181
6.V. CHILIN&ARIAN K.H. WOLF £EDITDRS1
ELSEVIER
DEVELOPMENTS IN SEDIMENTOLOGY 18B
COMPACTION OF COARSE-GRAINED S...
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181
6.V. CHILIN&ARIAN K.H. WOLF £EDITDRS1
ELSEVIER
DEVELOPMENTS IN SEDIMENTOLOGY 18B
COMPACTION OF COARSE-GRAINED SEDIMENTS, II
FURTHER TITLES IN THIS SERIES
1. L.M.J. U. VAN STRAATEN, Editor
DELTAIC AND SHALLOW MARINE DEPOSITS 2. G. C. AMSTUTZ, Editor SEDIMENTOLOGY AND ORE GENESIS 3. A. H. BOUMA and A. BROUWER, Editors TURBIDITES
4. F.G. TICKELL THE TECHNIQUES OF SEDIMENTARY MINERALOGY 5. J. C. INGLE Jr. THE MOVEMENT OF BEACH SAND 6. L. VANDER PLAS Jr. THE IDENTIFICATION OF DETRITAL FELDSPARS 7. S. DZULYNSKI and E. K. WALTON SEDIMENTARY FEATURES OF FLYSCH AND GREYWACKES 8. G. LARSEN and G. V. CHILINGAR, Editors DIAGENESIS IN SEDIMENTS 9. G. V CHILINGAR, H.J. BISSELL and R , W, FAIRBRIDGE, Editors CARBONATE ROCKS
10. P. MeL. D. DUFF, A. HALLAM and E.K. WALTON CYCLIC SEDIMENTATION 11. C. C. REEVES Jr. INTRODUCTION TO PALEOLIMNOLOGY
12. R .G.C. BATHURS T CARBONATE SEDIMENTS AND THEIR DIAGE~ESIS 13 A.A. MAN TEN SILURIAN REEFS OF GOTLAND 14. K. W. GLENNIE
DESERT SEDIMENTARY ENVIRONMENTS 15. C.E. WEAVER and L.D. POL LARD
THE CHEMISTRY OF CLAY MINERALS 16. H. H. RIEKE lil and G. V. CHILINGARIAN COMPACTION OF ARGILLACEOUS SEDIMENTS 17. M.D. PICARD and L.R. HIGH Jr. SEDIMENTARY STRUCTURES OF EPHEMERAL STREAMS
DEVELOPMENTS IN SEDIMENTOLOGY 18B
COMPACTION OF COARSE-GRAINED SEDIMENTS, I1 EDITED BY
GEORGE V. CHILINGARIAN His Imperial Majesty Shahanshah Arya Mehr Chair o f Petroleum Engineering, University o f Southern California, Los Angeles, Calif. (U.S.A.) and Abadan Institute o f Technology, Abadan (Iran)
AND
KARL H. WOLF Department o f Geology, Laurentian University, Sudbury, Ont. (Canada) Watts, Griffis and Me Quat Ltd., Consulting Geologists and Engineers, and Directorate Geneml of Mineral Resources, Jedda (Saudi Arabia)
ELSEVIER SCIENTIFIC PUBLISHING COMPANY Amsterdam - Oxford - New York 1976
ELSEVIER SCIENTIFIC PUBLISHING COMPANY 335 Jan van Galenstraat P.O. Box 211, Amsterdam, The Netherlands
Distributors for the United States and Canada: ELSEVIER/NORTH-HOLLAND INC. 52,Vanderbilt Avenue New York, N.Y. 10017
Library o f Congress Cataloging in Publication D a l s
Main entry under title:
(Revised)
Compaction of coarse-grained sediments. (Developments in sedimentology ; 18A-18B) Includes bibliographies and indexes. 1. Sediments (Geology) 2. Sand. 3. Grairel. 4. Diagenesis. I. Chilingarian, George V., 192911. Wolf, Karl H. 111. Series. m471.2.C65 551.3'04 73-85220 ISBN 0-444-41152-6 (American Elsevier)
Copyright @ 1976 by Elsevier Scientific Publishing Company, Amsterdam All rights reserved. No part of this publication may be reproduced, stored in a retrieval system, or transmitted in any form or by any means, electronic, mechanical, photocopying, recording, or otherwise, without the prior written permission of the publisher, Elsevier Scientific Publishing Company, Jan van Galenstraat 335,Amsterdam Printed in The Netherlands
This book is dedicated to His Excellency DR. MANOUCHEHR EGHBAL of the National Iranian Oil Company for his very important contributions to the world petroleum indus-
try, to our inspirers DRS. R.L. FOLK and H.D. HEDBERG and t o DRS. L.F. ATHY, J.M. WELLER, R.H. MEADE, G. DICKINSON, P.A. KRYUKOV, E.C. ROBERTSON, W. VAN DER KNAAP, K. TERZAGHI, G. RITTENHOUSE, H.J. FRASER, W. VON ENGELHARDT, E.L. HAMILTON, A.W. SKEMPTON, V.D. LOMTADZE and J. GEERTSMA for their important contributions to the field of compaction of sediments
CONTRIBUTORS F.A.F. BERRY Water Resources Division, US. Department of the Interior, Geological Survey, Menlo Park, Calif., U.S.A. D.F. BRANAGAN Department of Geology and Geophysics, The University of Sydney, Sydney, N.S. W., Australia G.V. CHILINGARIAN Petroleum Engineering Department, University of Southern California, Los Angeles, Calif., U.S.A. J.R. HAILS Department of Environmental Studies, University of Adelaide, Adelaide, S.A., Australia Y.K. KHARAKA Water Resources Division, U.S. Department of the Interior, Geological Survey, Menlo Park, Calif., U.S.A. D.M. RAGAN Department of Geology, Arizona State University, Tempe, Ariz., U.S.A.
M.F. SHERIDAN Department of Geology, Arizona State University, Tempe, Ariz., U.S.A. K.H. WOLF * Watts, Griffis and McQuat Ltd., and Directorate General of Mineral Resources, Jeddah, Saudi Arabia
* Formerly: Laurentian University, Sudbury, Ont., Canada.
CONTENTS
Chapter 1. INTRODUCTION by K.H. Wolf, G.V. Chilingarian and D.F. Branagan Sedimentology . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Sedimentary environments . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Conceptual model of sandstone sedimentology and petrology . . . . . . . . . . . . Provenance factors controlling types and compositional ranges of sandstones . . . Volcanic sandstones and tuffs . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Maturity concept in sandstone petrology . . . . . . . . . . . . . Possible loading conditions during sedimentation . . . . . . Sediments within the rock and geochemical cycle . . . . . . . . . . . . . . . . Bulk composition of sandstones . . . . . . . . . . . . . . . . .. . . . . . . . . . The effect of combined sedimentary and metamorphic processes in the origin graywackes and wackes . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Influence of mass properties on geophysical characteristics of sedimentary rocks References . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
. . . . . .
1 3 3 3 18 19 23 26 27
of 30 35 39
Chapter 2. CHEMISTRY OF WATERS EXPELLED FROM SANDS AND SAND· STONES by Y.K. Kharaka and F.A.F. Berry Introduction . . . . . . . . . . . . . . . . . . . . . . . . . . . . .. . Chemical controls on the composition of subsurface waters Mineral composition and porosity of arenaceous sediments . Composition of sandstones . . . . . . . , . . ....... : Porosities of sands and sandstones . . . . . Variation of shale porosity with depth . . . . . ... .. . . . . . . . . Membrane behavior of arenaceous sediments . . . . . . . . . . .. . . . Origin of the membrane properties of arenaceous sediments . . . . . The composition of the constituent minerals, 54 -The grain size of the con· stituent minerals, 54 - pH of the solution and type of t he exchangeable cations, 55 - The presence and nature of the organic matter in arenaceous sediments, 55 Ion exclusion and the double·layer theory . . . . . . . . . . . . . . .. Composition of solutions squeezed from membrane materials .. . Probable membrane effects in arenaceous sediments . .. .. . .. . Summary . . . . . . ................................. . References . . . . . . . . . . . . . . . . . . . . . . . . . . . .. .. . . . . . . . . . .. . .
41
42 47 47 49 52 54 54
55 58 61 64 64
Chapter 3. DIAGENESIS OF SANDSTONES AND COMPACTION by K.H. Wolf and G.V. Chilingarian General factors controlling compaction of sandstones Diagenesis in general : . . . . . . . . . . . . . . . . . . . . .
69
74
VIII
CONTENTS
Factors controlling diagenesis . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Diagenesis in sandstones . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Cementation of sandstones controlled by compaction fluids . . . . . . . . . . . . . . . . Monocrystalline grains . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Polycrystalline granular material . . . . . . . . . . . . . . . . . • . . . . . . . . . . . . . . Grains of hybrid composition . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Geologic applications . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . The studies of Levandowski and other investigators . . . . . . . . . . . . . . . . . . . , Textures resulting from compaction . . . . . • . . . . . . . . . . . . . . , . . . . . . . . . . . The concept of Morrow . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . The measuring technique of Kahn . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Depth of burial and diagenesis (Taylor, 1950) . . . . . . . . . . . . . . . . . . . . . . . Study on a carbonate-quartz system (Smith, 1969) . . . . . . . . . . . . . . . . . . . Complex diagenetic alterations (Aalto, 1972) . . • . . . . . . . . . . . . . . . . . . . . . Porosity and packing index. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Effects of textures and composition (Renton et al., 1969) . . . . . . . . . . . . . . . Observations on rims and coatings (Pittman and Lumsden, 1968) . . . . . . . . . . Study on the source of silica (Waugh, 1970) . . . . . . . . . . . . . . . . . . . . . . . . Study on grain contacts and packing parameters (Martin, 1972; Gaither, 1953, and others) . . . . . . . . . . . . . . . . . . . • . . . . . . . . . . . . . . . . . . . . . . . . . . The influence of physical compaction . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Formation of "glyptomorphs", 184 - 'T ransitional stages, 184 Porosity and permeability . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Studies by Rittenhouse and Fraser . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Plasticity, compressibility, density, and thixotropy of sandy sediments and sandstones Effects of compaction fluids on clay mineralogy in sandstones . . . . . . . . . . . . . . Effects of compaction fluids on trace-element and isotope composition of sediments
78 85 100 109 110 112 114 116 133 153 156 156 164 167 170 170 175 176
176 182 188 190 241 271 281
The trace·element budget .•..••.•..••...•. , . • • . . . . . . . . . • • . • . . . 281 The boron content . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Other trace-element contents . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Compaction fluids and isotope composition . . . . . . . . . . . . . . . . . . . . . . . . . Fluid-diagenesis in sandstones . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Regional features of compaction . . . . . . . • . . . . . . . . . . .. . . . . . . . . . . . . . . . Study by Fiichtbauer . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Study by Philipp and others . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Study by Adams . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .. . Study by Horn . . . . . . . . . . . . . . . . . . . . , . . . . . . . . . . . . . . . . . . . . . . . Diagenetic features and chemistry of pore fluids (Selley, 1966) . . . . . . . . . . . . The effects of compaction and depth of burial (Fiichtbauer, 1967) . . . . . . . . . Compaction and stratigraphic correlation (Conybeare, 1967) . . . . . . . . . . . . . Migration of compaction fluids (Magara, 1968) . . . . . . . . . . . . . . . . . . . . . . Study on g>"ain·contact types (Phipps, 1969) . . . . . . . . . . . . . . . . . . . . . . . . Alluvial fan deposits (Bull, 1972), . . . . . . . • . . . . . . . . . . . . . . . . . . . . . . . Experiments of Pryor (1973) . . . . . . . • . . . . . . . . . . . . . . . . . . . . . . . . . . Late diagenesis-epigenesis-burial metamorphism . . . . . . . . . . . . . . . . . . . . . . . Zones of secondary alterations . . . . . . • . . . . . . . . . . . . . . . . . . . . . . . . . . Zone of unaltered clay cement, 415- Zone of altered clay cement, 415Zone of quartz-like structures and illite--chlorite cement, 416 - Zone of spine-like aggregates and muscovite-chlorite cement, 416 Structures in sedimentary rocks as a result of compaction . . . . . . . . . . . . . . . . . References . . . . . . . . . . . . . . . . . . . . . • . . . . . . . . . . . . . . . . . . . . . . . . . .
281 287 288 290 309 311 3i8 332 335 342 342 365 372 379 385 390 395 415
422 423
CONTENTS
IX
Clulpter4. COMPACTION AND DIAGENESIS OF VERY COARSE-GRAINED
SEDIMENTS by J.R . Hails Introduction . . . . . . . . . . . . . . . . . .. . . • . . . . . . . . . . . Gravels . . . . . . . . . . . . . . . . . . . . . . . . • . . . . . . . . . . . Introduction . . . . . . . . . . . . . . . . . . • . . . . . . . . . . . Clean gravels . . . . . . . . . . . . . . . . . . • . . . . . . . . . . . Gravels with interstitial material and a high matrix content Gravelly soils and other very coarse-grained sediments . . . Laterite and silcrete gravels . . . . . . . . . • . . . . . . . . . . . Crushing strength of coarse-grained sands and gravels. . . . . . Conglomerates and breccias . . . . . . . . . . • . . . . . . . . . . . Summary and conclusions . . . . . . . . . . . . . . . . . . . . . . . Acknowledgements . . . . . . . . . . . . . . . . • . . . . . . . . . . . References . . . . . . . . . . . . . . . . . . . . . • . . . . . . . . . . .
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445 447 447 447 454 455 463 465 466 469 470 471
Introduction . . . . . . . . . . . . . . . . . . . . • . . . . . . . . . . . . . . . . . . . . . . . . . . Delineating and defining some compaction problems and concepts. . . . . . . . . . . . General comments . . . . . . . . . . . ,'.. . • . . . . . . . . • . . . . . . . . . . . . . . . . . Formation of ore deposits by water compaction . . . . . . . . . . . . . . . . . . . . . . Relationship between rheological properties of sediments and ore genesis . . . . . Relationships between the origin of petroleum and metalliferous concentrations
47 5 477 4 77 485 488 490
Chapter 5. ORE GENESIS INFLUENCED BY COMPACTION by K.H. Wolf
Examples of stratabound and stratiform ore deposits and their relationships to &edi· ment compaction . . . . . . . . . . . . . . . . . • . Missisaippi Valley-type zinc-lead deposits . Copper deposits in sedimentary rocks .. "" Uranium deposits in sediments . . . . . . . . • \ Sedimentary iron ores . . . . . . . . . . . . . . .. . Conclusions and summary . . . . . . . . . . . • : Acknowledgements . . . . . . . . . . . . . . . . . . References . . . . . . . . . . . . . . . . . . . . . .. .
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513 513 556 591 626 642 663 663
Introduction . . . . . . . . . . . . . . . . . .. .. . ... . ... .. ... . . Emplacement and deformation . . . . . . . .. . ... ... .. . ... .. Strain measurements . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Strain . . . . . . . . . . . . . . . . . . . . . . . • . . . . . . . . . . . . . . . . . . Bubbles . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Shards . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Pumice inclusions . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Variations in strain . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Compaction profiles . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Rigid inclusions . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Pumice . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Scoria and obsidian inclusions . . . . . . . . . . . . . . . . . . . . . . . . . .
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677 681 686 686 687 687 689 693 693 696 699 701
Chapter 6. COMPACTION OF ASH-FLOW TUFFS
by M.F. Sheridan and D.M. Ragan
CONTENTS
X
Basement topography Interpretations . . . . . . . Glossary of terms . . . . . Nomenclature . . . . . . . References . . . . . . . . .
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702 704 711 714 715
Appendix. COMPACTIONAL DIAGENESIS OF CARBONATE SEDIMENTS AND ROCKS by K.H. Wolf and G.V. Chilingarian
Compaction and the preservation of fossils Mass and petrophysical properties . . . . . . Compact ion and compaction fluids . . . . . Regional diagenesis . . . . . . . . . . . . . . . . References . . . . . . . . . . . . . . . . . . . . .
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719 7 36 7 54 761 767
REFERENCES INDEX . . . . . . . . • . . . . • . . . . . . . . . . . . . . . . . . . . . . . . . . 769 SUBJECT INDEX . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
781
Chapter 1
INTRODUCTION K.H. WOLF, G.V. CHILINGARIAN and D.F. BRANAGAN
This volume is a companion to Volume I of Compaction of Coarse-Grained Sediments, which contains subjects of interest to geologists in general, but in particular to sedimentologists, stratigraphers and petrologists. Those engaged in petroleum engineering, civil engineering, and geophysical studies (e.g., comparing the influence of sedimentary variables with the geophysical behavior of the rocks), should also have found numerous relevant chapters in Volume I. The present volume has gone one step further in including a topic of interest to "economic geologists"* or ore petrologists, inasmuch as the relationships among compaction, fluid movements, and the formation of ores in sedimentary (and volcanic) rocks have been examined and summarized. Although it was made clear in the Introduction of Volume I that it is beyond the scope of this publication to discuss in detail the complexities of the genesis of sedimentary rocks, some additional information is presented here introducing Volume II. This seems necessary, as in the present volume there is a strong shift to petrologic problems which suppose a fundamental knowledge of sedimentology. This chapter consists of two parts, namely, (a) additional data on the sedimentology of sandstones, and (b) data on the compaction of coarsegrained sediments that has not been covered elsewhere in the two volumes. In reference to (a), it is recommended that the reader familiarize himself with the introduction chapter in Volume I prior to reading the material below in this and later chapters. SEDIMENTOLOGY
Sedimentology and sedimentary petrology are becoming ever more multifaceted and, consequently, increasingly complex. Sandstone sedimentology, *In a way, the term "economic geologist" is unfortunate from the layman's point-ofview, and it is proposed that "ore petrologist" might be a better expression to be generally adopted. "Economic geology", as used in most cases, is not a combination of "economics" and "geology", but refers to the study of ores and their genesis, commonly without reference "to the "economic" details involved. In a strict sense, "economic geology" should include petroleum, metallic and nonmetallic industrial geology.
K.H. WOLF, G.V. CHILINGARIAN AND D.F.BRANAGAN
2
Petrography P e t r o l agy
HI nera logy I
Chemistry, Geochemistry
Biology--Paleobiology (e.g., b a c t e r i o l o g y i n diagenesisl
-I
GRAINED SEDIMENTS
Computer Science (i.e., handling of data)
Hvdrodvnami cs Akrodynamics (e.g., t r a n s p o r t a t i o n . textures. structures)
I-
Phy~its--Geophy5iCs (e .g., p a i e m a g n e t I c p r o p e r t i e s of h e m a t i t i c sandstones)
Geography--Paleogeography (e.g., ciimatologyl
Fig. 1-1. Diagram depicting the various disciplines that are utilized in sedimentology.
petrology, and stratigraphy are examples of a wide spectrum of subdisciplines of the so-called “soft rock” geology. Within the field of sandstone investigations, the examination of compaction forms one small part of this discipline of petrology. As shown in Figs. 1-1and 1-2, various scientific fields of studies supply the sedimentologist with concepts and techniques. The discipline of sedimentology, on the other hand, contributes to a number of fields (Fig. 1-2). The results obtained from the investigations of compaction of clastic and carbonate sediments and rocks are employed in one way or another in all fields listed in Fig. 1-2.
Economic Geology ( e . g . , u r a n i u m and copper i n sandstone)
t
COARSE GRAINED SEDIMENTS
* (e.g..
Oceanography r e c e n t sediments)
INTRODUCTION
3
SEDIMENTARY ENVIRONMENTS
An overall view of the major environments in which sediments can form is presented in Fig. 1-4,whereas Table 1-1 lists the more common types of sandstones (usually accompanied by claystones, mudstones, siltstones, conglomerates). Limestone-forming milieux, such as marine-shelf and biohermal reef complexes, are also depicted. Depending on the local conditions, the debris of reefs in addition to the oolites and skeletons from the shelf can be reworked, transported and deposited as beach, bar, and dune calcarenites (= sand-sized limestones), as mentioned in Chapter 3 in Vol. I. Consequently, they become closely associated with terrigenous sand deposits. Over a period of thousands of years, a thick pile of deposits may accumulate giving rise to a stratigraphic column several hundred feet thick. A hypothetical cross section of a possible stratigraphic section, composed of a terrigenous and a limestone complex, is given in Fig. 1-3. CONCEPTUAL MODEL OF SANDSTONE SEDIMENTOLOGY AND PETROLOGY
Figure 1-5 is a flow-chart diagram which includes all the processes and variables present in nature (from a regional macroscale down to the microscopic scale) that control the properties of sedimentary deposits. To read the model, one should start with the lowest numbered “box”, i.e., “Particle Properties” No. 1, then move to box No. 2 that includes “Textures” and “Fabrics”. All these are related to each other genetically so that they are part of box No. 3. All parameters in box No. 3 control the “structural” characteristics in box No. 4; these, in turn, control the “stratigraphic” properties in box No. 5, and so on. As the number of the boxes increases, the geologic scale of the investigation also increases and more and more variables and processes are included in the investigation of a particular area. (For details, see Wolf, 1973.) PROVENANCE FACTORS CONTROLLING TYPES AND COMPOSITIONAL RANGES OF SANDSTONES
Little has been said in Volume I about the geological variables in nature that influence the composition of terrigenous (i.e., clastic or detrital) * sedi*The terms “clastic” and “detrital” have been used in different ways so that they have lost a precise meanihg. They are used here synonymously with “terrigenous’’ or ‘‘continent-derived”.
K.H.WOLF,G.V.CHILINGARIANANDD.F.BRANAGAN
4
TABLE 1-1 Sandstone environments’ l a . Fluvial and alluvial l b . River delta 2. Watt (cf. No. 16) 3. Beach, tombolo, etc. 4. Bar 5. Dunes 6. Barrier Island 7. Chenier Island 8. Desert Refer to Fig. 1-4(pp. 5-7) vironments”.
9. Intramontane basins 10. Lake 11. Glacial 12. Volcanic terrain 13. Marine shelf (cf. No. 2 and 16) 14. Continental slope, e.g., turbidite environment 15. Bathyal (including terminal end of turbidite) 16. Tidal (cf. No. 2) 17. Deep-sea trench and correlate with “Conceptual model of sedimentary en-
COMPLEX NQI T E R R I G E N O U S ~ N T R lA -t
COMPLEX N0.2 (Mainly Limestone
-BASINAL
ENVIRONMENT^^
Fig. 1-3. Modalized stratigraphic section of two associated sedimentary complexes; one being terrigenous and the other mainly limestone in composition.
ments and rocks. Inasmuch as many classification schemes of sandstones are based on and, therefore, reflect these factors (see in particular Folk’s scheme: Fig. 1-15, Vol. I), an understanding of these variables is of utmost importance in environmental reconstructions. The details presented below are particularly significant in comparative studies that attempt to find reasons for local and regional sedimentological similarities and differences, which in turn influence differential compaction. Terrigenous sediments (in contrast to chemical sediments that form within the environment of deposition) originate in the source area, usually a land
pp. 5-7
Fig. 1-4. Conceptual model ofsedimentary environments. (After Kuenen, 1966, modified by Wolf,1973, fig. 6, pp. 164-166.)
-
pp. 8-10
~~
G L O B A L T E C T O N I C e.g ,continental d r i f t , p l a t e t e c t o n i c s
MASTER M OD E L O F SEDIMENTATION
------ -------------
------------.__.-_.-
REGIONAL
rn
S
ENVIRONMENTS
-
Miogeosyncline
Intra-rhontqn e
Cra ton
E u g e o s y n C Ii n e
b a s i n s
et c
Id
1ST
ORDER
1
16
.-
Y Sedimentary
-
8
I
SYNGENETIC- DlAGENETiC FACTORS 8 PROCESSES
I
Ti
-
REGIONAL FACTORS ( t o be considered I n addi ti on to those below) Geomorpholog IC depositional
Tectonism
environments
I I I
c g deIto.fIuvIaI.
I
I
I I
3rd order
I I
I
I I
2nd order
T l
7------------
I
environ rnent
I
v
HI
loke complex
L I
I I I
I I
PETROGRAPHY-
v
FACTORS WITHIN
I c
S E DI M E NTARY
C
I
1
1
c r
XI
I I I I
SOURCE AND
P,
E
I
C 0 L
-
I I
>
C
01
I
3
I
I@
+-@-
t
I
SedImerirory stuctllres
I
I I
I
I 1
7 a
1'
I
J
5
._--------------------------------Fig. 1-5. Master model of sedimentation; see text and original publication for details. (After Wolf. 1973, fig. 5. UP. 1617163.)
K H WOLF.f971-72 (fi rst revision)
I
J
INTRODUCTION
11
TABLE 1-11 Stages in the formation of a sedimentary rock GRADATION Bedrock
4
Soil
I. Degradation
1. Weathering 2. Mass wasting (= gravitational transfer) 3. Erosion and transportation by various agents (i.e., river water; ground
water; waves; currents; tides and tsunamis; wind; glaciers) Sediments Lithification Rock
11. Aggradation (by various agents - see erosion above)
111. Diugenesis to IV. Epigenesis (or catagenesis) and V. Burial metamorphism
mass or an island above sea level. The rock outcrops (i.e., “bed rocks”) undergo degradation (part I in Table 1-11), which includes mechanical and chemical weathering, that changes them to soil. The type of soil formed depends on the climate, topography, and numerous other variables, which determine the relative proportions of sand, silt, clay minerals, and organic content, as well as the amount and types of trace elements added and/or removed by solutions. Upon erosion and transportation, the particulate matter forms the clastic sediments. There are a number of different transportation mechanisms about which details are now available in many publications; these include mass transfer, rolling, sliding, saltation, and transportation in suspension or as a colloid. The transportation media are rivers and ground waters, waves and currents in lakes and oceans, tides and tsunamis (the misnomer “tidal wave” has been used for the latter), wind, and ice. The modes of transportation of the various components of sediments by stream waters are given in Fig. 1-6. After the particles reach a suitable site of accumulation, aggradation (part I1 in Table 1-11) commences. A number of cycles of deposition, re-erosion, and re-deposition can take place. Mechanical and chemical diagenesis, which is a complex field of investigation in itself (see Larsen and Chilingar, 1967, and Chapter 3 in this volume, for example) then changes the loose sediments into lithified or consolidated, solid rock over a period ranging from a few weeks to thousands or even millions of years. The final properties of the sediments are controlled by numerous interrelated variables, shown in simplified form in Table 1-111. As depicted in Table 1-IV, a crystalline igneous (or gneissic metamorphic)
12
K.H. WOLF, G.V. CHILINGARIAN AND D.F. BRANAGAN
True Solutions
True Solutionr
Fig. 1-6. Modes of transportation of principal components of sediments by stream waters. 1 = clay minerals; 2 = minerals of sands and rock fragments; 3 = other (a and a1 are only for mountain creeks in semi-arid climates). (After Strakhov, 1960, in: Perel’man, 1967, fig. 18.)
rock, for instance, undergoes various changes during weathering when the chemically unstable minerals (with a whole range of “relative stabilities”), such as orthoclase and plagioclase feldspar, change to clay minerals. This is accompanied by a relative increase in the percentage of stable quartz grains. Similarly, the proportions of the ferromagnesian minerals (amphibole, pyroxene, micas and olivines) are altered during weathering. As a result of differential erosion and transportation of the soil from the source areas, according to the grain size and shape and also due to the sorting and winnowing within the areas of accumulation, sediments are formed that range from clay-rich t o silt-rich muds and siltstones to sandstones and conglomer-
INTRODUCTION
13
TABLE 1-111 Factors determining character of sedimentary rocks
1. Nature of source rock (igneous, metamorphic, and/or sedimentary). 2. Topographic expression and relief of source area (high or low, slope characteristics,
a.
etc.).
Distribution of tectonic elements over source and depositional area (platform versus geosynclinal). 4. Intensity of tectonism in each tectonic element (slow to rapid subsidence or uplift, or stable). 5. Geology agents which weather, transport, and distribute sediments (e.g., arid versus humid weathering; slow versus fast transportation in suspension and by saltation and traction; sorting; abrasion; and turbidite versus slide deposition). 6. Patterns of environmental elements in the depositional area (e.g., presence or absence of lagoons, shelf, and delta). Example (numbers correspond to those above):
1 = plutonic, contact-metamorphic, and sedimentary source rocks; 2 = mountain; 3a = geanticline (= mountain source); 3 b = geosyncline; 40 = very active uplift; 4 b = very active subsidence; 5a = ice; 5 b = water and wind; 5c = water waves, currents; 5d = slides and turbidity currents; 6a = coastal plain, with or without a delta; 6 b = marine shelf; 6c = barrier reef or sand bar; and 6d = deep basin.
&.
The relationship between the grain size and grain composition is presented in Fig. 1-7 (Blatt et al., 1972). The petrographic identification of fine-grained material within sediments may offer difficulties. In particular the fine-grained rock fragments are hard to discriminate, because of the textural and compositional similarities and transitional possibilities, as illustrated in Fig. 1-8(Wolf, 1971). Figure 1-9 shows four source rocks that upon weathering, erosion and transportation of their sand-sized grains will form a sand deposit (the claysized particles would give rise to a mud): (a) plutonic and gneissic rocks supply feldspar,. quartz and ferromagnesian minerals; (b) terrigenous rocks, of which quartzites furnish mainly quartz, whereas arkoses supply quartz,
K.H.WOLF,G.V.CHILINGARIANANDD.F.BRANAGAN
14 TABLE 1-IV
Relationship between the minerals of igneous and sedimentary rocks (after Griffiths, 1967, table 10.1) Type of minerals
Essential
Varietal
Accessory Secondary (variable)
Igneous (%)
Minerals
Sedimentary (%)
orthoclase 0-30 0-50
0-20 0-20
<1
quartz amphibole { pyroxene 1 micas olivine 1 zircon tourmaline apatite garnet sphene, etc chlorite epidote
Type of minerals essential
0-9 5 accessory
<1
0-5
varietal
accessory
<1
In
L
0
Fig. 1-7. Relationship between grain size and composition of the detrital fraction in elastic silicate rocks. (After Blatt et al., 1972, fig. 8-19; courtesy of Prentice-Hall, Englewood Cliffs, New Jersey.) Cf. Strakhov, 1967-70, vol. 2, fig. 21, for similar data and for grain size distribution of heavy-mineral fraction. 1 = polymineralic rock fragments plus chert; 2 = monocrystalline quartz; 3 = clays and mica flakes; 4 = polycrystalline quartz; 5 = feldspars.
INTRODUCTION
15
meta-quartzite
Fig. 1-8. Circular diagram of ten lithologies between which textural and compositional transitions are possible that make identification and discrimination in sand grains and pebbles often difficult or impossible. (Modified after Wolf, 1971; courtesy of J. Sed. Petrol.)
K-feldspar and some ferromagnesian minerals; (c) volcanic rocks supply plagioclase feldspar and, occasionally, quartz; however, they mainly supply fragments of volcanic rocks that are often finely-crystalline and, therefore, very resistant t o breakdown; glass shards (=vitric grains) are also very common (see p. 19 for classification of tuffs); (d) older chemical sediments upon erosion supply a large variety of fragments, depending on their composition; carbonate fragments give rise to the calclithite debris (Folk, 1968; TERRIGENOUS SEDIMENTS
1
S A N D Fig. 1-9. Genesis of sand. Plutonic rocks supply mostly quartz and feldspar, terrigenous sediments mostly qpartz and rock fragments, volcanic rocks mostly rock fragments and glass, and chemical sediments mostly carbonate debris. (After Pettijohn et al., 1972, fig. 1-3,p. 3; courtesy of Springer-Verlag, New York.)
K.H. WOLF, G.V. CHILINGARIAN AND D.F. BRANAGAN
16
Folk et al., 1970); even glauconite has been shown to form second-cycle deposits. In so far as two or more source rocks can supply debris to one particular sedimentary environment, many sandstones and conglomerates are made up of a hybrid mixture of grain varieties. For this reason, one sandstone type can grade into all the other types, not only according to the grain-size fractions present, as depicted in the triangular diagrams in the Introduction chapter in Volume I, but also according to mineral composition of the grains shown in Figs. 1-15 to 1-17 of Vol. I. The geologic situation is, as depicted above, quite complicated. Figure 1-10 demonstrates that sandstones are not only composed of detrital grains that were derived from a terrigenous source, sometimes after longdistance transportation to the site of accumulation, but also of debris provided by the physical processes such as wave action in the depositional area. Chemical processes prevailing in the sedimentation milieux can also form particles that are incorporated in the sand deposits, Physicochemical processes can produce oolites and pellets of calcium carbonate, phosphate, and glauconite, for example, whereas biochemical processes may result in construction of protective and skeletal structures of animals and plants that later become detrital skeletal grains composed of whole shells or fragments thereof. Diagenetic processes at the water-sediment interface, or within the sedimentary unit itself, result in an alteration of detritus as well as neoformation of the so-called authigenic minerals, e.g., quartz, pyrite, calcite and dolomite. All these may be part of a sandy deposit. Figure 1-11depicts a flow-chart of
DETRITAL (TERRIGENOUS): FAR TRAVELLED
DETRITAL: FROM ENVIRONMENT
Labile
Srable
Quartz Feldspar Muscovite Amphiboles Tourmaline Pyroxenes Zircon I Clay minerals iinheritedl
primarily as skeletal fragments. oolites and
I I
1
CHEMICAL: DlAGENETlC ENVIRONMENT Quartz Tourmaline Chalcedony Zircon Feldspar Anatase Gypsum-anhydrite Pyrite Carbonates Petroleum Clay minerals (auihipenic+ inherited)
/
CHEMICAL: DEPOSITIONAL ENVIRONMENT Sluble
Lubile
Calcite Magnesian calcite Dolomite Protodolomite Phosphate Aragonite Clay minerals (mostly aurhigenic)
Fig. 1-10. Derivation of minerals in sandstones. (After Pettijohn et al., 1972, fig. 2-1, p. 28; courtesy of Springer-Verlag,New York.)
INTRODUCTION
17
-PHHYSICAL+CHEMICALEPICLASTIC (Weathering)
CATACLAsTrC-SAND
chemical- biochemical precipitation oolites -skeletal
SEAWATER
Fig. 1-11. Processes of sand formation. (After Pettijohn et al., 1972, fig. 8-1, p. 295; courtesy of Springer-Verlag, New York.)
the various processes involved, whereas Fig. 1-12 gives a useful birdseye view of the “evolution” of sandstones (and conglomerates): (1) formation of unconsolidated, loose sands through the derivation of the debris; and (2) transformation to lithified rocks through subsequent diagenesis.
18
K.H. WOLF, G.V. CHILINGARIAN AND D.F.BRANAGAN
Climatic controls on the origin of sediments are so distinct that it has been possible to establish useful correlation between the type of deposits formed and the climatic environmental variables, as done by Strakhov (1967-1970; see in particular figs. 170 and 171 in his volume 2, and fig. 2 in his volume 3). Many sandstones contain chemically precipitated cements and nodules or concretions, sulfide mineral concentrations, organic (= carbonaceous) matter, and shell (= fossil) fragments. Of these, the organic matter, together with clay minerals, may contain comparatively high amounts of adsorbed metals (e.g., Cu, Ni, Co, Pb, Zn, Mo, Mn), which upon release and subsequent biochemical or bacterial precipitation may give rise to ore minerals (e.g., see examples presented by Wedepohl, 1968). In some well-known instances, the latter are economically mined (see Chapter 5 in this book). VOLCANIC SANDSTONES AND TUFFS
In the discussions of sandstone classifications in the Introduction chapter of Volume I, little was said about volcanic sediments. As the models in the section on provenance clearly show, volcanic debris can form a major part of stratigraphic columns, especially in tectonic belts where volcanism was operative. Inasmuch as the present volume contains a chapter on the compaction of pyroclastic and related deposits, only a few additional comments are given here. The separation of tuffs from volcanic sandstones is, in practice, often very problematic. In theory, it is based on the presence or absence of reworking of the original volcanic detritus and erosion of volcanic flows. Volcanic debris dropping directly into water, however, may take on characteristics very similar to those of transported fragments of any origin. In addition, all gradations from true tuffs to volcanic sandstones can be expected within a single region. Volcanic fragments are commonly composed of: (1) glass shards (= vitric grains); (2) crystals of plagioclase feldspar and/or quartz, with some ferromagnesian minerals and mica booklets; and (3) various types of volcanic rock (= lithic) fragments. Accordingly, the tuffs are subdivided into the respective tuffs or tuffaceous sandstones, as shown in the triangular scheme of Fig. 1-13. All these are included in Folk et al.’s (1970) classification given in Fig. 1-15 of Vol. I. The volcanic lithic sandstones can be considered as a particular type of rock-fragment-rich sediments. Many sandstones, of course, contain small amounts of pyroclastic material among their lithic fractions. The possibility of pyroclastic sediments to act as source beds of metallic elements, which are released into subsurface fluids, has been discussed in Chapter 5 of this volume. These solutions, upon migration to geochemically
19
INTRODUCTION
GLASS
Fig. 1-13. Triangular diagram and end members of tuffs. (After Pettijohn et al., 1972, fig. 7-9, p. 269; see also O’Brien, 1963; Fisher, 1961, 1966; courtesy of Springer-Verlag, New York.)
suitable porous rocks, may precipitate minerals t o form stratabound ore deposits. MATURITY CONCEPT IN SANDSTONE PETROLOGY
The concept of textural maturity of sands was mentioned briefly in the Introduction chapter of Volume I. It requires further elaboration because of its fundamental control on the mass properties, such as porosity and permeability, and diagenetic behavior of sandstones (covered in Chapter 3), such as susceptibility to chemical and mechanical compaction. A clear distinction must be made between the various types of interstitial (= intergranular) infill in sands commonly referred to as “matrix” and “cement” (Fig. 1-14). Further discussion on this subject is provided in Chapter 3. The finely crystalline and occasionally amorphous components (usually clay minerals, chlorite, and micaceous minerals, with or without some silica, zeolites, and other rarer components) are termed “matrix”. It is now known that it can be of at least two different origins: (a) a fine detritus; &d (b) a secondary diagenetic and/or burial metamorphosed product (see section on graywackes, p. 30). The chemically precipitated crystals are known as “cement”, often composed of coarsely crystalline, blocky, drusy and/or acicular (= fibrous) calcite, aragonite, silica, and zeolite, for example. With a decrease in crystal size, the cement*gradually changes into what usually is termed “matrix”, so that a clear distinction between chemically-formed matrix and cement is
20
K.H. WOLF, G.V. CHILINGARIAN AND D.F. BRANAGAN
CLAY MATRIX
-
CHEMICALLY PRECIPITATED
ACICULARI(FIBROUS) ZEOLITE CEMENT
~‘LOCKY OR EQUID I MENS IONA L CEMENT
Fig. 1-14. Diagram showing the relationships between the framework of sand grains, on one hand, and a clay matrix and two types of chemically-precipitated cement, on the other.
difficult, even impossible. The terms “matrix” and “cement” are, therefore, somewhat unfortunate in transitional instances, because a fine-grained, chemically precipitated matrix lithifies or cements a sandstone just as well as does a coarsely crystalline cement. Nevertheless, the two terms are in wide use and can be employed unambiguously, if no genetic connotations are attached to them in doubtful cases. Those sandstones with a large percentage of clay or mud matrix are known as “wackes”. In studying the different textures, as outlined in Chapter 3, the different stages of both chemical and textural maturity have to be recorded. A useful visual diagram (Fig. 1-15) has been provided by Folk (1968),showing that any sandstone type * can range from immature to supermature, depending on the degree or intensity of winnowing (= removal of clay matrix), sorting (usually according to grain size, but also according to shape of the particles), and rounding (as a result of abrasion). Concomitant with the change of textural maturity, there is often a change in chemical maturity of the sand particles. If exposed long enough to the chemical and mechanical processes during weathering and transportation, the less stable minerals, such as feldspar, soft rock fragments, and ferromagnesian minerals, are destroyed preferentially and the stable minerals and quartz remain. The above-stated correlation between textural maturity of a sand and its
*According to the old definition of a graywacke, this clay-rich sedimentary rock does not fit this scheme. The new approach by Folk, however, appears to be more realistic in removing the clay matrix as an important variable in classifying sandstones. The matrix only indicates the stage of maturity.
INTRODUCTION
21
A
B
C
SUPERMATURE
(Rounding)
MATURE
(Sorting)
SUBMATURE
(Winnowing)
IMMATURE
QUARTZARENITE CHERTARENITE *%%Quartz
Abundont Chert
Fig. 1-15.Diagram illustrating the “maturity concept” of Folk (1968)as applied to the main sandstone types. The textural maturity ranges from immature, through submature and mature, to supermature depending on the degree of winnowing (which removes the fine-grained matrix), sorting, and rounding. With this change in textural maturity there may or may not be an increase in mineralogic maturity, depending on various environmental factors, such as climate and geologic time available for reworking. A. Igneous source; B. sedimentary source; C. metamorphic source. (Modified after Folk, 1968, p. 126;courtesy of Hemphill’s, Austin, Texas.)
mass properties becomes apparent on examining Fig. 1-15 and by considering the extreme cases. Porosity and permeability decrease with decreasing maturity. In the above discussion on employing textural maturity as one of the fundamental characteristics in classifying sandstones, the chemically precipitated cement within the sedimentary framework was ignored. Cementation is not a diagnostic criterion of the original depositional environmental param-
K.H. WOLF, G.V. CHILINGARIAN AND D.F. BRANAGAN
22
.A.
C
E
D
F
Fig. 1-16. Rock interstices and the relation of rock texture to porosity. Diagram showing several types of rock interstices and the relation of rock texture to porosity. A. Wellsorted sedimentary deposit having high porosity; B. poorly-sorted sedimentary deposit having low porosity; C. well-sorted sedimentary deposit consisting of pebbles that are themselves porous, so that the deposit as a whole has a very high porosity; D. well-sorted sedimentary deposit the porosity of which has been diminished by the deposition of mineral matter in the interstices; E. rock rendered porous by solution; F. rock rendered porous by fracturing. (After Meinzer, 1923,fig. 1;courtesy of U.S.Geol. Surv., Wash., in: Tolman, 1937,fig. 24, p. 98.)
50 Sedimentary
M Metamorphic
Fig. 1-17.Classification of terrigenous sandstones. (After Pettijohn et al., 1972,fig. 5-3, p. 158; modified from Dott, 1964,fig. 3; courtesy of Springer-Verlag, New York.)
INTRODUCTION
23
eters or variables, but is controlled by a number of secondary diagenetic factors operative from a few days to a few thousand years after deposition (in some instances up to a few million years later). On the other hand, both the original characteristics (e.g., textural maturity) and the secondary changes (e.g., degree of cementation, mode and degree of compaction, and the degree of fracturing and dissolution after lithification of the sediment) control the mass properties of rocks. Figure 1-16 illustrates the relationship between texture and porosity, for example. The amount of a clayey matrix has been considered in the sand and sandstone classification presented in Fig. 1-17. With an increase of the clay mineral proportion, the arenites (= matrix-free or matrix-poor sandstones) grade through wackes (= matrix-rich) and sandy mudstones into mudstones (= mainly composed of clay and, possibly, silty constituents). This is a useful approach in describing sandstones. Nevertheless, for detailed petrographic work, the approach developed by Folk (1968) and Folk et al. (1970), which includes both the composition of sand-sized grains and the textural maturity concept, is the most satisfactory so far (see Introduction chapter in Vol. I). POSSIBLE LOADING CONDITIONS DURING SEDIMENTATION
Possible loading conditions on a hypothetical sediment cube are presented in Fig. 1-18. The first condition presented (Fig. 1-18A) is polyuxiul loading, in which the three principal stresses are of different magnitude. Some investigators prefer to call this stress condition triaxial loading. The second loading condition presented (Fig. 1-18B) is hydrostatic, in which the three principal stresses applied are equal. The third type of loading (Fig. 1-18C) is triaxial, in which two of the three principal stresses are equal. Although some investigators justifiably refer t o it as biuxiul stress, the term triaxial is strongly imbedded in the earth sciences literature. In the uniaxial loading condition (Fig. 1-18D), the applied force acts in one direction only and is perpendicular to one surface of the sediment sample. The four faces of the cube parallel to the direction of the stress remain stationary. This arrangement can be achieved by placing the sample in a thick-walled, cylindrical chamber, the sides of which are stationary. The pressure can be applied with either one or two pistons, and the change in the volume of the sample is reflected by the change in the length of the sample. In the field of soil mechanics, this method is sometimes referred t o as triaxial testing. It should be mentioned also that some investigators reserve the term uniaxial for cases when there is a vertical stress with no lateral constraint and, therefore, lateral strain does occur. In biuxiul loading (Fig. 1-18E), the two principal stresses are equal, while two faces of the cube are held stationary. During the initial stages of
K.H. WOLF, G.V. CHILINGARIAN AND D.F. BRANAGAN
24
C
D
E
Fig. 1-18.Compaction loading classification. A. Polyaxial loading (p1 # p # p 3 ) , called “triaxial” by some investigators; B. hydrostatic loading (pl= p 2 = p g ) ; C. triaxial loading (pi = p 2 # p 3 ) , called “biaxial” loading by some investigators; D. uniaxial loading; four sides parallel to the stress p 3 are kept stationary; referred to as “biaxial” loading by some investigators; E. biaxial loading: p1 = p 2 and sides parallel to these two stresses are kept stationary. (After Sawabini et al., 1974,fig. 1,p. 133;courtesy of the Society of Petroleum Engineers of AIME.)
sedimentation, the loading condition is probably hydrostatic (Fig. 1-18B). As sedimentation progresses, the loading becomes triaxial (Fig. 1-18C). Finally, as the overburden load becomes large, the uniaxial loading condition (Fig. 1-18D) is approached. It seems also important to examine the frequently posed question, namely: Does the height of the overlying water column in a basin affect the compaction of coarse-grained sediments or not? It is commonly felt that compactive pressure at a certain depth below the sediment-water interface is higher if the depth of overlying water is greater. The compactive (or effective) pressure, p e , which actually causes compaction, is equal to the total overburden pressure, pt, minus the pore pressure, pp. The mathematical calculations presented here, however, show that the effective pressure at a certain depth below the sediment-water interface is independent of the depth of overlying basin water. If the depth of water is h, the specific weight of the hater is yw,the depth of point 1 (Fig. 1-19) below the sediment-water interface is h a , porosity of the sediment is Cp, and the
INTRODUCTION
25
Fig. 1-19. Schematic diagram used in calculating effective pressure at points 1 and 2 in a sediment. h3 = hq.
specific weight of solids is ys, then the total overburden pressure, pt, at point I is equal to (see Fig. 1-19):
(1-1) Ptl = hlYw + hsYs(1- $1 + h34YW Inasmuch as the pore pressure, pp,at point 1 is equal to the hydrostatic head: Ppl = hlY* +h3YW * the effective pressure, pe, at point 1 is equal to:
(1-2)
Pel =Ptl - p p 1 = h 3 r s ( l - @ ) +h3rw(@--1) Similarly, at point 2:
(1-3)
Pt2 = h 2 Y w
+
hsYs(1- 0) + h3dJYw
(1-4)
Inasmuch as th(s right-hand sides of the eqs. 1-3 and 1-6 are identical, effective pressure at point I is equal to the effective pressure at point 2. This is
26
K.H. WOLF, G.V. CHILINGARIAN AND D.F. BRANAGAN
true, however, only if the water above the watersediment interface is connected with the intergranular water (intercommunicating pore space). If there is an impermeable sediment layer above the coarse-grained sediment, then the effective (compactive) pressure at point 2 will be greater, if the pore pressures at points 1 and 2 are less than those indicated by eqs. 1-2 and 1-5, respectively. SEDIMENTS WITHIN THE ROCK AND GEOCHEMICAL CYCLE
The investigation of secondary mechanical and chemical changes of sedimentary rocks has to take into consideration the metamorphic alterations for reasons that are made clear in a section below, as well as in Chapter 3.
1I
INTERIOR OF EARTH
Fig. 1-20.The geochemical or rock cycle. (After Barth, 1962, fig. V-1, p. 378; courtesy of Wiley and Sons, New York.) Note: (1) Sediments are part of a cycle involving all major rock types. (2) There is no definite genetic separ?tion between sedimentary, igneous, and metamorphic rocks. ( 3 ) Considering geologic time, sediments (and all other materials) are not fixed in space and time.
INTRODUCTION
27
One has also to keep in mind that sediments are part of a complex cycle. As already discussed, sediments grade into tuffs and volcano-clastics and, through them, into other types of volcanic rocks. As the rock cycle in Fig. 1-20 illustrates, the sedimentary rocks also grade into burial-metamorphosed and higher-grade metamorphic products and, through the latter, into “granitic” gneisses. Some sandstones are composed of both purely sedimentary material and burial-metamorphic products; the wackes (i.e., graywackes or phyllarenites) serve as the best example (see pp. 22-23). Even high-grade metamorphic rocks in the Precambrian shields, in which all the original minerals have been altered, may retain certain features (e.g., pebbles, cross-bedding, size grading) that demonstrate their original sedimentary origin. Although such metamorphic rocks are often loosely referred to as “sediments” (e.g., graywackes, arkoses, and quartzites), yet it is understood by the petrologists working in the metamorphic belts that they are actually “meta-sediments” and “metavolcanics”. What is the significance of this to compaction? In future research, it will become important to extend the studies of compaction from sediments into burial-metamorphic products (as has been done occasionally already; see Chapter 6) and, eventually, into higher-grade metamorphic rocks, because porosity and permeability, among other mass properties, are functions of compaction and all concomitant processes. Detailed studies on diffusion in metamorphic rocks require a knowledge of the earlier processes prior to metamorphism (see example on p. 420). It is of some importance, therefore, to treat sedimentary and volcanic rocks as part of a rock or geochemical cycle, as depicted in Fig. 1-20. In the following sections, a few selectively chosen problems related to the extreme secondary changes of sandstones are briefly discussed to take the readers just one step beyond lower-grade alterations of which compaction is a part. Hopefully it will demonstrate that to fully comprehend the high-grade product it is necessary to attempt t o establish detailed petrologic models of all the lower-grade processes. BULK COMPOSITION OF SANDSTONES
Sandstone compositions (mineralogic, bulk chemical, minor and trace element, and isotopic) are important beyond sedimentary petrology and lowtemperature geochemistry. In the fields of igneous and metamorphic petrology and, especially, in the study of chemical differentiation since the Precambrian up to the Recent, it is often useful to compare the compositions of sedimentary rocks with those of high-temperature and high-pressure origins. Taylor (1968) showed how mixtures of certain sandstones plus basalt could
K.H. WOLF, G.V. CHILINGARIAN AND D.F. BRANAGAN
28
I
1
0
20
1
4 0 ' 6 0
80
BASALT IN MIX, 7'0 Fig. 1-21.Production of 60% Si02 by simple mixing of common rock types with basalt containing 50% Si02. A mixture of 20% quartzite With 80% basalt o r 50% granite and 50% basalt are examples of the quantities involved. It is clear that simple mixing requires large quantities of common rocks to be mixed with basalt. (After Taylor, 1968,fig. 5; courtesy of Pergamon Press, New York.)
5 Log ISiOZl Altos 1
Fig. 1-22.SiO2/Al2O3 versus Na20/K20 ratios of sandstones. Stipple outlines area of most analyses. Contours (heavy lines) show values of log [(SiOz + A1203)/(Na20 + KzO)]. Analyses drawn largely from Pettijohn (1963). (After Pettijohn et al., 1972,fig. 2-11,p. 62;courteky of Springer-Verlag, New York.) See also Garrels and Mackenzie (1971,figs. 8-4and 9-1,p. 227.)
29
INTRODUCTION TABLE 1-V Chemical classification of sandstones (after Pettijohn et al., 1972, fig. 2-10, p. 61.) Structural framework: anion
Exchangeable cations
Associated petrographic type
alkaline-earth-rich (carbonate cement)
quartz arenites
POUP
High SiOz /Alz O3 ratio (mature) (little clay or detrital Al-silicate)
alkaline-earth-poor (silica cement) Low SiOz /Alz O 3 ratio (immature) alkali-metal-rich (clay + detrital Al-silicate) (feldspar and clays)
alkali-metal-poor (aluminous clays)
Naz 0 > Kz 0
feldspathic graywackes
Naz 0 < Kz 0
arkoses; lithic graywackes lithic arenites
give rise to an igneous rock of andesite composition (Fig. 1-21),assuming that the mixture undergoes melting while a geosyncline, for instance, subsides into the earth’s crust. The compositions of sandstones can be made the basis for a chemical classification (Fig. 1-22and Table 1-V). The rock chemistry has a fundamental influence on sandstone diagenesis, including mechanical and chemical compaction. The style, degree, and rate of diagenesis will vary according to the chemical and mineralogic composition of the sediment. The similarities and differences between the various sedimentary and igneous rocks are demonstrated in Fig. 1-23.In order to predict the behavior of sandstones beyond diagenesis, in particular during low-grade (i.e., burial) and high-grade metamorphism, chemical classifications such as those given here are useful. The similarities between certain types of sedimentary and igneous rocks, shown in Figs. 1-21and 1-23,provide a partial explanation for the difficulty, even frequent impossibility, of distinguishing between these mdor rock types in metamorphic belts of the Precambrian Shields, because the similar original bulk composition is reconstituted to form identical metamorphic mineral assemblages. Unless some primary sedimentary and volcanic features, respectively, are preserved, the two rock varieties appear identical.
K.H. WOLF, G.V. CHILINGARIAN AND D.F. BRANAGAN
30 20
--
I
II
I
1' / Quartzites
I
16
Sandstones
B
12 /
5
0
Ark&
'" 0 -
// \
/
3
/ /
/
/
I
i
I
I
I
Sea water
(up 2 Log units)
I
I
I-
I
I
I
\
\
Average sandstone 0 Lithic sandstones
08
04
Average shale
0
I
I
0
10
log[(CaO
I 20
1
30
+ Na,O)/K,O]
Fig. 1-23. Relation between the compositions o f igneous rocks and those of sedimentary rocks. (After Garrels andMackenzie, 1971, fig. 9-1, p. 22'7; courtesy of W.W. Norton, New York.)
THE EFFECT OF COMBINED SEDIMENTARY AND METAMORPHIC PROCESSES IN THE ORIGIN OF GRAYWACKES AND WACKES
Matrix-rich (i.e., high clay mineral content) sandstones are discussed here as an illustration of the complexity of petrology, because a clear understanding of several processes is paramount in reading Chapter 3, for example. Until recently, it was assumed that the intergranular matrix of sands was of detrital (= clastic) Lorigin,i.e., it came from the source area and accumulated together with the sand grains, thus reflecting a lack of winnowing and sorting in the depositional environment. Detailed petrographic and geo-
I
INTRODUCTION
31
chemical studies have revealed, however, that much of the clay-sized fraction is of authigenic (= in situ) origin. It was formed later either diagenetically and/or through burial metamorphism as a secondary product within the sandstone framework, as a result of a chemical breakdown of unstable clastic mineral grains, t o form clay minerals, for example, often accompanied by chlorite and zeolites. A recent paper discussing the secondary alterations affecting graywackes, is that by Galloway (1974). Some details are given below as an introduction to Chapter 3. Figure 1-24illustrates three stages of chemical diagenesis that accompanied physical changes of the sand deposit. Grain deformation and
p q DEPOSITION
PHASE
PHYLLOSILICATE PORE FILL STAGE 3
ZEOLITE PORE FILL STAGE 3
ADVANCED BORlAL METAWRPHIC PHASE
I
COMPLEX REPLACEMENT G ALTERATION
TECTffllC PHASE
COMPACTION LUPLIFT
Fig. 1-24. Integration of data from the three sample suites shows that three stages of chemical diagenesis can be differentiated. Physical diagenesis, in the form of grain deformation and mechanical compaction, is especially pronounced in early to intermediate stages of burial when the fabric of the sandstone is open and cementation has not significantly filled-in available pore space. Formation of early calcite pore-filling cement (Stage 1) effectively insulates the sandstone from further diagenetic alteration or compaction until deep burial. Additional burial moves the sandstone into the realm of burial metamorphic grain ?alteration, which, if followed by a phase of tectonic deformation, can produce a graywacke grossly different in appearance from the original lithology. (After Galloway, 1974, fig. 8, p. 385; courtesy of Geological Society of America.)
K.H. WOLF, G.V. CHILINGARIAN AND D.F. BRANAGAN
32
mechanical compaction are particularly effective in the early and intermediate phases of burial as long as the framework of the sand is open and lithification is absent to poor. Upon precipitation of the first generation of cement (Stage l), compaction becomes ineffective until the rock subsides to deeper levels of the basin, where grain alteration due to the burial metamorphism commences. The minerals that form the “matrix” of this rock suite commence first as clay rims and coatings on the sand grains (Stage 2), which is followed by either phyllosilicate and/or zeolite pore filling (Stage 3). This variety of matrix is of secondary, chemical origin and, therefore, should not be confused with a detrital clay matrix. Laboratory experiments have confirmed the possibility of a chemical origin of matrix minerals by using the constituAMBIENT THERMAL GRADIENT PF/tOO‘)
3
(WtOOrnl
3.5
4
Fig. 1-25.Depth of burial at which Stage 2 and Stage 3 diagenetic facies are first encountered in fQur wells in the Queen Charlotte Basin, plotted against the present corrected thermal gradient at each well site. Length of bar denotes uncertainty resulting from variability of the sample intervals. Depth to the two diagenetic “tops” (below) increases significantly between wells characterized by relatively low ambient thermal gradients and wells with high thermal gradients (approximately 500-900 versus 1000-1200 m for Stage 2; 900-1800 versus 2400-2700 m for Stage 3), suggesting that temperature is a dominant factor in controlling stage of burial diagenesis. (After Galloway, 1974,fig. 9,p. 386;courtesy Geological Society of America.)
INTRODUCTION
33
r-BULK DENSITY (Gm/Ccl
I
SANDSTONE DENSITY / POROSITY
v S.
PRESENT DEPTH OF BURIAL
THERMAL
.........*
-~-Shll
-----
WELL GRADIENT Shrll Angb Horltquin 0-86 I.4°1F/100'
Anpb %CktYt E-66 f.9'F/fOOo0' Shtll Anglo Sacktyt 8-I0 WF/fM'
c___
Fig. 1-26.Relationships among bulk density, porosity and present depth of burial in three wells in the Queen Charlotte Basin. Sandstone porosity loss as a result of cementation is accelerated in the two wells with a high thermal gradient. Below 600 m (2000ft) sands in the hot wells have-considerably lower porosities than sands in the cool well at the equivalent depth of burial. (After Galloway, 1974, fig. 10, p. 386;courtesy Geological Society of America.)
K.H. WOLF, G.V. CHILINGARIAN AND D.F. BRANAGAN
34
1
2
5
3
4
I
6 I
Y
0 I
2
3 I
5
0
'
He - 9 10
I1 I2
I3 I4
15 I6
T
I I
Fig. 1-27.Summary plots showing average depth or stratigraphic interval in which diagenetic features first appear within each of the four basins. Once formed, such features are preserved over the entire depth-stratigraphic range investigated; thus, bars are continuous below point of first appearance. Where the sample grid is not continuous, spatial distribution is used to determine the relative age of various features. In the Bristol Basin, where depth of burial could bB only crudely approximated, the plot shows that successively older (and, on the average, more deeply buried) formations are characterized by types of cementation that occur at intermediate to deeper depths in the other basins. In all basins, early calcite predates all successive diagenetic phases. Clay rims and coats likewise consistently predate pore-fi!ling phyllosilicate or laumontite. (After Galloway, 1974,fig. 7,p. 348;courtesy Geological Society of America.)
INTRODUCTION
35
ents available from the sand grains (e.g., Whetten and Hawkins, 1970). It appears, however, that higher temperatures and pressures (i.e., burial diagenesis or burial metamorphism) are not absolutely necessary for the origin of phyllosilicate and zeolite minerals, as demonstrated by observations on Pliocene volcanic arenites and tuffs (Wolf and Ellison, 1971;Wolf and Surdam, 1971; as well as numerous other papers on zeolite genesis in tuffs). The Pliocene sandstones and conglomerates in question have not been buried by more than 400-500 f t and have not undergone appreciable increases in temperature and pressure, so that there is an insignificant amount of burial diagenesis, and burial metamorphism is absent. Yet, the intergranular spaces are filled with zeolites and phyllosilicates exhibiting the same textures as those described by Galloway (1974). Thus, it appears that both early diagenesis and burial metamorphism can give similar results. Only very detailed studies may prove the existence of differences, e.g. , certain varieties of zeolites form only under higher pressures and temperatures. Nevertheless, the problem outlined above makes it difficult to establish precise textural and mineralogical paragenetic sequences, unless the observed data can be correlated in boreholes with depth of burial, as done by Galloway (1974).In Fig. 1-25,he has plotted the first appearance of the indicators of Stages 2 and 3 against the presently observed corrected thermal gradient. The results obtained by Galloway suggest that temperature is a dominant factor in controlling the burial diagenetic processes. Under conditions of an increasing thermal gradient, the minerals indicative of alteration Stages 2 and 3 first appear and are encountered at greater depth. Both bulk density and porosity values undergo concomitant variations with depth (Fig. 1-26). Figure 1-27shows summary plots of the stratigraphic intervals indicating the first appearance of diagenetic products investigated by Galloway (1974) in four sedimentary basins. INFLUENCE OF MASS PROPERTIES ON GEOPHYSICAL CHARACTERISTICS OF SEDIMENTARY ROCKS
The geophysical behavior of sandstones is controlled by their mass properties, including those reflecting compaction processes. As illustrated in Fig. 1-28,the porosity, permeability, and the dielectrical, electrical, sonic-seismical, and thermal properties are a function of the total microscopic and megascopic makeup of the deposits. The field of geophysical behavior is a vast one and it is obvious that a separate book cmld be devoted to it. Nevertheless, two selectively chosen examples from the published literature are considered here to illustrate the
K.H. WOLF,G.V. CHILINGARIAN AND D.F.BRANAGAN
36
GEOPHYSICAL PROPERTIES:
SANDSTONE BODY Source rock environments
1
Factors during transportation
1
Factors in
Mm-
features
1
Mineralogic comwsit ion
Grain-size distribution 3 Textures 2
1
Depositional structures
2
Bedding relationships
I
Porosity and
1
Thermal properties
lithification 5 Degree
of
depositional environments
Fig. 1-28.Interrelationships among micro- and megafeatures of deposits, their genetic history and geophysical properties.
influence of various parameters of sedimentary rocks on their geophysical characteristics. Buchan et al. (1972) stated that the relationships between geophysical properties of sediments, e.g., acoustic and geotechnical, are a multivariate problem, and that multiple regression analysis may have to be used to sort out the variables that are redundant or less significant than
POROSITY, X
GRAPHIC MEAN, 8
Fig. 1-29.Relationship between sonic velocity and porosity. (After Buchan et al., 1972, fig. 2;courtesy of Q.J. Eng. Geol., Lond.) Fig. 1-30. Relationship between sonic velocity and graphic mean of grain diameter in phi units. (After Buchan et al., 1972,fig. 3; courtesy Q. J. Eng. Geol, Lond.)
INTRODUCTION
37
SAND FRACTION (>40), %
WET DENSITY, g / c d
Fig. 1-31.Relationship between sonic velocity and wet bulk density. (After Buchan et al., 1972,fig. 4;courtesy Q.J. Eng. Geol., Lond.) Fig. 1-32.Relationship between sonic velocity and percentage of sand fraction (>4 $J units in diameter). (After Buchan et al., 1972,fig. 6; courtesy Q. J. Eng. Geol., Lond.)
-
5ol.3
1.4
1.5
.
I6
1.7
I
WET DENSITY, g / c d
Fig. 1-33.Relationship between porosity and wet bulk density for various specific gravities of grains. (After Buchan et al., 1972,fig. 6;courtesy Q. J. Eng. Geol., Lond.)
K.H. WOLF, G.V. CHILINGARIAN AND D.F. BRANAGAN
38
others. Figure 1-29 prepared by these authors shows the relationship between sound velocity and porosity of sediments and also indicates that the Wood’s equation gives the lower limit of values (see original publication for details). The relationship between sound velocity and: (1)graphic mean of grain diameter in phi units, (2) wet density, and (3) sand fraction are presented in Figs. 1-30, 1-31, and 1-32, whereas Fig. 1-33 demonstrates a near straight-line relationship between porosity and wet bulk density. Zierfuss (1969) found that: (1) the heat conductivity values, as based on studies of clayey sandstones and compared with clean, clay-free sandstones, are appreciably lower in clayey sandstones than in clean sandstones having the same porosity; (2) correlations between heat conductivity and clay content in sandstones are poor; and (3) the heat conductivity of clayey sandstone can be related to the sum of porosity and clay content (expressed as a percentage of bulk volume). Zierfuss suggested that inasmuch as matrix conducts heat so much better than any pore fluid (oil, gas, or water) that may be present, the effect of the latter on the conductivity is small for rocks having moderate porosity. Figure 1-34a shows the relationship between heat Heat conductivity
u 20 LO 96
OO
a
Porosity
Fig. 1-34. a. Relationship between heat conductivity and porosity for West Netherlands clean sands (circles); Borneo clayey sands (triangles); and Bolivar Coast clayey sands (crosses). b. Relationship between heat conductivity and porosity plus clay content, expressed as percentage of bulk volume, for the same sands as above. (After Zierfuss, 1969, fig. 6, p. 259; courtesy of American Association of Petroleum Geologists.)
INTRODUCTION
39
conductivity and porosity, and it can be seen that the heat conductivities of clayey sandstones are less than those of clean sandstones. At first, Zierfuss did not find a clear relationship between increasing clay content and heat conductivity; however, his second attempt was more successful. Inasmuch as clay is a poor heat conductor in comparison with quartzitic material (see his table 111), Zierfuss considered it as part of the pore fluid. According to this reasoning, the actual porosities in effect were increased by the volumetric clay percentage. On plotting heat conductivity versus porosity-plus-clay content, a definite trend was established by him (Fig. 1-34b). In the new field of investigations on the use of nuclear explosion for practical purposes, such as stimulation of petroleum production, the numerous petrographic properties of the sedimentary rocks have to be considered as shown by the recent publication of Terman (1973). REFERENCES Barth, T.F.W., 1962. Theoretical Petrology. Wiley, New York, N.Y., 416 pp. Blatt, H.,Middleton, G. and Murray, R., 1972. Origin o f Sedimentary Rocks. PrenticeHall, Englewood Cliffs, N.J., 634 pp. Buchan, S., McCann, D.M. and Taylor Smith, D., 1972. Relations between the acoustic and geotechnical properties of marine sediments. Q. J. Eng. Geol. (Lond.), 5: 265-
284.
Cheng, J.T., 1973. The Effect o f Pressure and Temperature on Pore Volume Compressibility o f Reseruoir Rock. Thesis, Texas A & M Univ., Austin, 44 pp. Chu, H.H.H., 1973.Preliminary investigation for ground-water artificial recharge in Taipei Basin. In: Int. Symp. Underground Waste Managem. Artif. Recharge, New Orleans, La., 7 pp. Dott, R.H. Jr., 1964. Wacke, graywacke and matrix - what approach to immature sandstone classification? J. Sed. Petrol., 34: 625-632. Fisher, R.V., 1961. Proposed classification of volcaniclastic sediments and rocks. Bull. Geol. SOC.A m . , 72: 1409-1414. Fisher, R.V., 1966. Rocks composed of volcanic fragments and their classification. Earth-Sci. Rev., 1: 287-298. Folk, R.L., 1968.Petrology o f Sedimentary Rocks. Hemphill’s, Austin, Texas, 170 pp. Folk, R.L., Andrews, P.B.and Lewis, D.W., 1970.Detrital sedimentary rock classification and nomenclature for use in New Zealand, N . 2.J. Geol. Geophys., 13: 937-968. Galloway, W.E., 1974. Deposition and diagenetic alteration of sandstone in Northeast Pacific arc-related basins: Implications for graywacke genesis. Bull. Geol. SOC.A m . , 8 5 : 379-390. Garrels, R.M. and Mackenzie, F.T., 1971.Evolution of Sedimentary Rocks. W.W. Norton, New York, N.Y., 397 pp. Griffiths, J.C., 1967. Scientific ikfethod in Analysis of Sediments. McGraw-Hill, Toronto-New York, 508 pp. Inami, K., Iwamura, S. and Mitsui, S., 1972.Mechanical properHoshino, K., Koide, ties of Japanese Tertiary sedimentary rocks under high confining pressures. Geol. Surv. Japan, Rep., 244: 200 pp.
P.,
40
K.H. WOLF, G.V. CHILINGARIAN AND D.F. BRANAGAN
Kuenen, Ph.H., 1966. Geosynclinal sedimentation. GeoL Rundsch., 56: 1-19. Larsen, G. and Chilingar, G.V. (Editors), 1967.Diagenesis in Sediments. Elsevier, Amsterdam, 551 pp. O’Brien, R.T., 1963. Classification of tuffs. J. Sed. Petrol., 33: 234-235. Perel’man, A.I., 1967. Geochemistry of Epigenesis. Plenum Press, New York, N.Y., 266 PP. Pettijohn, F.J., 1963. Chemical composition of sandstones - excluding carbonate and volcanic sands. US.Geol. Surv. Prof. Pap., 440s: 19 pp. Pettijohn, F.J., Potter, P.E. and Siever, R., 1972. Sand and Sandstone. Springer, New York, N.Y., 618 pp. Sawabini, C.T., Chilingar, G.V. and Allen, D.R., 1974. Compressibility of unconsolidated, arkosic oil sands. SOC.Pet. Eng. J., 14 (3): 132-138. Strakhov, N.M., 1967-1970. Principles of Lithogenesis, 1, 2, 3. Consultants Bureau, New York, N.Y., 1: 245 pp., 2: 609 pp., 3: 577 pp. Taylor, S.R., 1968. Geochemistry of andesites. In: L.H. Ahrens (Editor), Origin and Distribution of the Elements. Pergamon Press, London-New York, pp. 559-584. Terman, M.J., 1973.Nuclear-explosion petroleumstimulation projects, United States and U.S.S.R. Bull. Am. Assoc. Pet. Geologists, 57: 990-1026. Tolman, C.F., 1937. Ground Water. McGraw-Hill, New York, N.Y., 1st ed., 593 pp. Wedepohl, K.H., 1968. Chemical fractionation in the sedimentary environment. In: L.H. Ahrens (Editor), Origin and Distribution of the Elements. Pergamon Press, LondonNew York, pp. 999-1016. Whetten, J.T. and Hawkins, J.W., Jr., 1970. Diagenetic origin of graywacke matrix minerals. Sedimentology, 15: 347-361. Wolf, K.H., 1971. Textural and compositional transitional stages between various lithic grain types (with a comment on “Interpreting detrital modes of graywacke and arkose”). J. Sed. Petrol., 41 : 328-332; 889. Wolf, K.H., 1973. Conceptual models, 1. Examples in sedimentary petrology, environmental and stratigraphic reconstruction, and soil, reef, chemical and placer sedimentary ore deposits. Sed. Geol., 9:153-193. Wolf, K.H. and Ellison, B., 1971. Sedimentary geology of the zeolitic volcanic lacustrine Pliocene Rome Beds, Oregon. J. Sed. Geol., 6: 271-302. Wolf, K.H. and Surdam, R., 1971.Diagenetic origin of zeolites, phyllosilicates and others in volcanic sediments, Oregon, U.S.A. Znt. Sed. Congr., 8th, Heidelberg, Abstr., p. 11. Zierfuss, H., 1969. Heat conductivity of some carbonate rocks and clayey sandstones. Bull. Am. Aesoc. Pet. Geogists, 53: 251-260.
Chapter 2 CHEMISTRY OF WATERS EXPELLED FROM SANDS AND SANDSTONES YOUSIF K. KHARAKA and FREDERICK A.F. BERRY INTRODUCTION
Almost all our knowledge about subsurface waters is derived from waters obtained from sand and sandstone beds (arenaceous sediments). These beds may, at shallow depths, serve as aquifers from which domestic or irrigation water is supplied; they may also serve as reservoirs from which oil,natural gas, and geothermal water and steam are obtained. Recently some of these beds have been used for the emplacement of liquid wastes obtained from industrial plants as well as oil-field brines produced with petroleum. There is essentially no field or experimental data directly concerning the chemistry of waters expelled from sands and sandstones as their porosity is reduced. The objective of this chapter is to investigate the various factors that might influence the chemical composition of such waters. Carbonate sands are specifically excluded from consideration. The chemistry of solutions expelled from arenaceous sediments by porevolume reduction is strongly influenced and sometimes totally controlled by the chemical composition of the original pore solutions. The origins and types of subsurface waters and the various factors that control the composition of interstitial solutions in subsurface environments are reviewed. Under the same physical conditions, the chemistry of expelled solutions will be different from the chemistry of pore solutions only to the degree that the arenaceous sediments exhibit membrane-type behavior. The mineral compositions and porosities of the principal sandstone types are briefly reviewed in this chapter, as are the factors that affect porosity reduction and thus water expulsion from various arenaceous sediments. (For details see Chapter 3.) The ion-exchange capacity of the various constituents present in sandstones is then discussed; the double-layer theory and the Teorell, Meyer and Sievers’ membrane model are used to develop some quantitative appraisal of the possible range of membrane influence on the chemistry of solutions expelled from arenaceous sediments of various composition. It is concluded that membrane phenomena may modify the chemistry of expelled pore waters from sands and sandstones providing that enough material with high ionexchange capacity - principally clays and/or kerogen - is present. The relative content of clay and kerogen is not nearly so important as the position in which that clay and/or kerogen is located.
Y.K. KHARAKA AND F.A.F. BERRY
42
CHEMICAL CONTROLS ON THE COMPOSITION OF SUBSURFACE WATERS
Waters that are present in shallow (< 1000 f t ) sand and sandstone beds are generally of meteoric origin and have relatively dilute chemical compositions. The solutes are derived from many different sources, including: (1) gases and aerosols from the atmosphere; (2) weathering and erosion of rocks and soil; (3) solution, precipitation and exchange reactions occurring along the flow path and in the aquifer; (4)man's activities; and (5) compaction of the underlying and adjacent sediments. Typical examples of chemical analyses of ground waters from such aquifers are given by White et al. (1963, table 4, p. F18). The waters generally contain less than 100 ppm dissolved solids and are dominated by Na+--.Ca2+--Mg2 cations and HCO;-SO$- C1- anions. The chemical composition and the salinity of interstitial waters from deeper formations are different from those of both surface and ocean waters. The electrolyte content of subsurface waters commonly increases with depth ranging from essentially that of fresh water to more than ten times higher than the concentration of sea water. Rogers (1917), Chebotarev (1955), White (1965), and others have noted that waters near the surface are dominated by sulfate ion. Bicarbonate waters tend to be intermediate both in their total salinity and depth of occurrence. C1- is the dominant anion in deeper waters, and its proportion to other anions generally increases with salinity. Recent data on the 'O/' 6O and deuterium/hydrogen ratios indicate that the major portion of almost all subsurface waters is meteoric in origin (Clayton et al., 1966; Hitchon and Friedman, 1969; Kharaka and Berry, 1973). The meteoric water commonly percolates from recharge areas located at higher elevations and is driven downdip by hydrodynamic forces. A few examples of connate waters (as redefined by White, 1965), however, are present in areas such as the Gulf Coast and California Coast Ranges where thick sequences of shales and siltstones are deposited rapidly and where the original high fluid potentials are maintained by continued compaction, dehydration reactions involving clay minerals, or tectonic phenomena (White et al., 1973; Kharaka et al., 1973). Connate waters as redefined by White (1965) are those waters that are generally similar in age with the associated rocks. Most connate waters are probably marine in origin and are associated with marine sediments. The chemistry of solutions expelled from arenaceous sediments will depend on the chemistry of their pore waters and their membrane behavior during reduction of pore volumes as a result of compaction. The knowledge of the different factors controlling the chemistry of pore solutions is variable. The origins of sulfate and bicarbonate waters are generally well under+
'
CHEMISTRY OF WATERS EXPELLED FROM SANDS AND SANDSTONES
43
stood. The sulfate in waters near the surface is derived from either the solution of gypsum and anhydrite or from the oxidation of pyrite present in the rocks (White, 1965). Pyrite oxidation by oxygenated water results in the liberation of Fe2+ as well as SO:-. The ferrous ion may move along the flow path in the now oxygendepleted ground water until it is precipitated or adsorbed on mineral surfaces. Also, the Fe 2+ cation may oxidize and precipitate as amorphous Fe(OH), that becomes more stable with age (Hem, 1967; Langmuir, 1969). The high bicarbonate content at intermediate depths is presumably due t o bacterial reduction of sulfate in the organic-rich environment. The reaction may be represented as follows: SO:- + CH4 + HS- + HCO; + H 2 0 The great differences in the salinity and ion proportions in the chloride waters, however, are not clearly defined. Several mechanisms (White et al., 1963; White, 1965) have been proposed to account for their origin. Evaporation and entrainment in expanding subsurface gas (Mills and Wells, 1919) and density stratification due t o molecular settling are unacceptable mechanisms because of the quantitative requirements involved (Russell, 1933; Mangelsdorf et al., 1970). Chloride waters can be derived from the solution of evaporite sequences in a few cases, such as the brines in the Devonian formations of Western Canada, but this is not possible in California and other sedimentary basins where evaporites are lacking (Chave, 1960). Reaction of water with minerals other than evaporites also contributes to increased salinities with depth, because the solubility and the rate of solution of aluminosilicates generally increase with increasing temperatures. Waterrock reactions also modify the chemical composition of subsurface waters. K/Na ratios in solution have been studied extensively both in natural and high-temperature experimental systems involving w a t e r rock interactions (Ellis and Mahon, 1964, 1967; White, 1965, 1970; Billings et al., 1969; and others). The K/Na ratios obtained are controlled mainly by exchange reactions with clays and feldspars. These reactions show such a strong temperature dependence that K/Na ratios are used in geothermal areas as a geochemical thermometer of subsurface temperatures (White, 1970; Fournier and Truesdell, 1973). K/Na ratios obtained from Kettleman North Dome Oil Field, California (Fig. 2-1) illustrate this temperature dependence. Kharaka and Berry (1973) concluded from a detailed study of geology, formation water chemistry, isotopic data, and regional hydrodynamics that the waters from the Temblor sandstones (Miocene) are meteoric in origin, whereas the waters in the McAdams sandstones (Eocene) represent the connate waters squeezed from the underlying shales and siltstones. The K/Na ratios in this field show no zonal or subzonal separation as shown by other chemical
Y.K.KHARAKAANDF.A.F.BERRY
*
PRODUCT ION ZONES *Temblor I 8 Vaqueros 0 Temblor II 0 Krcyenhegcn A Temblor Ill 0 Upper McAdams 0 Temblor IV e Lower McAdoms 0 Temblor V -t Mixed
0
0
0 0
0
OO
0
I
I
1
I
5
0
15
20
Q e e I
I
I
25
30
35
K / N a x lo3 Fig. 2-1. K/Na ratios (by weight) plotted against depth of midpoint of perforations below sea level of the production intervals at Kettleman North Dome, California.
ratios, but show a general increase with depth from Temblor I through Lower McAdams. The SiO concentrations in subsurface waters also increase with subsurface temperatures and are generally controlled by the solubility of quartz (White, 1970; Fournier and Truesdell, 1973). Ba and Sr concentrations in subsurface waters likewise are controlled mainly by the solubility of their respective sulfate and carbonate minerals (Kramer, 1969; Kharaka, 1971; and others) as in the following reactions:
CHEMISTRY OF WATERS EXPELLED FROM SANDS AND SANDSTONES
(Ba - Sr)SO, and:
* (Ba - Sr)"
45
+ SO:-
(Ba - Sr)C03 + H+ r+ (Ba - Sr)2 + HCO, +
Water-rock reactions control t o varying extents the concentrations of Ca2+ and Mg2+ in subsurface waters. Mg/Ca ratios obtained from subsurface waters are generally much lower than that of sea water. The increased loss of Mg with depth may be attributed to dolomitization. The loss of Mg in the diagenetic formation of chlorite, illite, Mg-montmorillonite, attapulgite and sepiolite, however, is more important in many sedimentary basins where carbonate sediments are not present in the geological section (Graf et al., 1966; Muffler and White, 1969). The general increase of Ca/Na ratios with depth is attributed by some investigators (Kramer, 1969; Hitchon et al., 1971) to a number of w a t e r r o c k reactions which release Ca to the fluid phase. These reactions include: dolomitization; solution of dolomite, calcite and gypsum; albitization; and diagenetic alteration of Ca-bearing silicate minerals to clays. Other investigators attribute most of the variations in Ca/Na ratios to membrane filtration (White, 1965; Kharaka and Berry, 1973). Diagenesis of organic matter incorporated in sediments also contributes to the geochemistry of subsurface waters. Studies have shown (White et al., 1963; Walters, 1967) that both I and Br are enriched in the organic phase, especially in some seaweeds and corals. Collins and Egleson (1967) obtained very high I and Br concentrations in oil-field brines from Anadarko Basin, Oklahoma, which they attributed t o diagenetic alterations involving marine organisms in sediments. Hyperfiltration through clays and shales acting as semipermeable membranes was first invoked by De Sitter (1947) t o explain the observed concentrations and the electrolyte content of the oil-field brines. Berry (1959), Berry and Hanshaw (1960), and Hanshaw and Hill (1969) invoked the membrane behavior of shales to explain the observed chemistry of the brines and anomalous low pressures within the various Cretaceous sands of the San Juan Basin. Berry and Hanshaw (1960) postulated that geological membranes were responsible for some of the fluid potential anomalies and the chemistry of their associated formation waters in western Canada and in the Wheeler Ridge Oil Field, California. Bailey et al. (1961) attributed the chemistry of the brines of t h e Cymric Oil Field, California, to diagenetic changes, modified by shales behaving as semipermeable membranes; Bredehoeft et al. (1963) developed mathematical models based on the membrane behavior of shales to obtain brines, with concentrations three to six times the concentration of present sea water, found at depth in the
Y.K.KHARAKAANDF.A.F.BERRY
46
Illinois Basin from fresh meteoric water containing small quantities of ions. Graf et al. (1966) also defined models and computed volumes of fresh and sea water needed to develop typical saline water compositions of the Illinois and Michigan basins by shale hyperfiltration. The contribution of geological membranes to the geochemistry of subsurface waters became widely accepted mainly due to White's (1965) classical paper which not only supported the concept of membrane filtration in the development of. saline formation waters, but also suggested specific mechanisms for providing differences in filtration rates of the major cations and anions commonly found in subsurface waters. Berry (1969) reviewed the laboratory and field evidence suggesting that shales serve as semipermeable membranes and discussed the relative importance of the various factors that influence the relative transport rates of dissolved species through geological membranes. Membrane behavior of shales was invoked by numerous other investigators (Weddle, 1967; Billings et al., 1969; Hitchon et al., 1971; Kharaka and Berry, 1973; and others) to explain the composition of formation waters in their respective areas. The chemistry of solutions squeezed from experimentally prepared monoionic (single cation and anion solutions) and multi-ionic clay-water systems (Kryukov and Komarova, 1954; Kryukov et al., 1962; Von Engelhardt and Gaida, 1963; Warner, 1964; Rieke et al., 1964; Chilingar and Rieke, 1968; and Kryukov, 1971) shows that the solutions become progressively more dilute in their electrolyte content and that the relative concentrations of the different species in the expelled solutions are different at different compaction pressures (Fig. 2-2). The chemical composition of solu-
-:
-
: I00 80
80
60
60
0 a 40
40
? 2
20
€ 2
5
.& c * u 0
u
\
ZEO
100,
rT
u
-c
120
120
0
0
; <
U
20
200
I50
water content
50
100 (In
percent
of
dry w e i g h t i
0
'0
-z -5 c
5
L
" c
10
0
V
u
0
60
40
20
0
Water Content (%dry w t 1
Fig. 2-2. Concentration o f solutes and conductivity (mhosicm) of solutions squeezed at different compactibn pressures from: a. bentonite; b. kaolinite. (After Kryukov, 1971, pp. 59 and 61.)
CHEMISTRY O F WATERS EXPELLED FROM SANDS AND SANDSTONES
47
tions squeezed from modern marine sediments is not significantly different from the overlying ocean waters (Shishkina, 1964; Siever et al., 1965; Manheim, 1966; Bischoff and Ku, 1970; and others). Some of these minor changes have been attributed t o the membrane behavior of natural materials (Siever et al., 1965); others have been attributed t o diagenetic water-rock reactions. The chemistry of pore waters in shales and the chemistry of waters originating from the natural subsurface compaction of shales and siltstones, however, are radically different from sea water and subsurface waters of other origins (Kharaka, 1971; Kharaka and Berry, 1973; Dickey et al., 1972; Schmidt, 1973; and others). These variations, which are generally attributed to the membrane properties of fine-grained sediments, parallel those reported by laboratory investigators. The chemistry of solutions expelled from arenaceous sediments will be different from the chemistry of their pore solutions only if arenaceous sediments act as membranes. The membrane behavior of arenaceous sediments will depend on their mineral composition, their degree of compaction and the chemistry of their pore solutions. MINERAL COMPOSITION AND POROSITY O F ARENACEOUS SEDIMENTS
Composition of sandstones Non-carbonate sandstones commonly are classified on the basis of the nature and relative abundance of their detrital constituents into four main types: orthoquartzites, lithic sandstones or subgraywackes, arkoses or subarkoses, and graywackes. (See Chapters 1and 3 for detailed discussion.) The relative abundance of these sandstone types is shown by the composite estimate from Krynine (1948), Middleton (1960), and Pettijohn (1963) given below: orthoquartzite lithic sandstone and subgraywacke arkose and subarkose graywacke
31% 27 22 20 100%
Pettijohn (1957, 1963) gave the following definitions for these principal rock types: Orthoquartzites consist of 90% or more detrital quartz. Lithic sandstones have an excess of rock fragments over detrital feldspar; if the non-quartz
Y.K.KHARAKAANDP.A.F.BERRY
48
TABLE 2-1 Composite average mineral composition (%) of various sandstone types (after Pettijohn, 1957,1963) Mineral composition
Rock type (No.of samples) orthoquartzite (10)
Quartz 96 Feldspar Mica Rock fragments (with shale) 4 Clay matrix Carbonates Minor constituents
arkose and sub- lithic sandstone arkose ( 7 ) and subgraywacke (54)
graywacke (35)
48 44 1
48
19 22 3
-
29 8 8
2
1 4
4 2
1
19 34
-
3
constituents (principally rock fragments and feldspar) exceed 25%, then the rock is a subgraywacke. Arkoses are sandstones where the rock fragments and feldspar comprise at least 25% of the rock and where feldspar comprises at least half of these non-quartz constituents. Subarkoses are sandstones where the rock fragments and feldspars constitute 10-2576 of the rock and again where feldspar constitutes at least half of these non-quartz constituents. Graywackes are sandstones that have sand-sized particles of quartz, feldspar and rock fragments embedded in a silt-to-clay-sized matrix that comprises 15%or more of the rock. Composite average mineral composition for these rock types is given in Table 2-1. Orthoquartzites, arkoses, and subarkoses form excellent reservoir rocks (high porosities and permeabilities). Lithic sandstones or subgraywackes are rocks with intermediate to poor reservoir properties; most “dirty sands” that serve as petroleum reservoir rocks are of this type. Graywackes have extraordinarily undesirable reservoir properties unless they are fractured. The clay content of arenaceous sediments may be present as shale fragments or as a clay matrix. The clay matrix may be present at deposition but oftentimes develops with depth as a result of alteration of unstable minerals. Rittenhouse (1973) indicated the variety of ways in which clay minerals may be present in sandstones. They: (1)may fill the pore space completely or partially; (2) may coat individual sand grains and thereby completely or partially block the smaller openings between the grains; and (3) may comprise the original detrkal sand grains that commonly are squeezed into the adjacent pore space upon subsequent burial.
CHEMISTRY OF WATERS EXPELLED FROM SANDS AND SANDSTONES
49
It is the presence of shales and clays in sandstones that must contribute principally to any membrane properties that sandstones might possess. Perusal of the compositions listed above clearly shows that membrane phenomena should be of considerable importance in graywackes and significant in clay-bearing lithic sandstones or subgraywackes. Membrane phenomena probably are of minor importance in,arkoses or subarkoses and are of negligible importance in orthoquartzites. Porosities of sands and sandstones Numerous porosity determinations have been made for unconsolidated
sands and various sandstones. Relatively unconsolidated sand deposits have relatively high porosities of 30-48% as shown in Table 2-11.
Most porosity determinations of sandstones have been made from petroleum reservoir rocks. There has been no attempt to measure and describe the porosities of sandstones, insofar as can be determined, by rock type. The majority of the sandstones for which porosities are available are orthoquartzites, subgraywackes, and subarkoses. The porosities of about 90 sandstones are listed by Birch et al. (1942). They range from 3.9% to 38%;the average is about 15-20%. The few porosity determinations of various sandstone types obtained from various sources are presented in Table 2-111. From this.limited list, there are no general conclusions which can be reached other than that the porosity variation for any given rock or rock type commonly is large and that the average porosity is, perhaps, 15%. The porosity of graywackes probably is highly variable depending princiTABLE 2-11 Porosities of relatively unconsolidated sands Formation name, age, and location
Porosity( %) range
Sand (Ft. Union Fm., Tertiary, Montana) 32.9-34.9 Sand (Lance Fm., Cretaceous, Montana) Sand (Kirkwood Fm., Miocene, New Jersey) 30.2-44.3 Sand (Mt. Laurel Fm., Cretaceous, New Jersey) Coarse sand (Rariton Fm., Cretaceous, New Jersey) 35.7-48.4 Fine sand (Quaternqry, California) Very fine sand (Quaternary, California)
Reference average
33.9 43.0
Birch et al. (1942) Birch et al. (1942) Birch et al. (1942)
44.4
Birch et al. (1942)
42.1 46.2 47.7
Birch et al. (1942) Clark (1966) Clark (1966)
Y.K. KHARAKA AND F.A.F. BERRY
50 TABLE 2-111 Porosities-of various sandstone types Sandstone type, formation name, age, and location
Orthoquartzites Berea Sandstone (Mississippian, Ohio and West Virginia) St. Peter Sandstone (Ordovician, Arkansas) Oriskany Sandstone (Devonian, West Virginia) Simpson Sandstone (Ordovician, Oklahoma)
Reference
Porosity (%) range
average
4.7-1 9
14.1
Clark (1966)
3.6-14
8.8
Clark (1966)
8.0-22
17.8
Birch et al. (1942)
16.0
Levorsen (1967)
15.0
Clark (1966)
Su bgray wackes
Bradford Sandstone (Devonian, Pennsylvania) 6.0-23.3 Temblor Sandstone (Miocene, California) 3.9-2 1.9 Bradford Sandstone (Devonian, , Pennsylvania) 4.6-1 8.6
Birch et al. (1942) Birch et al. (1942)
Su barkoses
Temblor Sandstone (Miocene, California)
3.9-21.9
13.0
Birch et al. (1942)
pally upon the degree of compaction of the fine-grained clay and siltstone matrix. Many graywackes have very low porosities: probably less than 5%" The amount of porosity that is developed in a particular sandstone and the subsequent modification of that porosity through various processes associated with burial are influenced by a large number of factors (Taylor, 1950; Weller, 1959; Rittenhouse, 1973). In general terms, burial of sandstones means porosity reduction which, in turn, means that pore water is expelled from water-saturated sandstones. Rittenhouse has found that the porosity of for each 1000 Gulf Coast Tertiary sandstones decreases on an average of 1% f t of burial, though the porosities at any given depth vary widely and the porosity decrease is not uniform. Porosity reduction in sandstone may be achieved not only by compaction phenomena involving an overall reduction -___
* Editorial note: 12 addition to the degree of compaction, the amount of primary matrix as well as of secondary-formed clays, zeolites and chlorite would be important (for details, cf. Chapters 1and 3).
CHEMISTRY OF WATERS EXPELLED FROM SANDS AND SANDSTONES
51
in the bulk rock volume, but also by various other processes while the bulk rock volume remains essentially unaltered. Thus the expression “porosity reduction” is the only overall term that describes the generally observed decrease in sandstone porosity with burial. (For details, see Chapter 3.) Several factors affect compaction in sandstones. Taylor (1950)described the importance of: (1)solution at intergranular boundaries; (2) distortion of weaker grains and minerals; and (3) crushing of sand grains in the compaction of sandstone. Rearrangement of the position and orientation of sand grains also may be an important factor in sandstone compaction. Porosity reduction and compaction can be modified by the nature of the packing. Equidimensional spherical grains cubically packed have a porosity of 47.676, whereas the same grains packed rhombohedrally have a porosity of only 25.9% (Graton and Fraser, 1935). Rittenhouse (1973)suggested that the natural processes that develop closer packing include: (1)alternate wetting and drying; (2) drying, particularly for clay-rich rocks; (3) downward infiltration of silt and clay; (4)increasing burial accompanied by pore-water escape; (5) earth tides; (6) microseisms and earthquakes; (7) intergranular flow of water; and (8) a decline in the ground-water level. As Rittenhouse indicated, a distinction as to effect must be made between the effect of solution of sand grains that become chemically unstable upon burial and that of the pressure-solution on grain boundaries owing to deep burial or tectonic stresses. The first type of solution phenomena may either decrease or increase the porosity, depending on whether or not the void space created is partially or completely filled by alteration of the packing and/or redeposition of the dissolved material in a more stable form. Pressuresolution at grain contacts, however, decreases both porosity and bulk volume. Plastic flow, crushing, or fracturing of sand grains also may decrease porosity and bulk volume. Plastic flow is particularly important when sand grains are composed of shale-type rock and micas. Crushing and fracturing are more important in the case of sand grains composed of rock fragments or feldspars. Porosity reduction in sandstones without reduction in the bulk volume is achieved through precipitation of crystalline material in the void space. Quartz, carbonates, anhydrite, and clays are the most common of these pore-filling materials. As Rittenhouse pointed out, pore filling of sandstones may occur any time after deposition. At the earlier stages, crystalline porefillers usually are calcite, aragonite or gypsum. Silica tends to be more important as a void filler at greater depth. Rittenhouse listed the following factors that control the amount of porosity and its subsequent alteration in a given sandstone: (1)the size, shape, original composition, and environment of deposition of the grains that com-
Y.K. KHARAKA A N D F.A.F. BERRY
52
prise the sandstone; (2)the composition and movement of the fluids in the pore space throughout the history of the sandstone; (3)the effective stress to which the sandstone is subjected, which is dependent upon the depth of burial, tectonic stress, and pore-fluid pressure; (4)temperature; (5)time; (6) mineral alteration; and (7) whether or not cementation of the sand grains, which converts them into a sandstone, occurs relatively early in the history of that rock. Porosities of sandstones also can be significantly reduced by man-made activities, particularly through the production of pore fluids, i.e., water, oil and/or gas. VARIATION OF SHALE POROSITY WITH DEPTH
Clays play such a significant role in the membrane behavior of arenaceous sediments that the variation of shale porosity with increasing depth of burial is briefly treated here (see also Chapter 3).Weller’s (1959)composite porosity-depth curve for shales to a depth of 10,000 ft, based on data from Terzaghi (1925),Athy (1930),and Hedberg (1936),is shown in Fig. 2-3. The greatest amount of porosity reduction in shales (80-45%) occurs within the first 50 f t of burial. The rate of porosity decrease with depth diminishes sharply below a depth of 50 f t and particularly below 100 ft. There are several qutstanding papers in which the various petrologic and geochemical factors that control the reduction of shale and/or clay porosities with depth and/or increased compaction pressure have been investigated (Terzaghi, 1925; Athy, 1930; Hedberg, 1936; Weller, 1969; Warner, 1964; Meade, 1964, 1966;Burst, 1969;Hiltabrand et al., 1973). (For their role in the release of ore-forming solutions, see Chapter 5 as well as the chapter by Mookherjee in Wolf, 1976,Vol. 4.)
DEPTH, F T . XI000
Fig. 2-3, Graph showing approximate relations between porosity and depth of burial for shale. (After Weller, 1959, p. 276.)
CHEMISTRY OF WATERS EXPELLED FROM SANDS AND SANDSTONES
53
The proper correlation between experimental compaction pressure and depth of burial is a complicated subject and is still not known. There probably is no one given correlation that is correct due to the different pore-pressure conditions to which different shales are subjected during their burial. Warner’s (1964)plot of his own and Skempton’s (1944,1953)experimental compaction data on clays with Weller’s (1959)composite curve for shale is shown in Fig. 2-4.The experimental data consistently indicate higher porosities for a given compaction pressure. This discrepancy is probably due mainly to the many reactions that take place under field conditions which are not duplicated by laboratory squeezing experiments. A significant part of the porosity decrease with depth of burial must be attributed to an increase in load pressure. Mineral-water reactions, the rates of which increase with increasing temperature of burial, also contribute to this decrease in porosity. Principally the temperature, pore-water chemistry, mineral composition, and time control the mineral-water reactions; the pressure environment is a modifying influence.
Fig. 2-4. Comparison between experimental and field compaction curves (Warner, 1964, p. 104): 1 = experimental curve, illite < 4 p (Warner, 1964,p. 104); 2 = experimental curve, highly colloidal slay (Skempton, 1953, p. 55); 3 = experimental curve, average clay (Skempton, 1944, p. 130); 4 = composite shale curve calculated from field data (Weller, 1959, fig. 3).
54
Y.K. KHARAKA AND F.A.F. BERRY
MEMBRANE BEHAVIOR OF ARENACEOUS SEDIMENTS
Origin of the membrane properties of arenaceous sediments The membrane properties of rocks result from charge deficiencies on the surfaces and edges of the constituent mineral particles. The charge deficiencies are balanced by exchangeable cations and anions. The concentration of exchangeable cations in rocks is generally higher than that of anions because of the net negative charges of the surfaces of rock particles. The cationic exchange capacity of sands and sandstones and, hence, their membrane properties will depend on: (1)the composition of the constituent minerals; (2) the grain size of the constituent minerals; (3) pH of the solution and type of the exchangeable cations; (4) the presence and nature of the organic matter in arenaceous sediments. The composition of the constituent minerals. Arenaceous sediments composed entirely of quartz will have very low, less than 0.05 mequiv./100 g, exchange capacity (Kennedy, 1965; Malcolm and Kennedy, 1970). The presence of unaltered feldspar will not increase the exchange capacity appreciably. Hem (1964) reported the exchange capacity of a beach sand sample composed of 45% quartz, 21% orthoclase feldspar, and 34% plagioclase feldspar to be 0.66 mequiv./100 g. The exchange capacity of the mineral portion of the sand fraction of Mattole River sediments (derived entirely from Franciscan eugeosynclinal rocks: predominantly graywackes, greenstones and cherts) ranged from 5.5 to 8.0 mequiv./100 g (Malcolm and Kennedy, 1970) and that of the sand fraction from twenty-one rivers across the United States ranged from 0.3 t o 1 7 mequiv./100 g (Kennedy, 1965). Kennedy attributed the observed exchange capacities mainly to those fragments composed of clay minerals. The presence of clay minerals, either as matrix or rock fragments, essentially will control the exchange capacity of sands and sandstones. The exchange capacity of the important clay minerals in mequiv./ 100 g at pH 7 are: (1)kaolinite: 3-15; (2) montmorillonite: 80-150; (3) illite: 10-40; (4) chlorite: 10-40 (Grim, 1968). Zeolite minerals, which are occasionally found in some sedimentary sequences, have high exchange capacities ranging from 230 to 620 mequiv./100 g. Feldspathoid minerals, which are probably not very important constituents of arenaceous sediments, possess very high exchange capacities up to 1000 mequiv./100 g (Carroll,1959). The grain size of the constituent minerals. The exchange capacity of a given mineral genekally increas$s as the particle size decreases. This is due to the fact that the broken bonds, which are responsible for almost
CHEMISTRY OF WATERS EXPELLED FROM SANDS AND SANDSTONES
55
all the exchange capacity of quartz, feldspars, and kaolinite, and part of the exchange capacity of other clay minerals, increase as the particle size decreases (Grim, 1968). In the case of kaolinite, Hendricks (1945) has shown that the exchange capacity of kaolinite increases from about 2 mequiv./ 100 g, with a specific surface area of 10 m2/g, to 8 mequiv./100 g, with a specific surface area of 40 m2/g. p H of the solution and type o f the exchangeable cations. The exchange capacity of a given mineral also may vary with the pH of the solution, as well as with the type of exchangeable cation used. The exchange capacity generally increases as the pH increases; the reported exchange capacities are generally for pH 7 (Grim, 1968). The presence and nature of the organic matter in arenaceous sediments. The exchange capacity of organic matter is high. Values of 150 to 500 mequiv./ 100 g have been reported for the organic fraction of some soils. The exchange capacity of particulate organic matter separated from the sand fraction of the Mattole River, California, was 94 mequiv./100 g (Malcolm and Kennedy, 1970). The contribution of the organic matter to the total exchange capacity of the sediments in this river was 15%. Exchange capacity of the organic matter constituted about 70% of the exchange capacity of the medium-grained sand (0.5--0.125 mm) fraction. The organic matter content of these sediments ranged from about 10 wt% in the medium-grained sands to about 3%in the clays and about 0.8%in other sands. The weighted-average organic matter content reported by Kennedy (1965) was 4.2% for the stream clays and about 0.5% for the sand fraction. The exchange capacity of the organic matter, principally kerogen, in shale and sandstone beds is not well known. Organic matter loses a large portion of the active groups (the hydroxyl, carboxyl, amine, and sulphur groups), that are responsible for its high exchange capacity in soils, upon burial. Some of ihese groups, however, are still present in kerogen, which comprises an average of 2% by weight of the shales. The clogging of the mineral exchange sites with kerogen also modifies the exchange properties of sediments. The exchange capacity of orthoquartzites and arkosic sandstones devoid of organic matter is very small and, hence, their membrane behavior is negligible. The exchange capacity of lithic sandstones and graywackes is appreciable and may exceed 1 0 mequiv./100 g in some sediments. Ion exclusion and the double-layer theory Because clay mjherals are responsible for the major portion of the membrane properties of arenaceous sediments, they are used in this section to
Y.K.KHARAKAAND F.A.F.BERRY
56 DOUBLE L A Y E R
D i s t a n c e f r o m C l a y SurfoceFig. 2-5. Concentration of cations and anions in the double-layer. The concentration of cations and anions increases and decreases, respectively, as the surface of the clay platelet is approached.
discuss the factors influencing the chemistry of water expelled from arenaC~OUS sediments. When a clay platelet is placed in a solution, the cations are attracted to its charged surface, whereas diffusion tends to carry them away towards the equilibrium solution where their concentration is lower. Anions are repelled by the surface, but diffusion counteracts the electric repulsion. Thus a “diffuse layer” is formed between the clay platelet and the equilibrium solution in which the concentrations of the cations and anions are higher and lower, respectively, than in the equilibrium solution (Fig. 2-5), and increase and decrease, respectively, towards the surface of the clay particle. A doublelayer, known as the Gouy-Chapman double-layer, is formed by the “diffuse layer” and the charged surface of the platelet. The thickness ( d ) in Angstroms of the electrical double-layer is given by Wiklander (1955)as:
(a)
where n is the dielectric constant of the solution, T is the temperature, u is the valence of the cation, and m is the molar concentration of the outside solution. Double-layer thicknesses for a clay platelet in 0.001 N and 0.1 N NaCl solutions at 25°C are approximately 100 and 10 A, respectively. Detailed discussions 01the double-layer theory are presented by Verwey and Overbeek (1948)and Van Olphen (1963).
CHEMISTRY OF WATERS EXPELLED FROM SANDS AND SANDSTONES
57
The concentration of cations (Ki) and anions (Ej)in the pore solution of a compacted'clay, which is in contact with an outside solution *, can be formulated from the thermodynamic equilibrium of salts in the outside and pore solutions. For a uniunivalent salt this is equivalent to: Pij = Pij
3
(2-2)
where pij and pij are chemical potentials of the ij salt in the outside and pore solutions, respectively. Also, the condition of electrical neutrality requires that for this uniunivalent salt in the pore solution: (2-3) mi =El + X where J f l is the concentration of negative sites in moles/l of pore space. Using appropriate units for the exchange capacity of clays, X- is then given by the following equation (Hanshaw, 1962): capacity X grain density x- = exchange percent (2-4) porosity Assuming the same standard state for the salt in the pore and outside solutions, eqs. 2-2 and 2-3 can be transformed (Marshall, 1948; Walton, 1958; Berry, 1959; Hanshaw, 1962; Kharaka, 1971; and others) to: (2-5) and: (2-6) where ai and uj are the activities, in the outside solution, of the cation i and anion j , and Ti and Tj are their activity coefficients in the pore solution. In as much as the values of Ti and Tj are not known, the values of yi and yj (the activity coefficients of i and j in the outside solution) are used in eq. 2-6 to illustrate the degree of anion exclusion that can occur in geological membranes. For example, with a bentonite sample compacted to 9000 psi, where is about 7, the concentration of C1- within the membrane would be 8* and 8 * when the concentration of the outside NaCl solution is 1Nand 1 10-1 N , respectively.
x-
* Solution surrounding the compacted clay, but which is not affected by the double-layer region of the cIay particles.
Y.K.KHARAKAANDF.A.F.BERRY
58
Composition of solutions squeezed from membrane materials The chemistry of solutions expelled from membrane materials is governed by a number of factors: (1)the concentration of ions in the pore solution (eqs. 2-5 and 2-6); (2) the velocity of the flowing water, which depends on the applied pressure that causes the flow; (3) the electrical interaction of ions with the negative sites on the clay particles; the negative sites not only control the distribution of the ions within the membrane, but also retard the flow of the cations with respect to water; and (4)the electrical interaction of the ions with the “streaming potential”. The “streaming potential” (Fig. 2-6a, b) is caused by the displacement of the double-layer (composed predominantly of cations) by water movement caused by the applied pressure. The effect of the “streaming potential” is the retardation of the flow of cations and the acceleration of the flow of anions within the membrane with respect to the flowing water. One overall effect of this potential is to cause the concentration of the anions in the effluent solution and, because of the requirements of electrical neutrality, the concentration of solutes is higher in the effluent solution than their concentration in the pore fluid. The “fine-pore membrane” model of Teorell, Meyer and Severs with some modifications is generally employed (Walton, 1958; Hanshaw, 1962; Kharaka, 1971; and others) for computing the concentration of solutions flowing through a membrane. This model, which is employed in this study, takes into consideration the concentration of the ions in the pore solution, Negotive poles
N e g o t i v e e l e c t r i c a l charges
Flow direction Fig. 2-6. The negative and positive poles of the “streaming potential” developed by displacement of the double-layer by the water flowing: a. from the middle part of the pore t o both ends; b. from one side of a pore to the other.
CHEMISTRY OF WATERS EXPELLED FROM SANDS AND SANDSTONES
59
the velocity of the flowing water, and the electrical interaction of the ions with the “streaming potential”. The electrical interactions of the ions with the negative sites on the clay particles are ignored in this model. Diffusion in and out of the membrane and the interference (coupling) between the various species in solution are also ignored. The concentration of ions in the effluent solution in this model is determined by their bulk transport by the flowing water, which depends on their concentration as well as on the velocity of the flowing water, and by the interaction with the “streaming potential”. The potential causes a flow of ions relative to the water, which sets up a “streaming current” I, given by Walton (1958)as: I, = x-Fkhdp (2-7) where kh is the hydrodynamic permeability, d p is the pressure differential, and F is the Faraday constant. Inasmuch as:
where d V / d t is the volume of solution flowing per unit of time, then:
Because both the cations and anions contribute to the “streaming current”, if one considers only a l : l electrolyte solution, then: (2-9) I, = I& + I,Ej The ni and Hi are the transference numbers (fraction of the current carried by that ion) of the cation and anion in the pore solution and are equal to:
miiii n. = _ _ q u ,+ m j i i j
(2-10)
and :
(2-11) where Tii and Tij are tLAe mobility of cation i and anion j in the pore solution. The concentrations of the ith and jth ions in the effluent solution due to the “streaming potential” are equal to:
(2-12)
60
Y.K.KHARAKAANDF.A.F.BERRY
and:
sdt n. ‘-(2-13) ’ FdV Concentration of ions in the effluent solution = their concentration in the pore fluid f the “streaming current” contribution, i.e.: (2-14) and :
(2-15) where mfr and mj* are the concentrations of ions in the effluent solution. Substituting I , from eq. 2-8in eqs. 2-14and 2-15,one obtains: (2-16) mp = - iiiX-
mi
and :
mT = mj + RjXElectrical neutrality in the effluent solution requires that: m$
=
mf
(2-17)
= m?
(2-18)
Substituting the values of Hi and Hi from eqs. 2-10and 2-11 in eq. 2-18, the following equation is obtained:
(2-19) The concentration of the effluent solution may be related to the concentration of an outside solution that is in equilibrium with the pore solution by substituting the values of mi and Xi from eqs. 2-5 and 2-6.The resulting equation can be transformed (Walton, 1958;Berry, 1959; Hanshaw, 1962; Kharaka, 1971)to yield: (2-20)
CHEMISTRY OF WATERS EXPELLED FROM SANDS AND SANDSTONES
61
Equation 2-20 is identical with equations relating the molality of effluent solutions ( m g ) with that of input solutions ( m i j )in the case of membrane filtration (Walton, 1958; and others). This indicates that the behavior of membrane materials is similar in cases dealing with filtration and compaction, Probable membrane effects in arenaceous sediments The theoretical concentration of solutions squeezed from membrane materials may be computed from eq. 2-20. For a Wyoming bentonite (composed of montmorillonite) with an exchange capacity of 88 mequiv./100 g and compacted at 10,000 psi pressure, the concentration of the effluent NaCl solution will be 8.5 * and 0.86 moles/l when the concentrations of the solutions (with which the compacted clay was equilibrated before compaction) are 0.16 and 1.65 moles/l, respectively. The membrane effect would be less for membrane materials having a lower exchange capacity. For an illite sample with an exchange capacity of 25 mequiv./100 g and subjected to the same physical and chemical conditions, the concentration of the effluent solutions would be 1.6 and 1.25 moles/l, respectively. The exchange capacity of orthoquartzitic and arkosic sandstones devoid of organic matter is probably less than 1mequiv./100 g. The value of computed from eq. 2-4 will be small and any membrane effects will be negligible in such sandstones. The exchange capacity of subgraywackes and graywackes will probably range from 5 to 20 mequiv./100 g. The value of X- for a graywacke with an exchange capacity of 1 0 mequiv./100 g and a porosity of 14%, is equal to about 2 moles/l of pore space. The concentration of solutions squeezed from such a graywacke will be 3.0 and 1.58 moles/l, when the concentrations of the equilibrated solutions are 0.16 and 1.65 moles/l, respectively. The concentration of solutions squeezed from membrane materials at equilibrium with the same outside solution will decrease with increasing compaction pressures. This decrease is due to the fact that the value of Z(eq. 2-4) for a given material is higher when its porosity is lower. The increased membrane efficiencies with increasing compaction pressures predicted from eq. 2-20, have been confirmed both in filtration experiments (Kharaka, 1971, for example) and in squeezing experiments (Kryukov, 1971; and others). Results from filtration and squeezing experiments both indicate that natural materials are selective in that the degree of retardation of different cations and anions is different in the filtration experiments, and that the relative concentration of cations and anions is also different in solutions expelled at different compaction pressures. Kharaka (1971) and Kharaka and
x-
62
Y.K.KHARAKAANDF.A.F.BERRY
Berry (1973) conducted filtration experiments at laboratory temperatures which show that the relative retardation of monovalent and divalent cations by different clays and a disaggregated shale were generally as follows: monovalent: Li < Na < NH, < K < Rb < Cs divalent: Mg < Ca < Sr < Ba. The relative retardation sequences obtained for anions at laboratory temperatures were variable, but at 70°C, the following sequence was obtained: HC03
CHEMISTRY OF WATERS EXPELLED FROM SANDS AND SANDSTONES
0
I
2
3
POTENTIAL
4
5
GRADIENT,
6
7
63
8
v/cm
Fig. 2-7. Normalized flow rates obtained with different materials as a function of electrical potential gradients. Ffowing solutions consisted of 0.5% by weight NaCl and 0.5%CaC12. (After Chilingaretal., 1970, p. 834.) qt = flow rate after application of direct electrical current; qi = initial flow rate.
the pore-volume reduction is achieved by compaction of those clays. The chemistry of the pore fluids expelled via the clay compaction must be controlled by the membrane environment of those clays. As a result, the degree of membrane influence on the expelled pore-water chemistry might be higher than would be deduced solely from a comparison of ion-exchange capacities. The degree t o which membrane phenomena affect the chemistry of the pore waters expelled from sandstones via porosity reduction will be dependent not only on the relative volume of clay minerals and/or organic kerogen present in the sandstone, but also on the position occupied by these clay minerals and/or kerogen. Membrane phenomena should be of greatest importance in sandstones where clays and/or kerogen are present as a coating on individual sand grains. In such a case, the majority of the water expelled from the sandstone via any mechanism that achieves porosity reduction would be subject t o membrane phenomena because of having to pass through the clay and/or kerogen passageways. Such a sandstone might have a relatively low total exchange capacity but a very large membrane-type effect on the chemistry of the expelled waters. It is probable that sands of this type
\
64
Y.K.KHARAKAAND F.A.F.BERRY
exist. Rittenhouse (1973)described sands in which the grains are surrounded by a clay coating. Sands deposited in waters with large concentrations of humic acid acquire humic acid coatings around individual sand grains. Sands of such origin are believed to have been very important in forming the host rock for the major sedimentary uranium deposits in the Morrison Formation (Jurassic) of the San Juan Basin, New Mexico (Kendall, 1971).
SUMMARY
The chemistry of solutions expelled from arenaceous sediments will depend on the chemistry of their pore solutions and their membrane behavior during porosity reduction. The chemistry of pore waters in arenaceous sediments is variable; water types range from essentially fresh potable waters devoid of dissolved salts in some shallow sand aquifers to highly saline formation waters with a content of total dissolved solids exceeding that of sea water by more than an order of magnitude. The proportion of different ions in the pore waters is also variable; the relative concentrations of C1-, Ca2+,and K" commonly increase with increasing depth of the reservoir from which the water is obtained. The chemistry of solutions expelled from orthoquartzites and arkosic sandstones devoid of organic matter will probably be close to the chemistry of their pore solutions. The membrane behavior of the constituent shales, clays and organic matter in lithic sandstones and, especially, in subgraywackes and graywackes will modify the chemistry of expelled solutions. The degree of this modification and the porosity at which this will occur will depend on the proportion and exchange capacity of their constituent clays and organic kerogen. The position where the clays and/or kerogen are located in the grain-void framework of a given sandstone probably is of critical importance. Finally, it should be stated again that direct experimental and/or field evidence is needed to test many of the conclusions reached in this chapter.
REFERENCES Athy, L.F., 1930. Density, porosity, and compaction of sedimentary rocks. Bull. A m . Assoc. Pet. Geologists, 14: 1-24. Bailey, E.H., Snavely, Jr., *P.D. and White, D.E., 1961. Chemical analysis of brines and crude oil, Cymric Field, Kern County, California. U.S. Geol. Surv. Prof. Pap., 398-D: 306-309.
CHEMISTRY OF WATERS EXPELLED FROM SANDS AND SANDSTONES
65
Berry, F.A.F., 1959. Hydrodynamics and Geochemistry of the Jurassic and Cretaceous Systems in the Son Juan Basin, Northwestern New Mexico and Southwestern Colomdo. Thesis, Stanford Univ., 213 pp. Berry, F.A.F., 1969. Relative factors influencing membrane-filtration effects in geologic environments. Chem. Geol., 4: 295-301. Berry, F.A.F. and Hanshaw, B.B., 1960. Geologic evidence suggesting membrane properties of shales. 21st Int. Geol. Congr., Copenhagen, Rep. Sess., Norden, p. 209. Billings, G.K., Hitchon, B. and Shawe, D.R., 1969.Geochemistry and origin of formation waters in the Western Canada sedimentary basin, 2. Alkali metals. Chem. Geol., 4: 211-223. Birch, F., Schairer, J.F. and Spicer, H.C.(Editors), 1942. Handbook of Physical Constants - Geol. Soc. A m . Spec. Pap., 36: 19-22. Bischoff, J.L. and Ku, T., 1970. Pore fluids of recent marine sediments, 1. Oxidizing sediments of 20° N Continental Rise to Mid-Atlantic Ridge. J. Sed. Petrol., 40: 960972. Bredehoeft, J.D., Blyth, C.R., White, W.A. and Maxey, G.B., 1963. Possible mechanism for concentration of brines in subsurface formations. Bull. A m . Assoc. Pet. Geologists, 47: 257-269. Burst, J.F., 1969. Diagenesis of Gulf Coast clayey sediments and its possible relation to petroleum migration. Bull. Am. Assoc. Pet. Geologists, 53: 73-93. Carroll, D., 1959. Ion exchange in clays and other minerals. Bull. Geol. SOC. Am., 70: 749-780. Chave, K.E., 1960. Evidence on history of sea water from chemistry of deeper subsurface waters of ancient basins. Bull. Am. Assoc. Pet. Geologists, 44:357-370. Chebotarev, I.I., 1955. Metamorphism of natural waters in the crust of weathering. Geochim. Cosmochim. Acta, 8: 22-48; 70-137; 198-212. Chilingar, G.V. and Rieke, H.H. 111, 1968. Data on consolidation of fine-grained sediments. J. Sed. Petrol., 38: 811-816. Chilingar, G.V., El-Nassir, A. and Stevens, R.G., 1970. Effect of direct electrical current on permeability of sandstone cores. J. Pet. Tech., 22: 830-836. Clark, S.P., Jr. (Editor), 1966. Handbook of Physical Constants - Geol. SOC.A m . Mem., 97: 23-24. Clayton, R.N., Friedman, I., Graf, D.L., Mayeda, T.K., Meents, W.F. and Shimp, N.F., 1966. The origin of saline formation waters, 1. Isotopic composition. J. Geophys. Res.. 71: 3869--3882. Collins, A.G. and Bgleson, G.C., 1967. Iodide abundance in oil field brines in Oklahoma. Science, 156: 934-935. De Sitter, L.U., 1947. Diagenesis of oil field brines. Bull. A m . Assoc. Pet. Geologists, 31: 2030-2040. Dickey, P.A., Collins, A.G. and Fajardo, I., 1972. Chemical composition of deep formation waters in southwestern Louisiana. Bull. Am. Assoc. Pet. Geologists, 55: 15301533. Ellis, A.J. and Mahon, W.A.J., 1964. Natural hydrothermal systems and experimental hot waterlrock interactions. Geochim. Cosmochim. Acta, 28: 1323-1357. Ellis, A.J. and Mahon, W.A.J., 1967. Natural hydrothermal systems and experimental hot waterlrock interactions, 2. Geochim. Cosmochim. Acta, 31 : 519-538. Fournier, R.O. and Truesdell, A.H., 1973. An empirical Na-K-Ca geothermometer for natural waters. Geochim. Cosmochim. Acta, 37 : 1255-1275. Graf, D.L., Meents,’ W.F., Friedman, I. and, Shimp, N.F., 1966. The origin of saline formation waters, 3. Calcium chloride waters. Ill. State Geol. Surv. Circ., 397 : 60 pp.
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Y.K. KHARAKA AND F.A.F. BERRY
Graton, L.C. and Fraser, H.J., 1935. Systematic packing of spheres, with particular relation to porosity and permeability. J. Geol., 43: 785-909. Grim, R.E., 1968. Clay Mineralogy. McGraw-Hill, New York, N.Y., 596 pp. Hanshaw, B.B., 1962. Membrane Properties of Compacted Clays. Thesis, Harvard Univ., 1 1 3 pp. Hanshaw, B.B. and Hill, G.A., 1969. Geochemistry and hydrodynamics of the Paradox Basin region, Utah, Colorado and New Mexico. Chem. Geol., 4: 264-294. Hedberg, H.D., 1936. Gravitational compaction of clays and shales. A m . J. Sci., 5th Ser., 31: 241-287. Hem, J.D., 1964. Deposition and solution of manganese oxides. U.S. Geol. Suru. Water Supply Pap., 1667-@:42 pp. Hem, J.D., 1967. Equilibrium chemistry of iron in ground water. In: S.D. Faust and J.V. Hunter (Editors), Principles and Applications o f Water Chemistry. Wiley, New York, N.Y., pp. 625-643. Hendricks, S.B., 1945. Base exchange of crystalline silicates. Ind. Eng. Chem., 37: 6 2 5-630. Hiltabrand, R.R., Ferrel, R.E. and Billings, G.K., 1973. Experimental diagenesis of Gulf Coast argillaceous sediments. Bull. A m . Assoc. Pet. Geologists, 57 : 338-348. Hitchon, B. and Friedman, I., 1969. Geochemistry and origin of formation waters in the Western Canada sedimentary basin, 1. Stable isotopes of hydrogen and oxygen. Geochim. Cosmochim. Acta, 33: 1321-1349. Hitchon, B., Billings, G.K. and Klovan, J.E., 1971. Geochemistry and origin of formation waters in the Western Canada sedimentary basin, 3. Factors controlling chemical composition. Geochim. Cosmochim. Acta, 35: 567-598. Kendall, E.W., 1971. Trend Orebodies o f the Section 27 Mines, Ambrosia Lake Uranium District, New Mexico. Thesis, Univ. California, Berkeley, 273 pp. (unpublished). Kennedy, V.C., 1965. Mineralogy and cation-exchange capacity of sediments from selected streams. U.S. Geol. Surv. Prof. Pap., 433-D: 28 pp. Kharaka, Y.K., 1971. Simultaneous Flow o f Water and Solutes through Geological Membranes: Experimental and Field Investigations. Thesis, Univ. of California, Berkeley, 274 pp. Kharaka, Y.K. and Berry, F.A.F., 1973. Simultaneous flow of water and solutes through geological membranes, 1. Experimental investigation. Geochim. Cosmochim. Acta, 37 : 257 7-260 3. Kharaka, Y.K., Berry, F.A.F. and Friedman, I., 1973. Isotopic composition of oil field brines from Kettleman North Dome, California, and their geokogic implications. Geochim. Cosmochim. Acta, 37 : 1899-1908. Kramer, J.R., 1969. Subsurface brines and mineral equilibria. Chem. Geol., 4: 37--50. Krynine, P.D., 1948. The megascopic study and field classification of the sedimentary rocks. J. Geol., 56: 130-165. Kryukov, P.A., 1971. Solutions Obtained from Rocks, Soils, and Silts. Acad. Sci. U.S.S.R., Siberian Branch, U.S.S.R., 21 8 pp. Kryukov, P.A. and Komarova, N.A., 1954. Concerning squeezing out of water from clays a t very high pressures. Dokl. Akad. Nauk. S.S.S.R., 99: 617-619. Kryukov, P.A., Zhuchkova, A.A. and Rengarten, E.V., 1962. Change in the composition of solutions pressed from clays and ion-exchange resins. Dokl. Akad. Nauk. S.S.S. R., Earth Sci. Sect., 144: 167-169. Langmuir, D., 1969. Geochemistry of iron in a coastal-plain ground water of the Camden, New Jersey, area. U S . Geoi. Surv. Prof. Pap., 650-C: 224-235. Levorsen, A.I., 1967. Geology o f Petroleum. F.A.F. Berry (Editor), W.H. Freeman & Co., San Francisco, 724 pp.
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Malcolm, R.L. and Kennedy, V.C., 1970. Variation of cation-exchange capacity and rate with particle size in stream sediments. J. Water Poll. Control. Fed., 42: R153-R160. Mangelsdorf, P.C., Manheim, F.T. and Gieskes, M.T.M., 1970. Role of gravity, temperature gradients and ion exchange media in formation of fossil brines. Bull. A m . Assoc. Pet. Geologists, 54: 617-626. Manheim, F.T., 1966. A hydraulic squeezer for obtaining interstitial water from consolidated and unconsolidated sediments. U S . Geol. Suru. Prof. Pap., 550-C: 256-261. Marshall, C.E., 1948. The electrochemical properties of mineral membranes, 8. The theory of selective membrane behavior. J. Phys. Colloid. Chem., 52: 1284-1295. Meade, R.H., 1964. Removal of water and rearrangement of particles during the compaction of clay sediments: review. U.S. Geol. Suru. Prof. Pap., 497-B: 23 pp. Meade, R.H., 1966. Factors influencing the early stages of the compaction of clays and sands: review. J. Sed. Petrol., 36: 1085-1101. Middleton, G.V., 1960. Chemical composition of sandstones. Bull. Geol. SOC.Am., 71: 1 0 11-1 026. Mills, R.V.A. and Wells, R.C., 1919. The evaporation and concentration of water associated with petroleum and natural gas. U S . Geol. Suru. Bull., 693: 104 pp. Muffler, L.J.P. and White, D.E., 1969. Active metamorphism of Upper Cenozoic sediments in the Salton Sea geothermal field and the Salton Trough, southern California. Bull. Geol. SOC.A m . , 8 0 : 157-182. Pettijohn, F.J., 1957. Sedimentary Rocks. Harper, New York, N.Y., 718 pp. Pettijobn, F.J., 1963. Chemical composition of sandstones, excluding carbonate and volcanic sands. In: Duta of Geochemistry, 6th ed. - U S . Geol. Suru. Prof. Pap., 440-S: 1 9 PP. Rieke, H.H. 111, Chilingar, G.V. and Robertson, J.O., Jr., 1964. High-pressure (up to 500,000 psi) compaction studies on various clays. 22nd Sess., Int. Geol. Congr., New Delhi, 15: 22-38. Rittenhouse, G., 1973. Pore-space reduction in sandstones: controlling factors and some engineering implications. Offshore Technol. Conf. A m . Inst. Min., Metall. Pet. Eng., Pap., OTC 1806: 1-683-1-692. Rogers, G.S., 1917. Chemical relations of the oil-field waters in San Joaquin Valley, California. U.S. Geol. Surv. Bull., 653: 119 pp. Russell, W.L., 1933. Subsurface concentration of chloride brines. Bull. A m . Assoc. Pet. Geologists, 17: 1213-1228. Schmidt, G.W., 1973. Interstitial water composition and geochemistry of deep Gulf Coast shales and sandstones. Bull. A m . Assoc. Pet. Geologists, 57: 321-337. Shishkina, O.V., 1964. Chemical composition of pore solutions in oceanic sediments. Geochem. Int., 3: 522-528. Siever, R., Kevin, C.B. and Berner, R.A., 1965. Composition of interstitial waters of modern sediments. J. Geol., 73: 39-73. Skempton, A.W., 1944. The consolidation of muddy sediments. Q. J. Geol. SOC.Lond., 100: 119-135. Skempton, A.W., 1953. Soil mechanics in relation to geology. Yorksh. Geol. SOC.Proc., 29: 33-62. Taylor, J.M., 1950. Pore-space reduction in sandstones. Bull. A m . Assoc. Pet. Geologists, 34: 701-716. Terzaghi, K., 1925. Principles of soil mechanics, 1. Phenomena of cohesion of clays; 2. Compressive strength of clays. Eng. News-Rec., 95: 742-746; 796-800. Van Olphen, H., 1W3. An Introduction to Clay Colloid Chemistry. Interscience, New York, N.Y., 301 pp.
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Verwey, E.J.W. and Overbeek, J.T.G., 1948. Theory of the Stability of Lyophobic Colloids. Elsevier, Amsterdam, 199 pp. Von Engelhardt, W. and Gaida, K.H., 1963. Concentration changes of pore solutions during the compaction of clay sediments. J. Sed. Petrol., 33: 919430. Walters, L.J., 1967. Bound Halogens in Sediments. Thesis, M.I.T.,Cambridge, Mass., 221 PP. Walton, H.F., 1958. Principles of osmosis applicable to oil hydrology. (Unpublished research report, Petroleum Res. Corp., Denver, Colo., 66 pp.) Warner, D.L., 1964. A n Analysis of the Znfluence of Physical-Chemical Factors upon the Consolidation of Fine-Grained Clastic Sediments. Thesis, Univ. California, Berkeley, 136 pp. Weddle, J.R., 1967. Oil field waters in southwestern San Joaquin valley, Kern County, California. Calif. Oil Fields, 53: 5-19. Weller, J.M., 1959. Compaction of sediments. Bull. Am. Assoc. Pet. Geologists, 43: 273-310. White, D.E., 1966. Saline waters of sedimentary rocks. In: A. Young and J.E. Galley (Editors), Fluids in Subsurface Environments - Am. Assoc. Pet. Geologists, Mem., 4: 342-366. White, D.E., 1970. Geochemistry Applied to the Discovery, Evaluation, and Exploitation of Geothermal Energy Resources. U.N. Symp., Geothermal Energy, 67 pp. White, D.E., Hem, J.D. and Waring, G.A., 1963. Chemical composition of subsurface waters. In: Data of Geochemistry - U.S. Geol. Sum. Prof. Pap., 440-F: 67 pp. White, D.E., Barnes, I. and O'Neil, J.R., 1973. Thermal and mineral waters of non-meteoric origin, California Coast Ranges. Bull. Geol. SOC.A m , 84: 547-560. Wiklander, L., 1955. Cation and anion exchange phenomena. In: F.E. Bear (Editor), Chemistry o f the Soil -Am. Chem. Soc., Monogr. Ser., 126: 107-148. Wolf, K.H. (Editor), 1976. Ores in Sedimentary and Volcanic Rocks, 1. Elsevier, Amsterdam (in press).
Chapter 3
DIAGENESIS OF SANDSTONES AND COMPACTION KARL H. WOLF and G.V. CHILINGARIAN
GENERAL FACTORS CONTROLLING COMPACTION OF SANDSTONES
Any discussion on the origin of sedimentary rocks should include references to compaction, lithification, and other diagenetic processes as shown in the diagrams of Figs. 3-1 and 3-2.These diagrams demonstrate that compaction is one process that determines the final property of a rock, and close scrutiny of Fig. 3-1,for example, will show that any primary sedimentological factor that controls the original characteristics of the sedimentary deposit
J I
I +
I
L_________________-__----------------J
Fig. 3-1. Theoretical system diagram showing principal processes that affect shallow-water marine sedimentation. Dashed outline separates exogenous processes that supply inputs to the system but do ?ot receive feedback from it. Endogenous processes are inside dashed outline. (After Harbaugh and Merriam, 1968; in Harbaugh and Bonham-Carter, 1970, fig. 7-1, p. 265; courtesy Wiley-Interscience, New York.)
70
K.H. WOLF AND G.V. CHILINGARIAN
Uplift
I
Fracturing
I
!
Fig. 3-2.Diagram of fluvial system. Uplift and climatic factors of temperature, precipitation, and wind are outside the system, feeding into the system, but not receiving feedback from the system. Lithification and compaction are also outside the system, but receive output from the system instead of feeding into the system. (After Harbaugh and Bonhamcarter, 1970, fig. 1-14, p. 22; courtesy Wiley-Interscience, New York.)
will also determine the compaction history of the accumulation (see Introduction chapter). A possible sequence of interdependent, large-scale geologic parameters, which influence compaction directly or indirectly, has been given in Fig. 3-3. Purely theoretical considerations or, as in the sequence in Fig. 3-3, mere “common sense”, may be sufficient to think of other similar mega-relationships. In the future, however, more subtle and quantitative interdependencies must be determined t o make progress in solving complex problems in sedimentology, as has been attempted by Wolf (1973a,b),
TECTONIC-~AO~~~~~~~~,~~-OR EXTERNAL INFLUENCES
A C C u M U L A T I O ~ ~ ~ ~ COMPACTION ~ ~ ~ ~ ~ - , EROSION C-DIAGENESIS
ADJUSTMENT
Fig. 3-3.Flow diagram showing the relationships between various large-scale geologic variables during the development bf a basin and sedimentation that influence compaction diagenesis. These, in turn, have a reciprocal control on the large-scale variables.
DIAGENESIS O F SANDSTONES AND COMPACTION
71
among others. To do this, it seems best to start not with the regional, but with the microscopic and mesoscopic processes and factors involved in compaction. These concepts should then be applied on a regional scale. Review of the literature shows that a large amount of data on compaction is already at hand, especially on the more local scale. The information on the vertical and horizontal variations of the effects of compactional diagenesis, on the other hand, is somewhat meagre, but several examples are available from the literature. In compaction studies of sandstones many factors are to be considered (Table 3-I), many of which are related to the presence or absence of components additional to the sand-sized fraction. For example, the compaction history of a well-sorted sandstone is different from that containing clay matrix. Also, a stratigraphic section composed of sandstone alone may have a much different history of compaction as compared with a section composed of sandstones interbedded with siltstones and/or shales. In an investiTABLE 3-1 Factors and processes controlling compaction of sandstones Group I (Inherited factors1 : individual properties of grains and fluids)
Group II (Inherited factors’ : mass properties)
grain size grain shape grain orientation grain surface features grain rounding grain sphericity grain electrostatic properties grain composition impurities on grain’s surface fluid composition
grainlmatrixlcement ratios shalelsiltstonelsandstone ratios in stratigraphic section absorbability adsorption capacity degree of cohesion size sorting compositional sorting thixotropic properties porosity and permeability grain stacking patterns packing packing heterogeneity grains’ frictional properties grains’ total surface area shear strength surface area configuration of deDosit topography (e.g., slope) of unit sedimentary structures
Group III (Dynamic factors) rate of fluid movement replenishing of fluid removal of fluid fluid pressure overburden pressure subsurface temperature rate of sedimentation (= rate of loading) earthquakes tectonism time Group IV (Inhibitory factors’ : reducing compaction) diagenetic changes neomorphism cementation recrystallization epigenesis (catagenesis) metamorphism
This term was adopted from Coogan and Manus, see Vol. I, Ch. 3.
72
K.H. WOLF AND G.V. CHILINGARIAN
TABLE 3-11 Factors and processes significant in controlling compaction of pyroclastics' Factors
Remarks
(1)Heat retained in the deposit
may cause welding of the particles, thus increasing the strength of the deposit controls the rate of increse in overburden
( 2 ) Rate of settling and accumulation of the pyroclastic debris (3) Distance from source and thus differential sorting of pyroclastic debris (4) Degree of reworking in the depositional environment (5) Amount of rainfall during volcanism (water precipitation is often reported during volcanicity) (6)Amount of ground water
controls thickness of deposit and, therefore, overburden pressure controls textures, fabrics, and structures of deposit controls initial compaction
influences diagenesis, e.g., controls cementation by zeolitization, which increases strength of deposit ( 7 ) Composition (mineralogic and bulk chemi- controls degree of reactibility (or chemi. cal) : cal stability) during diagenesis, as well as (a) acidic versus intermediate versus physical strength prior to and subsequent basic volcanic particles to compaction (b) ratio of phenocrysts/fine debris (c) bubbles versus groundmass in shards determine degree of resistance to mecha(8) Shape, internal textures, and fabrics of vitric shards nical compaction
Most or all of the factors listed in Table 3-1 are applicable here also, but have not been duplicated. Only those of specific importance in pyroclastics have been listed here.
gation of sandstone compaction, therefore, no clear separation can be made between the various lithologies, as illustrated by the many interdependent parameters in Table 3-1. It is also obvious that certain factors of importance in the compaction of pyroclastics either need not be considered in epiclastic (= terrigenous) sediments or are of relatively minor significance (Table 3-11). It may be advantageous to briefly explain compaction of sediments by using the model offered by Terzaghi and Peck (1968; see also Dickey, 1972) and illustrated in Fig. 3-4. A cylinder full of water, containing several perforated plates which are separated by springs, is used in the experiment. The small holes in the plates represent the low permeability of clays, whereas the springs imitate the strength of the clay-mineral accumulations that can sustain the weight of the burial or overburden pressure. At various levels between the plates, the cylinder is connected to manometers measuring the pressure of water. At the start of the experiment, i.e., before loading,
DIAGENESIS OF SANDSTONES AND COMPACTION
73
Fig. 3-4. Terzaghi's model of the compaction process. When a load is applied to the cylinder, the water first carries the full load. As water escapes through perforated plates, the springs sustain the load and water pressure drops t o normal hydrostatic. At time t = 0, pressure (p) is excess hydrostatic; pressure head in manometers ( h l ) in feet is equal to plyw, where p is in lb/sq. ft and 7,,, is the specific weight in lb/ft3. At time t l , pressure drops in the upper part of the model (first three manometers). At time t 2 , pressure drops in all four manometers. Finally, at time t = O 0 , the pressure is normal hydrostatic (level of fluid in all four manometers is at the top of model. (After Dickey, 1972, fig. 3, p. 6; courtesy Int. Geol. Congr., Montreal.)
-
the water in the manometers is at the same level as that in the cylinder (coincides with the horizontal line marked t = in Fig. 3-4).If pressure is suddenly applied to the cylinder (representing overburden pressure), the springs will compress and water will escape from between the plates. As the holes are small, the movement of the fluid will be slow so that compression of the springs is initially prevented and the load is carried by the water itself, leading to excessive hydrostatic pressure. This stage is presented by a horizontal line (t = 0), indicating that the water is at the same level in all the manometers. Subsequently, the fluid will pass out of the upper part of the cylinder and the pressure will decrease successively as time passes from the upper to the lower part, as indicated by the curves t l and t 2 to t,, until the hydrostatic pressure becomes very small and the full weight of the load is carried by the springs. The springs, in this case, represent the framework of the rock which would carry the overburden pressure. In sedimentary basins, the loss of large volumes of water takes a very long time, and excess hydrostatic pressures are common in younger formations, as described in numerous publications. One example, offered by Kidwell and Hunt (1958),is presented in Fig. 3-5: in the Recent sediments of the Orinoco delta at Pedernales, Venezuela, excess pressures increased to 8 psi and above at depths of 120 f t below the surface. The pressure decreased upward to zero at a permeable sandy bed, which allowed quick passage of intrastratal fluids,
74
K.H. WOLF AND G.V. CHILINGARIAN PCH-9
PCH-I1 20 L
60
% I l-
100
a
g
140
Fig. 3-5. Excess hydrostatic pressures (numbers in psi) in the recent sediments of the Orinoco River, Pedernales, Venezuela, As the muds compact, pure water is bleeding upward to a permeable bed at depth of about 30 f t , and also downward to the pre-Paria unconformity, which is also permeable laterally. (From Kidwell and Hunt, 1958; in Dickey, 1972, fig. 4,p. 6 ; courtesy Int. Geol. Congr., Montreal.)
and downward to about 4 psi at an unconformity. The latter permitted bleeding off of the water. DIAGENESIS IN GENERAL*
The widely-accepted division of sedimentary processes into syngenesis, diagenesis, and epigenesis is inadequate for detailed research. The stages at which most of the individual processes and products occur cannot be distinctly demarcated. Two or more processes may be active simultaneously. They may overlap or the termination of one may mark the commencement of another process; and still other alterations may occur independently in both space and time. Many of the interpretations depend on the scale - thin section, hand specimen or outcrop -- at which observations are made. Hence, “pigeonholing” of processes without contradictions is difficult, sometimes even impossible. Difficulties in genetic interpretations occur in particular in mono-mineralogic rocks, such as fine-grained quartzites. For these and other reasons, it is not surprising that no agreement has been reached on the definition and extent of diagenesis (see review by Teodorovich, 1961). Diagenesis has been restricted to those processes that cause lithification. *In dealing with the ge’neral aspects of diagenesis of sandstones, the section presented here has been adopted with minor changes from Chilingar et al. (1967).
DIAGENESIS OF SANDSTONES AND COMPACTION
75
Such a limited application, however, is arbitrary, artificial and impractical.
Not only are there several distinctly different lithification processes which are frequently difficult to recognize and separate, but they are so gradational as to defy precise definition and can occur at any stage during the early history of sediments. It is virtually impossible, therefore, to exclude from
diagenesis other early alteration. It is preferable to apply diagenesis in a wider sense to processes that affect a sediment after deposition and up to, but not beyond, lithification andfor filling of voids. Although these latter processes can, and usually do, take place at different times within a sedimentary formation, especially if composed of different facies, the final stage of lithification and/or filling of voids appears to be the most convenient time at which diagenesis can be terminated. * Hence, the following rather all-inclusive definition, in general agreement with the concepts of Ginsburg (1957) and Krumbein (1942),has been adopted here: diagenesis includes all physicochemical, biochemical, and physical processes modifying sediments between deposition and lithification at low temperatures and pressures characteristic of surface and near-surface environments. Postlithification processes grade into epigenesis, and epigenesis passes into metamorphism.** ’Epigenesis near the depositional environment is called juxta-epigenesis (“juxta-” meaning near), and epigenesis remote from the surface is named apo-epigenesis*** (“apo-” meaning far, remote). Most coarser-grained sediments have some small voids which have been partly or wholly filled by one or more generations of cement. Hence, it is possible in some cases to divide diagenesis into pre-, syn-, and postcementation stages. In other paragenetic investigations, however, the diageneticepigenetic boundary may have to be based on some other criterion to be determined by the individual investigator concerned. No definite rule is possible. As long as the boundary is precisely defined by certain fabric or structural relations, little confusion should occur. A diagenetic-epigenetic *This approach has been found to be of particular use in coarse-grained or open-textured sandstones, conglomerates and limestones, but may be more difficult to apply to finegrained rocks. Nevertheless, the paragenetic model used here is convenient, because after lithification (= infilling of voids by cement) of a rock, the intrastratal fluids related to surface conditions cannot penetrate readily the rock framework. In cases where sediments maintain their porosity and permeability for a long period of time, even after being far removed from the original depocenter, the intrastratal fluids occupying the cavities can also be looked upon as either of syngenetic, diagenetic or epigenetic origin. **Strictly speaking, epigenesis as defined here, passes into metamorphism only if an increase of pressure and/or temperature occurs. ***Weathering not related to the original depositional environment of the sediments is not included in diagenesis and epigenesis as defined here. The gradational frontiers between deposition (= syngedesis), metamorphism, and weathering have been illustrated by Dunoyer de Segonzac (1968) in Fig. 3-6.
K.H. WOLF AND G.V. CHILINGARIAN
76
boundary established on the basis of a few thin sections of a local outcrop, however, may have t o be revised and shifted up or down the paragenetic scale as soon as the petrologic and petrographic information of the whole formation is available. On the other hand, it may be found that the termination of diagenesis in one area may be completely unrelated to that of other localities. Not all diagenetic stages are present in sediments. For example, precementation dolomitization, with the formation of dense and relatively impermeable dolomite, may completely alter the calcareous components of a limy sandstone, resulting in an absence of the syn- and postcementation stages. Postcementation dolomitization may obscure the various cementation stages. The raw material of diagenesis, as Krumbein (1942) called it, consists of organic and inorganie sediment of allochthonous and/or autochthonous origin, interstitial fluids, and other components subsequently formed or introduced into the system. In general, it is possible t o subdivide the components that interact during the diagenetic processes into the following (Wolf, 1963): (1) diagenetic-endogenic; (2) diagenetic-exogenic: (a) supergenic-exogenic; (b) hypogenic-exogenic. This is merely an expansion of Amstutz’s (1959) division: syngeneticsupergenic, syngenetic-hypogenic, epigenetic-supergenic, and epigenetic-. hypogenic. In most cases, diagenesis derives its raw material from both endogenic (within the sediments) and exogenic-supergenic (outside source-from above) sources. One or the other may prevail. Under unusual conditions, however, a volcanic, i.e., exogenic-hypogenic, source may supply components for diagenesis without a marked increase in temperature. This would be particularly true for siliceous material introduced into a geosyncline (or other depocenter); during diagenesis of the sediments, quartzose arenites can WEATHERING
DEPOSITION
METAMO~~~ISM
Fig. 3-6. Diagenetic frontiers. (After Dunoyer de Segonzac, 1968, fig. 1.)
DIAGENESIS OF SANDSTONES AND COMPACTION
71
be converted t o orthoquartzites, carbonates can become siliceous, and fossils can be replaced prior to dolomitization. Diagenesis may express itself in a number of different ways. Krumbein (1942) mentioned that a total of about thirty separate diagenetic processes have been described in the literature. They may result in mineralogical changes, addition and removal of material, and textural and structural modifications and alterations ranging from slight to extensive or complete. Several generations of diagenesis may each leave evidence, or each successive one may obliterate or destroy the products of earlier processes. In many cases, however, diagenesis may appear to be absent if only visually-obtained information is considered. * The division of sedimentary processes into several subdivisions is suitable to our present purpose and state of knowledge of sandstone diagenesis, including compaction. It should be pointed out, however, that the pre-, syn-, and postcementation sub-groups will become arbitrary and artificial with an increase in our understanding of diagenesis. To ascribe to certain products merely a genetic term such as “precementation-diagenetic”, without relating it to the sediment’s history as a whole, invites criticism. It is more accurate to relate all processes and products t o a paragenetic sequence. In other words, a paragenetic scheme furnishes less ambiguous information in cases where it seems impossible to define exact syngenetic-diagenetic-epigenetic boundaries. The absolute time of formation may be impossible to determine, but the textural and structural relationships permit the interpretation of relative time of formation. The widely used terms “primary” and “secondary” have very little meaning in diagenetic investigations unless precisely defined, although they may be quite useful in a very general colloquial sense. Paragenetic interpretations are relatively easy and noncontroversial if the investigation is made on the scale of one thin section or hand specimen. Syn-, dia-, and epigenetic processes, however, are not only gradational, and overlap in time and space on a microscopic scale, but especially do so on a regional scale. Regional diagenetic studies may be rather tedious and resemble structural analyses, for example, in that micro-, meso-, and macroscopically examined features are assembled step by step. The so-called predepositional (or presyngenetic) processes and products can be deduced from rock fragments, which were derived from older sedi*It should be noted that diagenesis is not part of the petrographic (= descriptive) stage, but belongs to the subsequent stage of petrology and petrogenesis (= interpretive), although diagenetic data is often made part of petrographic descriptions. Reliable diagenetic reconstructions cannot be made, therefore, on the basis of the study of a few local thin sections, but must be based on as much geochemical, petrographic and stratigraphic information as circumstances permit to be obtained.
K.H. WOLF AND G.V. CHILINGARIAN
78
mentary, volcanic, plutonic, and metamorphic rocks. The sedimentary and volcanic rocks may have undergone diagenesis before erosion and transportation and, therefore, show features that suggest the conditions of secondary processes in the source-rock environment. Factors controlling diagenesis Certain factors will initiate diagenesis, and the same or other factors will perpetuate the old and/or cause commencement of new diagenetic processes. The sediments have a tendency to adjust to new physical and chemical conditions and would, theoretically, reach equilibrium. The micro- and macroenvironmental conditions above and within the sediments, however, change continuously. Sometimes, equilibrium may be established, as, for example, in cases where unstable feldspar grains are completely replaced by clay minerals. In many cases, however, the physical and chemical conditions shift so rapidly that only a small fraction of the reaction involving the sedimentary framework reaches equilibrium. In particular during the early diagenetic stages numerous successive and overlapping processes will be acting at a relatively fast rate on both micro- and macroscales, when movements of interstitial fluids are at a maximum, biological activity is producing chemically active substances, maximum pore space is available, and temperature change is more or less sudden due to diurnal exposure. The following factors influence diagenesis of sediments: (1)geographic factors (e.g., climate + humidity + rainfall + type of terrestrial weathering surface water chemistry); (2) geotectonism (e.g., rate of erosion and accumulation, coastal morphology, emergence and subsidence, whether eugeosynclinal or miogeosynclinal); (3) geomorphologic position (e.g., basinal versus lagoonal sediments + current velocity + particle size sorting + flushing of sediments); (4) geochemical factors in a regional sense (e.g., supersaline versus marine water and volcanic fluids and gases); (5) rate of sediment accumulation (e.g., halmyrolysis ion transfer + preservation of organic matter + biochemical zonation); ( 6 ) initial composition of the sediments (e.g., aragonite versus high-Mg and low-Mg calcite and isotope and trace-element content); (7) grain size (e.g., content of organic matter -+ number of bacteria + rates of diffusion); (8)purity of the sediments (e.g., percentage of clay and organic matter -, base exchange of clays altering interstitial fluids); (9) accessibility of salldstone framework to surface (e.g., cavity systems permit replacements) ; -+
-+
-+
DIAGENESIS OF SANDSTONES AND COMPACTION
79
(10) interstitial fluids and gases (e.g., composition, rate of flow, and exchange of ions); (11)physicochemical conditions (e.g., pH, Eh, partial pressures of gases, and COz content); (12) previous diagenetic history of the sediment (e.g., previous expulsion of trace elements will determine subsequent diagenesis). The numerous large-scale environmental parameters listed above influence in one way or another the more local environments and these, in turn, influence the microenvironments. There is a complete gradation and overlap of these macro- and microfactors as one example below illustrates: climate + geomorphology particle size
4
amount and type of bacteria
4
rate of diagenesis
4
pH and Eh
4
type of replacement The actual processes that lead t o diagenetic alterations and modifications of sediments can be divided as follows. (1) Physicochemical processes: solution, corrosion, leaching, bleaching, oxidation, reduction, reprecipitation, inversion, recrystallization, cementation, decementation, authigenic mineral genesis, crystal enlargement, replacements, chemical internal sedimentation, and aggregation and accretion. ( 2 ) Biochemical and organic processes: accretion and aggregation, particlesize reduction, corrosion, corrasion, mixing of sediments, boring, burrowing, gas-bubbling, breaking down and synthesizing of organic and inorganic compounds. ( 3 ) Physical processes: compaction, dessication, shrinkage, penecontemporaneous internal deformation and corrasion, and mechanical internal sedimentation. Many of the above processes are commonly considered syngenetic. As they can occur within the sediments and directly alter and influence diagenesis, however, they must be considered as part of diagenesis. It is the total or collective influence of all factors that must be examined in a final analysis. As Krumbein (1942)pointed out, variations in the diagenetic end-products may occur either-with different sediments in the same environment, or with the same kind of sediment in different environments.
TABLE 3-111 Various lithogenetic stages as defined by different authors in English-Americanliterature (after Dunoyer de Segonzac, 1968, table 4, p. 189)
-
T h e detrttal particles still in movement in the water
1 -*
a
in a sediment wlth o high water content but isolated from the environment of sedimentation
T h e sediment has became a'more ar less compact rock
TWENHOFEI
PETTIJOHN
1926. 1939.195
(1949.1957)
I
1
I
WILLIAMS TURNER GILBERT (1954)
I
iI
,-on,,y
I
(1960)
j
T Depor,r,on
DlAGENESlS
-I-----I
~
I
~
~
~
.
.
I
I
METAMORPHISM
. . . .
I lrrh,frcol,on
'
I' . . . . . 1: . . .. . .. . . . __._.... ~ ~ ._._____.__ ~ _ .~ . ~ ._... . . .. . . . . . . . . The sedimentary . ... . . . . . . . series finds itself . ,. . ,. . . . .. . .... .. ... . , . .. .. . . . .. .. . under metomarphic . ., . . . . .. . .. . ... .. .. ... conditions on account ......... : ... . . . . .. .. .. . .. .. .. . ..... .. .. ... ...... . of orogeny . . .. .. ... .. .. . .. . ............... . . .. . .. . .. . . . ,
-
'
'
_ _ _ ~ ~ ~ ~
'
-
Tectonic phenanena place the sediments r r r r p y w under conditians of Xofomorph,sm decompression and 'del!lk,ficoIron, leaching. 01 exposed in outcrops
t
* In reality parts 11 and 111 represent very unequal thicknesses of sediments, tens of meters for part 11, hundreds or even thousands of meters for part 111.
DAPPLES (1959)
I
DAPPLES (1962)
DIAGENESIS OF SANDSTONES AND COMPACTION
81
As pointed out already, the lithogenetic stages used in the study of sedimentary rocks have been established differently by various investigators, an excellent summary of which has been provided by Dunoyer de Segonzac (1968). His summary tables, giving the stages of lithogenesis, are presented here in Tables 3-111, 34V, and 3-V, as based on the English, German, and Russian literature, respectively. Compaction is a diagenetic process and one may even speak of “compaction diagenesis”. This book is devoted mainly to this particular process, which may occur either as a physical and/or physicochemical and biochemical phenomenon. Compaction of sediments is the process of volume reduction expressed either as a percentage of the original voids present or of the original bulk volume. Although the process affects mainly loose, unlithified sediments, it may also have profound influences on well-cemented deposits, TABLE 3-IV Various lithogenetic stages as defined by different German researchers (after Dunoyer de Segonzac, 1968, table I, p. 159)
*
In reality parts IT a d 1 1 1 represent very unequal thicknesses of sediments:tens of meters for part 11, hundreds o r even thousands of meters for part 111.
TABLE 3-V Russian nomenclature related to diagenesis (after Dunoyer de Sgonzac, 1968, table 2, p. 168) - -. ~~~~~
AUTHORS
I F LITHOGENESIS
I
F ERS MAN j1922)
PUSTOVALOV
S H V ET SOV
TEODOROVICH
RU K H I N
(1933, 1 9 4 0 )
(1934, 1957)
(1961)
(1961 )
The detrital particles stlll in movtrnent in the water Particles immobilized in o sediment with o high water content but isolated from the environment of sedimentaiion
.*.
................
T -
The sediment hos become a more ar less comDact r o c k
SEDIMENTS
T h e sedimentary series finds i t s e l f under metamorphic conditions on accouni of orogeny
v*
Tectonic phenomena place the sediments under conditions of decompression and leoching. or exposed in outcrops
I n reality parts I 1 and 111 represent very unequal thicknesses of sediments:trns of meters for part 11, hundreds or even thousands of meters for part 111.
VA S SOEV I C ti (1962)
1
(1958,1963)
I
DIAGENESIS OF SANDSTONES AND COMPACTION
83
as will be discussed below. The intergranular spaces of clastic and detrital sediments are eliminated by closer packing, crushing, deformation, expulsion of fluids, and, possibly, dissolving of grains. Cemented sediments may undergo compaction through solution along stylolites, for example. Krumbein (1942) gave the following average values of porosities of freshly-deposited material: sand = 45%, silt = 50-65%, mud = 80-90%, and colloids (less than 1mu in diameter) = approximately 98%water. According to Miiller (1967, p. 135), the initial water content of argillaceous muds is approximately 50-80%, which corresponds to a porosity of 70-90%; whereas porosity for sands is only 30-5096, which corresponds to 20-30% water content. S e m y a (1969), however, found an average of about 150% water* (on dry weight basis) for the upper 200 cm of the lemanic sediments of Lake Geneva. She also reported that the water content of the very first layer of sediments was 250% $ when measured on an undisturbed sample. The degree of compaction generally depends partly on the ratio of fine to coarse material and on the character of the sediment framework. Inasmuch as mixtures of fine and coarse clastic grains are quite common, in this chapter a reference will be made occasionally t o clay-sized material, though the book is devoted to the compaction of coarse-grained sediments. One should note here, that the behavior of a mixture of different grain-size classes may or may not lie somewhere between the behavior of the separate grain-size classes, but little is known about this subject and more research work is required. Although compaction is part of diagenesis, if one makes compaction the “center of consideration”, then diagenesis could be subdivided into pre-, syn-, and postcompaction stages (Table 3-VI). On first sight this may look to be an artificial division; however, in detailed work on, for example, the relationship between fluid movements during diagenesis and their influence on the origin of various types of chemical cements, including metalliferous ores (see Chapter 5), such an approach should prove to be useful. Little is known about the diagenetic features formed by “water of compaction”, in contrast to those produced by other varieties of intrastratal fluids and mixtures of solutions of different derivation. Consequently, in future investigations some type of genetic classification of cements and other diagenetic features formed by different types of solutions and a corresponding nomenclature will become necessary. The various sources of chemicals that form cements in sandstones are presented on p. 76 above; the reader is also referred to Chapter 5 on ore genesis in sediments by compaction fluids. Compaction, as illustrated and defined in Tables 3-111, 3-IV, 3-V, and 3-VI, is not clearly defined by an upper and lower time- and/or space-bound*Weight of water divided by the weight of dry solids.
03
TABLE 3-VI General relationship of the process of compaction to diagenesis DIAGENESIS
lb
1 - -+
EPIGENESIS I
J
METAMORPHISM
+
Precompaction
_2
-
syncompaction -
*
noncompactional fluid movements in general (i.e., fluids of various origins)
postcompaction
-
-1
WEATHERING
-
e--
<
“water of compaction” movements
-
-chemical precipitation from all varieties of fluids
,
x
I=
chemical precipitation from compaction fluids
DIAGENESIS OF SANDSTONES AND COMPACTION
85
ary, so that such tables must be used with extreme caution. The rate of
compaction and the decrease in porosity and permeability, as well as the rate of expulsion of fluids, will change with time both vertically and horizontally in sedimentary basins. In a study of release of chemical elements, e.g., lead, into compaction fluids and their migration from the fine-grained into coarser-grained sediments, one would have t o consider the changes in the properties of the sediments during compaction with geologic time, and any “definition” of compaction-diagenesis and its paragenesis is to be taken as a guide rather than a rule. The amount, rate, and mechanism of release of fluids are controlled by: (1)changes in permeability, which also controls the rate of flow of fluids; (2)temperature increase during burial; and (3) dissociation of water during temperature and pressure increases with increasing depth of burial (Blatt et al., 1972, fig. 7.1).Several case histories are presented in the various sections of this chapter. The study of diagenesis has been mainly an observational science until relatively recently when geologists and geochemists have obtained data on natural and artificial (= laboratory) diagenetic, chemical systems. More quantitative data is available now on diagenetic processes than ever before. The interpretation of these data will be assisted by the development of new theories. Berner (1971, 1972), for example, presented theoretical models, e.g., chemical kinetic models for steady and non-steady diagenetic processes. He also quantified, or offered formulae for, total compaction, rate of compaction, rate of flow through a horizon, and total volume of water passing through the sediments. DIAGENESIS IN SANDSTONES
The literature on the various aspects related to sandstone diagenesis is voluminous (e.g., see Pettijohn et al., 1972) and for the purposes of the present chapter, the publications by Dapples (1962,1971,1972)are particularly useful. On the basis of mineral associations, intergrowth, and replacements, he recognized that four oxide series among the sandstones are not purely composed of silica (or quartz) but have other constituents: (1)alumina, lime-magnesia, iron oxide series; (2) silica, lime-magnesia, iron oxide series; (3) silica, alumina, iron oxides series; and (4) silica, alumina, limemagnesia series. Based on simple mineralogy, he listed (see Table 3-VII) the recurrence of certain types of mineral associations showing secondary growth and mutual replacement. On using the four oxide series, Dapples (1962, figs. 1 to 4, pp. 915-929) showed that chemical components are combined to form specific minerals, which he discussed in detail. One should note that in his diagram, Dapples listed minerals (sometimes end-member
86
K.H. WOLF AND G.V. CHILINGARIAN
TABLE 3-VII Minerals showing mutual intergrowth and replacement in poorly-cemented sandstones of simple mineralogy (after Dapples, 1962, table 1, p. 915) ( 1 ) Quartz-hematite
(2) Quartz-calcite (3) Quartz-hematite-calcite-siderite ( 4 ) Quartz-chert-clay mineral (5) Quartz-chert-clay mineralsiderite (6) Quartz-glauconite-clay mineral (7)Quartz-calcite-clay mineral (8) Quartz-glauconite-clay mineral-calcite
minerals with minerals of transitional composition in between) that are stable under near-surface, early-burial, and late-burial conditions. Dapples (1962) explained that according t o the type of chemical reactions and the major sequence of occurrence, three diagenetic stages are recognizable (Table 3-VIII): Stage l-redoxomorphic, which comprises the episodes of sediment accumulation and early burial. The principal chemical reactions are reduction and oxidation. Stage 2-locomorphic, involving the replacement of one mineral by another without the minerals entering into mutual reactions, which is typical of the early-burial stage and forms an important part of the process of lithification. Stage 3-phyllomorphic, occurs subsequent to the locomorphic replacements and involves the origin of micas, mainly as a transformation product from clay minerals. In sandstone petrology, the progress of diagenesis may be considered t o be terminated at either the locomorphic or phyllomorphic stage. It must be recognized, however, that the phyllomorphic stage overlaps with the zeolite and chlorite stages of burial or low-grade metamorphism. Many other investigators have discussed the problem of gradational change from diagenesis into metamorphism. Blatt (1966) mentioned that, according to metamorphic petrologists, diagenesis per se stops short of the zeolite facies in tuffaceous sandstones and before recrystallization of quartz in orthoquartzites, arkoses and lithic sandstones. These limits imply temperatures of approximately 10°C to 200 f 50°C and pressures of one bar to about 2000 bars. Pettijohn et al. (1972) used the term diagenesis in a wider sense, namely, that it affects the sediments up to the lowest grade of metamorphism (the greenschist facies). There is no definite demarcation between diagenesis and metamorphism. As a sandstone is buried deeper and deeper in a sedimentary basin and heat and pressure increases, a stage will be reached where one can call the rock either sedimentary or metamorphic depending on the classification and nomenclature adopted. Both diagenetic and metamorphic stages
DIAGENESIS OF SANDSTONES AND COMPACTION
87
TABLE 3-VIII Reactions characterizing diagenesis (after Dapples, 1962, table 2, p. 931) ~
Redoxomorphic stage (reversible reactions)
~~~
Locomorphic stage (replace- Phyllomorphic stage (unidirecment reactions) tional reactions)
+ muscovite
aragonite. by calcite
“clay minerals” or biotite
Hematite + calcite T?.siderite
calcite by dolomite
montmorillonite -+ chlorite
Hematite + calcite + Mg”’ siderite + ferrodolomite + dolomite
carbonates by quartz or chert
“clay minerals”
*
-+
chlorite
Hematite + clay minerals + quartz or chert by carbonates “clay minerals” + Fe+* silica + chlorite + greenatite lite + stilpnomelane (?)
-+
bio-
Hematite + chlorite + chamo- feldspar by carbonates site (?)
“clay minerals” + chert rite
Hematite + illite + glauconite
quartz, chert, clay minerals by carbonates
“clay minerals” + quartz -+ sericite
“Bauxite” + silica nite
opal by chert or quartz
kaolinite + glauconite or. illite + Mg+2 -+ chlorite +
silica solution =+chert or quartz
glauconite
+ kaoli-
K+
-+
-+
chlo-
feldspar
Diaspore or boehmite + silica + clay minerals
illite or glauconite + muscovite
Diaspore (?) + silica + K+ + clay minerals (glauconite)
kaolinite + illite + glauconite + calcite + Mg+2 -+ muscovite + biotite + feldspar + dolomite + chlorite
Bauxite + hematite + silica (minor) =+clay minerals + pyrite
illite or glauconite + calcite + Mg+2 + micas + feldspar + dolomite
Kaolinite + K+f i illite
feldspar
Biotite
plagioclase -+ chlorite + chert
* glauconite
Feldspar chert
=+clay minerals +
Glass montmorillonite + chert -+
-+
-+
chlorite
sericite
88
K.H. WOLF AND G.V. CHILINGARIAN
leave imprints on rocks, e.g., on graywackes. As Pettijohn et al. indicated, although there is a diagenesis-metamorphism continuum in physical variables, e.g., temperature and pressure, there are important differences between the two stages. Metamorphic petrology involves the genetic interpretations of secondarily formed mineral assemblages, controlled by bulk chemical composition, as indicators of pressures and temperatures, assuming equilibrium or a close approach to it. Very often there is no original sedimentary feature left, except under certain circumstances, e-g., in the case of burialmetamorphosed graywackes. This approach results in an order of mineral assemblages in the metamorphic rocks expressed either by isogrades, metamorphic facies or petrogenetic grids. On the other hand, the sandstones, even after diagenesis, are composed of mineral assemblages that tend to reflect the composition of the original, obviously non-equilibrium detrital mixtures more than the effects of pressure and temperature, because the maximum values of the latter two variables involved in diagenesis are much lower than in metamorphism. In addition, the rates of reactions of mineral neomorphism are slow at low temperatures. Certain equilibrium assemblages of very low-grade burial metamorphism (considered late diagenetic or catagenetic stage by some investigators) may be the result of changing composition of intrastratal fluids rather than the increase in temperature and pressure, e.g., in pyrodastic rocks. Pettijohn et al. (1972) offered a diagrammatic representation of six stages t o which a sandstone is exposed during burial (Fig. 3-7). This type of scheme is particularly useful, because they attempted to assign clay-mineral reactions for each stage as shown in Table 3-IX. For a good summary of sandstone diagenesis the readers are also referred to an excellent book by Pettijohn (1972, pp. 383-437), who divided the evidence of diagenesis into: (1) textural, (2) mineralogical, (3) physical, and (4) chemical. As shown later in separate sections, compaction and/or compaction fluids are involved in these diagenetic modifications. Pettijohn et al. (1972) also discussed the composition of sandstones in some detail. They stated that the silica minerals in sedimentary rocks are either clastic or of a secondary chemical and/or biochemical origin. The secondary chemical and biochemical processes can also alter clastic silica grains by dissolving, corroding and etching them. As conceptually shown in Fig. 3-8, the chemically formed silica is represented by the following minerals: opal, chert, or quartz (see also Carozzi, 1960, pp. 291-292, among others, for details). It is important to know that the source of S i 0 2 is fourfold: biogenic silica, volcanic glass, various silicates, pore waters, and detrital quartz grains during pressure solution. All these may release SiOz into compaction fluids which, in turn, may precipitate silica (and other minerals) when exposed to a different chemical milieu. Similar considerations apply to the origin of authigenic feldspars in sandstones which require
DIAGENESIS OF SANDSTONES AND COMPACTION Depth of burial
89
Microscopic Appearance Immediately after deposition. Ex& Original detritus. hieh porosity.
to air or water
of depo%itmnalenvironment
Buried a few meters to tens of meters. Exposed to Interstitialwaters Some wmpaclion. some wrly chemical precipitates possible.
Buried to modcrate depths of about ID00 m. Pore water may be a brine Chemical Cements may reduce porosity. days may be altered.
5 q burial to thousands of meters perhaps with folding. Porosity may bc "cry low
from chemical osminl and pressure solution.
Incipient metamorphism. Growth of chlorite and other metamorphic minerals wnlh extensive pressure solution and quaNitic texture.
@in and erosion. within tens of meters ol land surface. Invasion by meteoric watei. dcccmcntation and "wuthcrin8" of clays may increasz porosity. After
Fig. 3-7. The stages of diagenesis in relation to depth of burial and increase of pressure and temperature. (After Pettijohn et al., 1972, fig. 10-1, p. 387; courtesy Springer, New York.)
a suitable solution, possibly mobilized by compaction, to form overgrowths of albite or K-feldspar on clastic feldspar grains, for example (Fig. 3-9). Chemical conditions for the precipitation of feldspar are determined by the pH and relative amounts of various components in solution, i.e., K+,Na+, Mg2+, Ca2+, and Si02. The alkali metal/hydrogen ion ratio, at a minimum amount of SiO14- in solution, is indicative of the stability of feldspar with respect to a solution and thus of the possibility of its precipitation (Garrels and Christ, 1965, pp. 359-363; Pettijohn et al., 1972, p. 38). Hemley and Jones (1964)suggested that slightly elevated temperatures, that are associated with moderate to deep burial, are important in precipitating feldspar. Diagenetic processes are not uniform and regular as demonstrated by sandstones that are hundreds of millions of years old and only slightly cemented in contrast to more recent, well-cemented sediments. Also, several sandstones with the same degree of lithification may have had very different post-lithification * histories. In general, during increasing alteration of different types of sandstones, there is an eventual convergence to a chemical
K.H. WOLF AND G.V. CHILINGARIAN
90 TABLE 3-IX
Some clay-mineral reactions during sandstone diagenesis (after Pettijohn et al., 1972, fig. 10-2, p. 431) Clay mineral formed
Precursor
Components added to (+) or subtracted from (-)
Stages of diagenesis (see Fig. 3-7)
Kaolinite
feldspar
1,2,6
Kaolinite Illite
pore space kaolinite
Muscovite
kaolinite
Illite
montmorillonite
--(K+,Si02) +HzO* +(A1203, SiOz, H2O) +(K+, Si02) -(A1203, HzO) +K+ -Hz 0
Chlorite
montmorillonite
Montmorillonite
volcanic glass
Glauconite
illite
+K+
-(SiOz, HzO, Na+, Ca2+, Mg2+, Fe2+, etc.) +(Fez+, Mg2+) -(SiOz, HzO, Na+, Ca2+)
+HzO
-(Na+, K+, Ca2+) +(Fez+, Fe3+) -(K+, A1203 1
* 2KAISi308 + 2H+ + 2HCOT + 9 H 2 0
+
2,6 3,4,5 5
3,4 3,4,5 1,2,3,4 192
AlzSizOs(OH)4 + 4H4Si04 + 2K+ + 2HC03.
equilibrium in both mineralogic composition and texture (as long as the original bulk chemical composition is not very different), but certain initial differences can persist into the higher grades of alterations. A fresh sand is a porous, non-equilibrium assemblage of clasts the composition of which is I
Pressure SOlUllO"
Volcan,c glass and other silicate\
iJIACIENF5IS
PROVFYAhCF
DIAGENESIS OF SANDSTONES AND COMPACTION
91
1
Fig. 3-9. The origin o f feldspar in sandstones. (After Pettijohn et al., 1972, fig. 2-4, p. 38; courtesy Springer, New York.)
“derived” from the source areas. During diagenesis, that includes compaction, cementation, and other alterations of various types, there is a loss of unstable detrital grains and an increase in stable authigenic components. After long and deep burial a quartzitic arenite would become a well-cemented quartzite (= quartz clasts cemented by quartz cement). Lithic arenites will show complete cementation by a combination of quartz, carbonate and clay minerals, the latter commonly represented by an illite-chlorite assemblage. (This complex material makes it difficult to draw a definite boundary between “cement” and “matrix”, as discussed below.) The final products of all the secondary alterations of sandstones reflect the temperature and pressure increases, composition of pore fluid chemistry, original composition of the sediment and its texture, and geologic time. The total post-depositional geologic history of the sediment was involved. Fiichtbauer and Muller (1969) presented the textural and mineralogical results of sandstone diagenesis (Table 3-X), including the influence of compaction. Each one of the diagenetic processes is treated at some length in subsequent sections of this book. Practical applications of detailed diagenetic and related studies of sandstones are exemplified by the work of Griffiths (1964, and his numerous other publications). Some examples are provided in this chapter. Griffiths (1964, p. 640) investigated several thousand core samples of barren and oil-producing sedimentary rocks, and showed that the latter exhibited specific characters which differed in degree from those of the oil-bearing but non-producing sands. Both producing and non-producing oil-bearing sands,
K.H. WOLF AND G.V. CHILINGARIAN
92 TABLE 3-X
Textural and chemical-mineralogical results of sandstone diagenesis (free translation of table 3-14,p. 105 in Fuchtbauer and Muller, 1970) Chemical process
Example
Influence on: porosity compaction
Dissolution (a) pressure-solution (b) intrastratal solution
pressure-quartzite dissolved heavy minerals
(-1 strongs ( + ) filledS
strong none
( + ) strong
none
Neoformation (partly cementation) (a) autochthonous1 sand grains grew at the expense of siliceous matrix homogene2 quartz grains dissolved at pressure points; SiOz precipitated in pressure shadows (b) autochthonous kaolinite formation in pores heterogene3 from feldspars, which were dissolved in other parts of the sandstone (c) a l l o c h t h ~ n o u s ~ silicification as a result of SiOz homogene supply from an external source cementation by anhydrite (d) allochthonous heterogene
,
Replacement (a) autochthonous (b) allochthonous
* Autochthonous
kaolinization of feldspar dolomitization of orthoclase
(-) doubled
( + ) strong
none
(-) filled
none
(-) filled
none
( + ) strong none
none none
= derived from the components of the sandstone (= internal source); homogene = neoformation of a mineral already present; heterogene = neoformation of a type of mineral not previously present in the sandstone; allochthonous = delivered from an outside source; ti (-) = reduction in porosity and (+) = increase in porosity
*
in turn differed from the barren sediments. Griffiths stated that this is an example of different textural properties of the sediments showing complex interdependent relationships (see also Griffiths, 1961). According to him (1964, p. 640), laboratory experiments have shown that sands having different grain sizes, exhibit different saturation characteristics (Griffiths, 1957):
DIAGENESIS OF SANDSTONES AND COMPACTION
93
“when a fluid-saturated sand is invaded by a second fluid, the second fluid is preferentially concentrated in the coarse-grained sands. For example, in the secondary uranium ores on the Colorado Plateau, when differential saturation occurs, the ore is selectively confined to the finer-grained sands. Since the ‘second fluid’ is expected to saturate the coarse-grained layers, this suggests that uranium saturation was original and the second fluid removed the ore from the coarser-grained sands.” Uranium concentrations have been found in coarser silts associated with very fine sands in the Entrada Formation, whereas they occur in very fine sands associated with fine to medium sands in the Salt Wash Member of the Morrison Formation. Uranium is also concentrated in fine-grained gravel to coarse-grained sands associated with coarser gravels in the Shinarump deposits. (Jobin, 1962, presented similar explanations for this distribution pattern.) In some cases reported by Griffiths (1964,p. 641), carbonate cement occurs in the coarser-grained sandstones, showing that the carbonate-bearing solutions invaded the sedimentary pile displacing earlier fluids. Griffiths stated: “It may well be argued that the controlling factor is not really grain size but some property which varies in a similar manner; however, since the variation in grain-size is closely associated with that in the controlling factor, grain-size variation is a convenient and efficient means of following the critical changes which are associated with, and in some complex way control, the differential saturation.” Of course, other variables are influential in addition to grain size, e.g., amounts and compositions of matrix and cement, sorting, orientation, packing, and degree of compaction, in determining porosity and permeability. But the variations in the several parameters, many of which are related to grain size, are interrelated and interdependent. As Griffiths (1964,p. 642) suggested, the question is how much additional effect, to that supplied by changes in grain size, is supplied by variation in the type and amount of cement and/or matrix and several other textural parameters. All of them should be considered step-by-step to determine their influences on porosity and permeability. Compaction, of course, which is our main interest in this book, will have to be considered here also. In order to evaluate the influences of several petrographic parameters on the mass properties of sediments, Griffiths (1964,p. 642) proposed to use statistical methods to determine: (a) the most important controlling property; (b) how much control this property exercises over the mass properties; (c) what additional information on control is supplied by adding new properties; and (d) what is the most parsimonious combination of properties which accounts for the greatest degree of control over porosity and permeability, for example. The degree of predictability and the interrelationships are determined by means of multiple regression and by using discriminant equations. The latter method was employed by Griffiths (1964,p. 649) in the
94
K.H. WOLF AND G.V. CHILINGARIAN
TABLE 3-XI Comparative measures of interrelationships between porosity and petrographic variables in some Berea oil-producing sands and in some barren Pocono sandstones (after Griffiths, 1964, table 1, p. 644) Variable
Degrees of freedom
~~
Berea
Pocono
0.7647*** 0.0265 0.15 08 * 0.0034 0.0011 0.16 31 0.1833*** 0.0149 0.08226
0.0062 0.4700 0.0055 0.4141 0.3995 0.4393 0.2891 0.1586 0.2108
Percent variation explainedl Berea
Pocono
62.0 2.2 12.2 0.3 0.1 12.4 89.2 10.8 100.0
0.2 14.9 0.2 13.1 12.6 13.9 54.9 .45.1 100.0
~
Packing Additional for grain size % Quartz Orientation Size sorting Shape Cumulative total Residual Total 1
Mean square
1 1 1 1 1 1 6 9 15
*
Percent variation = Sxf/Sx,2,t.
example given in Table 3-XI. Samples from the Berea oil-producing sandstone were compared with those of the barren Pocono sandstones, using the following petrographic properties: (1)mean proportion of quartz; (2) grain size (Le., the average length of the long axis of quartz grains); (3) size sorting (i.e., the standard deviation of long axis of quartz grains); (4) shape (i.e., the ratio of short over long axis of the quartz grains); (5) perfection of preferred orientation (i.e., standard deviation of axial inclination in degrees); and (6) packing of the quartz grains. Upon testing, Griffiths (1964, p. 643) was able to reduce the discriminant equation based on parameters 1 to 6 to an equation where packing, shape, and grain size led to maximum discrimination. Thus, the evaluation of these three parameters is adequate to differentiate the oil-bearing sandstones from the barren sandstones. The porosities differed in these two sets of samples and separate multiple regressions showed that the relationships between porosities and the petrographic variables also differed in the two cases: the porosity and petrographic variables of the Berea specimens showed a high degree of predictability, which was absent in the case of the Pocono sandstones (Table 3-XI). The data of the Berea sandstone indicated that the packing, proportion bf quartz, and grain shape “explains” or “accounts” for 88.9% of the variation in porosity, whereas in the Pocono specimens all above-mentioned six properties account for 54.9% variation. It is particularly interesting to note that no single variable controlled significantly the predictability of the porosity. Together with other
DIAGENESIS OF SANDSTONES AND COMPACTION
95
studies, it is evident that certain properties, in this case porosity, is controlled by other different properties in different sandstones. The above is a particular case of investigation, but Griffiths (1964) went beyond it to find a more general model, which is of special relevance to the present theme of the chapter. He pointed out (p. 647) that from studies of the petrology of detrital sediments, the dominant factors are source area, processes of sedimentation, and diagenesis. According t o him, a model can be established in which these three factors may be related to the measured properties of the sediment. This model can be used to compare roles played by different properties under varying genetic circumstances. Table 3-XI1 is an idealized model based on first approximation in which, for example, the percentage of clastic components is related t o source area and weathered source area materials, but is independent of other factors and properties. In reality, however, the simple relationships presented in Table 3-XI1 are more likely to be more complex, as shown in Table 3-XIII. In this table, the useful property of independence is largely lacking and, as Griffiths pointed out: “Effects of subsequent factors modify the relationships of properties to factors and may completely subdue the original relationships. Interdependencies among the properties are thus the rule rather than the exception. It would be useful, therefore, to find an analytical tool which could in some way separate out the various effects and, where appropriate, introduce indeTABLE 3-XI1 Idealized component analysis reflecting relationships among provenance, genetic process and properties of sediments (after Griffiths, 1964, table 2, p. 648) Property
Model character (C)
C1 source area Proportion of: Detritus Matrix Cement
c3
c2
weathering
xxxx xxxx
erosion
transportation
deposition
diagenesis
0 0 0
0 0 0
0 0 0
xxxx xxxx
0
0 0
0 0
Size Sorting
0 0
0 0
xxx
xxx 0
xxx xxxx
0 0
Orientation Packing
0 0 ‘
0 0
0 0
0 0
xxx xxx
xxx xxx
0
K.H.WOLF AND G.V. CHILINGARIAN
96
TABLE 3-XI11 Relationships among provenance, genetic process and properties of sediments (after Griffiths, 1964, table 3, p. 649) Property
Model character (C)
source area
c3
C1
c2
erosion
transportation
deposition
xxxx xxx xx xxx X
X X 0
X
X
X X 0
Size Sorting
X X
X X
xx xx
xxx xxx
Orientation Packing
0 0
0 0
0 0
X 0
xxxx xxxx xxxx xxx
Proportion of: Detritus Matrix Cement
weathering
xx X
diagenesis
X
xxx xxx x x xx xxx
pendence, so that the effects of the various factors on the variation in the measured properties may be separately evaluated.” Component analysis, which is one technique among many in factor analysis, is such a tool (Griffiths, 1962). In its use, the parameters are reduced to three for simplicity, i.e., C1,C2,and C3 in Tables 3-XI1 and 3-XIII. An example is given in Tables 3-XIV, 3-XV, and 3-XVI. The matrix of coefficients of correlation of zeroorder are given in Table 3-XIV for eleven properties of a Mississippian quartzose, low-rank graywacke (see Pettijohn’s classification in the Introduction chapter) of the Maxton Sandstone. Table 3-XIV indicates that there are: (a) a typical high correlation ( r > 0.95) between long “a” and “b” axes of quartz grains; (b) moderately high degree of association between the standard deviations (8, and 6,) of these axes, i.e., among size sorting of these axes; and (c) association between porosity and the log of permeability ( r = 0.85). Apart from these features, it is difficult to make an interpretation of the other data because of complex interdependencies present among the measured variables. When the data is treated by component analysis, the characteristic roots (last row in Table 3-XV) demonstrate that five components account for some 90% of the variation, i.e., the relationships among the eleven variables can be represented in terms of five components. As shown in Table 3-XV, the first component (C,)“explains” 50.6% of the variation in the relationships among the variables, the second component (C,) an
TABLE 3-XIV Matrix of correlation coefficients for eleven properties of the Maxton Sandstone, Mississippian, West Virginia (after Griffiths, 1964, table 4, p. 652) Mineral
Quartz Rock fragments Matrix Silica Xa
zb
aa &b
8
Quartz Rock frag- Matrix ments
i
Silica
Grain size
1
Size sorting
(%a
(jZb
(6,)
($1
Orientation (6")
Porosity
Log. permeability
~
-0.230
-0.441
-0.256
-0.186
-0.147
1
-0.512 1
-0.505 0.382 1
-0.580 0.811 0.342 1
0.581 0.781 -0.336 0.993 1
0.048
0.049
0.555 -0.347 -0.322 -0.263 -0.276 1
0.427 -0.274 -0.345 -0.255 -0.280 0.724 1
Porosity Log. permeability For n = 33, rij 2 0.344 significant at Po5 level; subscripts a and b refer to long and short axis, respectively.
-0.225
-0.044
0.534 -0.471 -0.190 -0.558 -0.589 0.241 0.172 1
0.702 -0.667 -0.606 -0.709 -0.700 0.416 0.185 0.477 1
0.096 0.575 -0.617 -0.678 -0.617 -0.605 0.386 0.217 0.306 0.850 1
98
K.H. WOLF AND G.V. CHILINGARIAN
TABLE 3-XV Component matrix of Maxton sandstones (after Griffiths, 1964, table 5, p. 654) Component
C1
Variable Quartz Rock fragments Matrix Silica Grain size a Grain size b Sorting a Sorting b Orientation Porosity Permeability Variance accounted for
103(rij‘) = relative loading -504 784 145 351 -229 7 95 341 -140 -8 29 -095 -379 -631 325 147 -867 287 177 -865 637 285 561 645 346 461 0 28 -554 608 -076 -076 880 -071 098 814
c3
c2
50.6
64.4
77.3
c4
c6
213 -119 -242 5 20 -232 -246 205 364 216 -321 -436
145 -01 5 022 -369 111 089 -161 062 458 -1 34 -184
86.5
90.8
additional 13.8%,the third component (C,) another additional 12.9%, and so forth. All components beyond the fifth (i.e., 6 to 11 or C6 to C l l ) are considered to represent random variation totalling some 9%. To attempt an interpretation of this data as based on the idealized models of Tables 3-XI1 and 3-XIII, Griffiths (1964) offered the simplified Table 3-XVI. According to him (p. 655): “The largest loading in the first component is average grain size and so this component is considered to represent effects of the factors erosion, transportation, and deposition. Since both proportion of rock-fragmen& and matrix show their heaviest loadings on this component also, it is deduced that variation in these detrital constituents is largely affected by selective sorting on the basis of size. If this interpretation is correct, then the matrix material is largely detrital in origin.” Griffiths pointed out further that inasmuch as porosity and permeability also show their heaviest loadings on this component, then these dependent properties are most closely associated with the average grain size. Thus, the porosity and permeability of this reservoir rock is largely dependent on the same factors that determined the size distribution, and the other properties are affected to some degree by the effects of the factors dominating this component. Griffiths also showed that the second component exhibits its heaviest loading on the two measures of size sorting. This component is independent of (and orthogonal to) the first component and, consequently, this may be interpreted as a subsidiary but important effect of selective sorting, i.e., the first two components both reflect selective sorting on the basis of size and shape, the one representing,
DIAGENESIS OF SANDSTONES AND COMPACTION
99
TABLE XVI Simplified component matrix of Maxton sandstones (after Griffiths, 1964, table 6, p. 656) Component +
1
4
5
X
+
X
+ +
+
2
3
xx
+++
Variable Quartz
+ +
+++
Rock fragments
xxx xx xxxx xxxx
Matrix Silica Size Z, fb
Porosity
++ + ++ ++++
Permeability
+++
Sorting ab
Orientation
+ + ++ ++
xx
+ X X
Explanation: (rij’)1o3
>850 700-850 500-700 250-500 <250
positive
++++
+++ ++ +
blank
negative
xxxx xxx xx X
for example, the average current velocity and the other the range in current velocities. Griffiths (p. 657) stated that ‘‘. . . the proportion of quartz possesses a heavy loading on this second component and so a large part of the variation in quartz proportion is a reflection of control by factors affecting fluctuations in current velocity. Again grain size and proportion of rock fragments and matrix are affected to some degree by the component. Since some two-thirds of the total variation (64.4%) is accounted for by these first two components, the remaining components are not nearly so distinct. It seems likely that as proportion of silica cement is dominant in the 4th component and present in the 3rd and 5th, diagenesis is the factor reflected in these components. The 3rd possesses its heaviest loading on quartz with a strong second on perfection of orientation, which suggests that the factors involved are processes operating at a late stage in deposition or early (physical?)
100
K.H. WOLF AND G.V. CHILINGARIAN
diagenesis. Porosity and permeability are also affected by the factors present in the 4th component, or essentially by diagenetic processes connected with the distribution of silica cements.” The reasons for presenting the above information from Griffiths’ publication is: (a) to demonstrate the complex interrelationships of numerous factors, which are discussed in this and other chapters; (b) to show the necessity of using statistical techniques; and (c) to point out that compaction will also have to be given consideration in the future investigations that range over thick vertical stratigraphic sections. The amount of matrix and cement and the degree of orientation and packing of grains, for example, may be controlled by compaction and interactions of sediments with compaction fluids. One useful approach may be that of combining Griffiths’ techniques with detailed statistical textural and fabric studies, e.g., determination of types of grain contacts (see pp. 133-163). CEMENTATION OF SANDSTONES CONTROLLED BY COMPACTION FLUIDS
The chemical precipitation of cements within the framework of sedimen-
tary deposits is dependent on a supply of chemical elements by intrastratal
solutions, usually, moving solution. Although it is conceded that compaction fluids are not the only types of fluids that can cause cementation, the processes of compaction and lithification must be fully comprehended in the investigation of diagenesis in general, in order to be able to determine not only the direct influences of compaction on cementation, but also to establish, for example, at what stage of compaction the sedimentary deposit was lithified. To provide some of the fundamental data on cementation processes, the following paragraphs will discuss a number of aspects. Although no direct reference may be made to compaction itself, it is understood that all the solutions thought to have been responsible for cementation in the cases treated, could be considered to have originated during compaction of a sedimentary unit under natural conditions. At the same time, all other conditions, such as an increase in pressure and temperature for example, can be visualized to be the consequence of burial in a sedimentary basin. One may justifiably inquire about the relationship between the amount of cement, porosity, and permeability, on one hand, and compaction, on the other. Although at the present time quantitative answers are not available, it is clear that the relationships between the amount of cement and porosity and between porosity and permeability are a function of the degree of compaction of the sediment. The stage of compaction when cementation occurred and the amount of porosity reduction as a result of compaction should be considered.
DIAGENESIS OF SANDSTONES AND COMPACTION
101
In a publication on carbonate cements in quartzose sandstones, Dapples (1971) listed two main groups, namely, (1)cement that did not cause destruction of the grain-supported framework, and (2) cement that did (Table 3-XVII). This subdivision indicates that in detailed studies of compaction features in sandstones it is of fundamental importance t o recognize the various textural relationships between the grains and the cement (and matrix, if present), because the time of cementation and the degree of compaction are some of the factors that control the fabric relationships between the framework and the secondary minerals. In particular one should note that Dapples’ group 11-A (Table 3-XVII) incorporates four varieties of cements that result in volume expansion during crystallization (see Spry, 1969, pp. 149-152). Such expansion has been noticed also by Wolf and Ellison (1971) in Pliocene volcanic arenites and rudites which were never covered by a thick overburden. It seems that this expansion as a result of precipitation of cement can only take place during early diagenesis, as long as the sedimentary unit is near the surface, i.e., when the overburden is relatively thin and compaction is at a minimum. Experimental work may eventually determine the exact values of pressures involved in crystallization and the maximum depth of burial at which expansion as a result of cementation can occur. Griffiths (1958) successfully attempted to relate porosity to petrography TABLE 3-XVII Physical classification of carbonate cement in quartzose sandstone (after Dapples, 1971, table 1,p. 197)
I. Cementation without destruction of the grain-supported framework: A. crystallization as a single event: (1) filling interstitial space as a simple adhesive: (a) without mineralogic reaction; (b) incorporating clay minerals already present in intergranular space B. crystallization as part of dual events: (1)filling pore space following partial welding of quartz grains ( 2 ) recrystallization into single or compound crystals following welding of cement to clastic carbonate grains in a mixture of quartz and carbonate grains
11. Cementation with destruction of the grain-supported framework:
A. by crystallization and volume expansion: (1)in interstitial pores to produce local expansion (2) in porous grains, causing rupture and rotation of such grains (3) around sand nuclei to form small concretions in situ ( 4 ) as large poikilitic crystals B. by replacement of silicate-mineral grains: (1)without destruction of bedding (2) by developmint of large concretions
K.H. WOLF AND G.V. CHILINGARIAN
102
t POROSITY, %
Fig. 3-10. Permeability (maximum) versus porosity in Cow Run Sand, West Virginia. (After Griffiths, 1958, fig. 1, p. 16; courtesy4 Sed. Petrol.)
and has found that in many oil sands porosity and permeability are related exponentially, so that when the latter is plotted on a log scale and porosity on an arithmetic scale, a linear relationship emerges (e.g., Fig. 3-10). Griffiths then concluded that as a first approximation any relationship between porosity and petrography also applies to the log of the permeability. Inasmuch as these interdependencies are not exact, they cause a scatter of the data around the trend line, as in Fig. 3-10, for instance. One should note that the relation breaks down in the 0-10% porosity and less than 2 millidarcies permeability ranges. As demonstrated in Figs. 3-llA,B, carbonate cementation caused the reduction in porosity, so that the degree of cementation would appear to be one of the most important petrographic parameters to consider. According to Griffiths (1958), there is an inverse relationship (approximately linear) between the porosity and the carbonate content (determined by employing an alkalimeter). It is also independent of geographic and stratigraphic location in this particular study; however, this need not be so in other cases, because the percentage of cement, as well as its type, often were directly controlled by regional variations in the sedimentary and diagenetic milieu. Below 5%, the porosity varies independently of the carbonate cement con-
DIAGENESIS OF SANDSTONES AND COMPACTION
I
1
1
1
x- Calculated o = Observed
,
103
E
Fig. 3-11. Relationship between porosity (Y)and percentage of carbonate cement ( X ) for Cow Run Sandstone, West Virginia. ( A after Griffitths, 1958, fig. 8 ; B after Griffiths, 1967a, fig. 21-5; courtesy McGraw-Hill, New York.)
tent and, in these instances, grain size, size sorting, and amount of silica cement are the main factors controlling the porosity of the sandstones. The maximum porosity of this sandstone is around 25%and when the carbonate
104
K.H. WOLF AND G.V. CHILINGARIAN
content exceeds this amount, there is no obvious relationship between carbonate content and porosity. The data discussed above was obtained from measurements on specimens from the same rock formation, which most likely have the same compactional history. Future investigations should include attempts to establish correlation among the degree of compaction, depth of burial, porosity (including “minus-cement” porosity; see Pettijohn et al., 1972,p. 424), and the amount of cement. Several examples of the regional distribution of cements in sandstones are presented here, which reflect upon the present state of knowledge. Mellon (1964)reported that discriminatory analysis of calcite- and silicacemented phases on the basis of their compositional and textural properties indicated that grain packing and/or pore size were the only variables associated with the distribution of cements. In his study, calcite-cemented phases were more loosely packed than adjacent silicate-cemented phases (i.e., chlorite, illite, quartz), but grain size, grain orientation, and detrital composition showed no significant association with the cement distribution independent of packing. Calcite-cemented sandstones tended to be coarser-grained on the average than silicate-cemented ones because of the marked correlation between pore size and grain size. The chlorite was formed first, entirely enveloping the more tightly packed constituents of the sandstone. Illite and quartz genesis subsequently filled the remaining parts of the larger pores. The passage of later carbonate-bearing fluids w a s controlled predominantly by the distribution pattern of the earlier cements. Of these, the calcite was preferentially precipitated in the more permeable, originally loosely packed framework of the sandstone, at the same time replacing some of the detrital and authigenic components. Thus, it appears that compaction has had an indirect influence on the distribution of the diagenetic cements, because it controlled the degree of grain packing and the porosity prior to the chemical precipitation of the calcite. But other factors also have to be considered. Garrison et al. (1969)showed that the composition of neither the Fraser River water nor the sea water of the same area could account for the precipitation of the early diagenetic carbonate cement in sandy sediments they investigated. They believed that the dissolution of carbonate shells by subsurface fluids and the precipitation higher in the stratigraphic section from the expressed water of compaction could have been responsible. As Pettijohn et al. (1972)mentioned, calcite cement replacing secondary quartz must be a later diagenetic product related to the redistribution of calcium carbonate within the sedimentary pile. The solubility of carbonates decreases with increasing temperature and increases with greater pressure, although the latter has a much smaller effect. Thus, the net effect of burial is to decrease the solubility of carbonates. This could account for some of the cement, but
DIAGENESIS OF SANDSTONES AND COMPACTION
105
unless large quantities of solutions are passed through the system, only small amounts of secondary carbonates would be produced. Burial results in an increase in pressure-solution, an explanation usually offered for the solution-deposition sequence of silica cement, but which can also be advocated for carbonate cementation. According to Pettijohn et al., the hydrostatic pressure effect on the solubility of carbonates is greater than that of quartz, which is reflected in the common occurrence of late carbonate cement preceding quartz cementation. One may mention here also that in coarse clastic units, higher salinity and higher temperatures and pressures favor the formation of anhydrite over gypsum cement. No data seems to be available on the direct control of compaction on the composition of evaporite cements. Heald et al. (1962) presented an interesting approach in their study of the origin of interstitial porosity in an ancient sandstone by testing the hypothesis, advanced by many investigators, that porosity is often of a secondary type as a result of leaching of carbonate (i.e., decementation) below an unconformity. The small calcite cement patches at corners in pores, observed by Heald et al., were believed to have been vestigial, i.e., residue left after dissolution. Evidence, however, was presented against the leaching concept, based on the presence next to the pore walls of quartz with unblemished secondary faces. Adjacent to calcite, the faces were imperfectly developed, so that one can conclude that leaching of this carbonate would have left voids lined by blemished quartz faces. Heald et al. gave several reasons to support their argument as stated above, including some based on laboratory experimental data. They treated the sandstone samples with acid to dissolve the carbonate before examining them to determine the proportion of the surface of each quartz grain covered by unblemished faces. The results are shown in Figs. 3-12 and 3-13,using a “crystal face index” from 0 to 10. 1
10,
CRYSTAL FACE INDEX (after artificial leaching)
Fig. 3-12. Relation between natural porosity and crystal face index after artificial leaching. I = Eastern outcrop area; 2 = Kanawha Co. subsurface samples; 3 = Tucker Co. subsurface sample. (After Heald et al., 1962, fig. 4, p. 294; courtesy J. Sed. Petrol.)
K.H. WOLF AND G.V. CHILINGARIAN
106
0
' 8
I 6
4
2
CARBONATE. % (before artificial leaching )
0
I
2
4
6
8
NATURAL POROSITY, %
Fig. 3-13. Relation between carbonate content (and natural porosity) and crystal face index. (After Heaid et al., 1962, fig. 5, p. 295; courtesy J. Sed. Petrol.)
Zero indicates that the unblemished quartz faces were absent, whereas 10 is based on the highest proportion. Figure 3-12demonstrates that the unblemished crystal faces increased with porosity, indicating that the quartz faces in openings are well formed in contrast to the quartz terminations in contact with other grains. Heald et al. reasoned that if the observable quartz faces had been cleaned by leaching under natural conditions, then artificial leaching in the laboratory would have revealed more of the faces. The samples which were calcareous prior t o artificial leaching, however, had no higher crystal face index after leaching in comparison to the low-carbonate samples of equivalent porosity (Fig. 3-13). To rule out leaching before cementation, which is much more difficult to accomplish, Heald et al. considered the fact that most of the sandstone samples contain clastic carbonate grains (e.g., fossil fragments). If the porosity of the sandstone had been the result of leaching of the carbonate clasts, dissolution would have been widespread; but no evidence was found to support this. Although the clastic carbonate grains were typically mediumsand sized, open or cement-filled voids of this size were lacking, so that an origin of pores by removal of the clastic grains was not acceptable. Compaction could not have changed the shape and size of the voids because the more resistant detrital quaftz grains of the framework would have prevented it. In their section on porosity reduction, Heald et al. reasoned that because
DIAGENESIS OF SANDSTONES AND COMPACTION
107 SW *o. UC
It
N I,
-c
22
** I-
.?)
L 0 0
24
2 IT
28
t*
0
I1 . i
Fig. 3-14. Clastic constitutents, cement and porosity in the Oriskany Sandstone. (After Heald et al., 1962, fig. 7 , p. 297; courtesy J. Sed. Petrol.)
leaching was not a factor in the origin of pores, the porosity was controlled by incomplete cementation and degree of compaction, and presented evidence for their conclusion. The highly calcareous samples had a high primary porosity and consisted of well-sorted quartz grains and clastic carbonate constituents. Secondary cementation by sparite and quartz eliminated the porosity. The proportion of carbonate cement (= sparite) increased with primary carbonate clast content, most likely a result of solution and redeposition of the original carbonate. Variations in secondary quartz and carbonate sparite cements and pressure solution account for variation in the porosity of the more quartzose sandstone samples. In Fig. 3-14,the differences in the ratio (clastic grains)/ (cement + porosity) seem to be mainly due to the differences in the degree of chemical, pressure-solution compaction. This is supported by the number of contacts per grain and the degree of suturing, which are higher in specimens with higher proportions of clastic grains. The samples with 6% total carbonate content have relatively good porosities (Fig. 3-15).As Heald et al., however, pointed out, the relative importance of the two diagenetic factors in reducing porosity was difficult to assess. Heald and Renton (1966)reported on experiments carried out on sandstone cementation and its relation t o porosity changes. The tests were performed at 225"--360°C and at pressures from 2000 to 11,000 psi in hydrothermal reactors. The results, however, may apply to lower temperature and pressure conditions, if one considers geologic time as an additional influen-
K.H. WOLF AND G.V. CHILINGARIAN
108
6
P 0 R 0 S I T I (XI
I
Fig. 3-15. Relation between carbonate content and porosity in the Oriskany Sandstone, Kanawha County, West Virginia. (After Heald e t al., 1962, fig. 8, p. 298; courtesy J. Sed. Petrol.)
tial parameter. Particularly interesting were the observations on the control of cementation by variations in size, shape, angularity, and composition of the sand grains used during the tests, which demonstrated that the amount of artificial cementation was influenced by the original properties of the clastic components as well as by the variation in influx of cementing material. The initial rates of precipitation of the cement in fine-grained sand were greater than the rates in coarse-grained samples. As cementation proceeded, however, the coarser-grained sand cemented faster than the finer ones, because the permeabilities of the latter were appreciably reduced after a moderate degree of cementation. Heald and Renton’s experiments were divided into cementation of monocrystalline grains, polycrystalline grains, and mixtures of grains (e.g., quartz grains; lithic fragments of fine-grained sandstones, quartzites and cherts; and grains of arkose and micaceous quartzite).
109
DIAGENESIS OF SANDSTONES AND COMPACTION
Monocrystalline grains The following results were obtained by Heald and Renton from their experiments performed on monocrystalline grains: (1)The experiments on the relative growth rates of precipitated quartz on round quartz grains of different sizes are shown in Fig. 3-16.The gain in weight percentage due to cementation increased with decreasing grain size, but the ratio of the growth rates remained essentially constant. The greater rate of growth on the smaller grains is a reflection of the larger specific-surface area of the sand sample. (2) The rate of addition of secondarily precipitated cement is a function of the concentration of the solvent, but the ratio of the growth rates is independent of the concentration of the solution. (3)As Fig. 3-17 demonstrates, the relative growth rates of secondary quartz was the same in Na2C03 and NaOH solutions.
m
W
0
10
20
30
40
PERCENT GROWTH -STANDARDS
50
50
PERCENT GROWTH -STANDARDS
Fig. 3-16. Growth relationship between 9-10 mesh, 12-16 mesh, and 21-32 round quartz grains. (After Heald and Renton, 1966, fig. 2, p. 979.)
mesh
Fig. 3-17. Growth pf 12-16 mesh grains compared to standard grains in solutions of 0.03 M NaZC03 and 0.25 M NaOH. (After Heald and Renton, 1966, fig. 3, p. 980; courtesy J. sed. Petrol. )
K.H. WOLF AND G.V. GHILINGAREAN
110
)
-
P E R C E N T GROWTH STANDARDS
PERCENT GROWTH - STANDARDS
Fig. 3-18. Growth relationship between angular quartz grains of 9-10 mesh, 12-16 mesh and 21-32 mesh size. The 9-10 mesh grains and standard grains were contained in separate baskets, whereas in the charges of smaller grains, the standard grains were dispersed through the samples as internal sbandards. (After Heald and Renton, 1966, fig. 5, p. 981; courtesy J. Sed. Petrol.) Fig. 3-19. Growth relationship between aitgular and round grains of 12-16 (After Heald and Renton, 1966, fig. 6, p . 981; courtesy J. Sed. P e t r d . )
mesh size.
(4) It is noteworthy, that when 0.25-M solutions of Na2C03 were used, cryptocrystalline quartz was deposited, but in a reduced, 0.03-M solution normal quartz overgrowths were the result. (5) Figure 3-18 indicates that the precipitation of cement was faster on smaller angular than on larger angular grains. (6) Figures 3-19 and 3-20 prove that the relative amount of cement is consistently greater in the case of angular in contrast to round grains, when variations in circulation of the solutions and temperatures remain small.
Polycrystalline granular material
The conclusions to be drawn from Heald and Renton’s experiments on polycrystalline granular material are:
bIAGENESIS OF SANDSTONES AND COMPACTION
80
111
1
j-
PERCENT GROWTH - SEEDS
PERCENT GROWTH -BEREA POLYCRYSTALLINE
Fig. 3-20. Growth relationship between angular and round grains of 21-32 mesh size. (After Heald and Renton, 1966, fig. 7 , p. 981; courtesy J. Sed. Petrol.) Fig. 3-21. Growth relationship between granules of fine polycrystalline quartz and granules of coarse quartz from the Berea Sandstone. (After Heald and Renton, 1966, fig. 10, p. 983; courtesy J. Sed. Petrol.)
(1)The growth rate of quartz cement on very coarse grained quartz was nearly twice that of the growth rate on grains of fine-grained polycrystalline quartz fragments, as shown by the curve in Fig. 3-21 with all values along the y-axis being nearly twice that of the x-axis. (2) In experiments on monocrystalline quartz, quartzite and polycrystalline chert fragments of the same size and shape (and surface area, therefore), the growth rates in Fig. 3-22 demonstrate that (a) the rate for quartzite was low but increased with time. The slow rate of growth seems to be a reflection of the small crystal size of the polycrystalline grains. As the cement enlarged the small crystals with a cansequent increase in surface area, faster growth was promoted. (b) The lower values of growth rate for monocrystalline quartz was- the result of reduced circulation of the solution, because of porespace reduction by cementation. (c) The chei-ts grew even more slowly during the cementation experiments, the finer chert showing the lowest rate.
112
K.H. WOLF AND G.V. CHILINGARIAN
PERCENT GROWTH-POLYCRYSTALLINE QUARTZ
DURATION, Hours
Fig. 3-22. Growth relationship between monocrystalline and polycrystalline quartz grains.
1 = fine chert; 2 = medium chert; 3 = quartzite; 4 = monocrystalline quartz. (After Heald
and Renton, 1966, fig. 13, p. 985; courtesy J. Sed. Petrol.,)
Fig. 3-23. Porosity change during cementation of grains of quartzite and monocrystalline quartz. 1 = quartzite; 2 = quartz. (After Heald and Renton, 1966, fig. 14, p, 985; courtesy J. s e d . Petrol.)
(3)The porosity decreases during quartz-cement precipitation are presented in Figs. 3-23 and 3-24. Again, there are marked deviations between cementation of monocrystalline quartz and polycrystalline grains. Grains of hybrid composition
The results of the experiments by Heald and Renton on the precipitation of cement on grains of hybrid composition (i.e., mixtures) were: (1) The relative rates of cementation of an arkosic sand (= mixture of quartz and feldspar), in comparison with pure quartz sand, are illustrated in Fig. 3-25.Only 1-2%.of secondary quartz grew as very slender pesmatic crystals on the feldspar grains. Most of the quartz was precipitated on the
1 288
I U
432
OURATION, Hours
Fig. 3-24. Porosity change during cementation of grains of chert and monocrystalline quartz. 1 = chert; 2 = quartz. (After Heald and Renton, 1966, fig. 15, p. 985; courtesy J. Sed. Petrol.) W
I
I
I
I
1
I
a u)
20
40
-
60
80
PERCENT GROWTH PURE QUARTZ
Fig. 3-25. Growth relationship of arkose and micaceous quartzite compared to pure quartz. 1 = pure quhrtz; 2 = arkose; 3 = micaceous quartzite. (After Heald and Renton, 1966, fig. 16, p. 987; courtesy J. Sed. Petrol.)
K.H. WOLF AND G.V CHILINGARIAN
114
DURATION, Hours
DURATION, Hours
Fig. 3-26. Porosity change during cementation of arkose and pure quartz. 1 = quartz and feldspar (arkose); 2 = quartz. (After Heald and Renton, 1966, fig. r7,p . 987; courtesy J. Sed. P e t r o l . ) Fig. 3-27. Porosity change during cementation of micaceous quartzite and pure quartz. I = micaceous quartzite; 2 = quartz. (After Heald and Renton, 1966, fig. 18, p. 987; courtesy J. S e d . P e t r o l . )
quartz sand particles and was approximately proportional to the amount of quartz grains. (2) The changes in porosity during cementation of pure. quartz, arkosic sand, and micaceous quartzite sand, are given in Figs. 3-26 and 3-27. (3) The rates of cementation of sand composed of micaceous quartzite are shown in Fig. 3-25and the rate was much slower than in the case of quartzite, seemingly a result of the smaller surface area of quartz grains exposed. Geologic applications
Heald and Renton offered some geologic applications of th’eir experimental data:
DIAGENESIS OF SANDSTONES AND COMPACTION
115
(1)Well-sorted, coarse-grained sand will be cemented faster than finergrained sands as a consequence of greater influx of solutions into the more permeable coarser material. Where the invasion of solution is the same in both, however, the fine-grained sand is cemented more rapidly. To extend this conclusion to the theme of the present book on compaction, one must deduce from the above data that differential or preferential compaction of different sediments will depend on the numerous parameters, and may be prevented or modified by varying rates and degrees of cementation from one unit t o the next. (2) Heald and Renton showed that the medium- to coarse-grained sand 2 times faster than very coarse grained sand. Theoretically, very cemented : fine sand will cement 16 times faster than very coarse grained sand, as based on surface area differences. Consequently, if a bed or lens of coarse-grained sand is surrounded by a finer-grained sand, porosity reduction due to the precipitation of intergranular cement from solution would be quicker in the coarser-grained sediment. The rate of cementation would depend on the relative differences between the sand-grain populations. (3) As demonstrated from the data given by Heald and Renton, the rates
DURATION, Hours
Fig. 3-28. Porosity change during cementation of round and angular sand. 1 = round quartz; 2 = angular quartz. (After Heald and Renton, 1966, fig. 21, p. 989; courtesy J . Sed. Petrol.)
116
K.H. WOLF AND G.V. CHILINGARIAN
of the precipitation of cement vary with angularity of the sand particles: it proceeds considerably faster in angular than in round granular material of the same size grade. Figure 3-28illustrates that the higher porosity which is characteristic of angular sands is reduced to the same value of rounder sands after some chemical cementation takes place. As a result of differences in packing and because of the normally greater degree of initial compaction of angular sands, the combined effect of compaction and subsequent chemical cementation causes a more rapid decrease in porosity in angular sands. (4) Due to the different rates of chemical cementation of monocrystalline quartz sands, polycrystalline lithic sands, and polymineralogic (e.g., arkosic, micaceous, etc.) sands, there may have been different degrees of “lithification” from one bed t o another in a vertical sequence, as well as in the same bed horizontally, depending on the mineralogic composition of the original detrital material. These variations in susceptibility to cementation, among others, must be considered in studies of mechanical and chemical compaction on a regional scale. The studies of Levandowski and other investigators Levandowski et al. (1973)demonstrated the important role of cementation in petroleum geology, especially with regard to migration, accumulation, and storage of hydrocarbons in sandstones, because it controls porosity and permeability. Their work can serve as an example of detailed petrologic and geochemical principles being applied to regional diagenetic studies, which are discussed in a separate section below. A distinction was made between early, late and “differential” cementation by employing a paragenetic approach based on the age relationships of the cements, and the diagenetic features were then related to the paleohydrology of the basin in which the Permian Lyons Sandstone of Colorado accumulated. Similar techniques will have t o be used in ore genesis investigations in sedimentary and volcanic piles. The sandstone is quartz-rich with about 75% quartz and is bonded by a matrix and cement. As to the mineralogic composition, the individual beds vary from orthoquartzites to subarkoses. The composition of intergranular material is as follows: clay and chlorite = 1-1076; iron oxide = trace to 10%; secondary quartz = 0-2896; solid organic matter (i.e., petroleum residue) = up to 25%; anhydrite, which occurs as a common cement = 0-25%. Carbonate cement (calcite and dolomite) is also present and pyrite occurs in minor quantities as cubes and small patches. The paragenetic sequence of the above constituents is given in Fig. 3-29.Interesting to note is that solution and overgrowths on grains is absent where organic, iron oxide, and clay coatings are thickest and, apparently, have prevented the chemical interaction between the solutions and granular components. Anhydrite is
117
DIAGENESIS OF SANDSTONES AND COMPACTION EARLY
- -
PTLtDETRITUS GRAINS CLAY IRON OXIDE CEMENT CUARTZ
* LATE
~
SOLID ORGANIC MATTER
POSTDEPOSITIONAL STAGE
D
ANHYDRITE CARBONATE PYRiTE
--
-
Fig. 3-29. Paragenetic sequence of minerals in Lyons Sandstone, Denver Basin, Colorado. (After Levandowski et al., 1973, fig. 12, p. 2228; courtesy Am. Assoc. Pet. Geologists.)
associated with secondary quartz and dolomite, but rarely with calcite replacing quartz. The vertical and regional distributions of the quantitatively important cements are shown by cementation fucies in a cross-section and in a series of maps (e.g., Fig. 3-30). The cross-section gives part of the Lyons Sandstone in sw
NE
50110 ORGANIC MATTER
OM 'ORE
~
:Lx\z
CARBONATE A N D / M
ANMVWITE
PORE SPACE
0
MILES
Fig. 3-30. Distribution of cements and porosity in Lyons Sandstone. (After Levandowski et al., 1973, fig. 13, p. 2228; courtesy Am. Assoc. Pet. Geologists.)
118
K.H. WOLF AND G.V. CHILINGARIAN
which various cementing agents have been found (Le., quartz, organic matter, carbonate, and/or anhydrite) and the porosity was larger than 50%prior t o cementation (= pre-cement porosity). The extent to which Levandowski et al. have gone in their detailed petrologic evaluation is further demonstrated by their figs. 5, 7, 14 t o 18, 20, and 21, to which the reader is referred. They have prepared separate maps and/or cross-sections for: (1)tne distribution of red and gray iron oxide-bearing sandstones; (2) the distribution of quartz (see also Fig. 3-30 here); (3) the geographic variations of organic matter; and (4)the carbonate and anhydrite cements. In Fig. 3-30, it is shown that the distribution of quartz cement is independent of depth, but tends to increase towards the west. This build-up of quartz largely accounts for the permeability barrier. In the eastern part of the area, quartz grains are strongly etched, but little or no deposition of secondary quartz took place, indicating a removal or loss of S O 2 . The solid organic matter is present in a considerably larger area than oil, the former being closely associated with the gray Lyons sandstones, so that its regional distribution pattern is similar to that of these sandstones. The carbonate and anhydrite cements are shown to be concentrated near the top of the Lyons Sandstone, below the organic interval and below intervals containing a relatively large amount of quartz. Figure 3-30 gives their regional vertical distribution demonstrating the cement’s concentration at the top and base of the Lyons Sandstone and below the solid-organic matter zone, as well as at the flanks. The general build-up of the carbonate and anhydrite cements in the west, just east of the quartz concentration, is equally clear. Spectrochemical analysis of anhydrite demonstrated a regional Sr content variation in anhydrite (as a substitute of Ca*+).According t o Levandowski et al., the Sr content can be used empirically as criteria of anhydrite origin or type: (a) “Normal” or “primary” anhydrite present in bedded evaporite sequences tend to have a nearly constant Sr 102/CaS04ratio, because they formed in equilibrium with brines of similar composition. (b) The “secondarily redistributed” anhydrite (= “leached anhydrite”) has an Sr content that suggests frequent mobilization: Sr content of the normal anhydrite is lowered appreciably by extensive contact with formation waters. (c) “Secondarily introduced” anhydrite (= “enriched anhydrite ”) in veins and solution cavities is distinguished by higher than normal Sr content. The data by Levandowski et al. (1973) illustrate the enrichment of anhydrite in a northwest-southeast trending zone lying adjacent to and roughly parallel with the zone of high-Si02 cementation. It is associated with the known Lyons oil fields and the greatest enrichment occurs near these fields. The beds below the Lyons Formation and much of the Lyons sandstones to the east appear to contain leached anhydrite. The anhydrite of the cap rock
DIAGENESIS OF SANDSTONES AND COMPACTION
119
overlying the Lyons Sandstone is uniformly composed of normal anhydrite, suggesting that it was not affected by solution. Based on these regional distribution patterns, the formation waters seem to have moved upward into the Lyons Sandstone as well as laterally toward the edge of the basin because of the impermeable cap rock. This is supported by the petrographic study of the anhydrite itself and the fact that the greatest concentration of anhydrite occurs parallel with, and just basinward of, the permeability barrier (Fig. 3-30). As to the origin of the cements that are related to pressure and temperature in the lower ranges of values and, thus, being dependent on burial in the sedimentary basin, Levandowski et al. provided the following information. It should be noted at the beginning that the fresh-looking plagioclase with coatings of clay and chlorite in the sandstone samples suggested that burial metamorphism was absent or at a minimum and that the minerals making up the coatings were the result of diagenesis.
i WO
t
-
0
IO’IT~K
2
4
6
8
0
12
PH
Fig. 3-31. Solubility of quartz and amorphous silica as function of temperature. (From Siever, 1962, in Levandowski et al., 1973, fig. 25, p. 2238; courtesy Am. Assoc. Pet. Geologists.)
Fig. 3-32. Relation between pH and solubilities of calcite (after Correns, 1950), amorphous silica (after Krauskopf, 1958), and quartz (from data of Van Lier, 1959). Six areas are delineated: C = calcite; A = amorphous silica; Q = quartz; p = precipitates; d = dissolves; 1 = calcite in sea water at 2OoC; 2 = amorphous silica at approx. 25°C; 3 = quartz at 25OC. (Fiom Blatt, 1966, in: Levandowski et al., 1973, fig. 26, p. 2238; courtesy Am. Assoc. Pet. Geologists.)
K.H. WOLF AND G.V. CHILINGARIAN
120
The source of the SiOz for the secondary quartz cement was fourfold: (1) solution of quartz sand grains; (2)deposition from sea water; (3)precipitation from connate water; (4) deposition from fluids derived from other beds. The fluids of types 3 and 4 sources could be wholly or partly made up of compaction fluids. As to the influence of overburden pressure, Kennedy (1950) showed that extremely high pressures, not usually encountered in sedimentary rock environments, are required to appreciably increase the solubility of quartz. On the other hand, Levandowski et al. (1973)found it reasonable to speculate that granulation and submicroscopic fracturing at grain contacts may produce some amorphous Si02, as does grinding of quartz in the laboratory (Siever, 1962). The interstitial water would preferentially dissolve the amorphous Si02 (Figs. 3-31and 3-32)and as the latter is removed in solution, the pressure along the grain contacts would produce new amorphous silica. Under natural conditions, other parameters in addition to the pressure would have to be considered, e.g., temperature and composition of pore fluids. Anhydrite and carbonate replacement of quartz releases Si02 to solutions that is then available for cementation. Figure 3-31 shows that the silica solubility increases with increasing temperature, whereas the CaCO solubility below 120°C and at constant Pcoz decreases with increasing tempera-
.20
c\
TEMPERATURE, O C
Fig. 3-33. Solubility of anhydrite in water, as function of temperature and pressure. 1 = pressure of 1000 bars; 2 = pressure of 500 bars; 3 = pressure of 1 0 0 bars; 4 = vapor pressure. (After Dickson et a t , 1963, in Levandowski et al., 1973, fig. 29, p. 2240; courtesy Am. Assoc. Pet. Geologists.)
DIAGENESIS OF SANDSTONES AND COMPACTION
121
ture. Anhydrite (= calcium sulfate) solubility also diminishes with increasing temperature (see discussion below and Fig. 3-33).If intrastratal fluids saturated with anhydrite and/or calcite are undergoing burial, temperature increases and CaCO and CaSO may precipitate, with concomittant dissolution of S O 2 . As Siever (1959)has mentioned, within the range of temperature t o which sedimentary rocks are subjected during subsidence of a basin, the temperature effect is much greater on the solubility of calcite and anhydrite than that of the pressure. The observation made by Levandowski et al. that S i 0 2 was moved from the basin center toward the margin where it was precipitated, indicates that the composition of the intrastratal fluids also influenced regional variations in cementation. As illustrated in Fig. 3-32 (e.g., Blatt, 1966), the calcitequartz replacement reactions are dependent on pH, e.g., dissolved SiOz tends to precipitate when the pH falls below 9.8. It may be proposed, therefore, that formation fluids in the sediments of the basin center were highly alkaline as a result of the presence of evaporites. The high alkalinity and the existing pressure and temperature conditions tended to increase the solubility of the quartz. Under late-diagenetic burial and compaction, the SiO2enriched waters moved horizontally underneath the tight anhydrite-rich unit. These basinal compaction fluids became mixed with less alkaline meteoric waters present along the basin margin. The resultant lowering in pH caused SiOz precipitation. As to the carbonate and sulfate cements, Levandowski et al. showed in their paragenetic sequence (Fig. 3-29)that they were the last to form, some evidence indicating that CaC03 may have been the latest of the two. The vertical and horizontal regional distribution of the last cement was at least partly controlled by the distribution of the earlier-formed intergranular pore fillings, e.g., quartz cement. The trace-element composition of the anhydrite suggested that the solutions moved from below and migrated laterally to the basin margins underneath the tight anhydrite cap rock. This took place during the compaction of the rock sequence. As shown in Fig. 3-33(Dickson et al., 1963), anhydrite solubility varies inversely with temperature and directly with pressure, whereas Fig. 3-34(Blount and Dickson, 1969)demonstrates that the anhydrite solubility is much greater in ionic NaCl solutions than in pure water at the same pressure and temperature. These investigators concluded that: (1) as a result of a combined temperature and pressure decrease that leads to a solubility increase, anhydrite will not precipitate from ascending saturated solutions; (2)at a constant temperature, anhydrite is precipitated from a saturated solution that migrated from a high-pressure to a low-pressure zone; and (3)anhydrite will precipitate from a high-ionic strength solutiov upon dilution by fresher water. Using the above data, Levandowski et al. reasoned that in the Lyons
K.H. WOLF AND G.V. CHILINGARIAN
122
to 0
250
1
1
500
750
I
1000
PRESSURE, Bars
Fig. 3-34.Anhydrite solubility as a function of pressure at several constant temperatures and NaCl concentrations. 1 = temperature = 100°C, NaCl concentration = 2M;2 = temperature = 100°C, NaCl concentration = 6 M ;3 = temperature = 15OoC,NaCl concentration = 6M; 4 = temperature = 2OO0C, NaCl concentration = 6M;5 = temperature = 2OO0C, NaCl concentration = 5M; 6 = temperature = 2OO0C, NaCl concentration = 4M;7 = temperature = 2OO0C, NaCl concentration = 2M; 8 = temperature = 100°C, in HzO;9 = temperature = 200°C, NaCl concentration = 1M;10 = temperature = 2OO0C, in HzO. (After Blount and Dickson, 1969,fig. 4,p. 235;courtesy Geochim. Cosmochim. A c t a . )
Sandstone sequence investigated by them, the formation fluids were highly alkaline, containing a large amount of sulfate. During compaction, the upward and, subsequently, laterally moving solutions moved towards the basin margins, where the pressure decrease resulted in precipitation of anhydrite and possibly some calcite. The theory of ionic impedance proposed by Fothergill (1955) was employed to explain the concentration of carbonate and anhydrite cement near a secondary quartz build-up that acted as a semi-permeable membrane behind which ions were concentrated on the high-pressure, basinward side. As compaction progressed, the formation fluids on the basinward side may have been subjected to higher pressures than on the landward side. Behind the semi-permeable membrane or barrier, ionic impedance progressively increased the concentration of Mg2+, Ca2+,and Sr2+ ions, until the solubility products of the respective minerals were reached and anhydrite and calcite were precipitated as open-space fillings. The sequence of cementation would have been: Sr-enriched anhydrite anhydrite -+ calcite. Figure 3-35 summarizes the geologic stages of the origin of the intergranular components present: --f
DIAGENESIS OF SANDSTONES AND COMPACTION R E PRESE N T A T I VE
GEOLOGIC
123
SETTING
STAGE 5 PRESENT- DAY
STAGE 4 ANHYDRITE AND CARBONATE
STAGE 3 PETROLEUM
STAGE I C L A Y AND IRON O X I D E
Fig. 3-35. Schematic diagrams illustrating cement development in Lyons Sandstone. (After Levandowski et al., 1973, fig. 31, p. 2242; courtesy Am. Assoc. Pet. Geologists.)
K.H. WOLF AND G.V. CHILINGARIAN
124
(1)Deposition of sands, with some detrital clay and iron oxide matrix. Basin*ard the Lyons sandstones are interbedded with limestones, dolostones, anhydrite (thick evaporite sequence) and shale. (2a) Accumulation of the anhydrite-rich unit (symbol: crosses) provided a relatively impermeable cap rock. (2b) Progressive accumulation of younger sediments increased the overburden, thus pressing out the connate water as compaction fluid. Migration was mainly lateral toward the basin margin within the sandstone underneath the anhydrite cap. The high alkalinity (pH) of the water caused SiOz dissolution and its subsequent reprecipitation as cement as pH was lowered. The silica solution predominated along the zone of the basinward sandstoneevaporite interbedding. On the other hand, near the western basin margin, the mixing of the alkaline with fresher water resulted in lowering of pH that caused SiO to precipitate. (3)Oil from the basinal shales and carbonate source rocks migrated into the Lyons Sandstone prior to complete compaction. The hydrocarbons moved laterally because of buoyant forces and the pressure gradient resulting from compaction. (4) With progressive burial, the overburden pressures led to the formation of solution effects, e.g., stylolites, with a consequent further expulsion of connate waters. These fluids dissolved CaS04 and CaC03 in older horizons and reprecipitated CaSO while migrating upwards and laterally. CARBONATE
4
A 0 0
LAUMONTITE
Jambwoo Sandstone Kioma Sandstone Westley Park Sondstonr
CHLORITE
LAUMONTITE
QUARTZ
CHLORITE
Fig. 3-36. Composition of the cement of sandstones from the Broughton Sandstone, Kiama, N.S.W.,Australia, based on carbonate, laumontite, chlorite, and quartz contents. Solid symbols represent micrometric analyses, whereas open symbols represent visual estimations; triangles = Jamberoo Sandstone; squares = Kiama Sandstone; circles = Westly Park Sandstone. (After Raam, 1968, fig. 6, p. 326; courtesy 3. Sed. Petrol.)
DIAGENESIS OF SANDSTONES AND COMPACTION
125
Raam (1968)described a terrigenous sequence in which the sandstones have a variety of matrices and cements of diagenetic origin (e.g., Fig. 3-36). Inasmuch as quartz and carbonates rarely are present together, two separate ternary diagrams were prepared. Zeolites (e.g., laumontite) and chlorite were frequently suggested to be of low-metamorphic or burial-metamorphic origin (e.g., Coombs, 1954, and Packham and Crook, 1960). Separate sections on burial metamorphism and related aspects are presented in this chapter. Raam attempted t o establish t o what degree the temperature and pressure increase was caused by overburden load, resulting in the formation of matrix and cement. He calculated that the unit was covered by sediments having a maximum thickness of only 2500-3000 ft, which is equivalent to a pressure of less than 200 atm and a temperature of about 35"C (assuming an average world geothermal gradient of loC/10O ft). Inasmuch as it was found elsewhere that laumontite forms at temperatures above 200" C and that although it can form in the presence of quartz at lower temperatures, i.e., 50--100"C, burial metamorphism could not have given rise to zeolites (e.g., laumontite), because they are present frequently in the absence of quartz. To find some additional support for his conclusion, he considered the rank of the coal in the rocks of the same basin and found that the maximum coal rank did not coincide with the general center of sediment accumulation, nor with the structural center of the basin. Raam, therefore, concluded that neither the authigenic minerals nor the high coal rank can be attributed entirely to burial metamorphism. He suggested that the physicochemical parameters of the early diagenetic environment played a more significant role. Hence, zeolitic reactions can occur at much lower temperatures and pressures than had previously been suggested. Hrabar and Potter (1969)gave an example of decementation. They found a sandstone body which was more permeable near the surface than its subsurface equivalents, because it had been decemented by ground water flowing downward through it into the underlying limestone (see Figs. 3-37A and 3-37B).The very high permeability values are due to the position of the sandstone body below the present erosion surface. Figure 3-38shows schematically the permeability distribution in an abandoned delta finger below an unconformity. The closer the sandstone to the unconformity, the greater its permeability. Important parameters controlling permeability were: (a) the length of time represented by the unconformity; (b) the intensity of decementation process; and (c) the extent of recementation after later burial. One might add another factor, i.e., any compaction that took place between cementation, decementation, and recementation could have reduced porosity and permeability. The formation of several generations of cement may be important in the .compaction history and should, therefore, be studied in detail.
K.H. WOLF AND G.V. CHILINGARIAN
126
Lithology and heah’ Mean Grain Size Bedding Facies (microns)
Horizontal Pernleability (100 millidarcies)
Percent Porosity
Fig. 3-37A. Vertical profile of bedding, grain size, permeability, and porosity, obtained from diamond drill cores. There is fining upward and vertical decline in permeability. Interval extends from Beech Creek Limestone to the top of Blue River Group. (After Hrabar and Potter, 1969,fig. 12;courtesy Am. Assoc. Pet, Geologists.)
Although the investigation of regional “diagenetic facies” is still in its infancy, a number of published examples clearly indicate a promising new field of petrology and sedimentary geochemistry. One variety of the diagenetic facies is based on the variation of different types of cements within sandstones (see also Wolf et al., 1967, pp. 128-130). Strakhov (1969, vol. 2,
OIO
20
50
too
zoo
500
low Moo
Jo
LOG PERMEABILITY, md
Fig. 3-37B. Stepwise cumulative curves of horizontal permeability values. There is a marked enhancement of permeability in outcrop. The 400 subsurface samples are from Bethel Sandstone in Midland field, Muhlenberg, Kentucky, whereas outcrop samples are from shallow drillhole in Owen County. There is an overall trend of permeability increase with depth. (After Hrabar and Potter, 1969,fig. 13, p. 2159;courtesy Am. Assoc. Pet. Geologists.)
DIAGENESIS OF SANDSTONES AND COMPACTION
2 .-
F
127
5600
c e .-
e 0
z-4000 0 t 3
I
I
I
\
\ - \
-
\
\
\
\
1
Fig. 3-39. Composition of the cementing minerals in various facies of Aalenian and Middle Jurassic deposits of Dagestan. Facies: A = alluvial; D = underwater deltaic; FB = foreset and bottomset beds of delta; M = marine; I = calcite; 2 = ankerite and siderite; 3 = quartz. (After G.N. Brovkov, 1958; in Strakhov, 1969, fig. 200, p. 513.)
pp. 513-515) mentioned the variations in distribution pattern of cement in sandstones and muddy sediments in different sedimentary depositional environmental facies (Fig. 3-39).He explained the variations to be the result of differences in fluid composition, differential migration of chemical elements during compaction, and related phenomena. Warner (1965, 1966) undertook a detailed study of the type, quantity, and distribution of cement, in addition to other parameters (Figs. 3-40,3-41, and 3-42),and demonstrated three stages of cementation by thin section analysis. He showed that the distribution pattern of the primary cement is related to the source rock and the distance of transportation from the source to the depositional milieu, whereas the secondary cements are more closely related to the structure of the region and the source rock, which suggests a ground water influence. Warner concluded that more investigations of this type are required to give information on sedimentary and tectonic structures, drainage direction of both surface and subsurface waters, permeability trends, distance of transportation influencing chemical differentiation, and source rock. The information gained would be of practical application not
1-1 1-
Outcrop limits of the Duchesne River Formation Contour representing percentage of cement Area of predominantly calcite cement Area of predominantly iron oxide. clayand silica cement
Fig. 3-40.Types of cements in the sandstones of the Duchesne River Formation. (After Warner, 1966,fig. 8, p. 952;courtesy Geol. SOC.Am.)
Contour representing percentage of grains showing secondary enlargement (silica cement) ...................... Synclinal axis in the Duchesne River Formation Fig. 3-41. Distribution of silica cement in Duchesne River Formation. (After Warner, 1966,fig. 9,p. 953;courtesy Geol. SOC.Am.)
DIAGENESIS OF SANDSTONES AND COMPACTION
-
129
Outcrop limits of the Duchesne River Formation
1-1
Contour representing the average number of trace elements
...
Sandstone faciesj 2.33;average
ratio
Mudstone facies,0.43:average
ratio
Fig. 3-42.Sandstone/mudstone ratio and facies of the Duchesne River Formation. (After Warner, 1966,fig. 13,p. 955;courtesy Geol. SOC.Am.)
Fig. 3-43. Carbonate mineral relationships in the CaC03-MgCOs-FeCO~ system. A. Composition of carbonate minerals in the Reedsville, Bald Eagle, and Juniata formations; two tie lines connecting mineral pairs are shown. B. Subsolidus relations at 400%. (After Rosenberg, 1960,1g67,based on laboratory study; in: Horowitz, 1971,fig. 1;courtesy J. Sed. Petrol.)
K.H. WOLF AND G.V. CHILINGARIAN
130
only in petroleum geology and hydrology, in general, but also in the understanding of the origin of copper and uranium "cements" in sandstones of the red-bed type (see Chapter 5 ) . The study of compactional history, together with textural and fabric investigations, may reveal the amount of fluids released, direction of fluid movement, and possible contribution of compaction fluids t o cement precipitation. Horowitz (1971) determined the stratigraphic variations in the ferrous iron content of the carbonates present in sandstones and used the data t o decipher the origin of these cements. Figure 3-43shows the general carbonsystem, whereas ate mineral relationships in the CaC0,-MgCO ,-FeCO Figs. 3-44 and 3-45illustrate the vertical changes in composition. Ankerite exhibited an upsection decline in ferrous iron content at four widely separated localities (Fig. 3-44),suggesting a vertical chemical gradient during its formation. Although Fez+ content in ankerite continued to increase downsection, total Fez+ in the carbonate system, i.e., ankerite plus calcite, actually decreased below the shoreline units (= Bald Eagle-Reedsville boundary),
,
REEDSVILLE
IOOO'
TYRONE
1' ' 1
8Od
I
600'
v,
LOYSBURG
1
0
t
I -
406
1
-
"
I
0
z
BEDFORD
P
n
u I
200'
u
100'
+ a
o
0,
a '3 a
Li
-100'
- 200'
NO ANK. ,
300
(Fe/Fe
I
30 0
+ Mg ) RATIO
Fig. 3-44. Stratigraphic variation in iron and magnesium content of ankerites at four localities in the Appalachian region of Pennsylvania. Color of rock units noted on side. Stratigraphic distance is shown in logarithmic intervals; zero mark shows the ReedsvilleBald Eagle boundary. (After Horowitz, 1971, fig. 2; courtesy J. Sed. Petrol.)
DIAGENESIS OF SANDSTONES AND COMPACTION
ca co,
131
Ca Mg (Cod.
Fig. 3-45. Approximate stratigraphic variation in bulk composition of carbonate cement. Letters relate positions in the stratigraphic column (left) to compositions on the ternary diagram. (After Horowitz, 1971, fig. 3; courtesy J. Sed. Petrol.)
because calcite became increasingly more abundant downsection (Fig. 3-45). Inasmuch as the decline of Fe2+ in carbonates above shoreline units corresponded to the successive upsection color change from dark grey to drab to red, it was concluded that the vertical chemical gradient reflected a change in the ogidation potential. Horowitz offered two hypotheses to explain his observations: (.1)successive facies differed in their Eh during sediment accumulation; or (2) diagenetic reducing solutions, originating from the lower marine units, reduced originally oxidized sediments with upsection effectiveness. Horowitn accepted the latter hypothesis as the more plausible one. The reducing fluids originated from the dark-colored marine Reedsville Shale and were expressed upward into continental red beds during compaction. Horor witz (1971) stated: “These solutions also may have carried the magnesium and calcium which combined with the reduced iron ,to form ankerite., since these two cations would. probably be more abundant in waters of marine origin. . . . That permeability variations influenced the migration paths of the reducing solutions can be inferred from the relationship between lithology and color. Proceeding westward from the depocenter, red beds are encountered first in the shalier (muddier) coastal plain facies (Fig. 3-46)and, more specifically, within shaly units of:this facies. Higher in the section, drab beds of the once more peimeable fluvial sandstane facies extend even further westward before passing into red beds. These observations suggested that reducing solutions favoured permeable sandstones. Displacement of most of the expressed water t o the west instead of eastward through the even ,more permeable conglomeratic facies wm atttibuted to the influx of meteoric water from the east and development of a hydrodynamic gradient to the west.” From the above discussions ‘one can conclude that the following factors must be considered in studying the cementation of sandstones by compaction fluids:
132 A
K.H. WOLF AND G.V. CHILINGARIAN A'
UNCONFORMITY
Fig. 3-46. Model proposed to show the migration paths of reducing fluids and explain the origin of the red-drab color boundary between the Bald Eagle and Juniata formations. 1 = reds beds, Juniata Formation; 2 = black shale, Reedsville Formation; 3 = drab beds, Bald Eagle Formation; 4 = direction of water movement. (After Horowitz, 1971, fig. 4; courtesy J. Sed. Petrol. )
(1)The distribution of sedimentary facies (i.e., conglomerates, sandstones, clayey deposits, limestones, dolomites, and evaporites), which deterniines the regional variations in primary porosity, original amounts of fluids trapped within the sediments and, therefore, the potential rate and degree of compaction to be expected on a theoretical basis. (2) The mineralogic composition of sedimentary particles, which determines the rate of release and composition of the fluids from clay minerals, for example, to form compaction solutions (see separate sections in this chapter, pp. 290-303). (3) The textures and labrics, which indicate the mechanisms, degree, and regional variation of compaction. (4) The diagenetic effects (e.g., genesis of cements and matrices; removal of heavy minerals by chemical dissolution; paragenetic relationships; alterations of clays). (5) The structural history of the sedimentary basin, which controls the amount and rate of subsidence and, therefore, the rate of sediment accumulation and the rate of compaction. The latter, in turn, controls the rate and direction of compaction fluid movements. The regional structure will also control later ground water. migration, which may cause decementation and recementation that could erase or obliterate the earlier-f ormed diagenetic features. To the knowledge of the writers, a total and all-inclusive regional
DIAGENESIS OF SANDSTONES AND COMPACTION
133
petrologic and geochemical investigation of the compaction history of a sedimentary basin has not been undertaken as yet. Most investigators, who have considered compaction in their stratigraphic and environmental examinations of sedimentary piles, have employed one or two particular techniques, such as studying textural changes in a vertical profile and/or variations in cementation, that resulted in data satisfactory for their particular needs and goals. Inasmuch as numerous different techniques and concepts have been developed and are available for compaction studies, they should be combined in the future in a deliberately comprehensive examination of the three-dimensional compaction history of a sedimentary basin. Such investigations could be particularly attractive because of their practical applicability in the petroliferous basins and sedimentary ore districts, and may also assist in unravelling the possible relationships between hydrocarbons and ores in the same basin (see Chapter 5). TEXTURES* RESULTING FROM COMPACTION
Geologists require details on the textures and their paragenetic relationships, in addition to data on composition, to be able to unfold the geologic history of sedimentary rocks. Much data are now available on the textures of the clastic rocks, especially the sandstones. The textures have their start with primary depositional characteristics and grade without definite boundaries into the secondary features, among which are those formed by compaction as a result of overburden pressures. The strictly primary and secondary-diagenetic textures, in turn, grade into catagenetic and metamorphic textures, fabrics, and structures. Thus, it is essential to be familiar with all types of petrographic characteristics in sedimentary petrology. There is hardly any concept in the interpretation of the origin of sediments that does not include their textural properties, so that for a proper comprehension of the various sections of this chapter it is necessary to have some understanding of the textures of sandstones. To name but one example, Modarresi and Griffiths (1963)have demonstrated by the use of statistical petrographic analyses of reservoir rocks that the textures of the framework of sandstones controlled subsequently developed silica and carbonate distribution. Although textures and fabrics are fundamental units of the larger-scale structures in sedimentary deposits, the latter are not considered here. Large features resulting from compaction and which, therefore, fall into the group of secondary structures, are also bypassed in this section, but a list with *The term “texture”’as used here includes those features that have been called “fabrics” by others, but excludes the large-scale structures.
K.H. WOLF AND G.V. CHILINGARIAN
134
pertaining references is provided at the end of this chapter. Most of the information obtained on textures and fabrics comes, from studies utilizing the petrographic microscope, but the use of the electron microscope and X-ray textuameter is rapidly increasing. Sippel (1968) employed luminescence petrography in the study of sandstone diagenesis and was able to: (a) distinguish detrital quartz from quartz overgrowths; (b) recognize fracture-healed quartz grains as well as features resulting from crushing; and (c) distinguish recrystallization from polycrystalline features inherited from themouroe area. All these textures are frequently not recognieable under a normally-equipped petrographic microscope. Pettijohn\et 4.(1972) presented a summary of the fabrics of sands and sandstones. They stated that the arrangement of grains to form an aggregate reflects: (a) manner of deposition at the time of sediment accumulation; (b) grain size; (c) sorting; (d) shape of the clasts; and (e) physical. and chemical compaction. The measurements of these parameters, together with other information, such as mineralogic composition, type and abundance of fossils, structures, and regional relationships of the various lithologic units, enables
SbTURED GRAINS
CONCAW-CONVEX CONTACT
POINT CONTACXLONG CONTACT
FLdATlMG GRAINS
LlNE .OF TRAVERSE
Fig. 3-47. Definition sketch of fabric terminology: quartz (white), mica (lined), and matrix (stippled). Illustration of the application of the quantitative fabric indices listed in Table 3-XVIIL. (After -Pe&ijohn et al., 1972, fig. .3-10;p. 91; courtesy Springer, New York.)
D M E N E S I S OF SANDSTONES AND COMPACTION
135
one to infer the primary and secondary rock formation regimes. The above information can also serve as a guide in estimating the crushing and bearing strength of sediments and rocks, as illustrated in some of the chapters of both Volumes I and I1 of this book. Pettijohn et al. (1972) offered a summary of the qualitative and quantitative teFms and indices that are used in describing the grain-to-grain relations. They have pointed out that the terminology and methodology are available now for detailed textural investigations, but there is a shortage of systematic mapping studies that could serve as models. Table 3-XVIII presents the TABLE 3-XVIII Terms used to specify sandstone fabric (after Pettijohn et al., 1972, table 3-4, p. 90) Qualitative Concauo-convex contact (Taylor, 1950, p. 707): one that appears as a curved line in the plane of section Fixed margin (Allen, 1962, p. 678): that part of a grain in contact with another in the plane of section Fixed grain (Allen, 1962, p. 678): fixed margin exceeds free margin Floating grain: no contacts with other grains in the plane of section Framework fraction: the stress-transmitting portion of a sand Free margin (Allen, 1962, p. 678): that part of grain not in contact with other grains in the plane of section Free grain (Allen, 1962, p. 678): free hargin exceeds fixed margin Long contact (Taylor, 1950, p. 707): a contact that appears as a straight line in the plane of section Packing (Kahn, 1956, p. 390): mutual spatial relationships among grains Sutured contact: mutual stylolitic interpenetration' of two or more grains Tangential contact (Taylor, 1950, p. 707): one that appears as a point in the plane of section Quantitative Condensation index (Allen, 1962, p. 678): ratio of percentage of fixed rock fragments t o percentage of free grains Contact index : number of contacts per grain Horizontal packing intercept (Mellon, 1964, fig. 7): average horizontal distance between framework grains Packing density (Kahn, 1956, p. 390): length of grains intercepted divided by length of traverse X 100 Packihg index (Emery and Griffiths, 1954, p. 71): the product of the number of quartz to quartz contacts per traverse and the average quartz diameter, the product being divided by the total length of traverse Packing proximity (Kahn, 1956, p. 390): number of grain-to-grain contacts divided by total number of contacts of all kinds (grain-to-matrix and grain-to-cement)X 100. Vertical packing intercept (Mellon, 1964, fig. 7): average vertical distance between framework grains
136
K.H. WOLF AND G.V. CHILINGARIAN
TABLE 3-XIX Sandstone fabric (modified from Adams, 1964, table 1, after Pettijohn e t al., 1972, table 3-5, p. 92) A. Grain-to-grain relations specified chiefly by qualitative observation supplemented by some quantitative indices on non-oriented samples; chiefly used to interpret and predict reservoir porosity ( 1 ) Much pressure solution: many sutured contacts, grain-to-grain contacts large, packing density high, and no porosity and cement (21 Moderate pressure solution : some sutured contacts, principally equidimensional grains, chiefly with concave-convex contacts; cement is mostly quartz overgrowths with or without minor amounts of carbonate or clay; little porosity (3) Minor pressure solution: mostly original grain outlines with long and tangential contacts; low t o moderate number of grain-to-grain contacts and moderate packing density; may be either well or poorly cemented by quartz overgrowths, clay, and carbonate; may have moderate porosity, if poorly cemented ( 4 ) No pressure solution: chiefly original grain outlines with either tangential contacts or floating grains. Number of grain-to-grain contacts is low as is packing density; cements are mostly carbonate or clay; porosity is high when cementation is limited B. Orientation of framework specified by quantitative measurement of oriented samples; porosity and cementation are independent of orientation; chiefly used t o determine current direction in undeformed sediments ( 1 ) Particulate methods: visual, direct measurement of either long axes or apparent long axes of framework grains, usually in thin section (2) Aggregate methods: measurement of a bulk geophysical property by an appropriate black box that can be correlated with the orientation,of the framework grains
qualitative and quantitative terms employed in textural investigations of sandstones, whereas Fig. 3-47 illustrates the quantitative terminology for textures. One should note that Griffiths (1967b) used a somewhat different nomenclature as given in Table 3-XX.’ One example, based on actual observations showing the relationships among fabric, pressure solution, and cementing agents, is presented in Table 3-XIX. Griffiths (1967a) and Blatt (1966) stated that packing, a parameter of particular importance in compaction, depends on: (1)mode of sediment deposition; (2) orientation of the particles, (3) grain size; (4) grain shape; (5) range in grain size (= sorting); (6) amount of clay matrix; (7) mineralogic composition; ( 8 ) overburden pressure and degree of compaction; and (9) time of precipitation of cementing agents. Both cementation and compaction convert loose sediments into consolidated rocks, and are part of what is usually called “lithification”. The amount of compaction of sand, according to Blatt (1966) depends on: (a) mean grain size; (b) sorting; (c) amount of argillaceous material present; and (d) other factors that are, however, believed to be of minor importance. In
DIAGENESIS OF SANDSTONES AND COMPACTION
137
TABLE 3-XX Classification of the configuration of contacts (after Griffiths, 196713, table 8.6, p. 173) Kind of contact
Symbol
Diagrammatic appearance
Class No.
~~~
Floating
F
0
Tangent
T
1
Long
L
2
Complete
C
3
Sutured or serrated
S
4
particular, the influence of a fine-grained matrix on compaction requires special attention in future research. The problem is accentuated by the fact that there are two types of matrices, namely, a detrital, primary variety and a diagenetic-metamorphic, secondary type. Determinations of the relationships among relative proportions and types of the grains, matrix and cement and their ratios, on one hand, and the style, amount and rate of compaction, on the other, necessitate a precise terminology in the study of the matrices. Dickinson's (1970) publication is a very significant one from this point of
K.H. WOLF AND G.V. CHILINGARIAN
138
.,....
. ......... . . . I.
...
,.
..........
.
\ . . . .. .\ . . . . . . .
.....
b
a
0.
.: . .
C
..'
. .. .
... ...
. .. . . . . . . . . / . . .. . .. . :
d
.
.
f
e
;:,*.;
......... ..
. . . . .: ..: .
,./_.
.. : . ..... . .:
-.
:,..... . '!>. .I_.
..
5
5.
. .
..
..
............... . ..............
h
I
k
I
...
.; . .:,..:.
..
%.
Fig. 3-48. Grain-enlargement, pressure-solution, and micro-drusy (open-space filling) textures, (After Glover, 1963, fig. 5, p.5l;courtesyJ. R . SOC.W. Aust.) 1. Enlargement textures. Simple enlargement textures: ( a ) Calcarenite showing sparry calcite in crywtallographic continuity with crinoid debris. Stippled fraghents are calcare-
DIAGENESIS OF SANDSTONES AND COMPACTION
139
view as it summarizes his own as well as the ideas of other investigators. He offered useful terms for matrices of four different origins: proto-, ortho-, epi-, and pseudo-matrix. (For a related terminology on carbonate sediments, see Wolf and Conolly, 1965.) The papers by Dapples (1962,1971, 1972) on the diagenesis of sandstones should also be consulted. As pointed out already, in the investigation of compaction, the petrologist must be familiar with all types of genetic textural varieties, i.e., both of non-compaction and compaction origin. For this reason, summaries on textures based on the studies by Glover (1963) and Strakhov (1957), are given here in the form of diagrams (Figs. 3-48, 3-49, 3-50 and 3-51). The former investigator offered five basic diagenetic textural groups of general application, some of which can be directly attributed to compaction, whereas others could be either indirectly the result of the interaction with compaction fluids or could be due to other phenomena. Strakhov’s diagrams are particularly useful in the investigation of cements and matrices in conjunction with the information supplied by Dapples (1962, 1971, 1972) and Dickinson (1970). Much information of interest t o sedimentary petrologists can be obtained
ous pellets. ( b ) Quartz sandstone with quartz outgrowths crystallographically continuous with clastic cores. Note faces. Lightly stippled areas represent pores. (c) Quartz sandstone with pores completely occupied with secondary quartz. Some outgrowths bounded by plane surfaces, some not. No sutured boundaries. Indentation textures: ( d )Dolomitic sandy marl in which dolomite is partly surrounded by, or has partly penetrated, quartz outgrowths. Texture does not reveal whether quartz or dolomite grew first. ( e ) Dolomitic sandy marl, same as ( d ) , except one of the dolomite grains is moulded onto a clastic quartz core. Dolomite, therefore, preceded secondary quartz. Enclosure texture: ( f ) Dolomitic sandstone with dolomite completely enclosed by quartz. Dolomite, therefore, formed first, and order is confirmed by moulded dolomite (upper center). Note how a moulded dolomite has retreated marginally (left center) due to slight solution during silicification. 2. Pressure-solution textures. ( g ) Quartz sandstone with clastic grains showing sutured boundaries due to compaction, deformation or both. Secondary quartz(s) which fills voids, may have come partly or completely from quartz dissolved along sutured contacts. (h) Calcarenite with microstylolite due to compaction. Microstylolite outlined by ironstained argillaceous matter and small quartz grains, both insoluble in the particular conditions of its formation here. Sparse distribution of quartz in rock suggests compaction equivalent to field of view. (i) Quartzite with sutured boundaries between outgrowths due to deformation after diagenesis. As much a metamorphic as a diagenetic texture. 3. Micro-drusy textures (or pore-filling textures). Simple micro-drusy textures: (j)Calcarenite partly cemented with fibrous calcite. Fibers are elongated normal to grain boundaries. (k) Calcarenite completely cemented with sparry calcite. Long axes of calcite crystals are normal to grain boundaries. ( 1 ) Lithic (volcanic) sandstone cemented by fibrous chlorite. Texture basically the same as in 0’) and (k).
K.H. WOLF AND G.V. CHILINGARIAN
140
a
C
f
b
d
e
n
Fig. 3-49. Composite micro-drusy, reorganization, and replacement textures. (After Glover, 1963, fig. 6, p. 53; cqurtesy J. R. SOC. W.A u s t . ) Composite micro-drusy textures: ( a ) Lithic (volcanic) sandstone with pores filled by three minerals which are, from the outside, a micaceous mineral, chlorite (stippled), and feldspar. Note how minute chlorite
DIAGENESIS OF SANDSTONES AND COMPACTION
141
from the literature on soils (e.g., Brewer, 1964), but only a few serious attempts have been made in this direction. The publications by civil engineers on compaction and compressibility of soils have also been only occasionally consulted by sedimentologists. Pedologists, for example, use textural terminologies and genetic concepts that may be applicable not only to paleosoil investigations (e.g., Yaalon, 1971),but also in sedimentological studies. The results of the work on artificial compaction of soils may throw some light on the compaction of sediments. The textural data supplied by soil specialists will eventually allow one to establish criteria enabling a petrologist to differentiate genuine sediments from paleosoil per se as well as permit a formulation of a list of textural differences and similarities. Soils are often composed of a whole gamut of individual grains of sand, silt, and flocculated clay minerals arranged in an arching skeleton (Fig. 3-52) enclosing large voids, termed “honey-comb” fabric (Casagrande, 1940, p. 85; see also Gillott, 1968, fig. 27). With the increasing use of electron-microscope techniques in determining the origin of the different types of matrices of sandstones, it is suggested that the petrologist may find the pedological literature a fruitful mine for ideas in attempts to differentiate between the clay matrices. Figure 3-53,for example, shows the various modes of textural associations of clay minerals as used by pedologists. Similar modes should be serrations are directed inward. The micaceous mineral may be a reconstituted clay film on the clastic grains. The order of formation was: (1)micaceous mineral (or its precursor), (2) chlorite, and (3) feldspar. (Based on a diagenetic sequence in Arrowsmith Sandstone.) (b) Lithic (volcanic) sandstone showing the following diagenetic sequence: (1)micaceous mineral; (2) quartz; (3) chlorite; (4) quartz (note euhedrism); (5) feldspar. (Based on a diagenetic sequence in Arrowsmith Sandstone.) Reorganization textures: ( c ) Claystone with vermicular kaolinite crystals. Fragility of the crystals is a proof of in situ formation. ( d ) Graywacke with chlorite-sericite matrix, formed from clay-sized detritus. The new flaky minerals penetrate margins of clastic fragments, making this partly a replacement texture also. (e) Fontainebleau Sandstone in which calcite has reorganized to form large crystals. Replacement textures: ( f )Calcilutite partly replaced by quartz euhedra with calcareous inclusions. The long fragment is calcite. ( g ) Shelly limestone with shells partly replaced by chalcedony; matrix is dolomitized. The dolomite has zonal inclusions, see 0’). Unreplaced shelly material has recrystallized. (h) Quartz sandstone cemented by sparry barite. The original pyritic and argillaceous matrix is represented by patches of argillaceous impurity rind isolated pyrite grains. (i) Ferruginous quartz sandstone in which quartz grains are apparently corroded by the ferruginous matrix. One quartz grain shows an outgrowth product of an earlier diagenetic phase. The texture resembles that where carbonate corrodes quartz; many such sandstones may originate as a result of replacement of calcite by iron oxide. 0’) A complex but fairly common texture, in which dolomite has partly replaced the matrix of a sandy marl. Carbonaceous and argillaceous inclusions form a dark zone in each dblomite rhomb. There has been later recrystallization of the matrix to sparry calcite, with expulsion of impurities.
K.H. WOLF AND G.V. CHILINGARIAN
142
. . .;
.
,.
Fig. 3-50. Authigenic enlargement of quartz in sandstone from the Birdrong Formation (Rough Range Bore No. 1, core 7 , 3,633-3,636 ft). The stippled areas are partly filled with clay-sized material. Quartz crystal faces are indicated by p (prism) and r (rhombohedron). The sketch is slightly idealized to show faces clearly. Width o f field = 0.6 mm. (After Glover, 1963, fig. 2, p. 39; courtesy J. R. SOC.W. Aust.)
expected in the matrices of sandstones and in mudstones, in particular if the clay is of diagenetic origin and in an uncompacted state. To better understand the very early compaction mechanisms of sediments (and physical and chemical diagenesis, in general), the textures of detrital, flocculated and chemically precipitated clay accumulations must be thoroughly studied using electron microscopy. Griffiths (1961) pointed out that some of the most important aspects of sedimentary petrography are definition and evaluation of procedure and an outline of a uniform methodology. He also stated that relatively few detailed analyses of sedimentary rocks exist and most of those that do exist are qualitative rather than quantitative. Griffiths offered a “conceptual definition” of a sedimentary rock by specifying five properties, which are to be considered also in detailed investigations related to physical and chemical compaction: (1)proportions of different kinds of grains (= m ) ; (2) sizes of grains (= s); (3) shapes of grains (= s h ) ; (4)orientation of grains (= 0);and (5) mutual arrangement or packing of grains ( = p ) . In several publications and books, Griffiths then has offered the following formula that depicts the properties of a sediment: P = f(m,s,sh,o,p), where P = unique index and f = function. Griffiths (1961, 1967a) discussed the complex interrelationships among these properties and described procedures for specification of each property, which when expressed as numbers, can be substituted in the above formula. The resulting value of P is a unique description of the rock specimen. In a multivariate system, the properties may be interdependent, dependent, or independent, as Shown by the three-variable system in Fig. 3-54 (A& and C ) : ( A ) size changes with shape, showing a direct linear depen-
143
DIAGENESIS OF SANDSTONES AND COMPACTION f
4
5
7
6 d
6 C
d
a
Y
€
4
71
Fig. 3-51. Textural types of cements. 1-8: types of cement association with detrital grains; a-e: argillaceous cements; and a-: cements of chemical origin. (After Strakhov, 1957, fig. 20.)
144
K.H. WOLF AND G.V. CHILINGARIAN
Fig. 3-52.Proposed fabric of clay soils. Honeycomb-structure flocculated soil. (After Casagrande, 1940,p. 85;in: Gillott, 1968,fig. 27;courtesy Elsevier, Amsterdam.)
Fig. 3-53.Modes of particle association in clay suspensions, and terminology. A. “Dispersed” and “deflocculated”; B. “aggregated” but “deflocculated” (face-to-face association, or parallel or oriented aggregation); C. edge-to-face flocculated but “dispersed”; D. edge-to-edge flocculated but “dispersed”; E. edge-to-face flocculated and “aggregated”; F. edge-to-edge” flocculated and “aggregated”; G. edge-to-face and edge-tosdge flocculated and “aggregated”. (After Van Olphen, 1963, p. 94; in Gillott, 1968, fig. 37;courtesy Elsevier, Amsterdam.)
*-I/
145
DIAGENESIS OF SANDSTONES AND COMPACTION A) DEPENDENCE
A
Y = a + p x
W
U
a
a
I
0
a
0
= comt k ; 8'1
0
0 0
SIZE,X Size changes with shape; lineor dependence
3 INTERDEPENDENCE
8 ) INDEPENDENCE
I
0
SI2E.X Size
0
0
0
independent of shape
t
to shope but relationship with size changes
SiZE,X Size reloted
Fig. 3-54. Interrelationship among the properties of sediments. (After Griffiths, 1961, fig. 1, by permission of The University of Chicago Press, copyright @ 1961 by the University of Chicago.)
Fig. 3-65.Geometrical illustration of three-factor experiment and its interactions. Main effect = solid lines; first-order interactions = broken lines; and second-order interaction = dotted line. (After'Griffiths, 1961,p. 493, by permission of The University of Chicago Press; copyright 0 1961 by the University of Chicago.)
146
K.H. WOLF AND G.V. CHILINGARIAN
TABLE 3-XXI Three-factor experimental design t o illustrate sources of variation (after Griffiths, 1961, table 1, p. 4 9 1 ) Source of variation: Size Shape Orientation
three main effects
Sizeahape Size-orientation Shape-orientation
three first-order interactions
Size-shape-orientation
one second-order interaction
dence; ( B ) size is independent of shape; and (C) size is related to shape, but the sizeshape relationship is not the same in all size ranges. Griffiths (1961) then proceeded to a three-variable system which is even more complex; the sources of variation are summarized in Table 3-XXI and graphically represented in Fig. 3-55. In a discussion of the application of his equation, Griffiths mentioned that he has used it successfully in: (a) detecting a favorable change from barren sandstones to potential oil reservoir rocks, and (b) differentiating ore-bearing from barren sandstones in the uranium-containing sandstones of the Colorado Plateau-type of mineralization. His approach should also find application in detailed work on compaction where precise description of all variables is fundamental.. Using “bedding” as an example, the problem of defining the microstructure of a sedimentary rock can be demonstrated. According to Griffiths, bedding represents heterogeneity or non-randomness of arrangement of the constituent elements. Bedding may be due to: (1) change of composition (Fig. 3-56,]); (2) to grain-size variation (Fig. 3-56,2) if composition is constant; (3) to grain-shape variation (Fig. 3-56,3) if both composition and grain size are constant; (4)to a change in orientation (imbrication) (Fig. 3-56,4), if composition, grain size and grain shape are maintained constant; and (5) to differential packing (Fig. 3-56,5), if all four properties remain constant. The question arises here as to how changes in one or all of these five parameters would affect chemical and physical compaction. Most studies made so far in this area have been rather simple and were based on one or two variables; however, in more recent studies, attempts are being made to consider more variables and their interdependencies. If subtle differences in the above-mentioned grain parameters influence the mode, degree, and rate of compaction of sediments, and if a set of more or less constant characteris-
DIAGENESIS OF SANDSTONES AND COMPACTION ~
G OF C OEM P O S ~ T ~ O N
I3.CHANGE
OF SHAPE
147
1 2 . CHANGE OF SIZE
I 4 . C H A N C E OF ORIENTATION
I
Fig. 3-56. The fundamental basis of “bedded” microstructure. (After Griffiths, 1961, fig. 3, p. 495; by permission of The University of Chicago Press, copyright @ 1961 by the University of Chicago.)
tics that vary between certain limits can be established for each sandstone group, then each type of sandstone can be characterized by certain differences in physical and chemical diagenesis, including compaction. Although a relatively abundant data is available already on the differences and similarities between arkoses, graywackes and quartzites (some of the differences are inherited from the definitions of these three petrographic terms and the arbitrary boundaries set by different researchers), many aspects remain to be investigated. One example is that described by Griffiths (1967b) on the mean and variance of roundness in the three basic sandstone types (Fig. 3-57): the arkoses show a wide variability in variance for a very small range in mean roundness, whereas the graywackes and quartzites exhibit a trend in which the variance increases with increase in mean roundness. The scatter of the data for the quartzites is largest, i.e., they differ in both sphericity and roundness considerably (see the very inclined line in Fig. 3-57). These differences are not the result of differences of the various rock types in the source area, but reflect the degree of selective shape-sorting of the grains: the arkoses are all near-source sediments, the graywackes vary in both texture and composition, and the quartzites are of several different textural types. The near-source sediments were subject t o relatively little selective sorting and, therefore, show wide variability within samples but little among sample means. On the other hand, well-transported sediments tend to exhibit wide variability in sample means and less variability within samples. One example
148
K.H. WOLF AND G.V. CHILINGARIAN
/-
---
TRENDS INSERTED BV EYE
@,a,@,ROCK-TYPE
MEANS
Fig. 3-57. Mean and variance of roundness measurements. (From Curray, 1949; in: Griffiths, 1967b, p. 139; by permission of McGraw-Hill, New York, 0 1967 McGraw-Hill, New York.)
of the change of sphericity and roundness with distance of transportation of the grains from the source area is presented in Fig. 3-58.The above geological occurrences are to be considered “ideal” situations and should be used as “models” in the search of a number of possible exceptions. In much of the theoretical considerations of the porosity, permeability, and depositional fabrics and textures of sandstones, most investigators use spherical particles as the fundamental, ideal constituents of the sediments. Inasmuch as compaction features and the behavior of sand grains under stress, for example, are directly related to the initial characteristics of the
F
i DISTANCE, Miles
Fig. 3-58. Form of the function representing the relationship of change in particle shape with distance (or time). 1 = analogous to .Spey River (after Mackie, 1897); 2 = analogous to Mississippi River (after Russell and Taylor, 1937). (Based on Krumbein, 1941; in: Griffiths, 1967b, fig. 6-16; by permission of McGraw-Hill, New York, copyright 0 1967 McGraw-Hill, New York.)
DIAGENESIS OF SANDSTONES AND COMPACTION
149
grains, the theoretical evaluations of the secondary modifications are also based on the assumption that constituent particles are ideal spheres. Allen (1969, 1970), however, has pointed out in several papers that theory and experiments based on the sphere as the ideal particle, fail to account satisfactorily for the concentration G f solids observed in unconsolidated natural sands. Allen (1969, p. 309)stated that “when the prolate spheroid of moderate axial ratio is substituted for the sphere, a theory which is satisfied by observation becomes possible, for the reason that equal spheroids can be regularly packed in both close and open ways”. Allen pointed out that the freshly deposited natural sands have a relatively low particle concentration, even though theory based on the sphere as the ideal sedimentary particle predicts for them a relatively high concentration of solids. Natural sands are chiefly deposited either as laminae on horizontal or nearly horizontal surfaces or as steeply-inclined laminae on slip faces or ripples and/or dunes (see results of numerous experiments by Jopling, 1963, 1965a,b, 1966, 1967, for example). The concentration of grains, C , which is equal to (space occupied by solids)/(total space), ranges between 0.50 and 0.65 in the above two cases. The variable C depends on the average grain size, degree of sorting, medium of deposition and intensity of deposition, and the porosity is equal to (1-C). According to Allen (1969, pp. 309-310), the range of possible particle concentration C, obtained experimentally by using haphazardly packed equal solid spheres is roughly between 0.60 and 0.64. Theory based on the sphere as the ideal sedimentary particle, as mentioned in most text and reference books, predicts C to increase with increasing range of sizes of different sphere populations present. Consequently, C of natural sands should be distinctly larger than the possible range of C for equal spheres, because natural sands are composed of grains having an appreciable range of sizes. Observations do not agree with this, because the lower limit of C for natural sands is about 0.5, which is below the lower possible limit of C for haphazardly
mm Cubic Packing
Orthorhombic PockinQ
PACKING Ill
Tetragonol Packing
Rhombohedra1 Pocking
Fig. 3-59. The four dose packings of equal prolate spheroids when the major axes lie in the planes of the layers. (After Allen, 1969, fig. 1, p. 312; courtesy Geol. Mag.)
K.H. WOLF AND G.V. CHILINGARIAN
150
packed equal spheres by 0.1. The upper limit of C for natural sands deviates little from the upper limit for haphazardly accumulated spheres. Allen attributed this to the use of the sphere as an ideal sedimentary particle. As shown by numerous investigations, the natural sand particle is approximated by a triaxial ellipsoid with a long axis about 1.5 times longer than the intermediate axis and about 2.0 times longer than the short axis. Allen (1969, p. 310) continued to explain that equal solid ellipsoids can be regularly packed in more different ways than equal solid spheres, because these particles can be differently oriented in space, which need not be the same for all the ellipsoids (Figs. 3-59, 3-60, 3-61). Some of their regular packing is extremely open, and it seems that the comparatively low C values of natural sands can be explained by the presence of only a small proportion of the grains in an open-packing pattern. When spheres are used, one can take into account only sorting or size variation and not orientation together with size, as spheres have no dimensional orientation. The closer regular packing of equal ellipsoids is highly anisotropic, whereas the open packing is perfectly isotropic. Inasmuch as natural sands possess a fabric which is neither perfectly anisotropic nor perfectly isotropic, a combination of both close- and open-grain packings may be expected to be present on a microscopic or a local scale. For these reasons, Allen (1969) explored the different regular packings of ellipsoids using not the triaxial but the prolate spheroids (= ellipsoid of revolution about the long axis). Figures 3-59 and 3-60 illustrate the nine different packing systems. Figure
@@Do3 Plan
PACKING
Elevat ion
PACKING YIII
Plan
Elevation
Elevation
Q9 Pion
Plan
Fig. 3-60. The five open paikings of equal prolate spheroids. (After Allen, 1969, fig. 2, p. 313; courtesy Geol. Mng.)
DIAGENESIS OF SANDSTONES AND COMPACTION 0.8
I
0
2-
0
C 0.6 L
,
Pocking
'm ' m
'
'
1
0
1
b
1
151
C OM1 0.698 0.605 0.555 0.524
LL I-
z
0.454 0.417
W 0 0.4
z
0
u W
I! 0.2
ka L
0
0
0.2
0.4
0.6
AXIAL RATIO, b/o
0.8
1.0
Fig. 3-61. Concentration as a function of axial ratio for the spheroid packings of Figs. 3-59 and 3-60. (After Allen, 1969, fig. 3, p . 314; courtesy Geol. Mag.)
3-61 shows that: (1)C, for example, decreases increasingly rapidly as the axial ratio falls below unity (case V); (2) C declines increasingly rapidly (more so than in case V) as the axial ratio decreases below unity (case VI); (3) C decreases rapidly with the reduction of the axial ratio below unity (case VII); (4) C at first increases with falling axial ratio, but thereafter decreases rapidly (case VIII); and ( 5 ) C at first increases slightly before starting to decrease rapidly as the axial ratio is reduced below unity (case IX). Aside from the regular, well-defined packings presented here, it is not clear how they are to be used to form a model of the concentration of real sedimentary particles arranged in a partially disordered manner. On the basis of theoretical considerations provided by Allen (1969), it seems that on combining the anisotropic rhombohedra1 arrangement of packing IV with the two-dimensional isotropic packing VI, a field that partly represents natural sands is obtained, as shown in Fig. 3-62. As pointed out by Allen (1965, pp. 317-318), only a comparatively small proportion of spheroids in an open, isotropic packing is necessary to give concentrations in the observed range of natural sands. Theoretically, the particles of a natural sand should show a substantial degree of dimensional ordering, though not a perfect one. The degree of ordering should increase with increasing grain concentration, C, for a constant axial ratio. After the theoretical studies, an attempt was made t o compare the results with natural sands, e.g., the grain relationships were investigated. In order to exclude the possibility of secondary changes of the original depositional
K.H. WOLF AND G.V. CHILINGARIAN
152
z-
s I- 0.6 U
a I2 W
0.4
0
u w
-J
I-
0 0.2
a
a!
0
0.4 M 0.8 I.o AXIAL RATIO, b/a Fig. 3-62. Concentration of a spheroid assemblage of mixed packings (IV and VI) as a function of axial ratio and proportion of packing VI given as (13). (After Allen, 1969, fig. 4, p. 317; courtesy Geol. Mag.)
0
Q2
fabric by compaction, for example, the samples collected had relatively high contents of secondarily introduced cement and lacked: (1) empty pore space; (2) evidence of recementation and corrosion of the detrital grains; and (3)evidence of physical distortion or breakage of particles. It was found that the natural sandstones had local patches of grains packed according to the V through IX patterns (Fig. 3-62),which although uncommon were not rare. In the future, maybe statistical analyses can be performed on the occurrence and abundance of these various packing patterns. The results shown in Fig. 3-63are as follows: a,b = long axes virtually at right angles, packing V; c-e = three elongated grains, in a manner as in packings VI and VII; f-i = trios of grains, similar to packings VIII and IX; j , h = two sets of four grains each arranged roughly along the edges of a square or rectangle, similar to packing VI; and 1-p = clusters of a large number of grains arranged in a manner similar to packing VIII. From these results, Allen (1969,p. 320) concluded that the “concentration of a partially disordered assemblage of spheroids can be represented by a model in which one open and one close regular packing are combined in a proportion to yield observed concentration in natural sands and sandstones. In terms of such a model, the concentration of a natural sand could be achieved if a comparatively small percentage of an open regular packing were combined with a comparatively large percentage of a close regular packing. Moreover, the percentage would be such that the sand would stiH display a substantial degree of dimensional ordering of the particles.”
DIAGENESIS OF SANDSTONES AND COMPACTION
153
Fig. 3-63. Cluster of real grains simulating the packings of Fig. 3-60, as observed in sandstones of Old Red Sandstone age. (After Allen, 1969, fig. 5, p. 319; courtesy Geol. Mag. 1
The concept of Morrow Morrow (1971) used a concept that is not always given sufficient consideration, possibly as a result of difficulties in measuring the parameter which he termed *‘packingheterogeneity”. He defined it as ‘‘local variation in sorting or, more strictly, local variation in pore-size distribution”. By employing Figs. 3-64to 3-67,he gave examples of homogeneity-heterogeneity determined by grain-size and packing variations. In Fig. 3-64there is no variation in grain-size distribution from region to region within the illustration, whereas there is such a variation in Fig. 3-65,which makes the texture homogeneous and heterogeneous, respectively. In Fig. 3-66the style of packing is random, whereas in Fig. 3-67part of the particles show cubic and other hexagonal cells, so that again the texture is homogeneous and heterogeneous, respectively. Suction (capillary pressure) required to drain the samples is inversely proportional to the pore size. The slope of the capillary pressure curve obtained in the laboratory experiment (e.g., see Morrow, 1971) re-
154
K.H. WOLF AND G.V. CHILINGARIAN
Fig. 3-64. Homogeneous distribution of rock grains illustrated as monolayer. (After Morrow, 1971,fig. 1, p. 515;courtesy Am. Assoc. Pet. Geologists.) Fig. 3-65.Heterogeneous distribution of rock grains illustrated as monolayer. (After Morrow, 1971,fig. 2, p . 515;courtesy Am. Assoc. Pet. Geologists.)
flects the pore-size distribution of the sediment, because the larger pores tend to drain before the smaller ones in a definite sequence based on pore size. The flatness of the capillary pressure curve, i.e., when the major portion of the curve tends t o be parallel t o the abscissa, is indicative of uniformity in pore size. The curve will show an “irreducible water saturation” (Fig. 3-68A) independent of further increase in the externally applied pressure, so that this property is a definitive, characteristic property of a rock (see Introduction chapter, Vol. 1).According to Morrow, fluid and rock variables (i.e., interfacial tension, viscosity, fluid density, visco-elasticity, particle shape,
Fig. 3-66.Homogeneous (random) packing of spheres illustrated as monolayer. (After Morrow, 1971, fig. 3, p. 516;courtesy Am. Assoc. Pet. Geologists.) Fig. 3-67.Heterogeneous packing of spheres containing high proportion of cubic and hexagonal cells illustrated as monolayer. (After Morrow, fig. 4, p. 516; courtesy Am. Assoc. Pet. Geologists.)
DIAGENESIS OF SANDSTONES AND COMPACTION
155
.ARREOUCIBLE SATURATION
SATURATION, %
Fig. 3-68A. Form of typical capillary pressure drainage curve. Irreducible saturation is dependent upon the pore geometry but independent of absolute pore size, whereas range of displacement pressures is dependent on absolute pore sizes. (After Morrow, 1971, fig. 5, p. 517; courtesy Am. Assoc. Pet. Geologists.) Fig. 3-68B. Retention of liquid at irreducible saturation; rock grains are shown as a monolayer. (After Morrow, 1971, fig. 6, p. 517; courtesy Am. Assoc. Pet. Geologists.)
size and distribution, and porosity) have little influence in themselves on the magnitude of irreducible saturation. Morrow’s experiments demonstrated that irreducible saturation can be an index of packing heterogeneity. He even suggested packing heterogeneity, which can be used as a supplementary characterization parameter of porous rocks, may be found to correlate with depositional environment and changes due to diagenetic processes, such as cementation and compaction. It has been found that irreducible saturation of sedimentary rocks increases as permeability and/or porosity decreases, for example. As to the relationship between irreducible saturation and packing, one should note that a variety of packings can give the same irreducible saturation, so that correlation is not a simple one. Figure 68B depicts the retention of liquid at irreducible saturation. The concept of “packing heterogeneity” must be given full consideration in the study of textures as a result of compaction, because any primary heterogeneities will control the subsequent style of grain movement during compaction.
156
K.H. WOLF AND G.V. CHILINGARIAN
The measuring technique of Kahn Kahn (1956) offered a method to measure packing and stated that packing is the mutual spatial relationship among the grains. Packing can be measured in thin sections by traversing the rock slice using a mechanical stage; the ocular must have a cross-hair and a micrometer to measure the length of the grains traversed (Fig. 3-69). As explained by Kahn and illustrated in Fig. 3-69, two concepts or measurements are used:
(1)Packing proximity, Pp = q/n X 100
(3-1) where q = number of grain-to-graincontacts, n = total number of contacts, as well as total number of grains, and 0 < q < n.
cgilt) n
(2) Packing density, Pd =(m
i=1
X 100
(3-2)
where m = correction term for various combinations of ocular, objective, and scale, t = total length of traverse, g = intercept values, n = the total number n t . of grains in a given traverse across the thin section, and 0 < i = 1 gi rn
c
Packing proximity, Pp, measures property of packing, expressed as the percentage of grain-to-grain contacts in a traverse of n contacts, whereas packing density, Pd, measures the aggregate property of packing. Kahn’s technique has been used with success, e.g., by Martini (1972).
Depth of burial and diagenesis (Taylor, 1950) Taylor (1950) studied several sequences of sandstones to determine what influence depth of burial has had on diagenesis, e.g., as reflected by graincontact varieties. The sandstones investigated by her cemented very early, and their textures and fabrics were controlled by depositional and accumulational factors. On the other hand, when induration occurred after additional beds of sediments accumulated (i.e., after burial), the role of many chemical and/or physical diagenetic factors may have been very important in determining the origin of secondary textures and fabrics. The controlling factors included: (a) mineralogic composition of the sand; (b) roundness and sphericity of grains; (c) sorting and grain-size distribution; (d) stratigraphic distribution of sand, i.e., whether a regional blanket or localized; (e) depth of burial; (f) tectonic or diastrophic processes; (g) temperature; (h) ground-water circulation condition; and (i) geologic time. These variables determine the type of pore-space reduction, which occurs as a result of: (1)pore filling; (2) elastic and plastic deformation of grains; (3)
DIAGENESIS O F SANDSTONES AND COMPACTION
c = c a r b o n a t e cement s = silica cement m =matrix .=void
157
-
ith grain 9,: i t h intercept value t = length of t r a v e r s e
ti
Fig. 3-69.Schematic representation of a sand-size sedimentary rock. The total number of grains in the traverse is nine. Of these nine grains there are nine contacts, i.e., grain 1 is in contact with matrix, grain 2 with a rock fragment, grain 3 with a void, grain 4 with matrix, grain 5 with carbonate cement, grain 6 with silica cement (an overgrowth), grain 7 with silica cement, grain 8 with matrix, and grain 9 with matrix. This results in, but one grain-to-grain contact. Packing proximity is, therefore, equal to 1/9 X 100 = 11.11%. Packing density is the sum of the intercept-size values for the nine grains, g, + g, + g, + g4 + gs + g6 + g7 + ge + gg = zr==tgi, divided by the length of transverse, t, and corrected for magnification. (After Kahn, 1956, fig. 2, p. 391; by permission of The Univ. of Chicago Press, copyright @ 1956 University ofChicago.)
solution and redeposition of dissolved material; and (4) plastic flow of material under pressure. Thus, pore-space reduction and compaction are of physical and/or chemical origin. Taylor found that in sandstones which have undergone early cementation (i.e., pressure effects are absent or at a minimum) there are three types of contacts: tangential, long, and concavo-convex. The shape of the grains and original packing control the grain-to-grain fabric. Tangential contacts (Table 3-XX) occurring as points in the plane of a thin section, have maximum development in deposits that were cemented early by pore infilling. Long contacts (relatively straight lines in the plane of thin section, see Fig. 3-47 and ,Table 3-XX) are more variable, just as the causes and controlling factors of their occurrence. A small edge of a grain or only part of an edge may be involved and thus form only a very short line. If grains are flat or have straight edges, then the nature of the final packing and orientation of the
K.H. WOLF AND G.V. CHILINGARIAN
158
grain accumulations control the proportion of long contacts. A sand with packing that results in maximum porosity has a minimum number of long contacts. With increasing degree of compaction, the number of long contacts increases. Concavo-convex contacts (curved line in the plane of thin section as shown in Fig. 3-47) are rarely found in sand in which the porosity has been reduced by monomineralic pore filling. In this case, the concavoconvex contacts are the result of the shape of the grains involved. The number of contacts per grain is small when early cementation took place. Random sections through haphazardly-packed spheres of uniform size have 0.63 contact/grain. Taylor has found 1.6 contacts/grain to exist in experimental sand deposits. Variations in size and shape are important in determining the number of contacts. Floating grains (no contact in the plane of thin sections as illustrated in Fig. 3-47) occur in sandstones that underwent early cementation. It should be pointed out, however, that in three dimensions these grains are usually supported. Chance packing of spheres gives rise to 47% floating grains, whereas the experimental sand prepared by Taylor had about 17% floating grains. In some cases, the power of crystallization of cement between the grains may have pushed them apart to form the “floating-grain” fabric. This may be indicative of either very early diagenetic precipitation of the cement and/or suggest a lack of thick overburden. This
EXPERIMENT&
SAND
FLOATING -TANGENTIAL=
LONG
CONCAVO-CONVEX- NONE
j
Fig. 3-70. Pressure phenomena, types. of contacts, and nature and percentage of secondary pore filling in Wyoming sands as related to depth. (After Taylor, 1950, fig. 1, p. 7 0 8 ; courtesy Am. Assoc. Pet. Geologists.)
DIAGENESIS OF SANDSTONES AND COMPACTION
159
interpretation applies to the volcanic arenites investigated by Wolf and Ellison (1971), which have a disrupted framework resulting from the calcite cement precipitation. Figure 3-70 indicates the general distribution of cement (<7%) in the sandstones studied by Taylor. In this figure, the dashed line represents less than 1%of the particular grain-contact type under consideration, whereas the full line depicts less than 10%. The various grain-to-grain contacts are also shown in Figs. 3-71 and 3-72. Pore-space reduction by plastic flow of material and by solution-recrystallization (or precipitation) of material occurs under pressure as a response t o burial and, perhaps, is controlled by the increase in geothermal gradient. Evidence of pressure is demonstrated by sandstones in which the volume of pore space plus the volume of chemical cement is less than the original pore volume of the sand. If a well-sorted sand has three or more contacts per grain, pressure probably has reduced the porosity. Pressure is also indicated where (1)a relatively large percentage of the contacts are long, concavoconvex, and sutured, and (2) grains are fractured, crushed, or bent. Although the above three contact types are chiefly due to pressure, concavo-convex and long contacts may also be due t o the nature of original packing and grain shape, as discussed above. Concavo-convex contacts are due to (a) plastic flow of the yielding grain, and (b) solution at points of contacts between grains. In the latter case, the dissolved material may be redeposited in nearby pores or removed in solution. DEPTH AND FORMATION 2885
4535'
' X P ~ ~ ~ ~ESAVERDE T A L SHANNON NUMBER OF CONTACTS PER GRAIN
"FLOAT IN G' GRAINS
CONCAVOCONVEX
SUTURED
4.4
I
4.9
1
5.2
' -
-
.3%
16 64h
I
9.6
~
191
.9
-
-
51.6
51.5
45 0
28 5
28 I
23 I
18 5
I9 7
31.8
Fig. 3-71. Number'and types of contacts in experimental and Wyoming sands. (After Taylor, 1950, fig. 2, p. 7 1 4 ; courtesy Am. Assoc. Pet. Geologists.)
K.H. WOLF AND G.V. CHILINGARIAN
160
+
-
3 60 v -
a
w40n
-
20
0
4 *.......... I
0
I
2000
....... .............. ” , 4.*:.. O-..-o...
3
-
1
......
I
I
4000
6000
8000
DEPTH Fig. 3-72. Graph showing number and type of contacts in Wyoming sands. 1 = tangential contacts; 2 = long contacts; 3 = number of contacts; 4 = concavo-convex contacts; 5 = sutured contacts. (After Taylor, 1950, fig. 3, p. 175; courtesy Am. Assoc. Pet. Geologists.)
Plastic flow has been observed in quartz, feldspar, and chert as well as in rock fragments, especially in shales, mudstones, phyllites, and schists. The numerous factors which determine the tendency of grains to yield include crystallographic orientation, chemical structure and composition, physical properties, original shape, amount and direction of pressure, amount of heat, amounts and types of matrix and interstitial fluids, and relative solubilities of the minerals. Taylor (1950, p. 710) stated that sutured contacts are due to local solution which produces wavy or zig-zag contact lines in the plane of thin sections. Impurities, such as clay, iron oxides, and bituminous material originally present as dust rings on the grains, are left behind as “insolubles” along the contacts. Extensive pressure solution may result in microstylolite contacts. Long contacts are due to (a) original packing of grains that have straight sides or edges, (b) precipitation of secondary cement, and (c) pressure. Consequently, great care must be taken in making an interpretation. In case a, some long contacts may result very early during compaction when grains rotate and adjust themselves. When plastic flow occurs in cases where grains have a straight edge, long contacts form. On the other hand, where adjacent grains are round, concavo-convex contacts will result. Long contacts between grains may represent a transitional stage between the original tangential and final sutured contacts. Figure 3-70shows the number &d types of grain-to-graincontacts, whereas in Fig. 3-72this data is presented graphically. There is a definite increase
DIAGENESIS OF SANDSTONES AND COMPACTION
161
in the number of contacts per grain with increasing depth, from an average of 2.5 contacts/grain in the Mesaverde Formation at 2885 f t to 5.2 contacts/ grain in the Morrison Formation at a depth of 8343 ft, with three intermediate values. The type of contacts shows a progressive change with depth, from those controlled by depositional packing to those resulting from pressure. Each curve in Fig. 3-72 has its own specific trend with depth. No relationship to depth of burial has been established for crushing, flowage, development of twin lamellae, cracking, and bending, which have also been recorded. As Taylor (1950,p. 716) has pointed out, the above-mentioned type of investigation is useful in studies of: (a) porosity and permeability reduction in vertical and horizontal sequences; (b) types and relative importance of processes that cause pore-space changes in sandstones that differ in composition, environment of deposition, and postdepositional history (i.e., diagenesis-burial epigenesis-metamorphism); (c) predicting porosity and permeability at different depths; and (d) determining the movements of hydrocarbons and ore fluids (see approach by Fuchtbauer, 1961, 1967a; and separate section on p. 311). Numerous investigators have employed the technique described by Taylor. Gaither (1953)presented some interesting results of investigations of packing in laboratory-sedimented, medium-grained St. Peter sand. The well-sorted sand with a porosity of 37% showed the following fabric: (a) 0.85 contacts per grain in the thin sections prepared; (b) 46% of the grains were in a floating position; (c) 31% of the grains were in contact with only one other grain; (d) 16% of the grains were in contact with two other grains; (e) 6%of the grains were in contact with three other grains; and (f) 1% of the grains were in contact with four other grains. More than three-quarters of the grain-to-grain contacts were of the tangential type, and most of the rest were long contacts. Concavo-convex contacts were uncommon, whereas sutured contacts were absent. Wright (1964)used Taylor’s (1950)approach in his investigation of clastic rocks. Figure 3-73 shows the results of a series of quantitative determinations of the number of grain contacts. Wright (p. 767) stated: “for most sediments the histogram roughly follows the shape of a probability curve, but in very poorly sorted sediments a tail occurs at one side of the histogram” (No. 15, Fig. 3-73).This tail represents the large grains, which are in contact with a large number of the surrounding smaller grains. All those histograms that have maxima around one grain contact (Nos. 6 and 8, Fig. 3-73),represent rocks that underwent early cementation. Many of the sediments studied contained clay minerals as a matrix that was not solid or resistant enough *toprevent compaction, but prevented compaction to proceed as far as in the clay-free deposits. Wright (1964,p. 758), therefore,
K.H. WOLF AND G.V. CHILINGARIAN
162
0
m ' 0 2 4 6 1
0 2 4
NUMBER OF GRAIN-CONTACTS PER GRAIN Fig. 3-73.Histograms prepared from determinations of the number of grain contacts per grain in fifteen thin sections of sandstones and siltstones. Sections 1-1 0 and 15 = Middle Grit Group of the Black Hill area, Holmfirth; 6 = siltstone below the Readycon Dean Series; 2 and 15 = Readycon Dean Series; 4 = siltstone below the Heyden Rock; 1, 3, 5 and 7 = from the Heyden Rock; 8, 10 and 11 = silts and sandstones above the Heyden Rock; 9 = Huddersfield White Rock; 12 = Rough Rock of Staffordshire; 13 = the Morridge Grit, Staffordshire; 14 = Elland Flags, Fagley Lane Quarry, Bradford. (After Wright, 1964,fig. 1, p. 757;courtesy ?J. Sed. Petrol.)
concluded that some account must be taken of the clay content before the significance of the average number of grain contacts can be assessed (Fig. 3-74),There are two factors which control the average number of grain contacts: depths of burial and proportion of non-quartzitic constituents in the quartz-rich deposits. Whether the number of grain contacts increases with depth in a proportional fashion or not has not been determined, neither by Taylor (1950)nor by Wright (1964).Figure 3-74,however, shows that there is a relationship between average number of grain contacts and percentage of quartz (or the percentage of non-quartz minerals). The small scatter may be due to variation in sorting. The straight line in Fig. 3-74cuts the grain-contact axis at a value of just over six, meaning that if the rocks were purely quartzitic (i.e., no day matrix), an average number of contacts per grain of about 6.2could be expected. Based on Taylor's observation that the
163
DIAGENESIS OF SANDSTONES AND COMPACTION
1
Lu
0
0
I
1
t
CEMENT
1
GRAINS
PERCENTAGE OF NON-QUARTZ CONSTITUENTS
I:) GroinlMatrix Ratio
MATRIX
Fig. 3-74.Modes of grain-contact histograms (Fig. 3-73)plotted against the percentage of non-quartzitic constituents. (After Wright, 1964,fig. 2,p. 758;courtesy J. Sed. Petrol.) Fig. 3-75. Ternary classification of sedimentary rocks using fundamental end members and showing range of samples studied. (After Smith, 1969, fig. 1, p. 261;courtesy Am. Assoc. Pet. Geologists.)
sandstone at a depth of 8300 f t in subsurface had 5.2 grain contacts per grain, an extrapolation was made by Wright for his own rocks, which suggested a burial depth of about 10,000 ft. A crucial question arises, of course, as to what extent extrapolations of this kind are reliable in other instances and an answer remains to be found by future investigators. Overburden pressure alone may not be the only parameter that controls the number of grain contacts and the type of contacts, because possibly the geothermal gradient and pore-solution chemistry, to name only two variables, may be effective also.
Puartz
I
95
75 25
F-Well Cores I
50 50
I
!
25 75
- I _Tyrone I
5 1 Dolomite 95 I ( O h )
Section
I
Fig. 3-76.End member dassification of dolomite-quartz system showing range of samples studied. (After Smith, 1969,fig. 2,p. 263;courtesy Am. Assoc. Pet. Geologists.)
K.H. WOLF AND G.V. CHILINGARIAN
164
Study o n a carbonate-quartz system (Smith, 1969) Smith (1969)investigated the relationships among petrography, porosity, and permeability in a carbonate-quartz system, which is particularly interesting as most investigations by other researchers were performed on monomineralic rock suites. Smith collected his specimens from two sedimentary rock suites of which one set was composed of nearly pure (quartz-poor) dolomite and the other of nearly pure (carbonate-poor) quartzite, as diagrammatically presented in Figs. 3-75 and 3-76.They are referred to as “F-well cores” and “Tyrone section specimens,” respectively. As illustrated in Fig. 3-76,the samples of the F-well cores range in lithology from 40% quartz (= sandy dolomite) to 100% quartz (= orthoquartzite). Figure 3-77shows a plot of porosity versus insoluble residue (I.R.) with a definite positive relationship. A few scattered points represent relatively higher porosity values than would be expected from their corresponding I.R. values. In general, porosity increases with increasing I.R. Smith mentioned several possible factors that may have resulted in this deviation, including compaction and cementation. There are two groups noticeable in Fig. 3-77, as shown, for example, by the relative flatness of the curve below 30--40% I.R. Fig. 3-78 shows the relationship between packing and porosity, which appears to be statistically significant and positively correlative. From the correlation in Fig. 3-77,one might expect a statistically significant positive relation between packing and I.R., which is indeed the case as shown in Fig. 3-79.One should note that between 80 and 100% I.R.,the degree of packing varies over a range of nearly 50%.
3
>:
“1 IS
::
INSOLUBLE RESIDUE, X
Fig. 3-77. Scatter diagram showing relation between percent insoluble residue and percent porosity for F-well ( 1 ) hnd Tyrone (2) samples. (Modified after Smith, 1969, fig. 6, p. 268; courtesy Am. Assoc. Pet. Geologists.)
DIAGENESIS OF SANDSTONES AND COMPACTION LO
1
1
I
I
I
I
1
I
165
I
0 0
0
I6
-
0
0
O
O 0
0
0
0
0
aQ
s 12-
0
5 P
o
2
-
0
0
~
0 0
0 8
0
o
o
! ? -
0
o
0
0
o
0
o
1 0
0
0
0
D
0
0
0 0
4 r
l o
0
0 - O '
0
'
20
I
I
40
60
I
80
I
~
K)o
0
40
80
PACKING, %
Fig. 3-79.Scatter diagram showing relation between percent packing and percent insoluble residue in F-well samples. n = 56; r = 0.489. (After Smith, 1969, fig. 8, p. 269; courtesy Am. Assoc. Pet. Geologista.)
166
K.H. WOLF AND G.V. CHILINGARIAN
Smith concluded that for this carbonate-quartz system an increase in porosity accompanies an increase in I.R. (quartz) between about 50 and 100% I.R. Part of the increase may be the result of changes in the degree of packing of the quartz grains. Intergranular porosity, which certainly is related to the packing (Fig. 3-78),was observed under the microscope in most of the samples containing 75-100% I.R. Based on statistical component analysis, Smith (p. 272) concluded that “the dependent property, porosity, is associated closely with the independent properties, I. R. and packing. The other independent properties, quartz grain size, sorting, shape, and standard deviation of shape, are unrelated to each other, and also are unrelated t o porosity, packing and I.R. The factors controlling the variability in porosity, packing, and I.R. therefore are unlike those controlling size, shape, and sorting, or the latter properties respond differently t o the same sedimentary processes.” “. . . porosity varies primarily as a function of the amount of I.R. and the degree of packing in the samples. There is no evidence to suggest that porosity is influenced appreciably by grain size in these sedimentary rocks, although grain size commonly has been shown to be one of the most important variables influencing porosity in other studies of this nature. Shape and sorting of the quartz grains in these samples also are unrelated to porosity.” In the section on the Framework Concept, Smith (p. 274) stated that the original porosity of the carbonate-quartz system was reduced as compaction continued until the framework of the grains began to support the overburden and thus preserved the remaining pore spaces. In the case of a pure limestone or dolomite composed of non-granular components (i.e., made up of lime mud or dolomite mud, which is more compactible than the deposits of oolites, skeletal fragments and pellets, for example), compaction reduces the pore space considerably more than it does in sediments with granular limy and quartz constituents. Continued addition of clastic or limy granular components to the non-clastic carbonate phase will eventually result in a sediment composed of enough grains in contact with each other to form a supporting framework. According to Smith, this type of granular framework will reduce or prevent compaction and generally preserve the larger pores in the sediment. As to a deposit composed of quartz grains that make up the framework, with a carbonate cement filling some or all of the intergranular pores, Smith (p. 275) stated that the final porosity in this type of carbonatequartz system depends on: (1)the original intergranular porosity which is a function of packing, for example; (2) the amount of secondary cement; and (3) the amount of cement removed by subsequent decementation. Smith compared the results of his work with those of Lucia (1962)and concluded that dolomitization of th*e rocks investigated by him occurred before compaction.
DIAGENESIS OF SANDSTONES AND COMPACTION
167
Complex diagenetic alterations (Aalto, 1972)
Aalto (1972)presented a fine example illustrating the complex interrelationships between textural, compositional and deformational factors for an orthoquartzitic or orthoquartzitic-sub-graywacke suite, with commonly uniform grain composition and complex diagenetic alterations among grains, cements, and clay matrices. In order to understand the terminology given in Figs. 3-80 and 3-81,the following information on the six grain types is necessary: (a) Undulose quartz (0-25" rotation) often has quartz overgrowths. Many grains contain bubble-like gas or liquid inclusions arranged in linear trains. In very deformed grains, these inclusions form parallel, closely-spaced trains
$ N ka 3 0 W
c
I
a
PHI MEAN SIZE
C
80 60
40
20
b
d
DEPOSITIONAL MATRIX (YO)
Fig. 3-80. Interrelationships of textural and deformational parameters. (a) Percentage of polycrystalline quartz present versus phi mean size; (b) percent of grains with overgrowths versus estimated percent of depositional matrix; (c) percent of grains with linear trains of inclusions and comb-like groups (clg) of trains versus phi mean size; and (d) percent of grains wit% linear trains of inclusions and comb-like groups of trains versus estimated percent of depositional matrix. (After Aalto, 1972, fig. 9; courtesy J. Sed. Pe tro 1. )
168
K.H. WOLF AND G.V. CHILINGARIAN CONCAVO-CONVEXlccl and SIMPLE LlNEIlbl
Fig. 3-81.Types of grain contacts. Letters presented here represent the samples studied and correspond to particular formations as follows: Q = Gutierrez-Quetame Sandstone; 2 = Cdqueza Sandstone; U = Une Sandstone; P = fine sandstones and siltstones of the lower part of the Guadalupe Group; D = Dura Sandstone; L = Labor Sandstone; T = Tierna Sandstone; G = La Guia Sandstone; C = El Cacho Sandstone; R = La Regadera Sandstone. The summary diagram (small triangle on left) portrays general relationships among textural, compositional and deformational parameters. (After Aalto, 1972, fig. 11, p. 338; courtesy J. Sed. Petrol.)
and resemble the teeth of a comb. Inter- and intragranular fractures are present in all sandstones studied by Aalto, but are most common in the oldest and most deformed rocks. Healing of these fractures by quartz gave rise to the linear trains of “comb-like” groups of bubbles. Microlites of micas, rutile, sphene, tourmaline and zircon are present in the quartz grains also.
DIAGENESIS OF SANDSTONES AND COMPACTION
169
(b) Semi-composite quartz grains have undulose extinction and overgrowths. The overgrowths and the deformational features are similar to those found in monocrystalline quartz, but are less common. (c) Composite quartz grains have undulose extinction, commonly sweeping over the entire grain. Overgrowths are uncommon, but where present they develop in optical continuity with different grains within the composite fragment. The inclusion trains and comb-like groups of trains are confined to individual grains. (d) Stretched quartz has many subparallel, elongate quartz members, and the boundaries are intensely crenulated, sutured or granulated. Strong undulose extinction commonly sweeps the entire grain, whereas the overgrowths, inclusion trains, and comb-like groups of trains are uncommon and only present in individual grains. (e) Polycrystalline quartz grains are present also. (f) Chert grains may be recrystallized, They may contain microfossils, glauconite, and silt and mud grains, all of which may be deformed and altered. Heavy minerals present as traces may show cleavage. (g) The matrix and cement is largely of diagenetic origin with kaolinite, illite, and sericite being common in older sandstones. Details on the diagenesis of “matrix-cement” and alterations of “grain-matrix-cement” and “grain-grain” are presented by Aalto. The interrelationships established by him among textural, compositional and deformational factors are shown in Figs. 3-80and 3-81.As pointed out by Aalto (p. 338), generally poor correlations exhibited in Figs. 3-79b,c, and d , and Fig. 3-81 are due to the presence of deformational features in grains formed prior to deposition by complex postdepositional deformation and by obliteration of grains by iron oxides during epidiagenesis. The data distribution in Fig. 3-81 suggest that sandstones characterized by a high percentage of simple irregular penetration contacts, microstylolite contacts, and pseudograins are commonly coarse-grained. These sandstones also contain more polycrystalline quartz, grains with inclusion trains, and grains with comb-like groups of trains. Aalto (p. 338) also stated: “Moreover, these sandstones have fewer floating grains and a clay or sericitic matrix. In contrast, sandstones with an abundance of simple line and concavo-convex contacts are more likely to be fine-grained and to have less polycrystalline quartz.” They also contain fewer grains with inclusion trains and comb-like groups of trains, more floating grains, and a greater variety of matrices and cements. Inasmuch as crystal face contacts are a product of overgrowth formation, sandstones with many such contacts commonly have more overgrowths. As shown in Fig. 3-80,with an increase in the amount of depositional matrix, sandstones commonly show a decrease in abundance of overgrowths on grains. With increasing amounts of depositional matrix, fine-
K.H. WOLF AND G.V. CHILINGARIAN
170
grained sandstones correspondingly have fewer grains with inclusion trains and comb-like groups of trains (Figs. 3-80 c , d ) . Aalto (p. 339) pointed out that exceptions have been observed for each one of the above generalizations, and that the observations described above reflect a complex interaction of several variables. Irregular and microstylolitic contacts, pseudograins, inclusion trains, and comb-like groups of trains (see Whisonant, 1970), associated with strong undulosity, all reflect increases in deformational stresses and presence of special chemical environment. Porosity and packing index Based on data presented by Masson (1951), Griffiths (1967b) related porosity to packing index (see definition on p. 135), as shown in Fig. 3-82. The artificially-packed sand had an index of 10, whereas the packing index of 32 sandstones from the Gulf Coast and U.S.A. Midwest oil fields varied from 18 to 62. Rosenfeld (1953, in Griffiths, 196713) using Masson’s technique, found that the average packing index of 6 1 orthoquartzite specimens (Fig. 3-83) was 41.4%, with a standard deviation of 17.7% and a range of 3 to 75%. Effects of textures and composition (Renton et al., 1969) Renton et al. (1969) studied the effects of textures and composition on compaction for various types of sands. One set of experiments was performed on quartz crystals surrounded by small grains but with no liquid, in order to simulate purely mechanical effects. After compaction, the quartz crystals showed visible fractures at points of contact and both radial and
0
PACKING INDEX
Fig. 3-82. Relation of porosity to packing index (After P.H. Masson, 1951, in Griffiths, 1967b, fig. 8-2, p. 1 6 9 ; copyright @ 1967, McGraw-Hill, New York.)
DIAGENESIS OF SANDSTONES AND COMPACTION
171
ui 12 w
J LL
5r n 8 LL
0
K :4
5z
0
0
20
40
60
PACK I NG I N D E X,
80
O '/
Fig. 3-83.Frequency histogram of grain-to-grain contacts in 61 specimens of Oriskany quartzite, f (mm arithmetic mean) = 41.4%; 8 (standard deviation) = 17.7%. (After Rosenfeld, 1953, in Griffiths, 1967b, fig. 8-3,p. 169, copyright 0 1967 McGraw-Hill, New York.)
TABLE 3-XXII Compaction of quartz and chert sands (after Renton et al., 1969,table 3, p. 1113) Expt.
Maximum Duraload (psi) tion (days)
Experimental material Initial Final Corn% void poros- poros- paction decrease ity(%) ity(%) (%)
7
6000
48
16-21 mesh round
6
6000
48
8
6000
47
14
6000
39
10
4000
23
11
4000
23
No.
quartz grains 60-80 mesh round quartz grains 60-80 mesh angular quartz grains angular quartzsilt (0.04mm) 60-80 mesh angular chert grains mixture 60-80 mesh angular quartz (50%) and angular chert (50%)
36.5
24.5
16.2
44.0
39.0
16.5
28.0
69.5
45.8
13.0
37.8
81.5
53.5
17.0
43.5
80.3
49.3
15.9
39.3
80.5
51.0
27.0
32.7
64.5
All experiments were conducted at 400"C with 0.5M NaZC03 solutions at a hydrostatic pressure of approximately 6000 psi. In experiments on quartz, load was padually increased to the maximum value over a period of 4 weeks and held at this value for the remainder of the experiment. Load gradually increased during the entire experiments on chert.
K.H. WOLF AND G.Y. CHILINGARIAN
172
annular fractures were present. The radial fractures were perpendicular to the surface of the crystals, whereas the annular ones formed nested cones. In the absence of intergranular liquid, which would have caused dissolution of the quartz at the points of stress, the quartz lattice failed as a brittle substance. Distinct pressure-solution pits were formed in distilled water, NaOH, Na2C03, and NaCl solutions, and natural brines. In a series of experiments, the relative rates of pressure solution and cementation in sands of different types were determined. The porosity, degree of compaction, and percent void reduction were measured and are given in Figs. 3-84 to 3-87. Porosity values before and after the compaction are presented in Table 3-XXII. According to Renton et al., compaction was predominantly the result of pressure solution at grain contacts, because no appreciable fracturing occurred and many concavo-convex contacts formed between grains. SiOz was precipitated as euhedral quartz overgrowths on the grains.
DURATION, days Fig. 3-84. Porosity reduction in various angular quartz and chert sands under t h e same conditions of compaction for each sample. On Figs. 3-84 to 3-87, solid symbols represent values based o n actual volume measurements, whereas open symbols represent values calculated from t h e amount of fluid removed during the experiments. As in Fig. 3-85, 5 = chert-quartz mixture; 4 = angular quartz silt (0.04 mm); 3 = 60-80 mesh angular quartz; 6 = 6--80 mesh chert. ( A f t e r Renton et al., 1969. Fig. 10,p. 1112; courtesy J. Sed. Petrol.)
DIAGENESIS OF SANDSTONES AND COMPACTION
173
Figure 3-86demonstrates the following facts: (a) The compaction rates decreased with time as the areas of contact surfaces of the grains increased and the load was distributed over larger areas. (b) Finer sands underwent greater compaction with resulting greater reduction in porosity, probably due to larger numbers of contact points. (c) Compaction of chert grains was more rapid than that of monocrystalline quartz. This may be attributed to the small crystal size of the chert or the presence of poorly-crystallized material very closely resembling a homogeneous chert, which facilitated more rapid pressure solution. (d) In mixtures of 50%chert and 50% quartz fragments, the initial com-
W-
v)
Q
W
a
50
0
I0
20
30
DURATION, doyr
40
50
0
10
20
30
40
50
DURATION, day8
Fig. 3-85.Variation in compaction rates of different types of sands under the same loading conditions that remained unchanged from sample to sample. 1 = 16-21 mesh round quartz; 2 = 60-80 mesh round quartz; 3 = 60-80 mesh angular quartz; 4 = angular quartz silt (0.04 mm); 5 = chert-quartz mixture; 6 = 60-80 mesh chert. (After Renton et al., 1969,fig. 12,p. 1113;courtesyJ. Sed. Petrol.) Fig. 3-86.Void reduction in sands of different types under the same experimental conditions that did not change from sample to sample. Legend as in Fig. 3-85.(After Renton et al., 1969,fig. 13,p. 1114;courtesy J. Sed. Petrol.)
K.H. WOLF AND G.V. CHILINGARIAN
174
paction occurred about as fast as in pure chert samples. The rate was similar to that of quartz, however, during the late stages of compaction. As shown in Figs. 3-84and 3-87, in the early stages of compaction, the pressure solution affected chert first. When the quartz grains were largely in contact, the rate of pressure solution was considerably less. Only 7% of the sample was composed of chert after 33%compaction. Another conclusion reached by Renton et al. (p. 1116) was that the rate of compaction and porosity reduction from pressure solution varies considerably depending on size, shape and type of sand grains (see Figs. 3-84 to 3-86). Pressure-solution compaction in samples having angular grains was considerably faster than in samples having round grains of the same size. Renton et al. (p. 1116) observed that the appearance of the angular grains after pressure solution and secondary quartz growth was not greatly different from that of round grains after comparable pressure solution. Thus, very little can be deduced in natural sands about the original grain shape after moderate pressure solution, unless remnants of clearly-defined dust rings are preserved.
DURATION, days Fig. 3-87. Porosity reduction in coarse-grained and fine-grained sand under the same conditions of compaction from sample to sample. Legend as in Fig. 3-86. (After Renton et al., 1969, fig. 11, p. 1113; courtesy J. Sed. Petrol.)
DIAGENESIS OF SANDSTONES AND COMPACTION
175
As shown in Fig. 3-86, although the fine angular sand grains had considerably greater initial porosity, their porosity became less than that of the fine round sands after about 25% volume compaction. The contrast between the compactability of angular silt and coarse sand (16-21 mesh) is striking. As pointed out by Renton et al. (p. 1116), the compaction of the silt was over 3 times greater than that of the coarse sand after the same length of time, i.e., 80%and 38%loss (by volume) of voids for silt and coarser sand, respectively. Porosity loss of samples of chert grains was considerably more rapid than that for monocrystalline quartz (Fig. 3-84). This diagram also shows (cf. Fig. 3-86) that there was a marked reduction in the proportion of chert, because in mixtures of chert and quartz grains much of the presolved chert was precipitated as overgrowths on the quartz particles. Hence, original chert contents in natural, compacted sands might be impossible to determine. From the above data, Renton et al. suggested that it is possible to “estimate porosity variation in certain natural sands where the degree of pressure solution was reasonably uniform and the dissolved silica was deposited locally. If the porosity reduction is known for a sand similar to one of the types shown in Figs. 3-84 and 3-85, the reduction can be estimated for the other types given. For instance, if a porosity reduction of 23% has occurred in 60-80-mesh round sand, only a 12% reduction would be expected in 16-21mesh sand. Exact porosity predictions are not possible in natural sands because of the large number of variables involved; however, these data should be of aid in estimating the relative magnitude of the porosity reduction’ from pressure solution.”
Observations on rims and coatings (Pittman and Lumsden, 1968) Pittman and Lumsden (1968) made observations, which were similar to those by Fuchtbauer et al. (see pp. 343-363), indicating that authigenic, fibrous chlorite rims on quartz sand grains inhibited pressure solution. They observed the following types of occurrences: (a) Thick continuous coatings of chlorite which preserve porosity, because pressure solution is retarded and quartz overgrowths are absent. (b) Sparse, thin, and discontinuous chlorite coatings; the sandstones are tight (= non-porous), because pressure solution and quartz overgrowth took place. (c) An intermediate stage between a and b where chlorite coating is thick and complete enough t o have retarded pressure solution and yet is discontinuous in places to allow some local quartz precipitation. Studies such as these have been used in petroleum reservoir investigations and may find application in th’e future in unravelling the origin of certain types of ores in sediments and pyroclastics.
K.H. WOLF AND G.V. CHILINGARIAN
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Study on the source of silica (Waugh, 1970) Waugh (1970)studied the source of silica, which forms quartz cement in continental red-bed sandstones deposited as barchan dunes in hot and arid desert environments. The cement is present as perfectly-formed bipyramidal quartz crystals. Waugh considered the more commonly proposed sources, including pressure solution, but found that none can adequately explain the origin of the secondary quartz in the sandstones he studied. By establishing that: (a) the number of contacts per grain averages approximately 1.8, (b) over 90% of grains have less than four contacts per grain, and (c) over 90% of the types of grain-to-grain contacts are either floating, tangential or long (Fig. 3-88),he demonstrated that pressure solution did not supply the silica. Waugh also discussed the possibility that silica dust, which is the result of desert abrasion, could have been the source for the quartz cement. Study on grain contacts and packing parameters (Martini, 1972; Gaither, 1953,and others) According to Martini (1972),the effects of compaction are revealed by the number and types of grain contacts and measurements of packing parameters, which describe the threedimensional distribution of the sedimentary particles. This data enables the estimation of the original porosity and permeability of the rock and assists in the interpretation of the secondary changes after deposition. In his investigation, Martini used the packing density, Pd (Kahn, 1956),and packing proximity, P p , as based on grain-to-grain contacts (Taylor, 1950). These packing parameters were measured in the hematitic quartzite to determine possible relationships between them and
A A
A) NUMBER OF CONTACTS PER GRAIN
8) TYPE
FLOATING 8 I
TANGENTIEL
.../
*:\ ,$A/
283
OF GRAIN CONTACTS
4-4’
48’4
LONG
CONCAVO-CONVEX 8 SUTURED
Fig. 3-88. Grain contact analysis illustrating lack of pressure solution effects in the Penrith Sandstone (Lower Permian) of north west England. (After Waugh, 1970, fig. 10,p. 1236; courtesy J. Sed. Petrol.)
DIAGENESIS OF SANDSTONES AND COMPACTION
177
how they relate to other sedimentary properties, e.g., grain size and grain orientation. The results indicated that degree and type of packing in this case correlated strongly with the grain-size variations and that pressure solution (as determined by grain-to-grain contact types) cannot wholly explain the presence of secondary silica cement in the sandstone samples. Inasmuch as the measured distributions of the two packing parameters deviated only slightly from a normal distribution, Martini used a normal statistical testing procedure in establishing relationship between these parameters. As shown in Fig. 3-89,a significant linear correlation is present at the traverse level between packing density and packing proximity and only 10.5% of the variability (scatter) is accounted for by the regression line. Based on the results obtained from the multiple regression analysis, Martini (1972,p. 418) found that better prediction of packing proximity can be made if the intercept size is not considered at the thin-section level. Although a statistical correlation does not imply a causal relationship, the results suggested that as grain size (i.e., intercept size) increases, the amount of space occupied by the particles in a given sample (Pd)tends to increase. Martini (p. 415) stated that “packing proximity was positively correlated with packing density at the traverse level of analysis (five traverses per sample), and that it was best predicted if only the average values of packing density of the samples were considered in a multiple regression analysis”. In relating packing and grain orientation (see Martini, p. 420, for method used), two major relationships were found among packing proximity, packing density, and grain orientation: (1)In sections perpendicular t o the bedding (for grain-imbrication determination) and in direction parallel to the vector mean of grain orientation, a
PACKING DENSITY, f
Fig. 3-89. Average relationships between packing density and packing proximity (circles = average per sample; triangles = average per stratigraphic section; solid line = regression at the traverse level; dotted line = regression at the thin-section level; dashed line = regression at the stratigraflhic section level; square = grand mean). (After Martini, 1972, fig. 2, p. 419; courtesy Int. Geol. Congr., Montreal.)
K.H. WOLF AND G.V. CHILINGARIAN
178
20
' 7 50
60
ro
a
PACKING DENSITY, % ,
PACKING PROXIMITY, %
Fig. 3-90. Relationships among packing density, packing proximity and imbrication of vertical thin sections cut parallel to the vector mean 6 of the grain orientations of the horizontal slides. Significant regression line: Pd = 70.608-1.365 8; r% = 88.67. (After Martini, 1972, fig. 3, p. 420; courtesy Int. Geol. Congr., Montreal.)
significant negative correlation was found to exist between the angle of imbrication and the packing density (Fig. 3-9042). Similar relationship existed between the packing proximity and angle of imbrication (Fig. 3-90,b). The relationships existed in the case of a more open-textured fabric when less space is occupied by the grains. This fabric is due to a higher angle of imbrication, because when the angle of imbrication is higher, relatively coarse grains obtain a better stable, static equilibrium. Also, when a relatively smooth lamina is formed, only the clastic particles with a high imbrication fabric may come t o a stop and be preserved in the microenvironment of sedimentation (Martini, 1972, p. 420). (2) Martini also fuund a significant negative correlation between the percent of vector magnitude of the grain orientation and the packing proximity in cases where P p is less or equal to 18%, whereas for higher values the two variables are independent of each other (Fig. 3-91). Similarly, as also shown in Fig. 3-91, the percentage of vector magnitude of the grain orientation is independent of the packing density. Martini concluded that the correlation between packing proximity and orientation implies that at high Pp values (i.e., when frequent interactions may occur between grains), the degree of preferred grain alignment is independent of the number of contacts between particles. On the other hand, when the Pp values are low (less than 18%),a better degree of grain orientation exists in the presence of fewer grain contacts, i.e., when fewer interactions took place. The grains, which are more elongated and streamlined, have their axes best aligned, and are better oriented in the direction of flow, have a higher probability of stopping when a smooth bed (i.e., with low roughness index) is being formed. Good static equilibrium requires also stronger upflow imbrications of the
DIAGENESIS OF SANDSTONES AND COMPACTION
179
PACKING DENSITY, 7.' 50
40
60
70
w ?
-I 1O 0
20
30
10
15
:
o
.
o
18
20
0 0
25
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PACKING PROXIMITY, f
Fig. 3-91. Relationships among packing density, packing proximity and vector magnitude per cent ( L )of the grain orientation. Significant regression line: Pp = 15.50-0.279 L ; r2% = 54.36. (After Martini, 1972, fig. 4, p. 4 2 1 ; courtesy Int. Geol. Congr., Montreal.)
particles. In naturally-accumulated sediments, therefore, the sand grains have a preferred orientation and imbrication. The erosion of these grains is more difficult than in instances where the particles are randomly oriented. Martini (1971) did not find a vertical variation in packing parameters in the 40-50 f t thick sandstone sequence, but noticed significant differences on a regional scale. Gaither (1953) presented the results of the investigations of several petrologists who have measured in thin sections the number and types of grain contacts and the porosity of natural sands of varying composition, texture, and postdepositional history. According to him (p. 182), Manry (1949) determined that compaction had reduced the pore space from an assumed original porosity of 38% to an average of 23% in a number of Paleozoic orthoquartzites. Simple pore filling by chemical precipitation of cement resulted in an additional reduction of an average of 7% porosity. In these specimens, the number of contacts per grain ranged from 1.2 to 3.1, which is considerably below the maximum value of 5.2 of Taylor (1950). This may be attributed to: (1)the different lithologies involved; (2) relatively early lithification of the quartzites with a consequent arresting of pore-space reduction; or (3) shallower burial of the quartzites. Beaudry (1950) investigated Pennsylvanian sandstones from depths ranging from 8353 to 13,096 ft, and found that the original porosity of about 37% was reduced to less than 9%. About 60-65% of this reduction in porosity was the result of simple pore filling by quartz and carbonate cement. The remaining 35--40% of pore space reduction was caused by: (1)
180
K.H. WOLF AND G.V. CHILINGAREAN
physical rearrangement of grains; (2) crushing and bending of grains; (3) solution and plastic flow; and (4)replacement and recrystallization. Low average number of contacts per grain (= 2.1-3.6), as compared with those reported by Taylor, and relatively small effects of overburden pressure, in general, were attributed by Beaudry to (1)more resistant composition, i.e., high percentage of quartz grains and low percentage of rock fragments and feldspar, and (2) early introduction of cement. There was no relationship between depth and (a) porosity, (b) number of contacts per grain, or (c) the type of contacts. There was, however, a relationship between the number of contacts per grain and composition, namely, the samples with the most contacts per grain had the largest percentage of calcite grains, whereas those with the fewest grain contacts per grain had the largest percentage of quartz. Hays (1951) studied samples from a depth'of 13,000-21,000 f t and observed that 43-75% of the original porosity was accounted for by simple pore filling, whereas the remainder was eliminated by pressure effects. Average number of contacts per grain ranged from 1.7 to 2.2. Hays concluded that in his samples the number and types of contacts were not related to depth. Samples with the largest number of carbonate grains had abnormally high percentages of sutured contacts. It seems then that, in some cases, composition may have a more important effect on the number and types of contacts than depth of burial. Figure 3-92 shows that when sands prepared in the laboratory are compared with natural, well-indurated sandstones, the percentages of grains with given numbers of contacts, as measured in polished and thin sections, widely
NUMBER OF CONTACTS
Fig. 3-92. Comparison of percentages of grains with given number of contacts as obtained from thin and polished sections of experimental and natural sands. 1 = thin sections; 2 = polished sections; 3 = implied-values for uncompacted sand; 4 = average of Beaudry's (1950) deeply buried sands (8,000-13,000 ft). (After Gaither, 1953, fig. 6, p. 191; courtesy J. sed. Petrol. )
DIAGENESIS OF SANDSTONES AND COMPACTION
181
diverge. In this diagram, the dotted curve 4 represents the distribution of grains with given number of contacts for Beaudry's (1950) deeply buried (8,000-13,000 ft) sandstones. According t o Gaither (1953),p. 193), compaction causes this curve t o differ considerably from the same kind of curve prepared for uncompacted material (heavy broken curve 3). The following four measurable variables can be employed to determine the presence or magnitude of compaction: (1)porosity; (2)average number of contacts per grain; (3) percentage of grains with given numbers of contacts; and (4) percentage of contact types. The first three parameters are closely related and variations in one will be attended by variations in the other two. The most important factors that can cause such variations in these parameters are compaction and sorting. Consequently, interrelationship between compaction and sorting must be thoroughly studied. In the study of undulatory quartz as related to deformation during burial,
q Y z
u
Fig. 3-93. Variation of the undulatory extinction of quartz in relation to the radiogenic age of different granitic rocks. (After De Hills and Corralin, 1964, fig. 1, p. 364; courtesy Geol. SOC.Am. Bull.)
K.H. WOLF AND G.V. CHILINGARIAN
182 100
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.
~
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FAULTED STRATA NON-FAULTED STRATA I
I
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d
the following factors have to be considered: (a) the derivation of undulatory quartz from the source rocks, e.g., granites and metamorphics; (b) differential transportation of undulatory versus non-undulatory quartz (see Conolly, 1965); (c) formation of undulatory quartz during burial; (d) formation of undulatory quartz by tectonism; and (e) conversion of undulatory quartz through recrystallization to less undulatory and/or non-undulatory quartz. By studying granitic rocks of three different ages, De Hills and Corvalh (1964) showed that the intensity of undulatory extinction increases with age as a result of increasing degree of tectonic deformation by successive orogenic movements. Their results are presented in Fig. 3-93. The publications by Blatt (1967) and Blatt and Christie (1963) contain information on properties of quartz from different source rocks, its undulatory extinction, deformation of quartz after deposition, and other related data. Conolly (1965) demonstrated that the percentage of undulatory quartz is greatest near faulted zones, which may also be true in folded rocks (Fig. 3-94). The influence of physical compaction Little information is available on the influence of physical compaction on surface features of grains, and considerable research work is required. It has
DIAGENESIS OF SANDSTONES AND COMPACTION
183
been also noted that certain chemical processes can form “frosting”, for example. These processes may or may not be related to the compaction fluids moving into and out of the sedimentary rock, and it may be impossible to determine the precise origin of the solutions that caused frosting. Surface alterations of grains occurring during calcite cementation or calcite replacement may result in textures that are easily confused with the frosting of eolian sand grains and, consequently, may lead to wrong paleoenvironmental interpretations. Krinsley and Donahue (1968), using an electron microscope, recognized four types of surface textures produced by the microenvironment during diagenesis: crystal surfaces, solution surfaces, pressure-solution striations, and fracture surfaces. One should note that these features are superimposed on surface features which were the result of the original sedimentary and depositional environment. Such studies are necessary t o determine eventually to what extent compaction gives rise t o different surface textures that can be distinguished from each other as to genesis. Margolis (1968) found that (1) sand grains from beaches of low-wave activity were almost all completely covered with crystallographically-oriented etch features; (2) sand grains from littoral zones with moderate wave activity show a combination of mechanical abrasion features and chemical etch features; and (3) quartz grains from beaches with high wave activity predominantly show impact V’s, breakage blocks, and scratches, with very few chemical etch triangles or rhombs. These findings are summarized in Fig. 3-95. Margolis (p. 255) stated that more investigations in other localities and in older sedimentary rocks are required to determine whether the abovedescribed findings d e generally applicable or not. Obviously, they could be applicable only in cases where diagenesis has not obliterated the original features. Margolis (p. 248) remarked: “The degree and the expression of this diagenetic pattern is a function of the length of time the grain has been exposed t o circulating meteoric and ground waters. Other factors that must be taken into consideration are climate, chemical composition of the local ground waters, pH and silica saturation of the waters, and the permeability of the sediment where the sand is located. A high clay content may protect a sand grain from diagenetic change and preserve the original surface textures.” Diagenesis in older sediments may have had a complex history, because the chemical composition and pH of interstitial fluids on the continents can vary greatly with time, from fresh ground water t o connate water during burial and then back to ground water after uplift. Hence, chemical corrosion of sand grains during burial, by compaction or other subsurface fluids, can either obliterate or destroy primary surface textures as mentioned by Margo-
184
K.H. WOLF AND G.V. CHILINGARIAN
v)
W IX
3
2
LOW
Energyl--Moderote I
1
Energy-I-High I
I
Energy I
-
I
50-
-
MEAN WAVE HEIGHT, cm Fig. 3-95. Correlation between mean wave height of the beaches sampled and the occurrence of mechanical features, chemical features, and combinations of both types of features on the surfaces of the sand grains. The ten beach samples were plotted according to the mean wave height (cf. table 1 of Margolis’ paper). The vertical axis indicates the number of grains, out of the total of 50, which exhibited the indicated features. (Energy classification of the beaches from Tanner, 1960). (After Margolis, 1968,fig. 10; courtesy Sed. Geol.)
lis (1968).It may also form new textures that can be easily misinterpreted as being of primary origin. Formation of ‘ ~ l y p t o m o r p h s ” .Rukhin (1958)stated’that the decrease in thickness of sediments during consolidation and compaction is favorable for the formation of “glyptomorphs”, i.e., crystalline aggregates of low solubility that can grow above the sediment’s upper surface. They can be composed of crystals of halite and gypsum, for example, forming within the sediments. Sometimes, these crystalline aggregates include particles of the sedimentary framework. Sand crystals formed by quick crystallization from solution constitute one well-known example. Due to the decrease in thickness of the sediments during compaction, the outline of these crystals may be impressed on the overlying and underlying sediments (Fig. 3-96).
Transitional stages. In the above discussion of sandstone textures and fabrics formed by compaction, no consideration was given t o those related to transitional stages between diagenesis and metamorphism, so that a separate, brief section is devoted to this particular subject. Skolnick (1965), for example, discussed the “quartzite problem” and
DIAGENESIS OF SANDSTONES AND COMPACTION
185
Fig. 3-96. Schematic sketch of the formation of preserved glyptomorphs or casts of halite crystals during compaction o f sediments. A = origin of crystals in newly-formed sediment; B = compaction of the sediment; and C = protruding glyptomorphs above bedding. (After Rukhin, 1958, fig. 56, p. 235.)
mentioned that many sedimentary quartzites are not true “orthoquartzites” (Len, quartz grains cemented by quartz infilling of original intergranular spaces), but have undergone pressure solution during compaction (= “pressolved quartzites”). The latter types are by far the most common quartzites and have certain characteristics that occur also in “metaquartzites”, i.e., formed or modified by metamorphism. Yet, little is known on the effects of time, depth of burial, tectonism, and the properties of the intrastratal fluids on the transition of quartz sand or sandstone to pressolved quartzite and, finally, to metaquartzite. Skolnick’s publication is essential for petrologists, because he described the origin of the ortho- and metaquartzites first and then outlined the terminology and classification of quartzites. In detailed work, it is necessary to distinguish among the three quartz-rich sandstones (i.e., pressolved, orthoand metaquartzites) with numerous possible transitions in between by using the types and frequency of grain contacts. Grain-contact types range from the original depositional varieties to those formed by different stages of compaction and, then, to those that resulted from metamorphism. Com-
186
K.H. WOLF AND G.V. CHILINGARIAN
parative studies of the textures of sandstones changing with depth in sedimentary basins may constitute a promising approach. During the past few years, data on the precise temperature and pressure conditions that cause changes in clays and coaly matter have been collected (for details see pp. 395-422). Thus, in a basin with a variable lithology, the types of quartzites could be compared with the variation in the types and amounts of various clay minerals and, possibly, organic matter with increasing depth. In the Precambrian and Phanerozoic metamorphic belts, the degree and type of metamorphism of clayey sediments and/or volcanics could be compared with the quartzite varieties. Whisonant (1970) reported on slightly metamorphosed sandstones and described postdepositional deformation of quartz-bearing sandstones that can result in the following changes: (a) straining of grains; (b) suturing of grain contacts; (c) solution and redeposition of silica as overgrowths and cement; (d) corrosion and replacement of quartz by sericite; (e) creation of microfaults; (f) formation of fractures either within or transecting the grains; and (g) creation of bubble-trains in healed fractures. The Cambrian rocks studied by Whisonant have undergone early stages of metamorphism as evidenced by secondary sericite, chlorite matrix, straining, suturing, crushing of quartz grains, abundant micro-cataclastic features, microfaults, and fractures. Whisonant (p. 1023) made the following observations on the effects of deformation: (1)As matrix quantity increases, the amount of overgrowth and cementation (by quartz precipitation) decreases; (2) as matrix quantity increases, degree of suturing of grain boundaries decreases; (3) as matrix quantity increases, the amount of straining (as shown by the degree of undulose extinction), which depends on whether the grain is single, semi-composite, or composite, decreases; (4)within the same sample, as grain size decreases, the amount of straining decreases; (5) when matrix is composed predominantly of sericite, as matrix quantity increases, sericitization of detrital quartz and K-feldspar increases; (6) as matrix quantity increases, the number of bubble-trains (i.e., healing microfractures) decreases. In regard to item 4 above, it can be said (see also Conolly, 1965, p. 126) that straining of quartz is a function of grain size and depends on sorting. Where both large and small clasts occur, the finer particles tend to be less strained than the coarser ones, whereas if only fine-grained particles are present, they show no significant difference in the amount of straining. This is due t o larger grains acting as “propsyyin the sandstone during stress application that results in strain shadows in the coarser fragments.
DIAGENESIS OF SANDSTONES AND COMPACTION
187
Whisonant also discussed the derivation of grains, with differing degrees of straining from a source rock area, which were subsequently differentially abraded and dissolved during transportation from the source to the depositional environment. In such cases, the final deposit would be composed of strained grains that differ in both proportion and degree of straining from the source rock material, as a result of the differential processes. Studies like the one mentioned above, will eventually lead to a better understanding of the vague transitional boundaries between diagenesis and catagenesis as well as between catagenesis and metamorphism. Maybe it is possible to make a number of precise subdivisions of the very low-grade metamorphic zones that would permit recognition of the various transitional stages from diagenesis into metamorphism. For this reason, a separate section in this chapter has been devoted to burial metamorphism and related topics (see pp. 395-422). As to using experimental results for the interpretation of natural geologic phenomena, certain difficulties have been encountered already. Blatt (1966), for example, pointed out that altho.u,gh intergranular suturing due to pressure solution is very common in quartz-rich sandstones even of shallow burial, it has not been duplicated in the laboratory in the absence of recrystallization. It seems that extended geologic time and intergranular solutions are also required. Granulation of quartz grains, according to Blatt, was produced experimentally under uniaxial stress with piston pressures of about 400 bars. This is equivalent t o a 6,000-ft overburden, if one assumes a 1bar/l5 f t gradient. As Blatt pointed out, these results are invalid as guides t o granulation of quartz, because grains in natural deposits are not confined laterally as they are in the cylinders. Review of the literature indicates that granulation is a relatively rare compaction feature in sandstones. The information on textures in sandstones presented above demonstrates that many petrologists have concerned themselves with both the nomenclature of textures and the chemical and physical variables that control their origin. Observations on natural sand deposits, as well as on those prepared by the free-fall method under laboratory conditions, have provided data on textures and mass properties of sands that are in the initial, uncompacted state. Although the degree and style of packing and the porosity of freshly accumulated sand range widely, depending on a number of variables that must be further investigated in the future, sufficient data is available at present t o permit the sedimentologist and petrologist to define with a certain degree of reliability the characteristics of uncompacted sands of comparatively simple primary texture and composition. The data gathered from investigations on naturally and artificially compacted sands also has reached a stage where the information can be successfully used in unravelling the compaction history, preferentially as part of the whole diagenetic history.
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K.H. WOLF AND G.V. CHILINGARIAN
POROSITY AND PERMEABILITY
The investigation of compaction of sand and sandstone, as well as any other sediment and sedimentary rock, rests partly on the study of porosity and permeability changes during the various stages of diagenesis. In this section, therefore, some fundamental data on porosity and permeability of coarser sediments, with occasional reference to clay- and silt-sized material, are presented. Although not all of this information has been directly related to compaction, a consideration of at least some selectively and preferentially chosen material on porosity and permeability is paramount in understanding and discussing compaction, and in planning future investigationsof the various compaction mechanisms and their effects. As Peck (1967;see also Pandey et al., 1974)has pointed out, there are no substances totally “non-porous” and “impermeable”, because there is no definite boundary between “diffusion”, on one hand, and “free flow”, on the other. The properties of sediments affecting diffusion and free flow are very important in the study of numerous aspects of sedimentology and petrology (e.g., diagenesis). No details on diffusion and free flow are discussed in this chapter. It is important to mention, however, that Darcy’s Law fails in extreme fluid-solids interactions, e.g., in the case of aqueous solutions of certain electrolytes flowing through some clays (Lutz and Kemper, 1959),or for fluid-solids systems characterized by Reynolds numbers greater than about 1 (Scheidegger, 1957,p. 124). Very promising seems to be a future study of the relationship between composition, texture, degree of compaction, etc., of sediments, on one hand, and the movement of fluids and ions by diffusion and free flow, on the other. In particular, the influence of the degree of compaction on porosity and permeability would be of interest as it would throw light on a number of genetic problems in the geochemistry of secondary products in sediments, including ores in detrital and carbonate rocks. As to the problems of diffusion, Pandey et al. (1974)presented fundamental theoretical concepts on pressure solution and movements of fluids. They pointed out that the migration and accumulation of solutions responsible for diagenesis and ore mineralization, as well as of oil and gas, is the result of a combination of upward and/or downward percolation, lateral movement, and diffusional processes. The reader should note that although little reference is being made to diffusion per se in this chapter, its fundamental significance is, of course, recognized. The lack of published data on the application of diffusion concepts in sedimentary petrology, is partly a reflection of the information available. As Pandey et al. mentioned, diffusion of fluid constituents is related to the configuration (e.g., size and tortuosity) of the pores within rocks. Thus, in petrologic investigations of diffusion, all
DIAGENESIS OF SANDSTONES AND COMPACTION
189
the factors discussed in the section on textures must be given due consideration. All mass properties of sandstones, including porosity and permeability, are a function of numerous variables most of which are, in turn, controlled by compactional diagenesis. Most of these variables have been treated in the different sections of this chapter. Inasmuch as porosity and permeability are among the most fundamental mass properties of sandstones, being also of considerable practical importance in the exploration for hydrocarbons, and, more recently, have been given increasing attention in ore petrology, some separate fundamental information is supplied below. Figure 3-97 (Pettijohn et al., 1972) indicates a wide variation of permeability, some in the order of 100,000 times, depending on the petrographic properties of the rock. Small-scale variations in the specific permeability values can be large in specific sandstones, especially in those that have been cemented, and usually this variability exceeds that of porosity. Tables 3-XXIII and 3-XXIV summarize the rock properties that influence permeability and flow response, wheras Table 3-XXV presents the hierarchical sequence of primary controls on permeability that can be represented also in a more complex conceptualized manner, as done by Wolf (1973a, p. 173, fig. 9), for example. The reader is referred to Pettijohn et al. (1972, pp. 93-97 and 523-533) for details on porosity and permeability as well as for examples discussed. The data on porosity and permeability is voluminous and is based on: (1) theoretical considerations; (2) experimental laboratory investigations using idealized models and natural sediments or sedimentary rocks; and (3) field studies. Some of the results are contradictory. A summary of the results of attempts to demonstrate a relationship between porosity and various fundamental textural properties is presented in Table 3-XXVI. The contradictory results, according to Griffiths (1967b, p. 230) “arise from the interdependencies between the measured properties of sediments, which, in turn, reflect the interactions between aggregates and environmental conditions in both laboratory and field observations”.
105
Specific Permeability, K Darcysl 10 1 10-1 10-2 1w3 10-4 10-5 Very fine rands silts. Clcon rand,, mixturesat sand silt bnd Unweathwed hxturssdckan
i o ~ 103 102
clay; glacial lillj&tmthed pmvcls clays; etc.
Flow charmkrirticlr
Gmd aquifers
Poor aquifers
clays
Imp.rvwus
K.H. WOLF AND G.V. CHILINGARIAN
190
TABLE 3-XXIII Permeabilities of various rock types - average values for k and K (after Davis and De West, 1966,p. 164) Type of rocks
h(mil1idarcys)
K(cm/sec)
Gravel Clean sands (good aquifers) Clayey sands, fine sands (poor aquifers) Representative values fork and K Argillaceous limestone (2%porosity) Limestone (16% porosity) Silty sandstone (12% porosity) Coarse sandstone (12% porosity) Sandstone (29% porosity) Very fine, well-sorted sand Medium, very well-sorted sand Coarse, very well-sorted sand Very well-sorted gravel Montmorillonite clay Kaolinite clay
106 - 108 103 - 106 1 -103 Wmd)
1 - 102 10-3 - 1 10-6 - 1 0 - ~ K (meinzers) * 1.8 . 10-3
* 1 meinzer = 4.72 .
cm/sec x 5.5
1.103
1.4 * 10' 2.6 1.1. 103 2.4.103 9.9.103 2.6.105 3.1 . lo6 4.3 . 107 10-2 1
-
2.50
4.74.10-2 19.90 43.60 18.00 . 10 4.6 . 103 5.8 . 104 7.88 . 105 10-4 10-2
darcys for water at 60°F.
Studies by Rittenhouse and Fraser Rittenhouse (1971a), who dealt with some of the most fundamental theoretical aspects related t o compaction of sand, has considered the amount of pore-space reduction as a result of pressure solution of quartz grains and the additional reduction of porosity caused by the chemical precipitation of the dissolved material. On considering single grains and spheres, and other idealized geometric forms, stacked according t o various packing arrangements, he found that the relative amounts of porosity loss due to solution and to cementation vary greatly, being dependent on grain shape and angularity, packing direction from which pressure was applied, and the amount of solution that had occurred. Using uniformly-sized spheres, and assuming a cubic and orthorhombic packing (as shown in Figs. 3-98a+), Rittenhouse calculated loss of pore space by solution and cementation as a result of overburden pressure (Figs. 3-98e-h). For various percentages of the radii of spheres lost by solution at points of contact, he presented (a) original and remaining porosity, (b) loss of porosity due t o closer packing, and (c) loss of porosity caused by chemical precipitation of the dissolved material as cement (Figs. 3-99a-d). The amount dissolved caused by assumed pressure solution is expressed by the percentage of the radii of the spheres lost during solution at points of contacts. For example, in fig. 3-99a, considering a loss of 28% of the
DIAGENESIS O F SANDSTONES AND COMPACTION
191
radii due t o solution, the removal of solid matter leads to a reduction in porosity of 13.3%.Inasmuch as the original porosity for this system of packing is 47.6%, and the 28% along the x-axis is equivalent to 34.3% on the y-axis, then 47.6 minus 34.3 gives the value of 13.3%lost. If the dissolved matter had been removed from the system, the only reduction in porosity would have been TABLE 3-XXIV Rock properties and flow response (after Pettijohn et al., 1972, table 11-9, p. 525) Rock property Effects on permeability and porosity Texture Grain size Sorting Packing Fabric Cement
-
permeability decreases with grain size; porosity unchanged permeability and porosity decrease as sorting becomes poorer although little data is available, tighter packing favors both lesser permeability and porosity in the absence of lamination, controls anisotropy of permeability; permeability is maximum parallel to the mean shape fabric the more cement, the less permeability and porosity
Sedimentary structures Parting lineation maximum permeability most probably parallels fabric in plane of bedding Cross-bedding scant available data suggest that horizontal permeability parallels direction of inclination and that the steeper the dip of the foreset, the weaker the horizontal vector of permeability little data, but fine grain size and more laminations combine to Ripple mark cause low permeability and, hence, ripple zones are commonly barriers to flow Grooves and flutes as judged by fabric, permeability should parallel long dimension Slump structures no data, but probably always greatly reduce horizontal permeability Biogenic structures destroy depositional fabric and bedding and, thus, drastically reduce permeability and cause minimal, if any, horizontal anisotropy of permeability; effect on porosity is unknown, but may be negligible Lithology Sandstone
Shale
thicker beds tend to be coarser grained and thus more permeable, i f cement is not a factor; if weakly cemented or uncemented, ratio of maximum to minimum permeability is perhaps less than 5 to 1; if cement controlled, ratio may reach 100 to 1 or more the prime barrier to flow that outshadows all others by far; thus it is the arrangement of sand and shale much more than permeability variation within the sand that controls flow in most reservoirs
TABLE 3-XXV Hierarchical sequence of primary controls on permeability (after Pettijohn et al., 1972, table 11-10, p. 526) Control
Remarks
Texture and fabric Defined by grain size, sorting, packing and shape orientation of framework grains Scale: 1 to a few cm3
fundamental “building blocks” that define the primary pore system; depositional fabric may be completely destroyed by burrowing organisms
Sedimentary structures Cross-bedding, ripple mark, and parting linea- directional structures consist of anisotroption are most common and nearly always have ic fabrics so that individual structures should behave as “flow packets” preferred orientation and anisotropic fabrics Scale: 1-102 m3 Bedding facies Defined by bed thickness, types and abundances of sedimentary structures and frequency of shale beds. Scale: 102-105 m3
probably the most important primary control on permeability distribution in a sandstone body; shale beds act as impermeable barriers to flow and are one of the more continuous lithologies
Composite sand bodies Superposition of one “cycle” of sand upon another, cycles commonly separated by unconformities Scale: I O ~ - I O ’ ~ m3
characteristic of many alluvial and deltaic sands; multilateral as well as multistory bodies possible
TABLE 3-XXVI Effect of fundamental properties on variation in porosity (after Rosenfeld, 1950; in: Griffiths, 1967b, table 11.4, p. 229) Rock property
Source of information* ~
theoretical analysis Coarse grain size Good size sorting Slight skewness toward finer sizes High “sphericity” and “roundness” (confused) Intermediate “roundness” Open packing Low chemical cement Low clay content
artificial mixtures
natural sediments
O+-
O+-
++
O+-
+ +
+
+
? ? ?
+ + ? ?
+
+
* 0, +, - indicate relative change in porosity for each rock property assuming the others are constant.
193
DIAGENESIS OF SANDSTONES AND COMPACTION
0
---- c u m PACKING
#--tunic PACKING, SOLUTION FROY 2 IUPPER e LOWER) CONTACTS EACH SPHERE
6 CONTACTS /SPHERE
r --cueic
PACKING, ROTATEO 45. SOLUTION FROY 4 CONTACTS OT EACH SPHERE
PACKING, ROTATLO 45. 6 CONTACTS / SPHERE
b ---CUBIC
c ---WITHORHOYBIC
OF
q --ORTHORHOYBIC PACKING, SOLUTION FROY 4 CONTACTS OF EACH SPWERE
PICKING
8 CONTACTS I SPHERE
[@ L-
_ A _ _ -
d--0RTHORHOMBIC PACKING, ROTATED 8 CONTACTS / SPHERE
WT
h --WTHORHOYBIC PACKING, ROTATE0 3 0' SOLUTION FROY 6 CONTACTS OF EACH
SPHERE
Fig. 3-98. Spheres in different packing and orientation relative to application of vertical pressure before and after solution from points of grain contacts. (After Rittenhouse, 1971a, fig. 1, p. 81; courtesy Am. Assoc. Pet. Geologists.)
194
K.H. WOLF AND G.V. CHILINGARIAN
13.376, but if all the dissolved material was precipitated in the pores, an additional loss of porosity of 9.1% (34.3 minus 25.2) would occur as calculated by Rittenhouse. Figure 3-100, on the other hand, shows that all porosity would have been lost before 30% of the spheres' radii had been dissolved. In this case, about half of the loss would be the result of solution and
47.6 X
F
F z
2
w
W
P t
!i
t > t
>
I-
a
0
P
B
a
P 0
0
IW PERCENT OF SPHERE RADIUS
LOST
ev
0
I5
SOLUTION
( b ) - U I B l C PICKING,ROTATEO 48 ' SOUITION FROM 4 CONTACTS FROM LACH SPHLRL.
( 0 )--CUBIC,VERTICAL PICKINO, SOLUTION fRw 2 COWTACTS OF EACH SPHERE.
LEGEND
LOSS 61 SOLUTlON
LOSS BY CEMENT
n
r J
50
REMAlNlNO n n E SPACE
F z W
.... , ..
X
P
W
0,
>
I-
8 B
I
0
0 PERCENT OF SPHERE RADIUS
(C
1--
OIITHORI(OMBIC PICKING; SOUITION FROM 4 CONTACTS Of EACH SPHERE.
Losr
m y SOLUTION
( d ) - - ORTHORHOMBIC PACKING; W T I O N
FROM 6 CONTACTS Of EACH SPHERE
Fig. 3-99. Original porosity and loss of pore space by solution and cementation for sphere packing and orientation presented in Fig. 3-98. (After Rittenhouse, 1971a, fig. 3, p. 83; courtesy Am. Assoc. Pet. Geologists.)
DIAGENESIS OF SANDSTONES AND COMPACTION
195
ORTHORHOMBIC, ROTATE0 30’
CUBIC,ROTATEO 45‘
POIOSITV
LOSS
DUE
TO SOLUTION
IPERCENTI
22.
.I .
OBLAlL 0 P R O L I T 1 SPHERIODS-TETRAGONAL PACKING ISPUASIIED.VERllClL,CUBlC~
20. O . O B t A l E WERMD K)*1015 SEMI A X E S - 4 C O N l A C l S 0 W L U i SPUEROIO l O l l O l 7 SEMI LXLS-4CONTACTS
I
5 W
Y
-
J
10. ’
0. (I.
a.
0
2
4
6
LORO.1TI
8 LOSS
K) DUE
I2 TO
I4
IS
SOLUTION
I8
20
22
24
ILEIISENTI
(bl
Fig. 3-100. Relationship between the loss of porosity by solution and the porosity loss resulting from the precipitation of cement. Curves in top figure are for perfect spheres, i.e., for particles with no angularity. In bottom figure, the graph for cubic-vertical packing is for oblate spheroids with perfect rounding, whereas the remaining three graphs are for prolate spheroids with angularity, the surface irregularities being expressed as minute “wedges”, “cones”,’and “pyramids” (see Fig. 3-102). (After Rittenhouse, 1971a, fig. 2, p. 82; courtesy Am. Assoc. Pet. Geologists.)
196
K.H. WOLF AND G.V. CHILINGARIAN
another half would be due to cementation. Similar calculations were made by Rittenhouse in preparing Fig. 3-99 for spheres of various packing arrangements, corresponding to those shown in Fig. 3-98.From the graphs, the cement-to-solution ratios (the relative losses by precipitation and by solution) can be determined. With greater amount of cement precipitated, the greater is the degree of pressure solution that took place. As recognized by Van Hise (1904,pp. 865-868), even the most concentrated subsurface brines can precipitate only small amounts of cement. Consequently, continued precipitation could not have come from a closed system without replenishment of solution and a steady supply of chemical elements is required. Figure 3-101 (Pettijohn et al., 1972, pp. 397-398) illustrates that with continued cementation pore spaces become progressively smaller, the permeability is steadily reduced, the flow rates decrease, and the rate of precipitation is slowed down. Adams (1964,pp. 1575-1577) indicated that the grain-size distribution of a sediment controls the initial permeability and its rate of decrease as a result of cementation. Fine-grained sands, with a lower primary permeability, undergo cementation prior to coarsegrained sediments. As indicated in Fig. 3-101,whatever flow model is proposed for the solutions causing the infilling of pore spaces by cementation, all have in common an exponential decrease in the amount of cement per
Fig. 3-101. Exponential decrease in porosity (A) and amount of cement precipitated (B), as cementation of a sandstone proceeds. The precise shapes of such curves depend on the rate of flow of solutions through the sand, the initial porosity and permeability, concentration of solutions, and the rate’of precipitation. (After Pettijohn et al., 1972, fig. 10-4, p. 398; courtesy Springer, New York.)
DIAGENESIS OF SANDSTONES AND COMPACTION
197
unit time, i.e., a long time is required for complete cementation in contrast to the time needed for partial cementation. As mentioned by Pettijohn et al. (1972, p. 398), the absolute length of time for complete cementation cannot be determined by any presently-known technique, but it has been observed that few Recent or Tertiary sandstones are completely cemented. A long period, of the order of lo8 years, seems t o be a requisite. The importance of packing in determining the cement-to-solution ratios is also shown by Fig. 3-99,a. If all four packing systems have a remaining porosity of 25%, the following ratios are applicable: cement-to-solution ratio packing system cubic-vertical 9.2113.4 cubic-rotated 4.2118.4 orthorhombic 4.2110.3 orthorhombic-rotated 4.5110.0 The calculations are based on the fact that in the first two packing systems the original porosity was 47.6% and in the other two, 39.5% (Fig. 3-100). If in all four systems only 25% porosity was left after diagenesis, then 22.6% (47.6 minus 25) and 14.5% (39.5 minus 25), respectively, were the corresponding reductions in porosity due to cementation and solution. Using the graphs in Fig. 3-994, the sum of loss of porosity by solution and that by precipitation of cement represents the diagenetic loss in pore space (e.g., for cubic-vertical packing system, as shown in the tabulation above, the total loss in pore space is equal to 9.2 + 13.4 = 22.6% and for orthorhombic packing it is equal to 4.2 + 10.3 = 14.5%). It can be seen then, as Rittenhouse pointed out, that in three of the four packing models, less than 10% cement is the result of pressure solution in cases where the quartzite has retained 25% porosity. Allen (1969, 1970) has pointed out that in theoretical considerations of problems related to textures and fabrics of natural sediments, the analogies to spherical particles of artificial sediments will give wrong results. The closest approximation can be found when oblate spheroids (or prolate spheroids) are used (Fig. 3-102), which more closely represent the spheroidal grains of the natural sands having sphericities of 0.7-0.8. Calculating the amounts dissolved by approximate methods and by using oblate spheroids of varied flatness in two arrangements, presented in Fig. 3-99,b, and a prolate spheroid in vertical arrangement, the reduction in pore space due to precipitation and solution has been approximated by Rittenhouse (1971a, p. 82) (see Fig. 3-103). These spheroid packings have the same original porosity values as those for spheres in similar arrangements. In his initial calculations, Rittenhouse only treated perfectly rounded grains; however, natural sand grains, although often well rounded as in the case of
K.H. WOLF AND G.V. CHILINGARIAN
198
Fig. 3-102. Two arrangements of oblate spheroids for which cement and solution relations were approximated. (After Rittenhouse, 1971a, fig. 4, p. 84;courtesy Am. Assoc. Pet. Geologists.)
multiple-cycle quartz particles, commonly have angularity. The “corners” and “edges” on surfaces of the grains can be visualized or expressed as minute comers or pyramids and as three-sided prisms. Results of Rittenhouse’s calculations of the cement-to-solution ratios for these types of grains 22
-
20-
I
ARE
POROSITY
LOSS
DUE
TO SOLUTION
(PERCENT)
Fig. 3-103. Relationship between maxima cement content and porosity loss due to solution for various types of sahds. (After Rittenhouse, 1971a, fig. 6, p. 85; courtesy Am. Assoc. Pet. Geologists.)
DIAGENESIS OF SANDSTONES AND COMPACTION
199
are given in Fig. 3-99,b. He pointed out that particles with rounded corners would give rise to graphs intermediate between those for spheres or spheroids and those for grains having angular corners and/or edges. Angularity decreases the cement-to-solution ratios. The original porosity increases with increasing angularity, but angular sands lose their porosity more rapidly by pressure solution at points of contact than by cementation. Rittenhouse (1971a) pointed out that pressure solution may be differential because of varying solubility from grain to grain and within individual grains. As a result, there may be penetration of the less soluble grains into the more soluble grains. In the idealized models described above, the grain and pore geometry were uniform in each case of packing; this was followed by considerations of nonsymmetrical packing and grain-size sorting and variations in other parameters. Rittenhouse (1971a, p. 86) concluded, after discussing the details, that “no type of packing and no variations in grain sphericity and roundness, sorting, or composition yield cement-to-solution ratios that are larger than found for equal-sized spheres in orthorhombic packing rotated 30”” (Fig. 3-99,a). This is in agreement with the results of Allen (1969). Thus, as pointed out by Rittenhouse (p. 87), “it appears that the ratios for this packing can be used t o indicate the maximum amount of ‘cement’ that can be derived from solution at points of grain contacts for any given amount of pore-space reduction.” If the original porosity of a rock is reduced diagenetically from 40 to 25% (a 15%reduction in porosity), the maximum amount of cement formed as a result of pressure solution (Fig. 3-104) will be 4.7%. According to Rittenhouse, his above-reviewed relationships for ideal systems can be applied to natural sandstones if certain adjustments are made. In ideal cases where the sand bodies are homogeneous and where all the dissolved material was precipitated in adjacent pores, “the cement-to-solution ratios for rotated orthorhombic packing appear to apply best to wellrounded, well-sorted sands that had original porosities of about 38%. For more angular sands, for example, for sands having higher or lower porosities, or for those having better or poorer sorting, these cement-to-solution ratios appear to be high.” Rittenhouse estimated maximum cement-to-solution ratios, as given in Fig. 3-103, for the following: (1)very poorly-sorted sand (porosity about 28%); (2) extremely well-sorted sands (porosity about 42.5%); and (3) well-sorted, very angular sands (porosity about 43%). The maximum values of cement-to-solution ratios for sands with intermediate values in porosity, sorting, and angularity would fall between the two curves plotted in Fig. 3-103. This “best estimate’’ curve of Rittenhouse is the same graph as the one labelled “wedges” in Fig. 3-99,b. Taylor (1950) already had observed the behavior of volcanic and other ductile grains under pressure (e.g., shale, anhydrite, and salt) and that their
K.H. WOLF AND G.V. CHILINGARIAN
200
nI2
-
10-
I-
well sorted rond.
6-
Very poorly sorted (2) Ettremely well rorlcd,of
(1 4-
(3) Very onqulor 29
0
2
4
6
8
10
ORIGINAL
12
I4
MINUS
16
18
20
22
24
26
28
30
32
34
P R E S E N T POROSITY [ P E R C E N T )
Fig. 3-104. Relationship between maxima cement content and difference between original and present porosity for various sands. (After Rittenhouse, 1971a, fig. 7,p. 86; courtesy Am. Assoc. Pet. Geologists.)
deformation increased with depth of burial. According to Rittenhouse (1971b, p. 92), the available experimental data suggest that sandstones containing ductile grains may undergo more intense mechanical compaction than those devoid of ductile grains and rich in quartz and/or feldspar and might lose porosity and permeability more rapidly. An example of this type of sands would be graywackes (Folk’s phyllarenites, i.e., a rock rich in phyllite rock fragments) containing relatively “soft” sedimentary and/or metamorphic rock fragments. Sawabini et al. (1974) also showed that compressibility increases with increasing feldspar/quartz ratio. It has been suggested by numerous investigators that postdepositional deformation due to compaction of the ductile fragments may give rise to a “matrix” (see Dickinson, 1970, for example). Rittenhouse (1971b) has offered an ideal model of a sand composed of nonductile and ductile spherical grains. The results of compactional deformation of the “soft” grains leading to various degrees of pore space and thickness reductions are presented in Figs. 3-105 and 3-106. Fraser (1935) discussed packing and its relationship to porosity and permeability in great detail (see also Graton and Fraser, 1935).When a number of spheres of constant size are uniformly packed, the remaining inter-sphere openings can be occupied by spheres that are of a particular smaller size. This can be repeated until the unit voids are filled with successively smaller
201
DIAGENESIS OF SANDSTONES AND COMPACTION
kEDUCTlON I N PORE SPACE,
Fig. 3-105. Relation between content of ductile grains and reduction in pore space caused by compaction. (After Rittenhouse, 1971b, fig. 2, p. 92; courtesy Am. Assoc. Pet. Geologists.)
r 45 40
a-' VI
3
35
30
4
f
25
W
10
5 0
0
5
10
I5
20
25
30
35
40
REDUCTION I N THICKNESS{COMPACTlON), o/'
Fig. 3-106. Relation between content of ductile grains and reduction in thickness caused by compaction. (After Rittenhouse, 1971b, fig. 3, p. 95; courtesy Am. Assoc. Pet. Geologists.)
202
K.H. WOLF AND G.V. CHILINGARIAN
20
24
28
32
36
40
POROSITY OF BINARY MIXTURE, YO
Fig. 3-107. Relationship between the matrix grain size (expressed as a fraction of diameter of large spheres) and the porosity of binary mixture. (After Fraser, 1935, fig. 1; by permission of The Univ. of Chicago Press, copyright @ Univ. of Chicago.)
and smaller spheres. Fraser stated (p. 921) that if the voids in an assemblage of spheres are filled successively with smaller spheres, the resulting porosity values for the combination (mixture of spheres) do not fall on a smooth curve. Instead, the porosity decreases by sudden steps, as shown in Fig. 3-107. As to mixtures of spheres of two sizes, Fraser stated that as long as the proportion of the smaller spheres is sufficient to keep the larger spheres separated from each other, the smaller spheres will dominate the general fabric or structure of the assemblage. The small spheres together with their own interparticle voids represent the “matrix”, whereas the larger spheres can be considered to be “foreign” particles floating in the matrix and disturbing the assemblage. Fraser noticed a fairly uniform decrease in porosity with increasing proportion of the large spheres (for an explanation, see p. 918 in his publication). According to Fraser, when the proportion of large spheres is increased beyond a certain limit, two alternative situations occur, depending on the relative diameters of the assemblages of spheres: (1)When the proportion of large spheres is just sufficient for them to be in contact with each other and to be self-supporting without requiring support from the small spheres, and when the smaller spheres have a diameter less than that corresponding to the “critical ratio* of occupation”, then the large spheres control the fabric or structure of the assemblage. The latter also holds true with further increase in the percentage of large spheres. At the particular point just mentioned, when the control changes from the smaller
* Ratio of the diameter of a small sphere, which can just pass through the pore throat between larger spheres into the interstitial void, to the diameter of the larger sphere.
DIAGENESIS OF SANDSTONES AND COMPACTION
203
to the larger spheres, the former just fill the voids between the latter ones without causing any distortion of the packing of the assemblage of the large spheres. Beyond this turning-point, as the proportion of large spheres increases and the smaller ones no longer can fill all the openings between the large particles, additional voids remain open. Consequently, beyond this turning-point, an increase in number of large spheres results in an increase in porosity. (2) In distinction from case 1, when the proportion of the large spheres increases beyond the limit of domination by the small particles, and when the diameters of the small spheres exceed the “critical ratio of occupation”, then and thereafter the two sets of spheres mutually interfere with each other. Neither set controls the fabric or structure of the assemblage until the content of large spheres reaches a total of 100%. Mutual interference raises the porosity; but, on the other hand, the porosity still decreases as the proportion of the large spheres increases. Certain of these relationships are shown in Fig. 3-108: Curve 1 represents porosities when the diameter of the smaller of the two series, d , is equal to 0.433 D,where D is the diameter of the larger spheres, and curve 2 depicts porosities when d is equal to 0.158 D. Curve 1 (“total porosity”) demonstrates a slower decrease in porosity with an increase in the proportion of large spheres than does curve 2, because the difference in the diameter of the spheres is greater in the latter case. When the content ratio of large to small spheres is 3/1 (= 72-7576 large spheres to 25-28% small spheres by volume), the large particles commence to control the fabric or structure of the assemblage and the porosity consequently abruptly increases (dashed lines). This porosity increase is particularly marked when the spheres differ widely in diameter. The two lines converge at the same porosity (39%) when the assemblage is composed of large spheres only (Fig. 3-108). Figure 3-108 also indicates that an increase in the proportion of large spheres results in increasing disturbance and looseness of packing of the small spheres which constitute the “matrix”. The two upper curves (“porosity in matrix”) present the proportion of remaining openings between the large spheres after their interparticle spaces have been occupied by the smaller spheres. As the number of large spheres increases, the porosity of the matrix also increases, because of growing disturbance of the fabric or structure and the loosening of the packing among the “matrix” as a result of increasing number and closeness to each other of the large spheres. The porosity of the assemblage of spheres changes with change in the ratio of the two diameters (Fig. 3-log), if the proportions of the two populations of spheres are constant. In Fraser’s experiments, the volumes of spheres of two different sizes were always equal; the diameter ratio of large to small spheres ranged from 1 : 1 (i.e., only one size of spheres) to 19 : 1. According to the data obtained by Fraser, the decrease in porosity is most marked in
K.H. WOLF AND G.V. CHILINGARIAN
204 LARGE SPHERES, %
t
t 0
d
4
so SMALL SPHERES, %
100
Y
POROSITY. %
Fig. 3-108.Relationship between the porosity and the relative proportions of large and small spheres. 1 and 3 = diameter ratio of large to small spheres is equal to 1 : 0.433;2 and 4 = diameter ratio is equal to 1 : 0.158. Curves 1 and 2 indicate porosity of matrix, whereas curves 3 and 4 show total porosity. (After Fraser, 1935, fig. 2;courtesy J. Geol.) Fig. 3-109.Relationship between porosity and the ratio of large to small sphere diameters for mixtures containing 50% of each size. (After Fraser, 1935,fig. 7; by permission of The Univ. of Chicago Press, copyright @ 1935 University of Chicago.)
the range of size ratios from approximately 2 : 1 to 6 : 1. Beyond a ratio of about 8 : 1, the growing difference in size of spheres of the two populations reduces the porosity only slowly. The effect of angularity on porosity was also considered by Fraser. He used a series of carefully sized materials, ranging in shape from spheres to flat plates, and maintained other parameters as constant as reasonably possible. The materials tested included: (1) the lead and sulphur shot (perfectly spherical sand-sized grains); (2) the marine and beach sand (partly rounded grains); (3) the crushed material (angular); and (4) the mica (flat, plate-like fragments) (Table 3-XXVII). The porosity was measured before and after the material was compacted by jarring, first when the constituents were dry and then when they were saturated with water. As shown m the Table, the angularity markedly affects the packing with a consequent increase in porosity. Assuming that the influence of grain shape on packing results in the variation of porosity, Fraser’s data showed a range of 8.73%(43.51-34.78) units (excluding mica) in porosity of compacted dry materials, i.e., from 34.78 to 43.51%. Of this range, 6.25 units are due to increase in porosity, because of irregular angularity, whereas 2.48 units are owing to porosity increase as a result of the ptesence of flattened particles (Table 3-XXVII). Moderately well-rounded sands exhibited only minor variations in porosity
DIAGENESIS OF SANDSTONES AND COMPACTION
20 5
TABLE 3-XXVII Effects of grain shape on porosity (after Fraser, 1935, p. 936) Material
Specific gravity
Porosity (%) type of packing dry loose*
~
Lead shot Sulphur shot Standard sand (marine) Beach sand Dune sand Crushed calcite Crushed quartz Crushed halite Crushed mica
11.21 2.024 2.681 2.658 2.681 2.665 2.650 2.180 2.837
wet compacted**
~~
40.06 43.38 38.52 41.17 41.17 50.50 48.13 52.05 93.53
loose
compacted
42.40 44.14 42.96 46.55 44.93 54.50 53.88
38.89 38.24 35.04 38.46 39.34 42.74 43.96
~~
37.18 37.35 34.78 36.55 37.60 40.76 41.20 43.51 86.62
~
~~
-
-
92.38
87.28
* and ** = before and after compaction by jarring. as a consequence of limited variation in the degree of rounding. The angularity led t o “bridging” and loose original packing of the grains. Independent of the grain shape, the wet constituents packed more loosely than the dry materials (Table 3-XXVII). This has also been observed in natural sand accumulations. The effect of angularity is more pronounced in the case of wet packs, as demonstrated by the differences in porosity of dry and wet packs for rounded versus angular materials (both loose and compacted). The accumulations of flat and needle-like components exhibit the greatest porosity (e.g., mica has over 90% porosity), which is also supported by observations on natural sediments. These high porosities cannot be reduced below 80%, even by prolonged shaking, as done by Fraser. An increase in pressure, causing mechanical compaction, was required to diminish the porosity of the wet mica accumulation to 67.4%. It should be pointed out, however, that the compressibility of sands increases with increasing mica and clay content (Sawabini et al., 1974). Fraser concluded that angularity usually increases porosity. He observed a decrease in porosity caused by “angularity” only in cases where the grains were mildly and uniformly disk-shaped. The control on permeability by the unformity of grain size, i.e., sorting, is very great. Just ‘as porosity is less for a population of mixed sizes, the permeability is also reduced within certain limits when particles of different
K.H. WOLF AND G.V. CHILINGARIAN
206
sizes are added. As to Fraser’s (1935) two-component populations, he remarked that a layer of openings, considered larger than those existing within the population of small spheres, exists around each large sphere. The small spheres adjacent to large ones can touch them only at one point, and the distance between these contact points is controlled by the size of the small spheres of the “matrix”. The relatively large voids around the large spheres connect freely in all directions, giving rise to a very permeable area or zone. As a result of disturbance of the packing of the “matrix” by the presence of a large “foreign” sphere, an area of looser packing extends for a distance equal to several large-sphere diameters away from these large particles. This also increases the porosity and permeability. Both of the influences mentioned above oppose the decrease in permeability caused by the presence of impermeable large spheres. The relative values of these opposing controls are given in Fig. 3-110 for a number of cases investigated by Fraser. The upper curve was obtained by adding larger spheres to a population of smaller ones with a diameter of 2.3 times smaller than that of the large spheres. With increasing proportion of the large spheres, initially the permeability increases slowly and then more rapidly. The reason for this behavior lies in the fact that the larger spheres are more effective in increasing the channelways by disturbing the packing of smaller spheres than in decreasing the permeability by blocking previously existing channels. It is rather interesting to observe that whereas the total pore space that remains after each addition of large spheres is reduced, the remaining pores form larger and more effective chanLARGE SPHERES, Yo
0;
‘
’
’
’
’
50
’
‘
SMALL SPHERES, %
’
‘
’
100
Fig. 3-11 0 . Relationship between the coefficient of permeability and the relative proportion of large and small spheres. 1 = ratio of the diameter of large sphere to that of the small one is equal to 3.61; 2 = large/small sphere diameter ratio is 6.28; 3 = large/small sphere diameter ratio is 2.30.. (After Fraser, 1935, fig. 11; by permission of The Univ. of Chicago Press, copyright 01935 Univ. of Chicago.)
DIAGENESIS OF SANDSTONES AND COMPACTION
20 7
nels. This creates the anomaly of increasing permeability accompanied by a decreasing porosity . As the lower curve in Fig. 3-110 indicates, the adding of large spheres with a relative diameter of 6.23 (i.e., diameter of large spheres is 6.23 times larger than that of small spheres) t o a population of smaller spheres initially causes a lowering of the permeability, which continues until the ratio of large to small spheres is about 50 : 50 by volume percent. Thereafter, the permeability increases until the content of the smaller spheres reaches about 31.5%by volume, when the permeability of the mixture is equal to that of the small spheres alone. From this point on, the permeability is greater. On the other hand, the porosity of the two-component system decreases until the ratio of the small to large spheres is 25 : 75. After that, the large spheres are in contact with each other and the spaces between them are completely filled with the small spheres. It is obvious then, that in mixtures containing from 25 to 50% of small spheres, the total pore space decreases whereas the permeability increases with increasing content of larger spheres. When the content of small spheres falls below 25%, the value of the porosity loses significance, because the small spheres are present in insignificant numbers to fill all the interstices between the large spheres. The central curve in Fig. 3-110 was obtained for a two-component system containing spheres, diameters of which were in the ratio of 3.6 : 1. Fraser (1935, pp. 941-946) discussed compaction of sands to some extent. He mentioned that it has been stated that sands show no shrinkage on drying and that this appears to be incorrect. As mentioned earlier, wet sand packs less tightly than dry sand by 1%or more percentage units of the total porosity (Table 3-XXVII). Experiments showed that wet sand had a porosity of 37.92%, whereas the same sand when dried had a porosity of 36.95%,and these results were reproducible. The difference may be due to the removal of a film of water that is present around each sand grain. Fraser discussed the degree of compaction of beach sands and stated that any accumulation of detrital components has a fairly definite range of porosity reflecting various degrees in the perfection of packing of the grains. He found that a sample of a moderately well-sized and well-rounded beach sand in the dry state exhibits a range in porosity from 37.8% to 46.576, depending on whether the fabric is that of a loose or tight packing. When wet, approximately the same range of porosity is present. The above change in porosity represents a change of 14%in the total volume of the sand. More research data is needed for better understanding of the influence of the numerous environmental parameters on the original porosity and the subsequent diagenetic compactional history as a result of increasing overburden. In one of his experiments, Fraser used actud beach sand that had an average natural porosity of 40.56%prior to sampling. Mere settling of the sand in the laboratory into a
K.H. WOLF AND G.V. CHILINGARIAN
208
container produced a porosity of 46.0% when wet and 46.3% when dry. Tapping produced settling or “compaction” until after 6 minutes the constant values of 38.26%and 37.82%in the wet and dry states, respectively, were obtained. These observations indicate that the naturally-accumulated sand on the beach with an average porosity of 40.56%had already undergone sufficient rearrangement to have lost most of the bridging effect. It has assumed such a fabric that further compaction while wet could take place only slowly, although considerable additional consolidation and loss of volume on drying still existed. Fraser also found that dry sand may be compacted much more quickly, especially during the early stages, than the wet sand (see table 6 in Fraser, 1935). Fraser further observed that when bridging of the sand grains is not present, no significant compaction can take place until the pressure applied to the system exceeds the crushing strength of the minerals. In the coarse sediments, the grain-size distribution and the matrix content are very important in controlling compaction. The degree of “compaction” mentioned above is, of course, related to the experimental conditions used by Fraser in attempting to cause the rearrangement of the fabric by tapping; therefore, the results cannot be easily extrapolated to natural or laboratory conditions where the sand is exposed to higher pressures and temperatures and to intrastratal fluids of varying composition. Experiments such as those done by Fraser should be performed at high temperatures and confining pressures. In addition, the influence of sorting, orientation, mineral composition, and grain shapes on compressibility, porosity, and permeability should be investigated in greater detail. The permeabilities of gravel packs having a porosity of about 35% are presented in Table 3-XXVIII.According to Hill (1941, p. 138), flow turbulence in part determines the permeability of gravels. Porosity also plays an important role, e.g., an increase in porosity from 35 to 40%results in an increase in permeability by as much as 60%. TABLE 3-XXVIII Permeability ranges of gravel packs (Q,% 35%) for gravels of different sizes (after Hill, 1941,table 3, p. 138) Gravel size (mesh)
Range in permeability (darcys)
Average permeability (darcys)
3- 4 4- 6 6-8 a-1 0 10-14
72 00-9000 3400-4000 1700-21 00 1000-1’300 700- 900
8100 3700 1900 1150 800
DIAGENESIS OF SANDSTONES AND COMPACTION
209
TABLE 3-XXIX Permeabilities o f various gravel-sand mixtures (after Hill, 1941, table 4, p. 129) ~
~~~
~~
Gravel size (mesh)
Gravel/sand size ratio
Average gravel permeability (darcys)
(Gravel + sand) permeability (darcys)
3- 4 4- 6 6- 8 8-1 0 10-14
15.0 10.6 7.5 5.3 3.7
8100 3700 1900 1150 800
40- 80 180-220 230-300 250-300 200-250
(avg. 60) (avg. 200) (avg. 265) (avg. 275) (avg. 225)
Hill (1941, p. 139) determined the permeability of these gravels after bridging them with a sand having the following properties: (1) logarithmic mean = 1.863 mm; (2) geometric mean = 0.275 mm; (3) standard deviation = 0.674; (4) skewness = 0.268; (5) kurtosis = 0.100; (6) ten percentile = 0.445 mm; (7) permeability = 6.6 darcys; (8) porosity = 35%.Sand was introduced into the gravel in a vertical-flow tube by flowing sand-water mixture. The results are presented in Table 3-XXIX. The permeability of the gravel-sand mixture was highest (average k = 275 darcys) in the case of gravel/sand size ratio of 5.3, whereas the lowest permeability (average h = 60 darcys) was present when the gravel/sand size ratio was 15. In the latter case there is an intensive migration of sand into the gravel caused by poor bridging action. Landreth (1969, p. 4) also presented permeabilities of packed gravels of certain size ranges (Table 3-XXX). Gaither (1953) considered the effects of sorting on porosity which can be quite complicated. Commonly, the porosity diminishes as the grains deviate from uniform size distribution, because: (a) the finer grains fill the voids TABLE 3-XXX Relationship between the particle size and permeability of gravels artificially packed in oil wells in order t o preclude sand production (after Landreth, 1969, p. 4) Gravel size range (inches)
U.S. Series Number (mesh)
Permeability (darcys)
0.023-0.032 0.032-0.046 0.046-0.065 0.065-0.093 0.093-0.1 31 0.131-0.18 5 0.185-0.2 63
20-30 16-20 12-16 8-1 2 6- 8 4- 6 3- 4
800 1100 1500 2100 2700 4000 6500
K.H. WOLF AND G.V. CHILINGARIAN
210
between the larger grains to form what is generally known as “matrixyy,and (b) the coarsest grains reduce the porosity by occupying a volume that would otherwise be occupied by the finer, porous material. These findings are in agreement with the observations made by Fraser (1935),as discussed above. Using King’s (1898)experimental data on porosities of sands, Gaither presented porosity curves for well-sorted sand and mixtures of sands (Fig. 3-111).Using a mathematical approach, Von Engelhardt (1960)presented an idealized (theoretically determined) curve. Examination of Fig. 3-111shows that (a) well-rounded, well-sorted sands have porosities ranging from about 34% for coarse sands to 38% for fine sands (curve 1); (b) decreasing the sorting by mixing different proportions of two different-sized sands results in a porosity decrease (curves 2 and 3 which are below curve 1); and (c) the lowest porosity of around 25% corresponds to a mixture of 60--70% sand finer than 0.096 mm in diameter and 30-40% of a well-sorted sand with an average diameter of 0.483 mm (curve 3). The value of 25% is approximately 12% below the average for a well-sorted, well-rounded, medium-sized sand. Morrow et al. (1969)presented information on the prediction of porosity and permeability from grain-size data. They pointed out that on the basis of a systematic study of unconsolidated material, Fumas (1929)showed that: (1) the porosity of a two-component sediment depends on both aggregate
0%
I
COARSE
100%
Fig. 3-111. Variations in the porosity with varying proportions of fine-grained and coarse-grained sands. 1 = change of porosity with average grain size for well-sorted sands (“simple” sands); 2 = change of porosity in mixtures of sands with average diameters of 0.611 mm and 0.152 mm; 3 = change of porosity in sand mixtures in which average diameter of coarse sands is 0.483 mm and maximum diameter of fine sand is 0.096 mm. (After King, 1898; in Gaither, 1953, fig. 2, p. 185; courtesy J. Sed. Petrol.)
DIAGENESIS OF SANDSTONES AND COMPACTION
211
40
8
*t
!
30
20
PERCENTAGE O F COARSE-GRAINED FRACTION
Fig. 3-112.Relationship between the porosity and the relative proportions of fine and coarse fractions in two-component sand aggregates for various size ratios. Percentages are by voIume. (After Furnas, 1929;in Morrow et al., 1969, fig. 1,p. 312;courtesy J. Sed. Petrol. )
I
0
0 c
10
-
*L
(L
I
.lo
x)
40
60
80
1 1
100
PERCENTAGE O F COARSE- GRAINED FRACTION
Fig. 3-113.Relationship between the ratio of (permeability of the mixture)/(permeability of the aggregate of small components) and the relative proportions of the fine and coarse fractions for two-component sand aggregates. Radius of coarse grains = 0.1 cm; porosity ofeither fine fraction or coarse fraction = 40%. Numbers on the curves designate the ratio of (permeability of the coarse fraction)/(permeability of the fine fraction). (After Furnas, 1929;in Morrow et al., 1969,fig. 2,p. 313;courtesy J. Sed. Petrol. )
212
K.H. WOLF AND G.V. CHILINGARIAN
composition and the size ratio of the particles (which has also been mentioned by Gaither, 1953, as discussed above), but is independent of particle size (Fig. 3-112); (2) the permeability of a two-component aggregate depends on relative proportions of fine-grained and coarse-grained fractions and their respective permeabilities (Fig. 3-113). In Figs. 3-112 and 3-113, the percentage of the coarse grains in an aggregate is plotted on the abscissa and ranges from 0 to 100%. At both extremes, of course, the system becomes a one-component sediment, composed either of fine or coarse grains, so that the lines converge to meet. The smalr t o large size ratio in Fig. 3-112 ranges from 0.5 to 0.05. The ratio of 1.0 would indicate no difference in size so that the curve would be a straight line parallel to the abscissa. With an increase in difference in grain diameter, greater volume of the intergranular space between the large grains can be filled by the smaller particles, giving rise t o a distinct decrease in porosity. The maximum decrease takes place when the two-component aggregate is composed of approximately 70% coarse material and 30% fine-grained components, for a size ratio of 0.2, 0.1 and 0.05. Morrow et al. (1969) determined the interrelationship among the grain-size distribution of unconsolidated sediments, porosity, and permeability. The mathematical formula that would describe the grain-size distribution best, according to them, is the Rosin-Rammler equation:
(3-3) where Y = cummulative weight percent of under-sized material, z = particle size, N = a measure of the narrowness of the particle size distribution, and D = a measure of the mean particle size. They also found that two other measurements, in addition to N and D , were useful in their correlations, i.e., fines = actual weight percent of material that passed a 325-mesh (44-micron) screen, and RR fines = weight percent of material less than 44 microns in size given by the fit to the Rosin-Rammler equation (eq. 3-3). These two measures enabled determination of the fine-end skewness. The findings of Morrow et al. can be summarized as follows: (1)The best correlation between the grain-size distribution and porosity was given by a plot of porosity as shown in Fig. 3-114. (2) Porosity versus log[lOON(RR fine~)l/~/fines], tends to increase with increasing closeness of the particle size distribution (i.e., N). (3) Porosity is independent of absolute particle size: there was a complete scatter on a plot of porosity versus D . (4) There is a reduction in porosity with increasing fractional weight of fines, which fill the interparticle spaces of a packing matrix formed by larger particles. ( 5 ) Porosity depends on both grain-size distribution and packing geometry. (6) The good correlation between the grain-size distribution and porosity suggests that the mode of packing is reasonably consistent for a given size distribution. (7) The
DIAGENESIS O F SANDSTONES AND COMPACTION
31
213
I
t
( 100xNx(RRFINES)”3 FINES
)
Fig. 3-114. Relationship between porosity and Rosin-Rammler size-distribution parameters. Root mean-square deviation = 1.65%; overlapping points = 7. Fines = actual weight percent of material that passes a 325-mesh (44 p ) screen; R R Fines = weight percent of material less than 44 p in size given by the fit to the Rosin-Rammler equation. (After Morrow et al., 1969, fig. 6, p. 317; courtesy J. Sed. Petrol.)
0
E
h
I
f
t
c
i I
W h
Y
4 LOG
(NrD),
microns
Fig. 3-115. Relationship between permeability and Rosin-Rammler size-distribution parameters. Root mean-square deviation = 0.32; overlapping points = 12. (After Morrow et al., 1969, fig. 7, p. 318; courtesy J. Sed. Petrol.)
K.H. WOLF AND G.V. CHILINGARIAN
214
LOG [ N X D X ( R R F I N E S ) ” ~ ] , ~ ~ C ~ O ~ ~
Fig, 3-116.Relationship between the permeability and the Rosin-Rammler size-distribution parameters. Root mean-square deviation = 0.345;overlapping points = 18. (After Morrow et al., 1969,fig. 8,p. 319;courtesy J. Sed. Petrol.)
MEDIAN DIAMETER,
/4
MEDIAN DIAMETER. mrn
Fig. 3-117A.Median diameters and porosity of recent North Sea sediments of Wilhelmshaven. (After Fuchtbauer and Reineck, in: Von Engelhardt, 1960, fig. 7; courtesy Springer, Berlin.) Fig. 3-117B.Median diameter and porosity of recent shelf sediments of the Californian coast (San Diego County). (After Hamilton and Menard, in: Von Engelhardt, 1960,fig. 8; courtesy Springer, Berlin.)
DIAGENESIS OF SANDSTONES AND COMPACTION
21 5
permeability of a sand increases with increasing mean particle diameter and with increasing porosity. Relationships between the permeability and the Rosin-Rammler size-distribution parameters are given in Figs. 3-115 and 3-116. ( 8 ) The porosity varies directly with the closeness of the size distribution. (9) The relative amount of fine particles affects the permeability through interstitial blockage (see also Fig. 3-113 of Furnas, 1929). Von Engelhardt (1960) observed that the primary porosity of freshlydeposited sands is dependent on the grain size down to the range of fine sand, as shown in Figs. 3-117 and 3-118. In one case, the fractional porosity between 0.40 and 0.44 is independent of grain size in the range of 120 to 240 p, whereas in the second example the fractional porosity between 0.38 and 0.45 is independent of grain size between 200 and 700 p. This indicates that the results of experiments with spheres apply only to coarser sands and are less applicable as the grain size decreases; this cannot be explained on the basis of purely geometric considerations. The examples of high-porosity sandstones obtained from the subsurface are presented in Tables 3-XXXI and 3-XXXII. The data indicates that inspite of overburden pressures, sandstones can retain a considerable amount of pore space, unless the pores have been filled as a result of chemical precipitation. Uncemented, very friable sediments have been found even at relatively great depth (approximately 4000 m), and it is in these cases where one can demonstrate that mechanical compression alone can reduce the pore space. Table 3-XXXII and Figs. 3-118 and 3-119 show examples of sands and sandstones from oil fields which have undergone mainly the mechanical
h OD6 @X12 .I3 .B .x)40 .M) I.Omm
Fig. 3-118. Grain-size distribution, porosity ($) and permeability (k in darcys) parallel to the bedding of unconsolidated to slightly consolidated sandstones of Valendis (Bentheimer Sandstone), Germany. From drillholes of the Ruhlermoor Oil Field near Meppen/Ems, Germany. (Laboratory of the Gewerkschaft Elwerath.) A = from depth of 782-788 m, 4 = 0.296, and k = 11.0 d; B = from depth of 782-788 m, $ = 0.297, and k = 8.70 d; C = from depth of 759’800 m, 4 = 0.325, and k = 2.10 d. (After Von Engelhardt, 1960, fig. 9; courtesy Springer, Berlin.)
K.H. WOLF AND G.V. CHILINGARIAN
216
TABLE 3-XXXI Examples of deeply-buried sandstones with relatively high porosities (after Von Engelhardt, 1960, table 3) Locality
Formation
Depth (m)
Porosity fraction
Weber, Pennsylvanian Cockfield, Eocene Miocene Tennsleep, Pennsylvanian Eocene Eocene Frio, Oligocene Oil Creek, Ordovician Pliocene Bromide, Ordovician
1860 2100 2160 2280 2320 2340 2740 3260 4300 4600
0.176 0.298 0.280 0.195 0.270 0.315 0.282 0.067 0.200 0.050
Lower Pliocene Lower Pliocene Lower Pliocene Lower Pliocene Lower Pliocene Upper Miocene
1555 1575 1695 1930 1960 2530
0.360 0.232 0.317 0.280 0.273 0.300
U.S.A. (1 ) Rangely, Colorado (2) Katy, Texas (3) University Field, Louisiana (4) Big Medicine Bow, Wyoming (51 Davis Lens, Texas (6) Liberty Co., Texas (7) Fishers Reef, Texas (8) Lindsay, Oklahoma (9) Fillmore, California (10) Carter Knox Field, Oklahoma
Italy (11) Cortemaggiore near Piacenza (12) Cortemaggiore near Piacenza (13) Cortemaggiore near Piacenza (14)Cortemaggiore near Piacenza (15) Cortemaggiore near Piacenza (16) Budrio East, near Bologna
NO. 1 , 2 , 6, 7, 8 after Winsauer, Shearin, Masson, Williams (1952); No. 3, 5 after Levorsen (1956);No. 4 after Waldschmidt (1941);No. 9 after Henriksen (1958); No. 10 after Stearns (1957); No. 11-16 after AGIP (1959). TABLE 3-XXXII Unconsolidated to slighlty consolidated sandstones from German oil fields (after Von Engelhardt, 1960, table 4) Locality Eldingen near Celle Eldingen near Celle Scheerhorn near Nordhorn Scheerhorn near Nordhorn Riihlermoor near Meppen Riihlermoor near Meppen
~
Formation
Depth (m)
Porosity
Lias (Y Lias (Y Valendis Valendis Valendis Valendis
1483 1463 1104 1120 842 853
28 29 23 27 30 33
* median diameter; ** sorting coefficient.
(%I
d50*
d75/d25
so**
1.60 1.20 1.49 1.60 1.67 1.59
1.12 1.11 1.04 1.23 1.01 1.10
(mm) 0.133 0.105 0.340 0.113 0.270 0.138
DIAGENESIS OF SANDSTONES AND COMPACTION
217
Fig. 3-119. Grain-size distribution, porosity (@) and permeability (k in darcys) parallel to bedding of slightly consolidated sandstones of the Lias (11. From drillholes in the Eldingen Oil Field near Celle, Germany. (Laboratory of the Gewerkschaft Elwerath.) A = from depth of 1490 m, # = 0.31, and k = 0.950 d; B = from depth of 1483 m, q5 = 0.28, and k = 0.420 d. (After Von Engelhardt, 1960, fig. 10; courtesy Springer, Berlin.)
compaction and only little chemical cementation. The sands from Ruhlermoor (Germany) are very loose, those from Scheerhorn (Germany) only weakly cemented, and those from the Eldingen oil field (Germany) are slightly more consolidated. These sands show good sorting and are of marine origin. The median diameter lies between 0.1 and 0.34 mm and the distribution curve is symmetrical. According to practical and theoretical considerations, the true primary porosity is about 40%. Thus, without chemical infilling of pores, the porosity was reduced from 40% to 23-33%, i.e., reduction of about 20-40%. Inasmuch as the plastic deformation and breakage of quartz is not present, the reduction in pore space is due to the rearrangement of the grains, but the movement was only minor. Primary textures are either modified or destroyed. In sands in which the primary orientation of grains was still observable after pore-space reduction, rotation along the long axes must have constituted the movements during compaction. According to Von Engelhardt (1960),if sand deposits of different sizes and shapes are undergoing compaction under the same overburden pressure, the mechanical reduction of the primary pore spaces will be different: the fine sands will experience less reduction in porosity than the coarser-grained ones, as the number of contacts per unit volume in fine-grained sands is larger and the resistance to compression is greater. The shape of the grains also plays a role in controlling degree of compaction. The sands composed of grains having irregularly-shaped surfaces maintain a higher porosity in contrast to sands having smooth, spherical grains, as the former have greater
K.H. WOLF AND G.V. CHILINGARIAN
218 30
I
I
I
I
I
I
I
I
-
28
2 24v)
2
22 20
0 00
-
100
180
280
I
#
340
1
1
420
500
MEDIAN DIAMETER,/
Fig. 3-120. Porosity of different beds of the Bentheimer Sandstone (Valendis) from a drillhole of the Scheerhorn Oil Field (Lingen, Germany). The porosity is controlled by the median diameter as shown by 190 measurements. (After Von Engelhardt, 1960, fig. 13; courtesy Springer, Berlin.)
numbers of long and concavo-convex contacts. An example illustrating the influence of grain size in a case where sands of different grain sizes are present is given in Fig. 3-120. A 25 m thick sandstone unit is coarser at the top and becomes finer towards the bottom. The distribution curve, e.g., sorting, is the same, so that only grain size varies. Figure 3-120 shows that the porosity of the fine sands is on the average greater than that of the coarse sands. Pettijohn et al. (1972, pp. 93-97) mentioned that the permeability of sandstones, especially from the same lithologic bed, often show a lognormal statistical distribution, whereas porosity is characteristically normally distributed. Variations in permeability are much greater than that of porosity as shown by Fiichtbauer (1967a) in his log permeability-versus-porosity plots. The same plots reveal an obvious correlation between effective porosity and permeability. The Kozeny-Carman equation shows that permeability is inversely proportional to the second power of specific surface (see Langnes et al., 1972). As demonstrated in Fig. 3-121, as grain size decreases, specific surface of the sediment increases and the resistance to flow increases. The permeability of fine-grained sand (or silt) is smaller than that of a coarsegrained one, with the effective porosity of both sands being identical. Von Engelhardt (1960) pointed out that the nature of sediments is such that permeability is never isotropic. The analyses of fluid movements in the pore spaces must take into consideration the fact that permeability changes with changes in the direction of fluid movement. Most important are the differences between the permeability parallel and normal to the bedding. In all examined cases, the permeability of sandstones normal to the bedding is less, as shown in Fig. 3-1122. Two important questions arise here: (1)What are the changes of permeability parallel and across the bedding during me-
DIAGENESIS OF SANDSTONES AND COMPACTION
219
I'
0
.o I
LOG GRAIN SlZE,h
..a"
' I
' 0
I
100
10
I"
I
Kx)
loo0
_
I IQoOord
kll
Fig. 3-121.Relationship between the permeability ( k ) and grain size ( d ) for the Bentheimer Sandstone of the Scheerhorn Oil Field. The regression equation is log k = -2.1007 + 2.221 loglod, where k is permeability in millidarcys and d is grain size in millimeters. Scatter diagram is based on random selection of data from fig. 49 of Von Engelhardt (1960).(After Pettijohn et al., 1971,fig. 3-14,p. 97;coueesy Springer, New York.) Fig. 3-122.Dependency of the ratio of the permeabilities (vertical (kl)/parallel (kII) to the bedding on kll as illustrated by measurements of the Lias, Dogger, and Valendis sandstones of the Gilhorn Basin in northwest Germany. Near the center, k l approaches the value of 1, while kll increases. (After Riihl and Schmid, in: Von Engelhardt, 1960,fig. 48;courtesy Springer, Berlin.)
chanical and chemical compaction? (2) What are these changes relative to each other? It has been shown by Jobin (1962) that the magnitude of permeability of sandstones controls the localization of uranium precipitation, so that answers to these and similar questions may be of assistance in the investigations of secondary changes of sediments and in exploration for ores. The control of various parameters on porosity and permeability were discussed in numerous other publications, of which a few are considered below. Rogers and Head (1961) offered a composite plot of groups of samples with four different median diameters. Figure 3-123 shows a general decrease in porosity with an increase in sorting coefficient (= an increase in spread of grain sizes). For groups of sands with the same median diameter, the relationship between porosity and sorting coefficient is approximately linear except for very well-sorted sands. The highest porosities in Fig. 3-123
K.H. WOLF AND G.V. CHILINGARIAN
220
approach the theoretical maximum of 46.7% for uniformly sized spheres with cubic packing. Where sorting is very good (i.e., where the coefficient is very low), the curves of different median diameters may coincide. When the sorting becomes poorer, the curves diverge from one another. Also, as the median size decreases, the porosity increases for the same sorting coefficient. In Fig. 3-124,these relationships are more readily shown when the median diameter is plotted against porosity for various sorting coefficients. When the sediment is well sorted, the porosity is independent of the actual grain size and is only a function of sorting coefficient. In cases where the sorting is poorer, the porosity is increasingly dependent on median diameter as well as on sorting. Rogers and Head (1961)also noted that sphericity decreases slightly with decreasing grain size and concluded that the increase in porosity with decreasing grain size is due to a decrease in sphericity and consequent poorer packing. Where all sand grains are perfect spheres, porosity is independent of absolute grain size. Beard and Weyl (1973)investigated the relationship between porosity, permeability, and texture of artificially mixed, dry- and wet-packed sands using eight grain-size subclasses and six sorting groups. Their conclusions can be summarized as follows: (1)In unconsolidated, artificially packed natural sands, the porosity is independent of grain size for sands of the same sorting, but porosity varies with sorting. (2) Wet-packed porosity probably represents the minimum porosity of an unconsolidated, clay-free sand following mechanical rearrangement of the particles before burial. (3)In compaction studies, one should always compare sandstones of the same sorting range, especially when compaction gradients in vertical and horizontal stratigraphic sequences are to be established, because of the variation of porosity 2.2,
1
I
I
I
I
,
1
1
, I
POROSITY, oh
Fig. 3-123. Relationship between porosity and sorting coefficient of sands with various median sizes. A = median diameter = 0.106 mm;B = median diameter = 0.151 mm; C = median diameter = 0.213 mm;fl = median diameter = 0.335 mm. (After Rogers and Head, 1961, fig. 1, p. 469; courtesy J. Sed. Petrol.)
DIAGENESIS OF SANDSTONES AND COMPACTION
221
POROSITY, %
Fig. 3-124. Relationship between porosity and median size of sands with various sorting coefficients. A = S o = -2.086;B=So = - 1 . 6 2 6 ; C = S o = ~ 1 . 2 7 9 ; D = S 0 = - 1 . 1 2 8 ; E = So = "1.061. (After Rogers and Head, 1961, fig. 2, p. 470; courtesy J. Sed. Petrol.)
with sorting. The same applies to many other studies of mechanical and chemical diagenesis. (4) Permeability decreases as grain size decreases and as sorting becomes poorer. (5) Low sphericity and high angularity probably
.. .
I
I
20
I
I
,
30
MEAN GRAIN SIZE, Bunitr
Fig. 3-125. Relationship between the porosity and mean grain size for all sandstone samples of Paluxy Formation, North Texas. (After Dodge et al., 1971, fig. 8 , p. 1826; courtesy Am. Assoc. Pet. Geologists.)
222
K.H. WOLF AND G.V. CHILINGARIAN
increase porosity and permeability. As to a review of the absolute porosity and permeability values in sandstones, as well as for newly obtained values by Beard and Weyl (1973),the reader is referred to their original publication. Dodge et al. (1971)also investigated the various petrographic and directional parameters of Cretaceous, shallow-marine, blanket sandstones that control both porosity and permeability. Mean grain size is plotted versus porosity in Fig. 3-125.These sediments have a high range of both porosity and grain size and lack a cement. There is an obvious absence of any correlation, probably as a result of an insufficient grain-size range in this particular study. Usually, porosity increases with decreasing mean grain size. On the other hand, Fig. 3-126 presents relationship between the permeability and mean grain size of the same sediment and shows a definite correlation: the permeability decreases with a decrease in the mean grain size. In other cases, where cement is present, the data shows a wider spread. The relationship between porosity and percentage of cement is shown in Fig. 3-127.The spread in the area of 0--5.0% cement and 25.0-40.0% porosity may be the result of one or more factors, e.g., variation in the grain size, sorting, and packing. It seems that up to 5% cement may be present before cementation starts to affect the porosity. The cement content is plotted versus the permeability in Fig. 3-128and the spread may be explained by factors mentioned above for Fig. 3-127,e.g., grain size. As shown in both figures, the permeability and porosity decrease with increasing cement content, as would be expected. But again, as mentioned earlier, in detailed investigations of such a
1.0
2.0
3.0
4.0
MEAN GRAIN SIZE, Bunits
Fig. 3-126. Relationship between the permeability and mean grain size for all Paluxy sandstone samples, North Texas. (After Dodge et al., 1971, fig. 9, p. 1826; courtesy Am. Assoc. Pet. Geologists.)
DIAGENESIS OF SANDSTONES AND COMPACTION
01
I
I
I
K)
I
20
223
I
CEMENT CONTENT, %
Fig. 3-127. Relationship between the porosity and cement content for D-1 outcrop, Dexter Sandstone member of the Woodbine Formation of the Gulf Series (Upper Cretaceous), Denton County, North Texas. (After Dodge et al., 1971, fig. 10, p. 1826; courtesy Am. Assoc. Pet. Geologists.)
a m
e
.a 0
2-
t
4 m 4 W
I
a W
a
0
3
10
15
20
25
CEMENT CONTENT, %
Fig. 3-128. Relationship between permeability and cement content for G-6 outcrop, Grayson County, Dexter Sandstone member of the Woodbine Formation of the Gulf Series (Upper Cretateous), North Texas. (After Dodge et al., 1971, fig. 11, p. 1826; courtesy Am. Assoc. Pet. Geologists.)
224
K.H. WOLF AND G.V. CHILINGARIAN
Fig. 3-129. Isometric diagrams showing the variation in cross-bedding dip direction and dip angle and their theoretical influence on permeability. A = low-angle cross-bedding. B = high-angle cross-bedding. Solid arrows indicate maximum directional permeability, whereas dashed arrows indicate minimum directional permeability within a given crossbed set. (After Dodge et al., 1971, fig. 12, p. 1827; courtesy Am. Assoc. Pet. Geologists.)
nature, directional variations in the rock’s mass properties must be considered. As shown in Fig. 3-129,the cross-bedding dip angles in a sandstone may affect fluid flow. If the dip angle is low, the flow will tend to occur along the laminae, whereas if the angle is high, various laminae will have to be crossed by the fluid and the flow will be impeded, depending on the various characteristics of the laminae, e.g., grain size. It should be noted that differential packing, i.e., varying degrees of packing from layer to layer, was not the result of burial in these particular sands, but was the result of the changing transportational and depositional hydrodynamic conditions. Only occasional reference is made in publications to the relationship between porosity and permeability, even though there is a large amount of data available on both porosity and permeability. It seems that quantitative relationships between the two properties should have been given more attention. In the example presented by Dupuy et al. (1963),there is a straight-line relationship when the permeability is plotted on the logarithmic and the porosity on the arithmetic scales, respectively (Fig. 3-130).(See also the data by Shenhav, 1971,given below.) In environmental reconstructions and stratigraphic correlations, it is particularly important to study the grain size, because it reflects the energy conditions. Also, mean grain size may be related to the mineral composition of the sands. Berg and Davies (1968),for example, showed that the quartz content increases with increasing mean grain size of quartz (Fig. 3-131).Also of interest is the relationship between the depositional environment and size and quantity of quartz (Fig. 3-132).In their example, Berg and Davies were able to establish four distinct environmental groups, although for clear field discrimination the knowledge of sedimentary structures and stratigraphic relationships is often required. Inasmuch as both porosity and permeability are a function of the quartz grain size and quartz content, these two parameters should also be related to the original sedimentary environment. This was
DIAGENESIS OF SANDSTONES AND COMPACTION
225
POROSITY, %
Fig. 3-130. Relationship between porosity and permeability for sandstones from Champ de Cazaux (Albian). (After Dupuy et al., 1963, fig. 17; courtesy 6th World Pet. Congr., Frankfurt am Main.)
indeed found to be the case by Berg and Davies (Fig. 3-132).The finest sediments have the lowest permeability, whereas the coarser sandstones have higher permeabilities. Similar results were obtained for porosity. Plots of permeability and porosity versus depth may also reveal vertical changes (through geologic time) in textures, which may occur in sediments of the same type of environment. Inasmuch as the primary, depositional environments control the textures, fabrics, structures, composition, and mass properties of sandstones, which, in turn, may influence subsequent diagenesis, including compaction, then they should be taken into consideration especially in regional and vertical stratigraphic compactional investigations. An extention of the above study can be found in the work of Shenhav (1971).He investigated the petrography, porosity and permeability of Cretaceous sandstone oil reservoirs, and found that the reservoir characteristics are determined by depositional environment and post-depositional changes. Shenhav recognized three main sandstone types each having its own distinct mineralogical, textural and stratigraphic features. Within these sandstones, four main varieties of porosities were defined : (1)Intergranular porosity, with
K.H. WOLF AND G.V. CHILINGARIAN
226
-90-
I-LAGOON
80-
\
X - BEACH OR UPPER SHOREFACE
a- MIDDLE
SHOREFACE I i N c L WASHOVER 5 s )
0- LOWER SHOREFACE
*-
\
I
L
BEACH AND UPPER SHOREFACE
\
(WASHOVER)
70-
\\ *
\
\
A\.
\
\
\
\
\
\
ri090
\
rXy2.096
\
\ I
M
0.05 SILT
0.10 VERY FINE SAND
0.1s
0 20
I
5
FINE SAND
MEAN GRAIN SIZE OF QUARTZ ( i t , r n r n
Fig. 3-131. Relationship between quartz mean grain size and quartz content in Muddy Sandstone, Bell Creek Field, Powder River County, Montana. (After Berg and Davies, 1968, fig 6, p. 1895; courtesy Am. Assoc. Pet. Geologists.)
crystalline cement at aminimum, depends mainly on the packing, sorting and orientation of the grains. There is an absence of pressure solution and the porosity ranges up to 32%. (2) Intercrystalline porosity (i.e., porosity between the crystals of the cement), with relatively high crystalline cement content, reaches a value of up to 8%. Porosity is related to mineralogy of the cement, i.e., the dense, xenotopic calcite cementation gives rise to low values, whereas the hypidiotopic" or idiotopic** dolomite cementation results in larger values of intercrystalline porosity. (3) Intermediate porosity between 1 and 2 ranges from 8 to 20%. The pores are bounded by both crystalline cement and by grains. (4)Intracrystalline porosity results from leaching of the centers in sanidine crystals. Only a small percentage (less than l%), of sandstones have a porosity of this type, which is ineffective in nature. Shenhav, as others before him, found a direct relationship between permeability and porosity; however, in a plot of porosity (arithmetic scale) versus permeability (logarithmic scale), the points have a tendency to spread out over a large area because of variation in the size and shape of grains (e.g.,
* Majority of the constituent crjrstals are subhedral (Friedman, 1965, p. 648).
** Majority of the constituent crystals are euhedral (Friedman, 1965, p. 648).
DIAGENESIS OF SANDSTONES AND COMPACTION
227
50,001
I BEACH AND UPPER SHOREFACE
ro.ooc spoc
500
X
MIDDLE SHOREFACE
u) IOOC
w
-
X
. \
+\
2
2 I
>
t
i
IOC
z
50
4 W
A
LAGOONAL
...
A
10
5
\
AND LAGOO~AL O
I I
O
T
.
SILT
0
5
\
'.
0
1
0.10 VERY FINE
1
' *
0.15
I
0.20
0.25
FINE SeiND
Fig. 3-133). Both porosity and permeability depend on the average grain size, but the effect of grain size on the porosity decreases with decreasing porosity. This is the result of an increase in the cement content. Figure 3-133 presents the general trends (mainly above 8%porosity) of the porosity-versus-permeability curves for various grain sizes. In the case of sandstones with intergranular porosity, the coarser varieties are more permeable than the finer ones of equal porosity. In the case of very fine sandstones with a high clay matrix content, there is no distinct relation between porosity and permeability (Fig. 3-135).Shenhav plotted porosity versus permeability for each one of the sandstone types, as exemplified by Fig. 3-134for type I. He
K.H. WOLF AND G.V. CHILINGARIAN
228
HORIZONTAL PERMEABlLiTY, M D fLOG
SCALE)
Fig. 3-133.Relationship between porosity and horizontal permeability for the various sandstone types of Helez Formation, Israel. (After Shenhav, 1971, fig. 25, p. 2222; courtesy Am. Assoc. Pet. Geologists.)
001
01
OD
I
0
10
w
mo
100
HORIZONTAL PERMEABILITY, M D (LOG
SCPLE)
Fig. 3-134.Relationship between porosity and horizontal permeability of Type-I sandstone samples of Helez Formation, Israel. (After Shenhav, 1971, fig. 22, p. 2219;courtesy Am. Assoc. Pet. Geologists.)
229
DIAGENESIS OF SANDSTONES AND COMPACTION
HORIZONTAL PERMEABILITY. M 0
(LOG
SCALE1
Fig. 3-135.Plot of porosity versus horizontal permeability trends of Helez Formation sandstones having various grain sizes. (After Shenhav, 1971, fig. 21, p. 2219;courtesy Am. Assoc. Pet. Geologists.)
also presented a composite curve showing the fields of data distribution for various sandstone types (Fig. 3-133). According to Shenhav, petrographic considerations allowed the recognition of three different depositional environments, giving rise to offshore marine, tidal channel or lagoonal, and dune sandstones. These sandstones had the characteristic mass properties shown in Table 3-XXXIII. The amount TABLE 3-XXXIII Mass properties of various types of sandstones (after Shenhav, 1971, figs. 22-25, pp. 22192222) Type of sandstone
Dune (eolian coastal)
Mass properties porosity (%)
permeability (md)
24 (avg.) up to 32
200 (avg.), max. over 2000 50 (avg.) up to 2000 less than 30
Tidal channel and/or lagoonal
16 (avg.) up to 30
Offshore marine
very low up to 16
Reservoir characteristics good intermediate poor
230
K.H. WOLF AND G.V. CHILINGARIAN
and mineralogy of cement, which are determined by the sedimentary and diagenetic milieu, are the main factors controlling the porosity, whereas permeability is related only to the amount of cement. In the case of intergranular porosity, permeability depends on the average grain size. One of the well-known difficulties confronting sedimentologists and petrologists is the large amount of time involved in the determination of precise data of the various petrographic variables by employing instrumental techniques. In many instances, despite a reduction in precision and reproducibility, as well as possible variations between researchers (or “operators”), useful methods have been proposed to speed up the procedure for petrographic work in both field and laboratory investigations. The charts for quick determinations of roundness and sphericity of sand grains and pebbles have been available for some time, in addition t o the charts for visual percentage estimations. Beard and Weyl (1973)have presented a new set of visual textural comparators, i.e., eight sets of photographs of sands ranging from very fine to very coarse, each composed of six photographs depicting extremely good to very poor sorting range. In combination with the other charts mentioned above, they permit rapid estimation of a total of four textural variables. Beard and Weyl suggested that these charts are particularly useful in describing compacted and silica-cemented sandstones. Upon considering some examples of investigations on the porosity and permeability variations that appear to have been controlled by primary depositional factors, one should examine the relationship between the depth of burial and mass properties of sedimentary rocks. Maxwell (1964)has stated that it is a valid generalization that porosity varies inversely with depth of burial and, less certainly, with geologic age. These conclusions are well supported for shales. Although the absolute values differ, porosity of shales decreases rapidly in the first few thousand feet of burial and much more slowly thereafter (see Rieke and Chilingarian, 1974). Less information is available on the variation of sandstone porosity with depth, because of the extreme variability in various parameters. Even less information is available on the comparative relations between porosity and depth of burial for interbedded sandstones and shales of the same basin or stratigraphic section. Figure 3-136 shows such a comparative approach by F’roshlyakov (1960). The porosityaepth relationship for shale is better defined than that for the sandstone. Maxwell stated that neither the shape nor the trend of the sandstone curve could be anticipated from the published information. He collected data on the porosities of natural quartz sandstones and compared them with published data. Inasmuch as temperature is an important variable affecting the compaction of quartz sands, he also collected temperature data and compared it with porosities in natural sandstones. Although there are certain difficulties in obtaining accurate temperature data, and in most wells
DIAGENESIS OF SANDSTONES AND COMPACTION
2000
231
c
* = SAND 0
= SHALE
POROSITY '10
Fig. 3-136. Relationship between porosity and depth o f burial for Jurassic-Lower Cretaceous elastic sediments (sands and shales), Cis-Caucasus, U.S.S.R. (After Proshlyakov, 1960; in: Maxwell, 1964, fig. 1, p. 698; courtesy Am. Assoc. Pet. Geologists.)
the temperature gradient has not been measured, enough information is available, however, to show some significant relationships among porosity, depth, and temperature. Maxwell discussed eight examples, of which one is reproduced here (Fig. 3-137). The porosity data has been plotted as bar graphs, each bar summarizing the measurements of one individual sandstone section several tens of feet thick. The small solid rectangles indicate the average porosities. Each figure summarizes the data for sandstones of the same geologic province and approximately the same age t o enable comparison between the depth of burial and temperature variations. Figure 3-138 is an example where the decrease in both maximum and average porosities with depth is striking. Also, the change of maximum porosity is nearly linear with depth, which is distinctive for these particular sandstones. It should be pointed out, however, that Maxwell discussed some exceptions to that rule. Maxwell (1964) also reported on the results of experiments on the compaction and cementation of quartz sands under conditions simulating deep burial. Results of the short-time compression tests at room temperature suggest that high porosity can persist to great depths. At pressures simulating a depth of 40,000 ft, a 25% porosity may remain, and if the normal hydrostatic pressure is present, the porosity may be as high as 30%. As Maxwell pointed out, experimental conditions do not necessarily resemble natural
232 I5
5 1
I
I
I
25 I
I
i1 3
5
K.H. WOLF AND G.V. CHILINGARIAN ~
.
~
5.000
1
10.000
20.000
/
-
15.000
/
/
/
/
ym, /
/
l5.000
A
20.000
POROSITY,%
Fig. 3-137.Relationship between the porosity: and depth of burial for Pennsylvanian and for two shallow Permian sandstones (at about 3000 and 5000 ft), North Texas and Oklahoma. Porosities from depth of below 20,000 ft calculated from sonic and neutron logs. Data from Shell Oil Company, Socony-Mobil, and from f i l l and Taliaferro (1949). Small rectangles give average pay porosity, whereas figures in parentheses represent number of porosity determinations. Temperature gradient in "C/lOOOf t is noted along dashed line limiting maximum porosity. (After Maxwell, 1964, fig. 5, p. 701; courtesy Am. Assoc. Pet. Geologists.)
conditions of compaction, because many variables are not considered in the laboratory experiments. In his tests, Maxwell considered composition of the particles and pore fluids, temperature, overburden pressure, pore-fluid pressure, and time; however, the effects of grain size and sorting, for example, were not evaluated. Maxwell presented a summary of porosity-versus-depth data for the relatively pure, well-sorted quartz sandstones as shown in Fig. 3-138,which is based on figs. 4, 5, 7, 9 and 10 of his 1964 publication. Two experimental curves based on the experimental data published by him in 1960 are superimposed in this figure. Noteworthy is the variability of the curves for the Oligocene sandstones as the thermal gradient changes from 7"--1O0C/10O0 ft. The following observations can be made on examining Fig. 3-138:(a) the Oligocene sandstones have larger maximum porosities at
DIAGENESIS OF SANDSTONES AND COMPACTION DEPTH
oo oO
I
MAXIMUM POROSITY X 5
10 10
I5
20
25
30
35
40
45
1
5-,000
5.000
10.0001 10.000
15.000
- 15.000
20.000
- 20.000
25.000
26.500'
- 3o.OOo
30.000
35.0%
233
5
20
25
30 ..
. .
4.0 -
a5 ._
35'W
Fig. 3-138. Relationship between the maximum porosity and depth of burial for various sandstones. (After Maxwell, 1964, fig. 12, p. 706, based on figs. 4, 5, 7, 9, and 10 of Maxwell, 1964, and experimental data of Maxwell, 1960, figs. 6 and 11; courtesy Am. Assoc. Pet. Geologists.)
shallow depth than the Miocene ones, but the two depth-versus-porosity curves cross at about 13,000 ft, i.e., the porosity of Oligocene sandstones decreases more rapidly with depth than that of Miocene sandstones; (b) the Ordovician sandstones have larger maximum porosities than the much younger Pennsylvanian sandstones. This behavior can be predicted by using the equation developed by Maxwell (1964, p. 702) from the experimental data for quartz sand (see also description of Fig. 3-139). In Fig. 3-139, two sets of curves were plotted, i.e., curves of the observed data taken from Fig. 3-138 and curves based on the calculated data. Although the two sets of curves agree only in a general way, the calculated data enable prediction of the actual field observations to a certain degree. The calculated depth at which the Oligocene and Miocene sandstones are expected to have identical porosity values depends on the values of the assumed original depositional porosity for each sandstone and, therefore, on the slopes of the two curves. The porosity values in Fig. 3-139a are anomalous, because extrapolation of the curves to zero depth gives improbably large initial maximum porosity values of over 50%for both sandstones. Maxwell (1964) offered two possible explanations: (a) measurements are too high as the sediments were poorly consolidated and the samples were disturbed or (b) one could assume that initial porosities of over 40% were maintained to depths of 7,000-8,000 f t because of lack of compaction. In view of the straight-line character of the
K.H. WOLF AND G.V. CHILINGARIAN
234
5,000
-I
t
0
10
20
30 POROSITY,%
10
20 _.
5,000
30 .. J C M. ‘63
Fig. 3-139.Relationship between the maximum porosity and depth of burial for various sandstones. a. Observed curves of maximum porosity obtained from Fig. 3-138;b. porosi~-~ where Gi = ties calculated by means of the equation ( @ i / q ! ~ f )=~ t[e(T-540)/(0-0186T] porosity at time of deposition, @f = final porosity, t = time in days, T = temperature, degrees Kelvin. This equation is based on experimental data and predicts porosity of sandstones for simulated “depth” of 26,500 f t (Maxwell and Verrall, 1954).Slope determined by extending line from this point toward assumed value of initial porosity, @i. (After Maxwell, 1964,fig. 13,p. 707; courtesy Am. Assoc. Pet. Geologists.)
maximum porosity boundary curves, the latter explanation seems unlikely. As Maxwell pointed out, however, the available data are inconclusive. In the experiments reported by Maxwell (1960, 1964), compaction of the pure quartz sand occurred mainly by fracturing of grains, accompanied by rotation of grains and fragments. Silica cementation was noticeable, but its amount was very small. Both physical yielding and chemical precipitation of cement are involved in the compaction and induration of natural quartz sandstones. In cases of the presence of largest porosity, it can be assumed that the sandstone unit has undergone least amount of secondary quartz cementation. Under these conditions, compaction may occur by physical failure of grains, similar to that observed in the experiments, in addition to the occurrence of pressure solution. One should note here, however, that many researchers have suggested that such fracturing occurs only in the shorttime laboratory experiments in which the long geologic time, available during the natural compaction in most sedimentary basins, is not considered. In the presence of abnormally high pore pressures, the graiwto-grain pressure may be reduced sufficiently to diminish compaction so that abnormally large
DIAGENESIS OF SANDSTONES AND COMPACTION
235
porosities could persist to great depths. Although the available data is inadequate for reaching any definite conclusions, Maxwell suggested that abnormally high pore pressures do not guarantee persistence of unusually large porosities a t depth, and stated that this problem offers a fertile area for more intense research. Maxwell (1964)discussed the cause of linearity of his porosity-versusdepth curves. This linearity was quite unexpected, as theoretically one would expect a change in rate of reduction of porosity with depth. Maxwell, therefore, offered an explanation for the linearity on the basis of experimental results obtained by various investigators: (1)At room temperature and a gage pressure of 1atm, the short-time compaction experiment showed porosity reduction per unit of load which was largest at the small initial loads and diminished rapidly with progressive reduction in porosity under increasing pressure. The plotted results gave a porosity-“depth” curve with a shape similar to that of the shale curve presented in Fig. 3-136. (2) When both the temperature and “depth” (= pressure) were held constant, porosity first diminished rapidly during the first few days and then decreased at a decreasing rate throughout the experiments which lasted up to 100 days. (3)When dry quartz sand was tested at a pressure of 1atm, the grains became progressively and distinctly weaker as the temperature was raised. The physical failure, therefore, was shown t o be strongly dependent on temperature. (4)When sea water and oil-field brines were surrounding the grains, failure of the grains was again strongly temperature dependent; however, porosities were lower than in the experiments performed under conditions described under (3).The solutions seemingly caused an additional weakening of the quartz grains. Pressure solution was negligible. (5)The solubility of quartz in distilled water increases with increasing temperature, especially at higher pressures (Kennedy, 1950).On the other hand, the dissolved chemical components in sea water and brines have little effect on the solubility of silica (Krauskopf, 1950;Siever, 1962). The solubility of quartz in natural waters, therefore, should increase as temperature increases. From the above-presented information, it appears that: (a) compaction and consequent porosity decrease in quartz sands should proceed at a decreasing rate with burial if the temperature is presumed to be constant; (b) on the other hand, the increase in temperature will cause a loss of porosity with burial at an increased rate, due t o both decreasing physical strength of the grains and as A result of increasing solubility of the quartz grains. Maxwell then suggested that these two trends tend to balance one another, so
K.H. WOLF AND G.V. CHILINGARIAN
236
that the curves of maximum and average porosity may approximate straight lines. The tests performed by Maxwell indicated that compaction continues as long as porosity is present; there is no indication of an equilibrium porosity for a particular depth which would persist throughout geologic time. Teodorovich and Chernov (1968) developed a relationship between the depth of burial, D, in m and permeability, h, in md for sandstones at a depth range of 500-4500 m: log k = 2.803 X e4*000074Dor (approx.) log k = 2.88 - 0.0002D
(3-4)
For siltstones at the same depth range the formulae are: log h = 2.961 X e4.000147 or log k = 2.87 - 0.003D.
(3-5)
Variation in the permeability of sandstones is much greater than that for siltstones at the same depth (see Table 3-XXXIV), because of the higher heterogeneity of the former, e.g., at a depth of 3950-4450 m, permeability varies from 30 to 350 md and, sometimes, up to 800 md (average k = 114 md). When the cement content is less than 7%, porosity at a depth of 4250 m is higher than 20-25% and permeability is higher than 150-200 md. With increasing depth of burial from 4250 to 6000m, permeability decreases 2.2-2.4 times for siltstones and 1.5-1.8 times in the case of sandstones. Meade (1966) observed that variations in the porosity of sands and in their water content are not predictably related to depth and an increase in the overburden load. Figure 3-140 shows this for a selected group of sediments that were exposed to their approximate maximum overburden loads at the time of sampling. All these sediments are Cenozoic in age and most are inorganic and terrigenous clastics, except for sample I11 containing about 10% siliceous skeletal material, sample I1 which is clayey, and sample IV with 10-20% CaC03. The curves are drawn in such a way that the area TABLE 3-XXXIV Variation in porosity and permeability with depth for sandstones and siltstones of A p sheron oil- and gas-bearing province in U.S.S.R. (after Teodorovich and Chernov, 1968, P. 88) ~~
Depth (m)
650 4250 5000 6000
Sandstones
Siltstones
porosity (%)
permeability (md)
porosity (X)
permeability (md)
26 f 4 17 i 4 15.5f 4 14 f 4
460 f 110 i 87 i 63i
28.5 * 16 i 14.7 * 13 i
5 0 0 i 75 4 0 i 10 26+ 8 17 f 5
80 25 20 20
4 4 4 4
DIAGENESIS O F SANDSTONES AND COMPACTION
237
DEPTH OF BURfAL IMI
Fig. 3-140.Relation between porosity and depth of burial in meters in selected clays and claystones (A) and selected sands and sandstones (B). I = Recent, Lake Mead on Colorado River (Gould, 1960, p. 176, lower 3 graphs); II = Recent, Santa Barbara Basin off southern California (Emery and Rittenberg, 1952, p. 755.);III = Recent and older (?), western Bering Sea (Lisitsyn, 1956, 1959,fig. 16);IV = Recent and older (?), eastern Black Sea (data from Ostroumov and Volkov, 1964,pp. 94-95); V = Recent and older (?), continental slope off Nova Scotia (Richards and Keller, 1962); VZ = Recent, Orinoco River delta (Kidwell and Hunt, 1958,p. 808); VII = Pliocene and Pleistocene, central California (Meade, 1963a;curves adjusted for artesian pressure); VIZI = Pliocene to Recent, Baku Archipelago (Koperina and Dvoretskaya, 1965, fig. 1; plus data from Korobanova, Kovaleva, Kopylova, and Safokhina, 1965,pp. 128-130); I X = Tertiary, Venezuela (Hedberg, 1936,p. 256);X = Miocene and Pliocene, Po Valley (Storer, 1959, p. 523); XI = Miocene, southern Louisiana and southeast Texas (Maxwell, 1964, p. 704).Where bulk density or water content (by weight) were reported, porosity was computed assuming a particle density of 2.60g per cm3. (After Meade, 1966,fig. 1,p. 1086;courtesy J. Sed. Petrol.)
above the lines gives the volume of solid particles, whereas the area below represents the interstitial volume. As to the influence of particle size on water content and porosity (the two are numerically equal because both are expressed as percentage of bulk volume), Meade stated that particle size may be as important a control as overburden load on porosity, especially during the early stages of compaction. A comparison of Figs. 3-140Aand 3-141A shows that an order-of-magnitude difference in diameter between 0.001 and 1 mm may cause about the same amount of porosity variation as an order-of-magnitude difference in depth of burial. With an increase in pressure, the finer sediments are compacted more rapidly than the coarser ones, and the relationship between grain size and porosity becomes less pronounced (Fig. 3-141B). A significant influence of grain size on porosity is still apparent at overburden loads approaching 100 kg/cm2. For the sediments presented in Fig. 3-141B, the magnitude of change in porosity with changing grain size is about equal to that related to depth of burial (cf. the curves in Fig. 3-141B with curve VII in Figs. 3-140A
K.H. WOLF AND G.V. CHILINGARIAN
238
50-
5
1
P MEDIAN DIAMETER fMMJ
Fig. 3-141.Relations between porosity and median particle diameter in selected Recent surface or near-surface sediments under overburden loads less than 1 kg/cm2 (A), and in sediments under overburden loads between 7 and 70 kg/cm2 (B), I = Lake Mead on Colorado River (Sherman, 1953, p. 399);ZZ = Lake Maracaibo, Venezuela (Sarmiento and Kirby, 1962, p. 719); IZZ = reservoirs in western United States (Hembree, Colby, Swenson, and Davis, 1952,p. 39);ZV = Gulf of Paria (van Andel and Postma, 1954, p. 108); V = North Sea (Fuchtbauer and Reineck, cited by Von Engelhardt, 1960,p. 15); VZ = continental shelf off southern California (Hamilton and Menard, 1956,p. 757); VZZ = San Diego Bay and adjacent continental shelf (data from Shumway, 1960,pp. 454-457); VZZZ = Rivers in Japan (Komura, 1963,p. 266);ZX = Pliocene and Pleistocene alluvium, central California (Meade). (After Meade, 1966, fig. 2, p. 1087; courtesy Am. Assoc. Pet. Geologists.)
and 3-140B). Another group of sediments in California exposed to overburden pressure of 5-60 kg/cm2 (Meade, 1963b), demonstrated that their downward decrease in particle size (together with changes in other factors) reversed the expected effect of increasing load and caused a systematic downward increase in porosity. With a further increase in pressure as a result of overburden, an inverse relation between the grain size and porosity changed to a more direct correlation. According to Meade (p. 1087), “the more rapid decrease in the water content of clays eventually overtakes the slower compaction of sands, perhaps at some depth of burial near 1km”. This is supported by the observations that deeply buried coarser sediments and their equivalent rocks are commonly more porous than adjoining claystones and shales. In his section on compaction of sands, Meade (p. 1096) stated that most sands are only slowly compacted during the early stages of compaction; however, relatively few unequivocal data is available to demonstrate this, because of the difficulty in collecting unconsolidated sands without disturbing them. Curve VII in Fig. 3-140B represents silty sands, and it seems that the silt gave the sand enough cohesion to withstand coring and sampling without too much distortion. Curve X I in Fig. 3-140B represents quartzose sandstones in approximate equilibrium with their present overburden pres-
239
DIAGENESIS O F SANDSTONES AND COMPACTION
sure. During the early stages of compaction, the parameters that control porosity are mainly textural, i.e., grain size, sorting, roundness, shape, and flexibility of certain types of grains (Von Engelhardt, 1960,pp. 3-16; Fraser, 1935; Gaither, 1953; Hamilton and Menard, 1956). The results of experiments on the influence of sorting and roundness on compaction of pure quartz sands are given in Figs. 3-142A,B.In the pressure range up to 100 kg/cm2, compaction occurs as a result of rearrangement of sand grains into a more dense packing system and, to a minor degree, by elastic compression of the individual grains. At pressures above 100 kg/cm2, compaction increased as a result of cracking and shattering of the grains in the compression apparatus. In agreement with other investigators, however, Meade doubted whether this cracking and shattering would occur under natural conditions of com-
i
A -
50i u.
- 0 1
t
0.4
10
100
1000
1
10
3 1000
100
: 4oor C
50
1-
MICA
20%
I
PRESSURE
1
to
IKG PER CM'J
Fig. 3-142.Influence of different factors on relations between porosity and pressure in sands, as determined in laboratory experiments. A. Influence of sorting in well-rounded quartz sands (Roberts and de Souza, 1958);sorting index (a@)defined by Inman (1952, pp. 135-136); median diameter of two better sorted sands = 0.60 mm; median diameter of sand with poorer sorting = 0.48 mm. B. Influence of rounding of quartz sands, 0.42-0.84 mm in size (Roberts and de Souza, 1958). C. Influence of mica particles mixed in different proportions with rounded quartz sands (Gilboy, 1928,p. 560);particPes of both constituents are 0.42-0.59 mm in size. (After Meade, 1966,fig. 10,p. 1097; courtesy A.m. h s o c . Pet. Geologists.)
240
K.H. WOLF AND G.V. CHILINGARIAN
paction of sands. The slow rate of pressure increase in nature, extending over thousands and millions of years, permits other processes to be operative, such as plastic flow, pressure solution, and reprecipitation. Sands that are composed of minerals other than quartz may respond differently to pressure, with the softer grains being more readily deformed. That the greater porosity of well-sorted sands, in contrast to more poorly-sorted ones, persists during early compaction is indicated by the results presented in Fig. 3-142A. The influence of particle roundness is demonstrated in Fig. 3-142B, where angular sands show greater initial porosities. This reflects the instability of the initial packing of the angular grains. The angular particles are more compactible than rounded ones of the same grain size. The pronounced influence of platy and flexible mica particles on the behavior of sands in compaction experiments, is demonstrated in the five graphs of Fig. 3-142C. The porosity, compressibility, and elasticity of sand increase with increasing mica content. Experiments by McCarthy and Leonard (1963) support these results and suggest further that the finer the mica flakes among the sand grains, the greater the increase in porosity per unit increase in mica content. It should be noted, however, that a permanent, rather than elastic, deformation of mica plates may also take place during the early stages of compaction of micaceous sands. At pressures between 0 and 100 kg/cm2, grain size is the most important parameter influencing compaction. Not only is it inversely related to water content and porosity of the sands, but the size of the particles influences most other factors that control mechanical and chemical compaction. The effect of grain size may be so strong that the expected decrease in porosity with depth of burial may be obscured. The presence of mica may have a greater influence on the mass properties of sands than textural variations, such as those of rounding and sorting, of the non-mica grains. It is important to stress here, that more complex combinations of sand, silt, and clay should be experimentally examined in order to determine the influence of their relative proportions on the rate and degree of compaction. In many laboratory investigations on the changes of mass properties with increasing pressure (simulating burial of sediments in basins), the possible effects of temperature were not taken into account. It is, therefore, of particular interest to note that Somerton and El-Shaarani (1974) found that the compressibility of sandstones increases with increasing temperature. The effect of temperature was more pronounced at the lower effective pressures. Sawabini et al. (1974) determined relationship between the void ratio (= volume of pores/volume solids) and effective pressure, pe (Pe = pt -pp, where pt is the total overburden pressure and pp is the pore pressure) for unconsolidated sandstones from .a depth of 3000 f t (Fig. 3-143). In the effective pressure range of (F3000psi, the void ratios varied from 0.85 to
DIAGENESIS OF SANDSTONES AND COMPACTION
I
O m
L
,
1
1
1
1
1
HXK)
I
I
1
L
I
24 1
I
QW
EFFECTIVE PRESSURE, psig
Fig. 3-143. Experimental relationship between void ratio and effective pressure for unconsolidated sandstones. (After Sawabini et al., 1974, fig. 9, p. 136; courtesy SOC.Pet. Eng. AIME.)
0.19. These authors tested unconsolidated, medium- to fine-grained, arkosic sand cores obtained from oil-producing formations of Pliocene and Upper Miocene age in the Los Angeles Basin, California. The overburden (external) pressure was held constant at 3000 psi in a hydrostatic (three principal stresses are equal) compaction apparatus, at a temperature of 140"F,while producing the interstitial fluids and thus reducing the pore pressure and increasing the effective (grain-to-grain) pressure. It is the latter stress that causes compaction and, consequently, the one that should be used in plotting porosity-versus-pressurecurves whenever possible. PLASTICITY, COMPRESSIBILITY, DENSITY, AND THIXOTROPY OF SANDY SEDIMENTS AND SANDSTONES
A number of physical and chemical properties of sandy sediments are related to compaction in that (a) the degree and rate of compaction controls these properties, and (b) vice versa, these properties themselves influence the style, rate and degree of compaction. The most important properties referred to here are plasticity, compressibility, density and thixotropy* of sandy sediments (with silt and/or clay), sands, and sandstones. These topics are only briefly discussed in this chapter and only some selectively chosen mate-
* Thixotropy is the ability to gel (become firm) upon quiescence and to become fluid upon agitation.
242
K.H. WOLF AND G.V. CHILINGARIAN
rial is presented. From the outset it should be pointed out that to confine the discussion to sand-sized components is really meaningless, because the four properties listed above are a direct function of purity, or impurity, of the sediments, i.e., the amount of clay and silt admixed with the sandy constituents is extremely important in determining the behavior of the sediment. To extrapolate this argument, deflocculated clay minerals may fall into the group of clay-sized components, although the largest particles are up to 5 p in size and may fall into the silt-sized group. Once the clays are flocculated and form larger aggregate particles, however, they may belong to the silt- and sand-sized classes. It should be remembered, therefore, that even though this book is devoted t o sands, the data on clays given below aids in better understanding the behavior of clayey siltstones and clayey and silty sandstones. This has been pointed out also in Chapter 2. The bulk density* of sedimentary rocks depends on: (a) the density of individual types of minerals present; (b) the proportions of different minerals with varying densities; (c) the porosity; and (d) the amount of fluids in the intergranular pores. Inasmuch as the porosity depends on the degree of compaction and cementation, for example, one can speak of original and secondary densities of the sedimentary rocks. The densities usually increase with increasing degree of compaction and lithification, and details are given below. Figure 3-144presents the hydraulic equivalents of light and heavy minerals commonly found in sediments. In a general way, particles with similar hydraulic properties have a tendency to accumulate together t o form sedimentary laminae or beds as a result of both shape and size sorting, during transportation to and within the depositional environment. The actual density of sedimentary particles is, of course, important, but inasmuch as the shape of the grains controls their behavior during erosion, transportation and settling in fluid media (i.e., water and wind), it is the combination of both density and shape plus size that determines the hydraulic properties of particles. To be able to compare these properties, hydraulic equivalents, like those in Fig. 3-144,have been determined. Although a number of variables control the density of rocks, particular
* Bulk density of sediments has been defined as the weight of the sediment per unit of bulk volume (bulk volume, V , = pore volume, V,, + volume of solid grains, V,) or the mass of the sediment per unit of bulk volume. The mass p is attracted by the earth with a force (weight) having magnitude p g , where g is the gravitational acceleration. For example, pure water, which has a specific weight of 62.4 lb/cu ft, has a density of 1.94 (= 62.4/32.2) slugs/cu ft (gravitational acceleration = 32.174 ft/sec/sec). If density is reported in g/cm3, then it is equal to specific gravity.
243
DIAGENESIS OF SANDSTONES AND COMPACTION
Hobit Aciculor
I [Gypsum IUI
Cordierite
Prisrnotic
[Sillimonite ApOaTourmaline
e-
zircon
Rutilei
Equant
To bu lo r
I
Anotose Feldsporbmmui
Barite Kyonite
Andolusite UIlU mnmnmmmChlorlte mTopoz BIOtiternmmmm mnmmOChloritoid mum Mvscovite ontrnorillonite Im Dlospore olinite
Platy
Density
20
25
30
35
40
Brookite
45
Hematite Ill
5.0
5
Fig. 3-144. Diagrammatic representation of hydraulic equivalence of the commoner light and heavy minerals. (After Griffiths, 1967b, fig. 10-5; copyright 0 1967 McGraw-Hill, New York.)
lithologic units have been found to have a mean bulk density that can vary distinctly from unit to unit (cf. Fig. 3-145, for example). Before entering the subject matter of plasticity, compressibility and re-
Oriskany
Venango
Bradford Miocene
Sespe
Fig. 3-145. Mean bulk'density for rock types, (After Griffiths, 1967b, fig. 18-3; copyright
@ 1967 McGraw-Hill, New York.)
K.H. WOLF AND G.V. CHILINGARIAN
244 Inviscid fluid
- - -
Viscous 1Iuid
IPascalian)--IStokesiani
/I
Visco-elastic fluid IMaxwelii
Elastlco-vIscOus solid (Kelvin)
Elastic solid Rigid solid IHookeanl-IEuclidea~)
Plastics IBoltzrnannl Plastics ilinghaml
Fig. 3-146. Rheological classification of materials. (After Fredrickson, 1964, fig. 55; courtesy F’rentice-Hall, Englewood Cliffs, N.J.)
lated phenomena, it should be pointed out that these concepts imply flow under pressure. The field of rheology is concerned with such phenomena and includes a whole range of “fluid” and “solid” types of materials, as shown in the classification in Fig. 3-146.The flow phenomena as related to one particular type of geologic or natural system, i.e., muddy sediments including sandy deposits, are presented in Fig. 3-147.A comparison between the flow behaviors of Newtonian (viscous) (called Stoke&an fluids in Fig. 3-146)and Bingham fluids is presented in Fig. 3-148. D i spers ions
Shearhardening di spers ionS
V i s c o s i t y dependent on r a t e o f shear
V i s c o s i t y independent of r a t e of shear
O i ~ p e r s i o n sw i t h v i s c o s i t y dependent on tinw o f r e s t and time of shear
D i spe r s Ions w i t h Y Is cos It y independent of time o f r e s t and t i m e of shear
False-body dispersions
Rheopectic d i spe rs i oils
Thixotropic dispersions
Plastic d I spe r s ions
Pseudo-plastic d i s p e r s ions
Diiatant dispersions
Non- rheopect ic
dispersions
Fig. 3-1 47. Classification table summarizing the relationships of the flow phenomena in muddy sediments. (After PryceJones, in: Boswell, 1963; courtesy W.H.Heffer and Sons, London.)
DIAGENESIS OF SANDSTONES AND COMPACTION
I
A
245
B
Fig. 3-148. Schematic diagram of flow behavior of Newtonian (viscous) ( A ) and Bingham ( B ) fluids. Below point 1 on curve A , the flow is viscous, whereas above it the flow is turbulent. Pressure p t is necessary to start the fluid movement (yield point). 1 ’-2 ’: region of plastic flow; above point 2 ’, the flow is turbulent.
Newtonian fluids are the simplest fluids from the standpoint of viscosity considerations, because the viscosity coefficient, p , remains constant at all rates of shear. Viscosity 1.1 is equal to K ( d q / d p ) , where K is a constant and d p / d q is the slope of the straight-line portion of curve A in Fig. 3-148. Curve B in Fig. 3-148illustrates the flow behavior of Bingham fluids. As shown in this figure, a finite, minimum pressure pt must be applied to start the movement of the fluid and is referred to as the yield point. The straightline portion of curve B represents the plastic flow region (line 1’-2’). The ~ equal to K ‘ ( d q / d p ) ,where K‘ is a constant and dqldp is plastic viscocity 1 . 1is the slope of the 1’--2‘ line. Some of the aspects related to the flow phenomena of sediments is determined by their rheology or rheologic properties. These properties, which determine the flow of materials in general, have been investigated in detail by physical scientists and engineers, but only occasionally have been applied to geologic problems (e.g., Boswell, 1961;Elliston, 1963;McNeill, 1963, 1966). The type of flow phenomenon in which the change from a rigid to a fluid condition is produced under ordinary temperature by mechanical action, without the application of heat, is called “thixotropy” (part of the encompassing field of rheology). After being undisturbed for a while, the material will set again to a jelly-like mass, in some cases at once and in others slowly. A water-plus-clay system can change from a gel to either a sol or a more fluid gel as a-response to mechanical disturbance, and then change back to gel when left at rest.
K.H. WOLF AND G.V. CHILINGARIAN
246
As Boswell (1963),for example, has pointed out, differences in texture (angular vs. rounded grains, grain-size distribution, etc.) and composition (e.g., percentage and types of clays in the matrix of sands) result in variations in the rheologic, especially thixotropic, properties of sediments. Differential compaction may depend on these properties, because the capacity to retain moisture and the tensile and shear strengths determine the sediments’ ability to change in response to overburden load. Many intraformational structures are evidence of this, although the present knowledge of rheological properties of sediments in controlling diagenetic disturbances and reworking of unconsolidated deposits is only in its infancy (e.g., Elliott, 1965). Figure 3-149A1is the more commonly accepted way of representing the composition of sediments, i.e., the interstitial water is ignored. If one wishes to consider the rheological properties of sediments, however, then the presFigure ence of water has to be considered, as shown in Fig. 3-149A2,B. 3-149A2also illustrates the changes in properties from “permanent fluid”, through “thixotropic”, to crumbly, plastic and cloddy, depending on the composition. Figure 3-150shows the rheotropic zones in relation to depth below the surface, as one might expect in a sedimentary basin, as a function of grain size. This diagram by Boswell (1963)has been applied to practical geologic problems by McNeill (1963,1966). Below the surface of sediment accumulation (i.e., “belt of variables” or zone of physical, biological, and chemical variables), the following zones succeed each other: (1)permanently fluid muds (clay and/or silt); (2) thixotropic sediments; (3)plastic deposits; (4)“indurated” and compacted sediments. In the thixotropic zone, enough water is present in the sediments to allow them t o flow under shock or repeated vibration. The sediments, however, become firm again when the disturbance ceases. The plastic sediments can flow under stress, but as soon SAND
SILT
WATER
SILT
and
SAND
CLAY
Fig. 3-149A. Diagrammatic representation of varying composition and characters of noncalcareous muds. A1 = dry; A2 = wet. (After Boswell, 1963, fig. 1 ; courtesy W.H. Heffer and Sons, London.)
247
DIAGENESIS OF SANDSTONES AND COMPACTION
SAND and SILT
Fig. 3-149B. Diagram showing relationships of calcareous and argillaceous muds. (After Boswell, 1963, fig. 2; courtesy W.H.Heffer and Sons, London.)
as the pressure is released, the material reverts to a rigid state. The formation
of small faults may be the result of plastic flow. The tendency for creation of fractures, faults, and brecciation increases with depth towards the indurated zones and towards areas where sediments are well cemented. Cementation has been ignored in Fig. 3-150 to simplify the presentation. It is edaphic deposits Belt of Variables
CLAY
I
SILT
hypedaphic deposits
I
0mm
SAND
Fig. 3-150. Diagrammatic representation of diagenetic zones in relation to depth below sea floor and grading. The Belt of Variables and edaphic and hypedaphic deposits extend across the figure. (Cf. also McNeill, 1963, in Elliston and Carey, 1963; and McNeill, 1966.) (After Boswell, 1961, p. 100, in Boswell, 1963, fig. 5; courtesy W.H. Heffer and Sons, London.)
248
K.H. WOLF AND G.V. CHILINGARIAN
significant to consider Boswell’s (1963)statement that compaction of a stratigraphic section of sediment will not proceed evenly with geologic time, but in stages when rheotropic yield values are reached for each sedimentary deposit as burial continues. The resistance to deformation is related to the various index properties of the sediments (including rheotropy or thixotropy, for example), and inasmuch as these properties change from one sedimentary unit to another, compaction will proceed irregularly. In some instances, older beds may require longer time than the overlying younger ones to reach maximum compaction. Sposmatic differential compaction or settling may, therefore, be a normal feature of mechanical diagenesis. As Boswell pointed out, there are various ways in which sediments can offer resistance to stresses and, consequently, the style of the disturbance can vary accordingly. Four examples of this are presented here: (1)“Shearrate blockage” (i.e., dilatancy) in the case of clean, well-sorted, and wet sands, silts, and muds, when the disturbance occurs rapidly. This leads to the formation of a more dense and rigid mass of sediment accompanied by the expulsion of water. As Boswell pointed out (1963,p. 107), originally bedded or laminated sedimentary units that underwent “shock” during an earthquake, when the deposits were still thixotropic, can become unstratified or homogeneous as a result of an obliteration of the bedding. (2) “False-body thixotropy” in systems where clay minerals and electrolytes are present, results in a “rigid” gel persisting until the yield point is reached, after which “liquid flow” occurs. (3)“Plasticity”, where the water content is within certain limits, causes the material to be a firm gel until the yield point is reached; thereafter, the sediment will fold, shear, microfault and, ultimately, when the water content is sufficiently reduced, it will brecciate. (4) Induration (i.e., cementation) at a comparatively early stage if certain material, such as carbonate, has been precipitated interstitially. This results in fracturing, faulting or brecciation when the stress is rapidly applied, or will flow, fold or flex if the stress is slowly applied. In a water-saturated, lithologically fairly uniform stratigraphic sequence, the value of “m”* (Boswell, 1961) decreases steadily as the overburden pressure increases, so that at any particular locality, “m” decreases with increasing age of the deposits. Exceptions, however, do occur. Table 3-XXXV lists some values of “m” from which the “porosity” or void ratios also have been calculated. There is some correlation between the geologic age of the strata and moisture content, with considerable overlap. Table 3-XXXVI illustrates the general porosity decrease with age of the rocks. In coarse-grained sediments and sedimentary rocks, porosity can be measured
* “m” is conventionally expressed as g of,water per 100 g of solid dried to constant weight at 105-110°C.
249
DIAGENESIS OF SANDSTONES AND COMPACTION TABLE 3-XXXV
Variation in moisture content, porosity, and void ratio with age of argillaceous rocks (after Boswell, 1961, table 1 ) m*
(range)
Devonian, Upper and Middle, marls Carboniferous, Middle, Coal Measures, shales Triassic, Keuper Marl Jurassic, Lias clay Jurassic, Oxford Clay Cretaceous, Gault clay Eocene, Lower Red Beds Eocene, London Clay
Porosity (%)
16
31
9-18
13
26
13 18 22 29 19 24.4
2.5-20 14-22 18-24 28-30 12-27 18-33
25-42
very variable
Void-ratio (e) calculated from m
calculated S and M from rn
9-24
Eocene, Upper 15-29 Oligocene, Upper 19-29 Pleistocene, bedded clays 16-48 usually Postglacial and Recent
rn
(mean)
0.44
(M)
0.35
26 33 38 44 34 40
4.8-20.2 (M) 22.5-27.7 (S)
0.35 0.49 0.61 0.78 0.52 0.66
39 40 41-54
29 (MI
40-82
30-50
4.0-14.6
24 (MI
mode 24.5 23 24
0.64 0.66 0.69-1.17 (S)
0.66-4.6
24-170
* rn = g water/l00 g solid; drying temperature = 1O5-11O0C;values of void ratio, e, and porosity were calculated from rn ; the values from Sorby (S)and Moore (M) are included. either directly by standard techniques such as those used by civil and petroleum engineers, or by indirect methods, such as by determining the natural moisture content “m”, as long as the rock sample is saturated with water when collected. In the latter case, however, the mineralogic composition of the sample should be known, because in the case of clayey rocks, the “porosity” values calculated from “m” will have a different connotation as “m” includes water in the lattices of the clays as well as that adsorbed on their surfaces. Drying temperature will also affect the results (see Cebell and Chilingarian, 1972). Susceptibility to thixotropic behavior increases if the system is essentially loosely packed and a large proportion of the particles is of a diameter less than 1p. The thixotropic properties of sand and silt are greatly enhanced by
K.H. WOLF AND G.V. CHILINGARIAN
250 TABLE 3-XXXVI
Relationship between porosity and age of rocks (after Boswell, 1961, table 2) Age
Porosity (%)
Pleistocene Mio-Pliocene Miocene Upper Cretaceous Permian Pennsylvanian Mississippian Devonian
36-41 31 33-40 24-25 15 15-18 10-1 1 11
the presence of clays. The addition of electrolytes may result either in an increase or decrease in thixotropy, depending on the type of electrolyte and its concentration. An addition of a small amount of clay or other colloidal material, which is strongly hydrophilic, would make a dilatant system more thixotropic. Whereas thixotropy and plasticity are related, dilatancy is antipathetic t o thixotropy. A dilatant state is typified by uniform, close packing of the system’s constituents. Certain natural or artificial systems are either thixotropic or dilatant, depending on the make-up of the system, e.g., percentage of water. The rate at which a disturbed thixotropic system sets varies considerably : some, such as suspensions of montmorillonite clays, can become a gel almost instantaneously, whereas others require minutes, hours, or even days. This time factor should be considered in natural geologic systems, because of the different rates of setting of sediments with a variable composition from one unit to the next, after they were disturbed by an earthquake. Both Boswell (1961) and Niggli (1952), among others, considered the base exchange properties of clay minerals in a system as related to its thixotropy. There is a decrease in thixotropy when electrolytes, such as NaC1, KC1, NaOH, and KOH, are added in certain concentrations to montmorillonite plus water systems, whereas the thixotropy of kaolinite and other clays and powdered minerals increase when the same concentrations of electrolytes are used. One of the authors of this chapter (G.V.C.) found that very small (<1%) concentrations of sodium hydroxide increase thixotropy of both montmorillonite and kaolinite muds, probably because of adsorption of hydroxyl ions on clays. High concentrations of electrolytes usually flocculate the system, greatly increase the viscosity, and decrease thixotropy of clay plus water systems. Upon shaking, the water and clay components in these systems may separate from each other. Much research, however, remains to
DIAGENESIS OF SANDSTONES AND COMPACTION
251
be done on the numerous factors that determine thixotropy. For example, some organic liquids greatly increase the thixotropy of fine sediments, whereas others reduce it. Some hydrophilic, natural organic colloids increase thixotropy by plating the mineral particles with colloidal sheaths. The three main factors which control the degree of thixotropy are: (1)electrostatic or ionic attractions; (2) Van der Waals forces of attraction between molecules; and (3) hydrogen bonds or bridges. In the thixotropic state, the energy of attraction between the particles exceeds that of repulsion, whereas when the mass of material is in the dilatant state, the reverse is true. The surface area of solid components in a system also plays an important role in determining the degree of thixotropy. The area increases as the particle size decreases and as the shape of the constituents departs increasingly from the spherical form. I t should be noted, however, that a thixotropic state can be present in a mass of equidimensional and even spherical grains. The surface area decreases as a result of flocculation of muds (i.e., clay plus water systems) by electrolytes, for example. Creation of gels can be explained by the formation of surface-to-edge contacts in clays, because surfaces of clay plates are negatively charged, whereas the edges are positively charged. Upon shaking, the surface-to-edge contacts are broken and the mass becomes fluid again. Niggli (1952) discussed some properties of sediments which determine compressibility. As pointed out previously, not much detailed information is available on the compressibility and related properties of sediments composed of various proportions of clay, silt, and sand. As the following information demonstrates, it is not sufficient in future research to merely add a certain amount of clay to silt- and sand-sized material for experimental purposes. The exact mineralogical composition of the clay-sized components have t o be known before the results of the experiments can be meaningfully interpreted. Both the adsorbed water content and the extent and type of base exchange that occurs, which mainly determine the compressibility of clayey sediments and, therefore, also of clayey sandstones, also must be known. Figure 3-151 shows diagrammatically the relationship between the range of plasticity (expressed as the percentage of water present in the system) and the type and amount of three different clay minerals (kaolinite, Ca-montmorillonite, and Na-montmorillonite) in clay plus quartz mixtures. With an increase in the content of quartz grains, the plasticity decreases. Figure 3-152 presents the compressibility of quartz, muscovite, kaolinite, and bentonite. One should note, however, that the curve for “bentonite” (an altered tuff; Wolf, 1959) is to be accepted only with care, because the properties of montmorillonite minerals, of which bentonite is commonly composed, varies according to the base exchange cation present (i.e., Caversus Na-montmorillonite). The amount of compaction (in percent, at a pressure of 1kg/cm2) increases with increasing pressure (see Table 3-XXXVII).
K.H. WOLF AND G.V. CHILINGARIAN
252
I
QUARTZ CONTENT,?,
U
u
Fig. 3-151.Plasticity range for various mixtures of quartz with: ( I ) kaolinite, ( 2 ) Camontmorillonite, and (3)Na-montmorillonite. Quartz content diminishes (= reduces) the plasticity range. (After Niggli, 1952,fig. 93,p. 269;courtesy Birkhauser, Basel.) Fig. 3-152.Dependence of the degree of compression on the type of pelitic (= clayey) minerals based on experiments of Haefeli and Von Moos, 1938. A, B, and C are >0.002 mm in size. A = quartz; B = muscovite; C = kaolinite; D = bentonite. (After Niggli, 1952, fig. 94,p. 269 ;courtesy Birkhauser, Basel.)
Table 3-XXXVIII presents the amount of water that can be adsorbed on finegrained materials and the time required to do so. Future comparative studies should determine to what extent the values obtained from laboratory work will apply t o newly-accumulated natural sediments in a number of different sedimentary environments. Table 3-XXXIX demonstrates the dependency of plasticity values on the base exchange capacity of clay minerals, whereas in Table 3-XL another variable, namely quartz content, has been considered. In Table 3-XLI, the grain size of the quartz grains was varied, whereas the amount and type of clay mineral remained constant. There is a general decrease in the lithification, roll, and plasticity values with an increase in grain size. The presence of quartz having diameters less than 2 p changes the. plasticity only little in contrast to when diameters of quartz grains are larger than 2p. In instances
253
DIAGENESIS OF SANDSTONES AND COMPACTION TABLE 3-XXXVII
Specific compressibility (in %) of compressibility a t 1 kg/cm2 as a starting point (after Haefeli and Von Moos, 1938, in Niggli, 1952, table 34, p. 268; courtesy of Verlag Birkhauser, Basel) Material
Pressure (kg/cm2) ~~
~
0.5
1
2
4
-1.51
0
2
4.35
8.3
-3.3
0
3.5
7.4
11.7
-2.39
0
2.7
5.5
8.6
-0.37
0
1.1
2.0
3.1
Platy minerals in low proportions in CaCO3-rich delta sand
0
1.41
3.0
3.5
Platy minerals in high proportions in CaCO3poor delta sand
0
2.6
5.5
8.7
Muscovite, pure
< 2p
Kaoline from Zettlitz, Germany + quartz, 2/3 of the mass of kaoline 200-5OOp in diameter + quartz, 1 / 2 of the mass of kaoline 200-500~ in diameter + quartz, 1/3 of the mass of kaoline 20b-800p in diameter
8
where the components used had diameters less than 2p, however, the plasticity values depended on the types of minerals present (Table 3-XLI1,a).Values for liquid and plastic limits and plastic index for naturally-occurring clay samples are given in Table 3-XLII,b. Sawabini et al. (1974) studied the compressibilities of unconsolidated, fine- to medium-grained arkosic sand cores, lf inches in diameter and 3-4 TABLE 3-XXXVIII Adsorption of water under normal conditions (after Endell, in: Niggli, 1952, table 36, p. 271; courtesy of Verlag Birkhauser, Basel) Type of material
Maximum water content in % dry weight
Fine quartz sand-quartz pelite 27-32 Mica pelite 125 90 Kablinite pelite Ca-montmorillonite 300 Na- montmorillonite 700
Required time in seconds 10-2 3 25 1200 2400 36,000
K.H. WOLF AND G.V. CHILINGARIAN
254
TABLE 3-XXXIX Dependence of plasticity range of the same soil (black soil of India, with 60% particles with d < 2p) after base exchange with different cations (after Oakley, in Niggli, 1952, table 37, p. 271; courtesy of Verlag Birkhauser, Basel) Li-clay Na-clay Mg-clay Water adsorption from N/10 chloride solution; water content (index) a t 110" C 46 Plasticity range (index) in % HzO 82
35 60
17 56
Ca-clay K-clay 17 42
15 22
TABLE 3-XL Dependence of plastic property on the quartz content and type of clay minerals (after Niggli, 1952, table 32, p. 268; courtesy of Verlag Birkhauser, Basel) Ratio of quartz to clay mineral
9:l 7 :3 1:l
0:l
Water content (%) of kaolinite consistency limits
Water content (%) of Ca-montmorillonite
lower
upper
plastici- consistency ty (plas- limits tic index) lower upper
15.5 12.8 12.8 35.7
16.3 19.0 21.6 65.3
0.8 6.2 8.8 29.6
17.8 21.9 49.5
52.5 75.3 140.6
Water content (%) of Na-montmorillonite
plastici- consistency ty (plas- limits tic index) lower gpper 34.7 53.4 91.1
18.5 23.5 47.0
47.5 122.3 214.0 475
plasticity (plastic index) 103.8 190.5 428
TABLE 3-XLI
>
Influence of grain size of quartz on the consistency properties of 50% kaoline (30% 2p; 14% 2-5p; 3% 5-10p; 3% < l o p ) + 50% quartz; quartz grain size was varied from sample to sample as given in the table (after Niggli, 1952, p. 268; courtesy of Verlag Birkhauser, Basel) Grain size of the quartz
Liquid limit Plastic limit Plasticity range (index)
<2p
5-lop
20-5Op
100-2OOp
200-500p
60 40.9 19.1
43 28.5 14.5
36.5 21 15.5
36.8 20.7 16.1
35.1 19.1 16.1
255
DIAGENESIS OF SANDSTONES AND COMPACTION TABLE 3-XLII
<
Plasticity range of material with grain size d 21.1 and naturally occurring clays (after Niggli, 1952, tables 35a,b, p. 270; courtesy of Verlag Birkhauser, Basel) Type of material
Plastic limit
Quartz (Dorentrup, Germany) approx. 28 Calcite, grain size < 21.1, artificially prepared 30 55 Muscovite, grain size < 2p, artificially prepared Kaoline (Zettlitz, Germany) 42 Illite (?) or montmorillonite containing kaoline (Sarospatak) 43 Ca-bentonite with montmorillonite, after Endell 50 47 Na-bentonite with montmorillonite, after Endell Kaoline (Union Co., Illinois): Total sample Lop 0.51.1
36.3 37.1 39.3
Liquid limit approx. 28 48 78 84
18 23 42
120 141 475
77 91 428
58.3 64.2 71.6
Na-montmorillonite (Belle Fourche, South Dakota): 97 Total sample
625-700
Attapulgite (Quincy, Florida): Total sample
177.8
116.6
Plastic index
0
32.3 528-603 61.2
j6
inches long. The cores were obtained from the oil-producing sands of Pliocene age from the Los Angeles Basin, California. Direct measurements of the pore fluid pressure and bulk volume change of each sample were made in a hydrostatic compaction apparatus as the pore fluids were expelled. A t a constant overburden pressure of 3000 psi and a temperature of 140"F, the calculated to 3 * l o d 5 psi-', bulk volume compressibilities ranged from 7.4 * whereas the pore volume compressibilities varied from to psi-' in the 0-3000 psi effective pressure range. The effective bulk and pore volume compressibilities were defined by Sawabini et al. (1974) as follows:
and :
K.H. WOLF AND G.V. CHILINGARIAN
256
where C b e is the effective bulk volume compressibility in psi-'; cpe is the effective pore volume compressibility in psi-'; vbi and Vpi are the initial bulk volume and the initial pore volume of the sample, respectively, in cc; p t is the total overburden pressure in psig; p e is the effective pressure, which is equal to the difference between the total overburden pressure and the pore pressure, in psig; and T is the temperature during the test, in F. Some of the results obtained by Sawabini et al. (1974)are presented in Figs. 3-153,3-154,and 3-155.These authors found that compressibility increases with increasing feldspar content. Somerton and El-Shaarani (1974) studied a group of sandstone and a group of siltstone cores obtained from wells drilled in the Imperial Valley, California, at temperatures as high as 200°C and pressures up to 16,000 psi. Bulk compressibilities of all rocks increased with increased temperature, but the effect of temperature was more pronounced at the lower stress revel (up to =5000 psi). Insufficient data were obtained t o evaluate the effect of temperature on the compressibility of liquid-saturated rocks. Somerton and El-Shaarani (1974),however, stated that the effect appears to be about the same as for dry rocks. Maxwell and Verrall (1954)experimentally investigated consolidation of quartz sand to sandstone, simulating deep burial. Compaction was under5 .-
n
'
c
a
-
K
W
I 3 J
9 0.2 o.z
Y
J
t
3
m W
t
1 u
ld4 01
I '
EFFECTIVE P~ESSURE,p i g
lo4
lo*
b
lo3
EFFECTIVE PRESSURE, psig
10'
Fig. 3-153. Relationship between void ratio and effective pressure for unconsolidated, arkosic oil sands. (After Sawabini et al., 1974, fig. 10, p. 136; courtesy SOC. Pet. Eng. AIME.) Fig. 3-154. Relationship between effective pressure and effective bulk volume compressibility for unconsolidated, arkosic oil sands. (After Sawabini et al., 1974, fig. 6, p. 135; courtesy SOC.Pet. Eng. AIME.)
DIAGENESIS OF SANDSTONES AND COMPACTION
0
s 4
257
J DEFORMATION
Fig. 3-155. Relationship between effective pore volume compressibility and effective pressure for unconsolidated, arkosic oil sands. (After Sawabini et al., 1974, fig. 8, p. 136; courtesy Soc. Pet. Eng. AIME.) Fig. 3-156. Shape of load-versus-deformation curve for sands compacted by direct pressure only. (After Smalley, 1963, fig. 1; courtesy Nature, London.)
taken at room temperature and at 1atm pore-fluid pressure. There was a steady decrease in porosity from about 35%at a total overburden pressure of 10,000 lb/inch2 to about 25%at 50,000 lb/inch2. At a pressure of 50,000 lb/inch ', which is equivalent to 45,000 ft of overburden, the rate of decrease of porosity with increasing pressure was found to be zero. Borg and Maxwell (1956) examined the nature of deformed material, without considering the particle size variation caused by the various pressures. The experimentally simulated high overburden pressure caused complete deformation of samples, which were initially incompletely compacted tind uncemented. On the other hand, Smalley (1963) designed experiments to show the nature of the deformation process as obtained in the laboratory (Fig. 3-156) by measuring the grain size alteration during compression (Fig. 3-157 and Table 3-XLIII), rather than measuring the porosity change. The three stages of deformation observed by Smalley (1963) in a piston compaction apparatus are presented in Fig. 3-156, where load is plotted versus deformation: Stage I depicts elastic compression and the beginning of yield, whereas Stage I1 represents the continued breaking of the sand grains and reduction of the pore space by filling of pores with the broken fragments. Stage 111. sets in when the pores have been filled and the sample becomes in effect a solid, incompressible mass. A 2-inch-diameter steel cylin-
K.H. WOLF AND G.V. CHILINGARIAN
258
LOAD,
tOnS
Fig. 3-157. Variation in particle size percentage in sands compacted at different pressures. (After Smalley, 1963, fig. 2 ; courtesy Nature, London.) Numbers on curves are B.S. sieve sizes.
der with a steel piston was employed t o compress 100-g sand samples at various pressures up to 20,000 lb/inch2. The quartz sand was graded prior to experiments so that 95.9 k 2% passed through a No. 16 and was retained by a No. 18 B.S. sieve. The samples were compressed at a chosen pressure and then analyzed again as to their three-component grain-size distribution, as shown in Table 3-XLIII and plotted in Fig. 3-157. The point of inflexion at the load value along the 18-sieve curve, represents the transition from the Stage I deformation to the Stage 11. The amount of material composed of quartz grains passing the No. 100 sieve steadily increases with compaction pressure. The 22-sieve curve flattens out very early, indicating that there was no increase in this fraction with increasing compression and that the percentTABLE 3-XLIII Size distribution of compacted sands (see Fig. 3-157) (after Smalley, 1963, table 1) Compacting load Compacting pres- % on No.18 B.S. % on No.22 B.S. % passing No.100 (tons) sure (lb/in.2) sieve sieve B.S. sieve
0 2 4 6 8 10 14 26 29
0 1,426 2,852 4,278 5,704 7,130 9,982 18,538 20,677
’
95.5 91.0 61.0 45.5 37.0 32.0 28.0 20.5 20.5
i:
2
4.0 f 2 5.5 15.5 18.5 17.5 18.0 16.5 15.5 15.5
0 0.5 2.5 7.5 10.5 13.5 17.5 25.5 27.5
DIAGENESIS OF SANDSTONES AND COMPACTION
259
age remained the same. The curves suggest that the very fine particles resulting from the breaking of the sand grains fill the pores during Stage I1 of the deformation. Also, it seems that 20% is a limiting value of the No. 18-sieve curve (Le., portion of the sample retained by a No. 18 sieve) and that a further increase in pressure will not change the amount of this size of the sand. Measurements after compression indicated that the assumed original porosity of 35% was reduced to 25% by application of a pressure of 20,000 lb/inch '. Additional experiments were carried out to establish whether or not differently-sized sands would cause any variations in the transition pressures for the various deformation stages. In Table 3-XLIV, Smalley (1963) showed the pressures at the Stage I-Stage I1 transition points in sands having different starting grain sizes. There is a steady increase in this pressure with decreasing particle size. As pointed out by Brooks (1966), the bulk densities of rock masses are important for interpretations of gravity anomalies. In geophysical work, the density variations with depth in sedimentary basins have to be known for precise interpretations and extrapolations of the results obtained. Numerous studies on density variations of fine sediments with depth are available (see Parasnis, 1960, for example), but at the time Brooks (1966) undertook his investigation, data on sandstones were less extensive. He referred to Taylor (1950) who offered some information on the relationship between density and porosity changes with increasing burial. Measurements on natural sandstones have to be compared with the results obtained from laboratory compaction tests on sands, so that one can assess the density distribution in ancient sandstone basins. The necessity of such a comparison becomes even critical, because many investigators claim that fracturing and fragmentation of sands, which occur in the laboratory compaction tests, do not occur in nature. It would be particularly valuable to assess the density distributions, mentioned above, in several basins that contain sediments ranging in age from the Recent through the Precambrian. Inasmuch as this range is TABLE 3-XLIV 1-11 transition pressures in compacted sands (see Fig. 3-156) (after Smalley, 1963, table 2)
B.S. sieve number
1-11 transition pressure (psi)
14-16 16-18 22-25 25-30 30-36 44-52
2570 2670 3200 3750 4050 4800
K.H. WOLF AND G.V. CHILINGARIAN
,
POROSITY %
Fig. 3-158. Relationship between porosity and saturation density (i.e., density of a rock when pore space is filled with water) for sands having different grain densities. (After Brooks, 1966, fig. 1, p. 63; courtesy Geol. Mag.)
rarely available in one basin, by investigating several basins the data can then be obtained step-by-step. Brooks collected samples across the rock sequence, as well as along the dip section, and performed compaction tests on disaggregated sandstone (Bunter Sandstone) to determine the effect of applied pressure on density. The information was used to prepare the density versus depth curve and to assess the bulk density variations in a sedimentary pile. A modified method of Parasnis (1960)for measuring densities was used by Brooks. A rapid method for determination of mineral and grain densities involves the use of a simple linear relationship between the saturation density of a rock and its porosity (= volume of pore space/total bulk volume = vp/vb, in percent), for rocks having the same grain density (see series of lines in Fig. 3-158).This relationship is used because most minerals associated with quartz in sandstones have densities significantly different from that of quartz (2.65g/cm3), e.g., 2.71 for calcite, 2.8-2.9 for dolomite, 2.76 for muscovite, 2.57 for orthoclase, 2.61 for albite, and higher than 5.0 g/cm3 for iron oxides. The densities of plagioclases, oligoclase, andesine, and labradorite, are close to that of quartz. A relationship between saturation density and porosity was shown to exist as follows: If ps = saturation density of rock (pore space filled with water) in g/cm3; pd = dry density (pore spaces empty) in g/cm3; pg = grain (or mineral) density in g/cm3; and r$ = porosity in percent = vp/vb x 100, where V, is the pore volume and vb is the bulk volume, then: r$ = (ps - Pd)
x 100
(3-8)
DIAGENESIS OF SANDSTONES AND COMPACTION
261
Dry density Pd, weight of grains/total volume, is equal to: (3-9) Combining eqs. 3-8and 3-9: =
441 -P g ) 100
+P g
(3-10)
For rocks of the same grain density, therefore, pa is directly proportional 6.The slope of the line is (l-pg)/lOOand the intercept on the pa-axisis pe. The family of curves in Fig. 3-158enables one to determine the average grain density of a rock if ps and 4 are known. Brooks used this diagram in his investigations to detect any significant differences in pg between specimens that could account for any observed density trends. Brooks had to make certain assumptions and mentioned that other factors affect density distribution in a sedimentary sequence, e.g., (a) irregular distribution of cement, (b) neomorphism of minerals during weathering at the surface, (c) presence of unconnected pore spaces, and (d) different grain-size distributions (see also Smalley, 1963). The influence of fluids on density has been discussed by McCulloh (1967)and some of the pertaining information is presented near the end of this section. Brooks' (1966)results, obtained from borehole specimens across the bedding planes and covering a depth range of nearly 1000 ft, are graphically presented in Fig. 3-158.Grain densities range from 2.64 to 2.68 g/cm3, with resulting variation of approximately 25% in saturation density. To reduce this variation to 1096, the grains with densities falling outside the range of 2.66-2.66 g/cm3 were not used. The remaining grains were tested for a possible relationship between ps and depth and a value of r = 0.281 was obtained. This indicates a probability of occurrence by chance of greater than 1 in 10,i.e., the correlation is inadequate. Hence, there is no significant increase of grain density (pg)with depth over the range exhibited by the borehole specimens. The samples collected along the dip and their graphical analysis indicated that the sandstones had a range of grain densities ( p e ) similar to that of the borehole specimens, i.e., across the bedding planes. No significant relationship between ps and the depth along the dip was found. Brooks (1966)performed compaction tests to examine further the possibility of significant density changes with depth by using disaggregated dune sandstones samples. The results of a number of compaction tests are presented in Fig. 3-159,in which density is plotted versus the applied pressure. The dry density values were converted into equivalent saturation densities before plotting, for direct comparison with densities on undisturbed samples. to
262
K.H. WOLF AND G.V. CHILINGARIAN
2.30
ESTIMATED DEPTH OF BURlAL,ft (p=2.25g/cm3)
p
5 m . I
law0 I
15,oOO
I
I
I
1
. l
cE
Pl
c
i
t cn z
2.10
W
n
z 0 Ia
a
2 1.90 I-
a
ul
I
0
1
2.0
4.0
I
SD
1
8.0
PRESSURE, tons/in'
Fig. 3-159. Relationship between pressure (or depth of burial) and saturation density for disaggregated Bunter Sandstone. Curves 1-3 = samples deposited through gauze; curve 4 = sample vibrated to minimum porosity. (After Brooks, 1966, fig. 2, p. 65; courtesy Geol.
Msg. 1
Equivalent depth of burial values are shown on the pressure-axis, representing the approximate thickness of overlying sandstones assuming a bulk density of 2.25 g/cm3. In one set of experiments, the sand was deposited by a free fall from a funnel into the cylinder, resulting in a high initial porosity (40--50%). The results of three compaction experiments are shown as curves I, 2, and 3 in Fig. 3-159. After a preliminary sharp increase in density, the density gradient was reduced to less than 0.04 g/cm3/1000 ft during further compaction. This low value may suggest why significant density increases were not apparent in the natural rock samples collected from the 800 and 600 f t of original depth range, in particular when the influence of grain size and sorting is not known. At a pressure equivalent to an overburden load of about 10,000 ft, the saturation density of samples compacted in the laboratory was always substantially less than the mean value of 2.26 g/cm3 for the natural samples collected. Also, an unloading curve (curve I , Fig. 3-159)shows an elastic relaxation of grains, which resulted in a reduction in density of about 9% of the total density increase due to compaction. Consequently, these curves do not represent the natural compaction of natural rock samples during burial. Inasmuch as naturally-deposited sands usually have a much lower porosity than those obtained by free fall in the laboratory, Brooks (1966) assumed that the unsatisfactory etperimental compaction curves may have been due t o unnatural, initial grain-accumulation packing. Hence, the samples, after
DIAGENESIS OF SANDSTONES AND COMPACTION
263
being dropped into the cylinder, were subjected to intense vibration under water to obtain the minimum porosity. The compression curve 4 (Fig. 3-159) shows that the initial density of the sample approached the natural density. After further compaction and a certain amount of crushing, the sample attained the mean natural density at a pressure equivalent to an overburden of 10,000 ft. Thus, curve 4 represents a more plausible curve of compaction for the natural rock specimens, with a field bulk density of 2.26 g/cm3. The original low porosity of the natural specimens were explained to be the result of wind accumulation, because the sediments of this origin have been shown to have always very low porosities. As Brooks pointed out, water-laid sandstones have different accumulation textures and, therefore, obey different compaction laws, especially if their original porosity was significantly higher than that of the natural sandstone samples used above. Compaction curves I, 2, and 3 (Fig. 3-159) could apply t o some actual water-laid sandstones with sufficiently low densities. These curves indicate that burial to 10,000 f t would result in deformation that would cause a 10% increase in density (not considering long-time effects of compression on pressure solution, for example). On the other hand, density gradients below 1000 f t are very low and could easily be obscured, over limited depth ranges, by other factors not discussed by Brooks. From his tests and theoretical considerations, Brooks concluded that his results are useful in indicating that the averaged densities of surface rock samples are likely to give close approximation to the bulk densities at depth, as long as no unexpected changes in lithology take place. McCulloh (1967), in his important contribution on the volumetric mass properties of sedimentary rocks in situ and the gravimetric effects of petroleum and natural gas reservoirs, has assembled data that is useful for geological and stratigraphic analyses of rock characteristics and in the field of petroleum exploration. Particularly interesting is the fact that he has given consideration to the properties of the various possible subsurface fluids that can alter some of the physical rock properties. McCulloh investigated variations in densities and porosities with depth, and stated that the first-order factors controlling density are mineralogic composition, porosity, composition of the pore-filling fluid, temperature, and fluid pressure. Second-order factors that determine the densities of sedimentary rocks are mainly those that influence the total porosity, e.g., grain size, sorting, depositional environment, depth of burial, postdepositional cementation and recrystallization, deformational history, pore-fluid pressure history, and geologic age. In a Paleozoic section, a few thousand feet thick, composed of consolidated and compacted sandstones, shales, and limestones, an intimate relationship between the dry-bulk density, total porosity, and density profiles over a vertical depth was determined by McCulloh (1967). In a rock with a
K.H. WOLF AND G.V. CHILINGARIAN
264
particular constant grain density, porosity is a straight-line function of drybulk density. The latter ranges from a density of the grains when the porosity is zero to a density approaching zero at 100% porosity. The variations of porosity and density of sedimentary rocks are a function of depth of burial, because gravitational consolidation reduces porosity and simultaneously increases the density of the rocks. These changes are mainly the function of maximum net overburden pressure, degree of compaction, and geologic time. Numerous second-order effects mentioned above, however, also play a role. The interplay of these complex relationships is sufficient to result in extensive variations in porosity versus depth curves from basin to basin and from locality to locality. Such is the case, for example, in Fig. 3-160(Dickey, 1972), where a number of curves for different sandstone units of varying geologic age indicate diverse bulk density and porosity versus depth relationships. McCulloh (1967)obtained a curve of maximum porosity versus depth as shown in Fig. 3-161,by plotting total porosity versus depth for a large number of sedimentary rock samples from a wide variety of geologic environments, represented by a broad range in lithologies of all geologic ages, and from an extreme range of depths. As more data is being accumulated, this curve will probably undergo revision in the future. The greatest variation in porosity is to be expected among the youngest rocks and, especially, newly-accumulated sediments. There is a tendency for later (diagenetic and epigenetic) processes to reduce or even eliminate the minor primary differences in porosity of otherwise litholpgically similar deDEPTH, m
“
T x x ) 2 m 3 o O 0 4 o O 0 ~
0 1
1
0
1
I
1
’
1
UMO
I
I
1
1
8ooo DEPTH, ft
I
“
I
I
‘
12.OOo
I
I
I
0
1 4 m
Fig. 3-160. Relationship between depth o burial and density (as well as porosity) of sediments from various localities. (After Dickey, 1972, fig. 5, p. 7; courtesy Int. Geol. Congr., Montreal.) 1 = Athy (1930), Pennsylvanian, Oklahoma; 2 = Dallmus (1958), Tertiary, Venezuela; 3 = Eaton (1969), Tertiary, California;4 = Dickinson (1953), Tertiary, Gulf Coast;5 = Boatman, Louisiana; 6 = Griffin and Bazer (1969); 7 = Rochon (1967), Louisiana; 8 = Reynolds; 9 = Hedberg (1926, 1936), Tertiary, Venezuela; 10 = Philipp et al. (1963), Germany; 11 = Atwater and Miller (1965), average sandstone, Louisiana.
0
TOTAL POROSITY. IN PERCENT
0
I
jl If
jf iI il
--Probabla maxlmum averaga porosltv of nr arvolr sandstone
if
iI il
probable I---Maximum rosity for most sedC mantary rocks I1 ;I
il
’
if I ;I ;I
-Probabia upper lltnlt of porosity for virtually all sedlmantary rock8
il
il
a of porosities of f I , A a n gmost redlments ynd
-$I I
sedimentary rocks
Fig. 3-161.Relationship between the total porosity of sedimentary rocks and depth of burial, based on empirical data from laboratory measurements of more than 4000 samples of conventional cores from (1)the Los Angeles and Ventura basins of California, (2) other scattered localities in the United States, and (3) the Po Basin of Italy. (After McCulloh, 1967, fig. 4, p. A9;courtesy U.S.Geol. Sum.)
266
K.H. WOLF AND G.V. CHILINGARIAN
posits. The finest-grained and best-sorted sediments tend to be the most porous and least dense upon accumulation, and may contain over 90% interstitial water. In the subsurface, under a few hundred feet of overburden, porosity varies over a wide range, and the coarse-grained and best-sorted sediments, e.g., sandstones and conglomerates, commonly are the most porous and the least dense. At depth beyond a few hundred feet, associated rock units of different lithologies have distinctly different porosities and densities just as they do upon accumulation. Based on extensive research work done recently by numerous investigators (e.g., Roberts, 1969;Chilingarian et al., 1973) one cannot state that clays are more compressible than sands (see Chapter 2, Vol. I). According to McCulloh’s data, the porosity may vary from about 10% at a depth of 20,000 f t to about 30%, and 0.0
DENSITY, IN GRAMS PER CUBIC CENTIMETf 0.5 -.
f t per bbl
cu f t per bbl
I
0-100.000 porn NaCl
I I I
I
I J ,
,
,
,
,
,
,
, I
Fig. 3-162. Variations in density of interstitial fluids as functions of depth, three sets of assumed temperature and pressure gradients, and fluid composition. Curves for interstitial waters from fig. 5 in original paper. (After McCulloh, 1967, fig. 6, p. A12; courtesy U.S. Geol. Surv.)
DIAGENESIS OF SANDSTONES AND COMPACTION
267
occasionally more, at 5000 ft. The greatest ranges in porosity occur in the Pliocene and Quaternary rocks, i.e., the youngest deposits, whereas the porosity of older rocks tend to vary less at all depths. The limits of the range of average porosities in spatially associated clastic marine rocks (i.e., sandstones and argillaceous rocks) with maximum porosities, are given in Fig. 3-161. As a result of variations in mineralogy, the dry-bulk density of sedimentary rocks with zero porosity varies from 2.65 g/cm3 to 2.87 g/cm3. As porosity increases, the variation in grain density will have a decreasing effect
PmbabI. marlmum averam
ecrasltY at l.m n
I-
630
eL
im B
-1
5 10
2 1.6 1.7 Id 1.9 2.0 2.1 2.2 2.3 2.4 2.5 2.6 2.7 FLUIWSATURATED ROCK OENSITY, IN QRAMS PER c u n i c CENTIMETER
0 1.S 1.7 1.8 1.9 2.0 2.1 2.2 2.3 2.4 2.5 2.6 2.7 FLUIO-SATURATED ROCK OENSITY. I N GRAMS PER CUBIC CENTIMETER
8. DEPTH. APPROXIMATELY 4,000 FEET; TEMPERATURE, 1ZO.F; PRESSURE. 1.795 POUNDS PER SQUARf INCH
A. DEPTH. APPROXIMATELY 1.000 FEET: TEMPERATURE. 81’F; PRESSURE,460 POUNDS PER.SQUARE INCH
I-
530
f
I” 2
2r
10
01.6
1.7
1.1
1.9
2.0
2.1
2.2
2.3
2.4
2.5
2.6
2.7
FLUIWSATURATED ROCK DENSITY. I N GRAMS PER CUBIC CEhTlMETER
C. DEPTH. APPROXIMATELY 6.000 FEET: TEMPERATURE. 150’F; PRESSURE. 2.M5 POUNDS PER SQUARE INCH
1.6
1.7
1.8
1.9
2.0
2.1
2.2
2.3
2.4
2.5
2.6
2.7
FLUID-SATURATED ROCK DENSIlY. IN GRAMS PER CUBIC CENTIMETER
D. DEPTH, APPROXIMATELY 10,oOO FEET; TEMPERATURE, 209.F; PRESSURE, 4,465 POUNDS PER SQUARE INCH
Fig. 3-163. Rock density in situ as a function of total porosity and fluid composition at various temperatures and pressures assumed to be prevalent in young deep marine sedimentary basins. Assttmed grain density = 2.67 g/cm3. (After McCulloh, 1967, fig. 8, p. A17; courtesy U.S.Geol. Sum.)
K.H. WOLF AND G.V. CHILINGARIAN
268
on the dry-bulk density, so that two rocks with a 40% porosity and differing in grain density by 0.1 g/cm3, will differ in dry-bulk density by only 0.06 g/cm If these rocks are saturated with the densest of all possible interstitial fluids (namely, a brine), the water-saturated, subsurface densities would also differ by only 0.06 g/cm3. If a rock with 40% porosity is filled with petroleum, instead of salt water, there will be a decrease of 0.2-0.4 g/cm3 in the in-situ bulk density. Consequently, inasmuch as the underground densities of natural interstitial fluids vary widely, their densities must be considered in calculating rock bulk densities. McCulloh (1967)has discussed this aspect at some length, and some of his data in the form of curves are presented here (Figs. 3-162,3-163,3-164,3-165and 3-166). Assuming fixed dry-bulk and grain densities, the density of a rock with a certain porosity is a function only of the density of the fluid filling the pores. Inasmuch as grain densities vary with lithology and fluid densities vary with composition, temperature and pressure, a simple straight-line relationship between porosity and rock density cannot be expected. If one assumes, however, that (a) the porous sandstone is composed of grains having the same density, and (b) the temperature, pressure, and composition of the fluid are constant, then a straight-line graph of rock density versus porosity can be obtained (Fig. 3-163A for a sandstone with grain density of 2.67 g/cm3). The data in Fig. 3-163is applicable to a hypothetical deep marine sedimentary basin having a temperature gradient of 1"F/67f t and a pressure
'.
F
40 Ic W z
35
0
a 30 W
a z 25 W a $20
88 2
W
& \
52
6000
*<
8500
*%
15
< 10
I- -
20,000
7 -
?
:!5 0
1.7 1.8 1.9 2.0 2.1 2.2 2.3 2.4 2.5 2.6 2.7 FLUID-SATURATED SANDSTONE DENSITY. I N GRAMS PER CUBIC CENTIMETER
1.6
Fig. 3-164. Relationship among average density in situ of reservoir sandstone, its fluid composition, and total porosity. Assumptions: a. probable maximum average porosity for depth, temperature, and pressure gradients prevalent in young marine sedimentary basins; b. gas-oil ratios as given in Figs. 3-162 and 3-163. (After McCulloh, 1967, fig. 9, p. A18; courtesy U.S.Geol. Sum.)
DIAGENESIS OF SANDSTONES AND COMPACTION
269
DENSIM, IN G R A M S PER CUBIC CENTIMETER
0
POROSITY, IN PERCEM
5
0
0
EXPLANATION
Densityfidd br many cornrnercially important ptmleurn fluids
Density field for most interstitial waters
Density field of sedirnentory Density field d sedimentary rocks saturated with wrocks saturated with troleum fluids woter
POCOSity fieU of sediMntS and sedimentary rocks
Fig. 3-165. Densities and porosities of sedimentary rocks and their constituents as functions of depth and of temperature and pressure gradients prevalent in young deep marine sedimentary basins. (After McCulloh, 1967, fig. 10, p. A19;courtesy U.S.Geol. Sum.)
gradient of 0.445 psi/ft. Similar density-depth curves can be calculated for other temperature and pressure gradients. In Fig. 3-163A,the sandstone is assumed to be saturated at different times with each of the four principal interstitial fluids. The densities of these fluids are plotted versus depth in Fig. 3-162,under temperature and pressure conditions common at a depth of 1000 f t in young and deep marine sedimentary basins (indicated as “normal” in Fig. 3-162).The four lines in Fig. 3-164 converge on the grain density at zero percent porosity, as the lower porosity results in the reduction of density contrast between rocks saturated with water and petroleum. According to Fig. 3-161,the maximum sandstone porosity in young’sedimentary basins is 50% at a depth of 1000 ft, whereas the average porosity of these sandstones at shallow depth is about 40% or
K.H. WOLF AND G.V. CHILINGARIAN
270
5 ;
40
0
35
'
w' 30
z
n
-2
a
v)
25 20
LL
15 l-
a
a 0 10 0 " 5
-1
2O
n
+
-0
-0.1 -0.2 -0.3 -0.4 DENSITY CONTRAST, IN GRAMS PER CUBIC CENTIMETER
Fig. 3-166. Density contrasts between reservoir rocks saturated with water and those saturated with petroleum fluids, as functions of hydrocarbon composition and average total porosity, assuming temperature and pressure gradients prevalent in deep young marine sedimentary basins. (After McCulloh, 1967, fig. 11, p. A20; courtesy U.S.Geol. Surv.)
less. The horizontal line at a 40% porosity intersects the sloping straight line of methane-saturated rock at the lowest density of 1.61 g/cm3 (Fig. 3-163A). This suggests that the rock density lower than 1.6 g/cm3 at a depth of 1000 f t is probably due t o the presence of coal, diatomites, or rocks of unusual composition. Figure 3-163B is analogous to Fig. 3-163A, except that the temperature and pressure are those common at a depth of 4000 f t , according t o Fig. 3-162.According to Fig. 3-3.61, the maximum sandstone porosity at a depth of 4000 f t is less than 40% and averages about 30%. If a rock at that depth has a density less than 1.90 g/cm3, it must be due to the presence of coal, a rock of unusual composition, or an overpressured formation (a rock dilated by abnormal fluid pressure). The graphs in Figs. 3-163C and 3-163D supply similar data for depths of 6000 and 10,000 ft, respectively. Relationship between the average sandstone density and depth is presented in Fig. 3-164. The temperature, pressure, and porosity gradients used in plotting this figure are nearly limiting values, i.e., temperature and pressure at any depth are near the minimum values found in boreholes, whereas porosity at any depth is near the maximum. Although an increase in temperature would result in a less dense fluid, it would probably be accompanied by a decreased porosity, thus compensating for each other. Figure 3-165 summarizes the data assembled by McCulloh (1967) on the relationship between the depth of burial and densities of sedimentary rocks.
DIAGENESIS OF SANDSTONES AND COMPACTION
271
Fig. 3-167. Density contrasts between water-saturated argillaceous rocks of various maximum porosities and reservoir sandstones saturated with water or petroleum fluids, as functions of average total sandstone porosity and fluid composition. (After McCulloh, 1967, fig. 12, p. A21; courtesy U.S.Geol. Sum.)
In Fig. 3-166,the density data presented in Fig. 3-164have been replotted in the form of density contrasts. Water-saturated porous reservoir rocks are used as a standard against which the densities of same rocks saturated with various fluid hydrocarbons are contrasted. The “total” porosity is the average porosity and, as in Fig. 3-164,the values of average total porosity have been correlated with probable maximum depths by reference to the porosity data in Fig. 3-161.Figure 3-167demonstrates that the greatest density contrasts between water-saturated sandstones and water-saturated argillaceous rocks occur where the sandstone porosity is maximum, i.e., in the youngest, uppermost rocks in a stratigraphic section. EFFECTS OF COMPACTION FLUIDS ON CLAY MINERALOGY IN SANDSTONES
It is now an accepted fact, supported by laboratory and petrologic observations on natural rocks, that clay minerals within sandstones (e.g., graywackes, kaolinitic quartzites, montmorillonitic pyroclastics or volcanic arenites, and argillaceous arkoses) can be the product of: (a) detrital origin, with or without secondary alteration in situ; (b) direct precipitation from fluids within the sandstone framework; (c) neomorphic replacement of unstable minerals, such as feldspar; and (d) any combination of the above. Fluids, like compaction solutions, are most important in the processes in-
272
K.H. WOLF AND G.V. CHILINGARIAN
volving clays in sandstones, because they supply and remove chemical components and make the environment conducive for reactions requiring a certain pH and Eh (Dapples, 1967). In other sections of this chapter, several publications were discussed that present data on the release of fluids from the clays during diagenetic mineralogical transformations (e.g., Powers, 1967; Fyfe, 1973). Some of the data concerning the influence of fluids on the composition of clays, which is a somewhat different, though related, topic from the release of water by clays during diagenesis, are presented here. Hiltabrand et al. (1973)performed diagenesis experiments on argillaceous sediments in artificial sea water and recorded mineralogic and chemical elemental changes with variation in temperature, depth, and geologic age. They concluded that the mineralogy of mudstones and shales is the result of diagenesis rather than a reflection of the chemical milieu of the depositional environment. This conclusion also applies to clays present as interstitial matrix within sandstones, as long as the clays are detrital in origin and diagenesis was related to the near-surface conditions of the depositional milieu. It may not apply to neomorphic clays because they can originate much later through the interaction of unstable elastic grains and compaction fluids, for example. If ion exchange is one of the processes involved, then the reader is referred to the publication by Carroll (1959)who described ion exchange of all major clay mineral groups, some other minerals (chlorite, glauconite, allophane, feldspar, quartz, and zeolites), and some rocks, e.g., basalt, tuff, shale, and bentonite. Sarkisyan (1971)used a scanning microscope in the study of clay cements in elastics of an oil-bearing basin to determine reservoir properties. The reservoir rocks are present at different depths and form zones of postsedimentary (= epigenetic) alterations of rocks so that zones of kaolinite, kaolinite-chlorite, and chlorite cements were distinguished. The properties of the allogenic (= detrital) and authigenic (= neoformed) clay cements as determined by electron-microscope examination allowed differentiation between the two. Towe (1962)discussed the derivation of silica from clay deposits as a source of cement in sandstones. He stated that with an increase in depth of burial and under certain geochemical conditions, Si4+ can be expelled from the tetrahedral layers as it is being replaced by A13+ in the montmorillonitetype clays and thus cause their alteration to micas. This would result in an increase in illite content while silicon will be released during this transformation process. Inasmuch as it has been shown that both the depth of burial and the geothermal gradient appear to have an effect on such changes, it seems quite plausible to assume that compaction and compaction fluids are involved. Towe, using published data, showed that 3 g of silicon are released per 100 g
DIAGENESIS OF SANDSTONES AND COMPACTION
273
of pure clay in the transformation of a montmorillonite (bentonite) to an average illite, wheras 1 g/100 g is freed in a conversion of mixed-layer illite-montmorillonite to illite. Hence, there is a sufficient amount of silica available to account for the cement in sandstones, but the amount of silica available varies depending on several geologic conditions mentioned by Towe. One of the main controls would be the type of clay mineral supplied from the terrigenous source-rock area. If the latter had furnished montmorillonite to a marine environment, the resultant clay deposit should be a good source of silica, whereas if illite clay had been supplied, only a poor silica source would be available in the depositional environment. Whetten and Hiltabrand (1972) used recent, coarse, silt-free and clay-free sand from the Columbia River, composed largely of andesitic volcanic detritus, and performed a hydrothermal experiment at a temperature of 200°C and a pressure of 200 psi in a brine solution. After five months, the clastic grains were altered to a mixture of 82% (by weight) sand, 5% silt, and 13% montmorillonite clay. Although only the hypersthene grains were severely etched, other mineral grains and lithic fragments probably also participated in the reaction to form clay. The clay, which coated various parts of the apparatus, appeared to have at least partly precipitated from solution. The ' , a slight decrease in Ca fluid phase showed a slight increase in Na+ and K and Mg2+, and a great increase in Si4' content (saturation); pH remained constant at 3.5. These results suggest that the matrix of graywacke sandstones could, at least partly, be of secondary origin, as has been proposed by Cummins (1962) and is shown in Fig. 3-168. This supposition is also supported by the geochemical investigation of Reimer (1972), who found that the matrix in Precambrian graywackes formed through reassemblage of
'+
Stable sand grain
Rather unstable sand grain
Very unstable sand grain
Interstitial water
Fig. 3-168. Diagram showing presumed post-depositional origin of matrix of graywackes. (After C u m i n s , 1963, fig. 4; in: Pettijohn et al., fig. 6-20, p. 210; courtesy Springer, New York.)
K.H. WOLF AND G.V. CHILINGARIAN
274
chemical compounds derived from decayed labile clastic constituents such as feldspars, mafic minerals, and fragments of igneous and mafic volcanic rocks. The resulting matrix consists of chlorite, sericite, and dolomite in varying amounts. When the matrix plus dolomite content is plotted versus the content of labiles, a volume increase of 1.6 to 1.9 parts of matrix plus dolomite for 1 part of labiles is indicated. This suggests an outside source, and transportation of the constituents by pore fluids into the graywacke sandstone. The Ca2+ cation was partly derived from the decomposition and partial albitization of the plagioclase within the graywackes, whereas the Mg2+ was released from the basic igneous rock grains within the deposits. Most of the Ca2+ of the dolomite, however, was derived from the destruction of the plagioclase in shales alternating with the graywackes and was brought in by pore fluids. This is indicated by the peculiarities in the Rb/Sr ratios of both graywackes and shales. The CaO of the average shale was reduced from about 1.64% to 0.19% and the Sr content from about 81 ppm to 25 ppm, whereas at the same time the CaO content of the graywackes increased from about 1.64% to 1.89% and the Sr content from 81 to 91 ppm. The COz gas was almost totally introduced by pore fluids. Bucke and Mankin (1971;see also Wilson, 1971) studied the illite, chlorite, and kaolinite present in both sandstones and shales of mid-Pennsylvanian age and found consistent differences between the types and contents of clays in the shales and sandstones (Figs. 3-169,3-170,and 3-171).The shales Illite
0
Chlorite
Kaol i ni t e
Fig. 3-169. Clay mineral composition of Desmoinesian (Middle Pennsylvanian) sands and shales (finer than 4 p). Open ciicles = sand; solid circles = shale. (After Bucke and Mankin, 1971, fig. 2, p. 974;courtesy J. Sed. Petrol.)
DIAGENESIS OF SANDSTONES AND COMPACTION Sholt S o d
OILLITE ~'-JCHLORITE
275
Shale SDnd
KAOLINITE
Fig. 3-170. Comparison of clay-mineral content of adjacent sand-shale pairs (finer than 4 p). (After Bucke and Mankin, 1971, fig. 3, p. 974; courtesy J. Sed. Petrol.)
have a limited compositional range and illite dominates in all samples. Cornpositional variation is prevalent in sands, but illite is still the major constituent. Bucke and Mankin closely investigated associated sandstones and shales, because they considered this approach to be a prime starting point for establishing the relative importance of detrital versus diagenetic origin of clay
SAND untreated
+SHALES-
+SANDS-
Size in microns
SAND
solvatcd
SHALE untreated SHALE solvoted
0ILLITE
CHLORITE
~KAOLINITE
Fig. 3-171. Variation of Clay-mineral content in Desmoinesian sands and shales with size fraction. (After Bucke and Mankin, 1971, fig. 4, p. 975; courtesy J . Sed. Petrol.)
276
K.H. WOLF AND G.V. CHILINGARIAN
minerals. Various textural and other considerations led to the conclusion that kaolinite is authigenic, whereas illite and chlorite are detrital in origin. It should be noted here that the sandstones are much higher in kaolinite content than shales. Bucke and Mankin (1971,p. 975)stated: “The illite and chlorite were subjected only to some minor ‘repair work’ or reconstruction diagenesis . . . Those clays deposited in shales or mudstones have been preserved with only minor reconstruction by ion absorption and exchange. Any changes that may have occurred were probably physical changes, notably water expulsion during compaction as described by Burst (1969).Those weathered illites and chlorites deposited with coarser material remained in a post-depositional environment with continued relatively high porosity and permeability even after compaction. These clays had prolonged access to the required ions, notably potassium and magnesium, brought by circulating interstitial water thereby reversing some weathering effects.” Using the paragenesis of several authigenic minerals, i.e., quartz, kaolinite, calcite, dolomite, and pyrite-siderite, Bucke and Mankin suggested a trend toward a higher effective pH during diagenesis. It must be pointed out, however, that although Zimmerle (1963)and Sharma (1968,1970) based their interpretation of diagenetic alterations on pH, Bucke and Mankin stated that the actual change in pH may have been minor because an increase in temperature has much the same effect on solubility of quartz and calcite as an increase in pH. Carrigy and Mellon (1964)described authigenic clay mineral cements in sandstones from Alberta. Often no clear distinction can be made between detrital and authigenic clays; indeed no sharp division was possible between grains, cements, and matrices among the fine constituents. Nevertheless, certain useful results were obtained. The stratigraphic section studied is dominantly non-marine with a maximum thickness of about 25,000 f t prior to uplift and erosion. A generalized section is given in Fig. 3-172,which also shows the distribution of the authigenic clay minerals and zeolites. The strata markedly decrease in thickness to the east from the Foothills to the Plains, where they form a succession of nearly flat-lying, interbedded marine and non-marine strata, which are several thousand feet thick and extend north and south to the margin of the underlying Precambrian basement. The details related to the fine-grained minerals are as follows (Carrigy and Mellon, 1964,pp. 468-47 0): (1)Kaolinite is the most widely-distributed clay mineral cement. It is common both in marine and non-marine sandstones and is present in both the deeply buried, folded strata of the Foothills and in the shallow, flat-lying strata of the Plains. Kaolinite is the dnly common authigenic clay in quartzose sandstones low in feldspathic and volcanic detritus, where it is associated
DIAGENESIS OF SANDSTONES AND COMPACTION
277
LEGEND
non%&ne
and shale
morine sandstone
-
marine siltstone and shale unconformity
z - zeolites c -chlorite
k - koolinite i - illite m - montmorillonite
Fig. 3-172. Generalized columnar section of the Cretaceous-Tertiary strata of the Rocky Mountains and Plains regions of Alberta showing the known distribution of authigenic clay mineral and zeolite cements in sandstones. (After Carrigy and Mellon, 1964, fig. 4, p. 469; courtesy J. Sed. Petrol.)
with authigenic quartz and, generally, calcite. Hence, kaolinite in quartzose rocks originated from the chemical precipitation from the alumina-rich compaction solutions expelled from the adjacent shaly strata and is not the alteration product of feldspars. Kaolinite is also abundant in many sandstones with moderate amounts of volcanic detritus, where it is found with other cements, excluding zeolites. The chlorite-kaolinite assemblage is common in the fluviatile sandstones, but only kaolinite is abundant in sandstones interbedded with coal or marine shales in these units. Kaolinite is also present as the main clay cement in sandstones associated with marine and non-marine sandstones. Montmorillonite and glauconite are also common in many of the above sandstones; however, both minerals, especially glauconite, appear to be associated largely with altered clastic grains (probably pyroclastics?) in contrast to the intergranular occurrence of the authigenic kaolinite. (2) Authigenic illite is less widely distributed than kaolinite. It was found
K.H. WOLF AND G.V. CHILINGARIAN
278
only in the Foothills as interstitial cement in some non-marine fluviatile sandstones, which contain moderate t o abundant amounts of feldspar and volcanic clasts. In these sandstones, illite is commonly associated with other authigenic components: chlorite or kaolinite, quartz, and calcite. Illite is also present in one basal, marine, feldspathic sandstone member, cemented by a mixture of kaolinite and fibrous illite in combination with quartz and calcite. (3) Authigenic chlorite is abundant in many volcanic sandstones and is associated with other authigenic minerals: laumontite, illite, kaolinite, montmorillonite, quartz, and calcite. Chlorite is scarce or absent in coal-bearing or marine facies. It is also scarce or absent in the more deeply buried sandstones many of which contain abundant volcanic clastics, but which have kaolinite and/or montmorillonite as cements. Relation between types of clay minerals and depth of burial is probable and should be investigated. (4)Montmorillonite is present in many marine and non-marine sandstones; however, its distribution as an intergranular cement, in contrast to its presence as an alteration product within clastic grains, is limited. Where present, montmorillonite is invariably found in volcanic-rich sandstones, although these sandstones do not necessarily contain montmorillonite. The suggested origin of the four groups of authigenic minerals described above and correlations between the gross composition and inferred depositional environments, are presented in Fig. 3-173. Assemblages of authigenic silicates found in the shallow, undisturbed sandstones of the Plains are less diverse than those observed in the folded sandstones of the Foothills. In the opinion of the present writers, these differences in regional distribution of the authigenic clay cements have occasionally been attributed to differences in the physicochemical conditions and depth of burial to which the rocks were subjected, rather than to differences
MARINE AND NE AR-MARINE
illite
- kaolinite - q u a r t z
kaolinite - q u a r t z
kooliniie
-94
V O L C A N I C ROCK FRAGMENTS AND FELDSPARS-
F e l d s p a t h l c greywackes
P r o t o q u a r t z i tes
Fig. 3-17 3. Chart showing the relationship among authigenic silicate assemblages, bulk composition, and depositional environment of Cretaceous-Tertiary sandstones of the Alberta Foothills. (After Carrigy and Mellon, 1964, fig. 5, p. 471; courtesy J. Sed. Petrol.)
DIAGENESIS OF SANDSTONES AND COMPACTION
279
in the original composition and sedimentary environment. This assumption seems incorrect, however, because Carrigy and Mellon stated (p. 471): ". . . within the Foothills proper, where the sediments have been subjected to geosynclinal depths of burial and subsequent tectonism and, by inference, to widely varying conditions of pressure and temperature, authigenic silicate assemblages show no obvious correlation with original depths of burial or intensity of folding. Instead, they are intercalated throughout the stratigraphic succession in relation to sandstone composition and, to a lesser extent, the inferred depositional environments of the rocks, demonstrating the stability of all four major types of clay minerals over the range of temperatures and pressures prevailing during diagenesis. Presumably, the ultimate factor responsible for their formation was the composition of the pore fluids during cementation, as determined by the availability of the component chemical constituents from the surrounding detritus." Meade (1963, 1964) found an anomalous increase in pore space with increasing depth when one would expect a decrease in porosity. Grain-size variation could not explain this anomaly, so that Meade searched for additional factors through a literature survey and examination of cores. He concluded that the following variables, in addition to overburden pressure, have had an influence on pore volume in fine sediments: particle size, clay minerals and their adsorbed cations, concentration of interstitial electrolyte solutions and their acidity, clay-particle orientation, and presence or absence of microfossils (i.e., diatoms in this case). Figure 3-174 summarizes some of the factors of interest here in relationship to compaction and other diagenetic processes. Figures 3-174, a and b demonstrate the influence of particle size on poor volume: in both unconsolidated marine sediments, and sediments under overburden pressure, the pore volume increases with decreasing particle size. Figure 3-174,c shows that in the pressure of 1-100 kg/cm2, the sequence of increasing pore volume is kaolinite + illite + montmorillonite (see also Chilingar and Knight, 1960). This is probably the result of particle size, 'because the specific surface measurements show a decreasing order of particle size in the same sequence. Figures 3-174,d, e, and f are of particular interest to illustrate a possible relationship between chemical diagenesis and overburden pressure (i.e., compaction). In sandstones with clay minerals, either as a matrix or as lenses or beds, the exchangeable cations adsorbed by clays are also influencing pore volume under low-to-moderate burial pressures as shown in Fig. 3-174,d. Meade (1963) mentioned that the smaller the valence and the larger the hydration radius of the adsorbed cation, the greater the pore volume of the clay sediment. This does not seem to be applicable under large overburden loads (30-3200 kg/cm '), as demonstrated by experiments performed by
K.H. WOLF AND G.V. CHILINGARIAN
280 a
h
O
0 - O
C
10
Average partide size (9)
Kaolinite
l 100
10
Pressure
Effective overburden load ( kg/cm2)
(kg/cm2)
e
d
f
47
21
> 0 10
100
Pressure t kg/crn2 )
0
1 1
100
l b M , NaCl
,
10
100
Pressure (kg/cm2)
K 10-'M NaCl
0
10
1 100
Pressure !kg/cm2)
Fig. 3-174.Relations of void ratio to other factors, observed in natural sediments and in laboratory experiments. Void ratio is ordinate in all graphs; note different void-ratio scales. a. Relation to average particle size observed in unconsolidated sea-bottom sediments. Curve I modified after Von Engelhardt (1960,p. 15); curve I1 modified after Shumway (1960,p. 663). b. Generalized relation to effective overburden load and particle size in sediments. Modified after Skempton (1953,p. 55). c. Experimentally determined relation to pressure and clay-mineral species. Modified after Chilingar and Knight (1960,p. 104), to show their results to 100 kg/cm2. d. Experimentally determined relation to pressure and adsorbed cations in <0.2 p fraction of montmorillonite. Modified after Bolt (1956,p. 91). e. Experimentally determined relation to pressure and electrolyte concentration in unfractionated Fithian illite (about 60%by weight coarser than 2 p ) . Modified after Mitchell (1960,fig. M3). f. Experimentally determined relation to pressure and electrolyte concentration in < 0.2p fraction of Fithian illite. Modified after Bolt (1956,p. 92). (After Meade, 1963,fig. 2, p. 237;courtesy Sedimentology. )
Von Engelhardt and Gaida (1963). Figures 3-174,e and f illustrate the interrelationships among overburden pressure, degree of compaction, particle size, and electrolyte concentration. Further research work is needed in this area.
DIAGENESIS OF SANDSTONES AND COMPACTION
281
EFFECT OF COMPACTION FLUIDS ON TRACE-ELEMENT AND ISOTOPE COMPOSITION OF SEDIMENTS
The control of compaction fluids on trace-element and isotope composition of sediments is a highly diversified field of investigation. It can only be treated briefly here by considering a few specific examples and by some discussions related to the numerous complexities involved in genetic interpretations. The trace-element budget in a sandstone, that is related to the intergranular clay matrix and can, therefore, be influenced by compaction (or any other) fluids, is considered first.
The trace-element budget It has been shown by numerous investigators that: (1)detrital clays can undergo diagenetic ion exchange at the site of sedimentation; (2) detrital clays can be diagenetically altered partly to completely to new (neomorphic) clay minerals and release certain ions (this process may continue into the metamorphic stage); (3) labile clastic grains can alter into clay minerals during diagenesis and burial metamorphism; and (4) neomorphic clay minerals can be the result of chemical precipitation from solution, both during diagenetic and metamorphic stages. Any of the above clays may undergo more than one generation of changes in elemental composition, depending on various physical and chemical factors changing through geologic time. Compaction fluids may cause ion exchange and bring metallic ions that were not available earlier, or they may remove elements that were released into the fluid. In the latter case, diminution or depletion of ions would occur. The composition of the clays in cases 2 to 4 can be influenced by the compaction solutions. The boron content Keeping in mind the above discussion, the boron content of sediments is considered next. One could have selected a number of other elements, but the geochemical behavior of boron is particularly important because it is used as paleosalinity and paleotemperature indicator. Harder (1970) stated that the geochemical distribution of elements in sedimentary rocks is genetically complex as they occur in the following forms: (1)in clastic fractions in the sand, silt or clay; (2) adsorbed on fine particles, e.g., clays; (3) chemically precipitated fractions as newly-formed minerals; and
282
K.H. WOLF AND G.V. CHILINGARIAN
TABLE 3-XLV Boron content of oxides and silicates (after Harder, 1959, in Harder, 1970, table V, p. 162) Mineral
Boron content (ppm)
Remarks
Quartz Agate
0-35 2-90
Nesosilicates Cyclosilicates
several ppm, sometimes under the detection limit
Phyllosilicates
distinctly higher than in the other see Table 3-XLVI minerals, especially the pure, clean micas, serpentine and montmorillonite
Tectosilicates
several ppm, the higher values in sericitized plagioclases and in minerals of the sodalite and scapolite groups
approximately, mean value distinctly higher than quartz
(4) organically precipitated fractions. Boron is found in all four cases. The boron content in different sediments, different grain-size fractions, and minerals is presented in Tables 3-XLV, 3-XLVI, and 3-XLVII. Illites contain the most boron, kaolinites the least, and montmorillonites and, probably, chlorites have intermediate contents TABLE 3-XLVI Boron content of phyllosilicates (after Harder, 1959, in Harder, 1970, table VI, p. 162) Mineral
Boron content (ppm)
Muscovite Paragonite Biotite Sericite Illite Glauconite Montmorillonite Kaolini te Serpentine Chlorite
10-500, very variable values 50-250 1-6 40-2000 100-2000 or more several thousands
5-200
10-30 low or very high values, according to conditions of formation around 50, but generally lower
DIAGENESIS OF SANDSTONES AND COMPACTION
283
TABLE 3-XLVII Boron content of sediments (after Harder, 1970,table I, p. 155) Rock
Clays and shales Sandstones and sands Limestones Dolomites Ironstones Glauconitic rocks Possible average for sedimentary formations
Boron content (ppm) range
mean value
25-800 5-70 2-95 10-70 20-200 350-2000
100 35 27 28 85
(see also Couch, 1971). Boron content is greatly dependent on the type of facies, with clays being the richest in boron. The boron content in limestones is controlled by the amount and type of clays present, whereas in sandstones it depends on the amount of clay and the presence of tourmaline. Especially micas and sericitized plagioclase are rich in boron. Fine-grained iron oxide or hydroxide precipitated in a marine environment can contain up to 300 ppm boron, whereas only a few ppm are present in lacustrine deposits. Chamosite has similar boron contents, but glauconitic rocks are richer in boron. Glauconite, which is a diagenetic neomorphic mineral produced through recrystallization, has a crystal structure that can easily accommodate boron. Its boron content is higher than that in chamosite and iron ores. In iron oxide, iron hydroxide, and iron silicate (except glauconite), the boron is probably adsorbed at the surface. This weak type of fixation is responsible for the sharp decrease in boron content after even surficial weathering. The effects of diagenesis and metamorphism are even more pronounced, as discussed below. The boron content of clay minerals depends on a number of variables: (1) boron content of the fluids, e.g., continental and marine waters, volcanic solutions, and surface and subsurface fluids (Fleet, 1965); (2) time of exposure of clays to the fluids (Fleet, 1965); (3)rate of sedimentation (related to 2 in early stages of diagenesis; Fleet, 1965); (4)salinity (Fleet, 1965; Couch, 1971); (5) water temperature (Fleet, 1965); (6) common-ion effect; (7) specific-surface area (controlled by grain size; Couch, 1971); (8)type of mineral (Couch, 1971); (9) crystallinity (related to 8;Couch, 1971); (10) pH of the interstitial solution (Fleet, 1965); (11)inherited boron or boron on recycled clays, which will retain boron from its earlier history (Fleet, 1965; Perry, 1972; Couch, 1971); (12) organic content of the sediments (Fleet,
284
K.H. WOLF AND G.V. CHILINGARIAN
1965; Walker, 1964). (The above selected references are among the many publications available on these topics.) Although many investigators discussed the conditions close to the surface of sedimentation that control the boron content in clays, many of the same factors should be influential in the subsurface during higher-grade diagenesis and metamorphism. The following factors control the release or take-up of boron by clays: (1)time of exposure available (depending on rate and degree of compaction which, in turn, influence porosity and permeability); (2) vertical changes in composition of the interstitial and compaction fluids; (3) temperature; (4) pressure; (5) recrystallization; and (6)pressure solution (which may release boron from quartz). Based on experimental results, Harder (1970)showed that the uptake of boron by illite, for instance, occurs in two steps: (1)adsorption of boron on the surface of the clay mineral which proceeds very quickly; and (2) subsequent boron incorporation in the tetrahedral sites of illite structure, which is a very slow reaction. Release of the adsorped boron is easier, therefore, than that of the absorbed boron within the structure. Harder (1961)stated that diagenesis and metamorphism through recrystallization and neomorphism decrease the boron content of sediments. This boron is added to the “hydrosphere” and returns to the sea, thus completing one part of the boron geochemical cycle. Diagenesis generally reduces the initial boron content of sediments with time, because of the increase in grain or crystal size through recrystallization and neomorphism. This has been observed in the case ‘of recent, volcanic, submarine iron-hydroxide precipitates that contain 300 to 400g of boron per ton obtained from the sea water, whereas in Devonian iron ores of a similar genetic type, the boron content is only 18 g B/ton of precipitates. Crystallization and recrystallization of the ferric hydroxide to hematite releases most of the boron. Similar trends of boron decrease have been observed in purely sedimentary, syngenetic oolitic iron ores, as well as in iron silicate minerals. The latter were formed through diagenetic reduction of ferric hydroxide to ferrous hydroxide and subsequent reaction of the Fez+ with aluminium and silica. The geologically young chamosite contains 60 g B/ton of ore, the Devonian ores contain 50 g B/ton, and the partly chloritized ores of Ordovician age have only 13 g B/ton. In the latter case there is a change from the kaolinitetype structure of the chamosite to the chlorite structure. If the iron ores underwent higher-grade metamorphism so that the whole rock has been chloritized and contains garnet, feldspar, actinolite and magnetite, then the boron content is very low (e.g., less than 5 g/ton). Where the boron is diagenetically released into circulating subsurface fluids, it may form authigenic tourmaline in unmetamorphosed sediments. Perry (1972) studied boron uptake into the lattice of clays during the
DIAGENESIS OF SANDSTONES AND COMPACTION
285
deep-burial diagenesis stage. Such a study seems reasonable, according to him, because experimental results have shown increased boron fixation at higher temperatures (Couch and Grim, 1968) and that many illites are diagenetic in origin. Perry examined samples from deeply buried Gulf Coast sediments and determined that the starting clay mineral assemblage was composed mainly of a highly expandable, mixed-layer illite-montmorillonite with lesser amounts of kaolinite, discrete illite, and sometimes chlorite. With increasing depth of burial, the montmorillonite layers were converted to illite layers and the proportion of illite layers in the mixed-layered illite-montmorillonite increases from 30 to 80% with increasing depth of burial. The boron content of the less than 1-mu fraction from two wells showed only a vague trend of increasing boron concentration with depth. In another approach, Perry plotted the boron concentration in the less than 1-mu fraction versus the proportion of illite layers in the mixed-layer illitemontmorillonite. He found that the relationship is much more clearly defined in this case and that the amount of boron fixation in the fraction less than 1mu in size is best related to the diagenetic formation of new illitic layers. According to E. Couch (personal communication, in: Perry, 1972, p. 156), clay minerals being supplied by the Mississippi River contain 50-80 ppm boron. These concentrations are close to the lower values found by Perry for the highly expandable clays. Through diagenesis, the boron content for the illite-montmorillonite increased to nearly 200 ppm. By extrapolation of this data, a boron value of 270 ppm was obtained for a non-expandable illite, which falls well within the range of boron contents of the marine illites of sedimentary rocks (Reynolds, 1965; Harder, 1970). Perry (1972, p. 156) concluded that ". . .the diagenetic illite formation and its concomitant boron uptake therefore can account adequately for the boron increase necessary to indicate a marine environment. With increased time under diagenetic conditions, boron would continue to diffuse into the illite structure in order to reach higher boron values found in some older illites. The higher temperatures and more concentrated pore fluids encountered in diagenetic conditions provide more reasonable conditions for boron incorporation in the lattice than the conditions at the watersediment interface, and the diagenetic formation of new illitic material is certainly more favorable." Perry (p. 156) continued to state that ". . . pore water flow could increase the amount of possible boron gain from pore waters, but only if the flow is from a clay-deficient area where boron has not already been removed in significant amounts. This situation requires flow into a shale section, probably from a sandy, more porous medium, and is not a very likely hydrodynamic situation, especially for the amount of flow needed. If boron is not primarily gained from an external source (diffusion also does not appear to
286
K.H. WOLF AND G.V. CHILINGARIAN
be reasonable mechanism on any large scale), one must look for a redistribution within the shale so that boron can be added to the lattice of the illitic material forming diagenetically. A newly deposited sediment will have boron located in two regimes: boron adsorbed on the clay surfaces at the depositional site and previously inherited ‘detrital’ boron in the clay lattice. The adsorbed boron would then be retained through burial and finally incorporated into the lattice of the new illitic material formed during diagenesis.” Perry (p. 159) finally concluded that the boron-paleosalinity technique may only be useful as a very general indicator of depositional environment, and only after careful clay mineral determination and petrologic interpretations. Couch also stated in his personal communication that certain corrections explained by him have to be applied in order to give paleosalinity values which are in good agreement with other evidence. He concluded that as long as one is concerned with rocks of low permeability, postdepositional processes apparently do not alter the original boronsalinity relation. Walker (1964)discussed the use of the boron content of sediments as a paleosalinity criterion and concluded that, in spite of numerous complexities which demand a careful scrutiny and adjustment of the boron data obtained, this geochemical tool does supply a reliable measurement of paleosalinity. Some of these complexities described by Walker are related to the postdepositional enrichment of boron. He remarked (p. 211 and pp. 215-216) that the coarse-grained sandstones are commonly porous and have an anomalously high boron content. If compaction fluids from shales in more saline parts of the basin moved into the porous sandstone bodies, a reaction between the clay minerals (especially illite) in the sandstones and the boron might occur. The boron content of normal sea water (salinity of 30-40 parts per thousand) is approximately 50 ppm, whereas the observed boron content of marine illites is 200-300 ppm. This requires an equilibrium in an open system between the sea water and the illite and, of course, a slow rate of sediment accumulation. On the other hand, in a closed system where the clays are exposed to trapped interstitial fluids with a low boron content of around 5 ppm, a small enrichment of boron in the illite within the porous sandstone would result in a large relative decrease in the boron concentration of the interstitial solution. Consequently, there would be no appreciable enrichment of the boron in the illite. All of the above-considered publications refer to the important influence subsurface fluids may have on boron content. Some authors even mentioned compaction fluids as a possible variable directly controlling the boron concentration. In general, whenever various researchers refer to interstitial waters or subsurface fluids without being more specific, it is safe to assume that at least part of these solutions may be compaction waters. Whenever the data allows, one should be more specific as to the precise origin of the fluids.
DIAGENESIS OF SANDSTONES AND COMPACTION
287
Other trace-element contents Although the above considerations are confined to boron in sediments, similar allowances must be made t o the possible presence or absence of the effects of compaction fluids on the content of any other trace element, e.g., potassium, sodium, uranium, copper, lead, and zinc, to name only a few. Ericsson (1972),in his publication on the chlorinity of clays as a criterion for paleosalinity, remarked that the measured chlorine contents often varied greatly. He stated (p. 5) that grain size and organic carbon content have some influence on the vertical chlorinity changes within sediments. The vertical salinity variation in a particular area may also be influenced by diffusion and leaching, either by surface or subsurface water. These processes are controlled, for example, by compaction, but the magnitudes of various effects are not known. In addition, salinity of compaction fluids changes upon migration (Rieke and Chilingarian, 1974,pp. 239-272). Hartmann (1963)described, among other interdependencies of chemical elements, the relationship between iron and vanadium in sandstones. It is reviewed here as an example of an element-element correlation. Both Fe and V behave similarly geochemically, which results in their parallel occurrence in (a) hematite-rich, clay-poor sandstones (Fig. 3-175),and (b) clayrich, hematite-free sandstones (Fig. 3-176).The former case is more or less self-explanatory, because the vanadium is associated with the iron in hematite as a consequence of their similar geochemical behavior. In case b, the iron associated with the clay minerals was at the same time accompanied by vanadium, as has been demonstrated in the laboratory. On assuming that fluids, compaction waters or otherwise, can precipitate authigenically new
Fe,O,
CONTENT, %
Fig. 3-175. Graph demonstrating the dependency of the vanadium content (ppm V) on the total F e z 0 3 content (%) in red, hematitic, clay-poor sandstones. (After Hartmann, 1963, fig. 6; courtesy Geochim. Cosmochim. Acta.)
K.H. WOLF AND G.V. CHILINGARIAN
288
h,O,
CONTENT, Vn
Fig. 3-176. Graph demonstrating the dependency of the vanadium content (ppm V) in the total F e z 0 3 content ( W )in hematite-free sandstones and claystones. (After Hartmann, 1963, fig. 5; courtesy Geochim. Cosmochim. Acta.)
minerals (such as hematite), can cause alterations (e.g., hematite to goethite, etc.), and can dissolve minerals, then the occurrence of any accompanying minor and trace elements will exhibit certain regularities, whether they occur adsorbed on the surfaces of minerals or form part of their internal crystalline structure. Compaction fluids and isotope composition
In regard to the possibility that compaction fluids may be responsible for determining the final isotope composition of sediments, one example is considered here. Monster (1972,’ p. 941) stated that “. . . there is no apparent reason to assume that in a particular crude oil the sulfur in the sulfur-bearing organic compounds has a uniform isotopic composition. On the contrary, from what is known of the sulfur isotopic composition of organically bound sulfur in marine sediments, there are good reasons to expect considerable heterogeneity in sulfur isotope ratios.” Studies of recent sediments have revealed a large isotopic heterogeneity for sulfur, including organically bound sulfur, with depth and location within a particular basin. On the other hand, there was a certain S-isotope homogeneity for various oil fractions obtained by Monster. This is in strong contrast to the wide variability in isotopic content of the organic sulfur compounds in recent marine sediments, which are supposedly the precursors of petroleum in reservoir rocks. There are two processes, in particular, that can effectively increase the S-isotopic homogeneity: (1) primary migrational mixing, and (2) S-isotope exchange reactions.
DIAGENESIS OF SANDSTONES AND COMPACTION
289
(1)It is reasonable to assume that any particular basin has a wide range of isotope ratios for organically bound sulfur, whereas the range in ratios will be relatively small at any particular location. The mixing of the various organic sulfur compounds in the fluids that undergo either lateral or vertical migration, commonly caused by compaction, may result in a reduction of the range of isotope ratios. At the final location, each individual chemical compound will have an averaged isotope composition resulting from the contributions of various locations. (2) Exchange reactions between different compounds can only take place if they are in proximity to each other. Again, compaction fluids can achieve transportation of various compounds to bring them together for chemical reactions. In this regard, Monster (1972,p. 946) stated that many chemical reactions will take place during diagenesis, migration, and maturation. According to him, "addition of extraneous sulfur compounds at some stage of migration will, in general, change the 34S/32S ratio of total sulfur of the oil, and might extend the range of isotopic ratios". Similar arguments can be applied to C-isotopes. Figures 3-177 and 3-178 are schematic diagrams indicating the major processes affecting the isotopic composition of petroleum moving from a primary source to a reservoir rock. The values presented apply only to a particular marine sedimentary basin; -40
GAS PHASE -30 to-38
-?lo
&
6c'3%0
- 20
-I
I0
,
0 I
I
,
I
,d
LlOUlD PHASE - 2 0 to -23
Fig. 3-177. 3C/ 2C in petroleum from primary source to reservoir rock in typical marine sedimentary basin. (After Monster, 1972, fig. 3, p. 947; courtesy Am. Assoc. Pet. Geologists.)
290
K.H. WOLF AND G.V. CHILINGARIAN
they are not general, world-wide ranges, so that similar diagrams for other marine basins may vary in range. Additional complexities can be visualized if one considers isotope variations in the (a) normal, (b) evaporitic, ( c ) barred (i.e., "Black-Sea" type), (d) brackish-water, and (e) fresh-water basins. Certain geologic conditions could result in mixing of two or more of the fluids derived from these environments during movements of compaction fluids in the subsurface. FLUID-DIAGENESIS IN SANDSTONES
The field of fluid-diagenesis is a comparatively new one in contrast to the study of diagenetic textures and minerals. The latter may reflect chemical changes in intrastratal fluids, which caused the precipitation of cements and led t o other secondary alterations. It is easier to investigate the visual effects of diagenetic processes and use them for theoretical interpretation of the causes. With so many new precise instruments and techniques now at hand, however, there is a trend in investigating the geochemical and physical con-
DIAGENESIS OF SANDSTONES AND COMPACTION
291
ditions that resulted in diagenetic modifications. This can be done by studying recent sedimentary environments and by laboratory experiments and then using the results for establishing “diagenetic conceptual models” for a spectrum of environments in nature. For example, precise investigations of fluids under natural conditions as well as in vitro promise to provide data on the interactions of fluids and sedimentary particles. Inasmuch as compaction fluids are believed to constitute one major group of solutions responsible for diagenesis, a brief discussion is presented here. In general, fluid-diagenesis in sediments has been conveniently divided according to the solid component involved in the chemical reactions, e.g., fluid-diagenesis in clayey, sandy, and pyroclastic deposits. As has been pointed out in other sections, as a result of the hybrid origin and composition of sandy sediments, the data available on fluid-diagenesis in clayey sediments, for example, is applicable also to clayey sandstones. Runnels (1969)discussed an interesting aspect of fluid-diagenesis, namely, the precipitation of minerals as a result of mixing of two different solutions. Numerous varieties of natural waters are summarized in Table 3-XLVIII, and inasmuch as sediments accumulate in each one of these environments, the trapped primary interstitial fluids vary in composition accordingly. The secondary differential fluid alterations during burial complicate the situation further. Many of these sedimentary facies are closely associated in three dimensions, so that during migration as a result of compaction the various types of fluids will mix (see also Shelton, 1964). Runnels (p. 1188)mentioned some of the consequences: (a) precipitation of minerals; (b) production of porosity through mixing of sodium chloride waters (he plotted the solubility of calcite as a function of the content of neutral electrolyte, NaC1, TABLE 3-XLVIII Classification of natural waters (after Runnels, 1969, table 1 , p; 1189) I. Meteoric waters
11. Surface waters A. Lakes: (1) fresh, ( 2 ) saline B. Streams: ( 1 ) fresh, ( 2 ) saline C. Swamps: (1) fresh, ( 2 ) saline D. Oceans 111. Subsurface waters A. Vadose: ( 1 ) fresh, ( 2 ) saline B. Phreatic (1) normal formation water: (a) fresh, (b) saline ( 2 ) ascending thermal water: (a) fresh, (b) saline
K.H. WOLF AND G.V. CHILINGARIAN
292
in solution); and (c) dissolution (e.g., decementation) and precipitation (e.g., cementation). He showed that an understanding of these processes is very important in petroleum-reservoir and ore-genesis studies (e.g., Colorado Plateau uranium ores; fluorspar of the Kentucky-Illinois district; and Mississippi-Valley-type Pb-Zn ores). The simple classification of natural waters, as given in Table 3-XLVIII, incorporates 14 separate categories of waters, which in turn can give 91 combinations on mixing any pair of fluids. When one considers, however, that in nature three, four or even more fluids can mix in both surface and subsurface environments, then many more combinations can be envisaged. The complexity of fluid geochemistry is even greater when other parameters are considered, e.g., temperature and pressure differences and chemical variations with depth, which have been established for many sedimentary basins. Runnels considered the upper limit of diagenesis to roughly correspond to conditions of temperatures and lithostatic pressures in the deepest wells drilled (20,000-25,000 f t depth), i.e., temperature of about 200°C, pressure of 1500 atm, temperature gradient of loC/10O ft, and pressure gradient of 100 psi/lOO f t for water-saturated sedimentary rocks. Under the more extreme conditions, the dissociation of water as a function of temperature and pressure (Blatt et al., 1972) must be considered in geochemical investigations on fluiddiagenesis (Fig. 3-179). Sharma (1968, p. 232) stated that petrographic investigations are the study of the end results or effects of chemical reactions involving pore fluids and the solid matter of the host material, so that petrographic data can give
5e 0
80
40
I20
TEMPERATURE, I
!moo
I
~qaoo
200
160
1
1
240
OC 1
1
2 o w 3opoo
DEPTH, ft
Fig. 3-179. Variation in the dissociation constant of water as a function of temperature and depth in a geosyncline. (After Blatt et al., 1972, fig. 7-1; copyright 0 1972 PrenticeHall, Englewood Cliffs, N.J.)
DIAGENESIS OF SANDSTONES AND COMPACTION
293 THIN PLASTIC
Fig. 3-180. Sharma’s sediment model. (After Sharma, 1968, fig. 1, p. 233; Minemliu m Deposita. )
rtesy
only a partial answer to the questions on the origin of diagenetic products. To fill this gap, he performed experiments to determine: (1)the relationship between the chemistry of the water in the sediments and the mineralogy of the sedimentary particles; (2)the factors controlling the processes of calcite and silica precipitation in sediments; (3) the processes responsible for transforming sediments into hard rocks; and (4)the paragenesis involved. Sharma used a box containing various sediments, as shown in Fig. 3-180, and pumped fluids of known composition (i.e., artificial sea water of different compositions) through the sediments. Chemical analyses were performed on the effluent waters collected at regular intervals, whereas petrographic and X-ray analyses were done on the chemical precipitates formed during the experiments. Table 3-XLIX shows the chemical changes that took place. Both aragonite and silica were precipitated as cements of various forms. Sharma utilized a total volume of solutions of approx. 4500 ml in his model experiments with a consequent flow of 35.5 ml/cm2. This requires a burial depth of 47 cm, if the data by Emery and Rittenberg (1952,p. 755)is accepted. These authors stated that a burial of 100 cm of sediments will result in a flow of interstitial fluids of 75 ml/cm2 due t o compaction. Thus, the experiments performed were comparable to situations of early diagenesis during shallow burial. Sharma (1970) discussed: (1)diagenetic alterations of montmorillonite and degraded illite to chlorite, possibly by ionic influence in the marine
K.H. WOLF AND G.V. CHILINGARIAN
294
TABLE 3-XLIX Chemical composition of influent1 and effluent water injected through the model in run No. 2 (after Sharma, 1968, table 2, p. 235) 7/25/63 Influent water ~
Potassium Magnesium Calcium Strontium Lithium Chloride Sulfate Bicarbonate Carbonate Specific gravity Resistivity PH
310 ppm 651 ppm 128 ppm 11 PPm 4 PPm 21,780 ppm 1240 ppm 24 PPm 48 PPm 1.028 0.312 ohm/m 9.01
7/29/63 Effluent water
Gain
Loss
~
290 ppm 749 ppm 528 ppm 32 PPm 3 ppm 21,080 ppm 1359 ppm 1579 ppm 0 PPm
98 PPm 400 ppm 21 PPm 119 ppm 1555 ppm
1.027 0.312 ohm/m 6.95
20 PPm
1 PPm 700 ppm
48 PPm
0.001
The influent water was injected with a 90-psi carbon dioxide pressure at 72"F.
environment, and (2) changes of kaolinite t o illite or chlorite, which are often present in sandstone matrices. These mineral changes distinctly alter the chemistry of the trapped pore sea water, which often exhibits vertical variations in composition in sedimentary sections. As pointed out by Sharma (1970, p. 722), under ideal conditions the intrastratal fluid composition may be controlled by the physical characteristics (grain size, packing, porosity, permeability, etc.) and the composition of the mineral phases with which the fluids are in contact. The degree and rate of compaction are also very important. Buried waters are usually different, in composition, from the overlying sea water as has been demonstrated by numerous investigators, e.g., by Friedman et al. (1968) and Friedman and Gavish (1970). Hydrodiagenesis depends on: (1) changes in the depositional environment; (2) rate of sediment accumulation; (3) length of exposure of detritus to fluids; (4) mineralogy of the sediments; ( 5 ) differential ionic movements due to diffusion and semi-permeable effects of clays; and (6) cation exchange. All of these factors are related to compaction and compaction fluids. For example, slow rate of deposition prolongs watersediment interaction and reduces the rate of compaction. There is also a longer interaction at the sediment-water interface as well as a prolonged subsurface watersediment interaction under these conditions.
DIAGENESIS OF SANDSTONES AND COMPACTION
295
Sharma’s (1970) investigation concentrated on the glacial sediments that were transported from Alaska t o the sea. The glacial material was mechanically eroded and abraded, but underwent a minimum amount of chemical alteration (i.e., chemical weathering was absent) on land prior t o being deposited in the sea. Inasmuch as these glacial sediments are chemically the most immature, the chemical interaction between the sedimentary particles and the sea water and pore fluids would be expected t o be particularly distinct. For example, hornblende and biotite release K+ and Mg2+ ions into the fluids. Sharma (1970, p. 727) stated that the initial composition of intrastratal fluids is controlled by that of the depositional milieu, such as type of surface water, presence and degree of flocculation of clays and organic matter, and ion exchange. Alterations in both surface and shallow interstitial fluids can be seasonal in fjords, estuaries, saline lagoons, and lakes, among others. The mineralogy of sediments, proportions of sand, silt, and clay fractions, and the chemistry of the interstitial fluids would determine the types of chemical reactions. For that reason, it was so important to use very immature mineral particles in Sharma’s investigation to eliminate at least one variable, namely, chemical alterations of the solid particles as a result of surface weathering prior to erosion and transportation. Grain size and degree of compaction is significant in controlling diffusion. According to Sharma (p. 729), the diffusion coefficient of dissolved ions or molecules in water is of the order of l o v 5 cm2/sec., but when the free ions or molecules encounter sediments, their movement through the pore fluids is impeded by the sedimentary particles. (See also Peck, 1967, for discussion on diffusion coefficients.) Duursma (1966a,b), among others, showed that many heavy metals, alkaline earths, and even some alkali metals, have a great affinity for clayey calcareous sediments, so that the ionic diffusion coefficient is reduced by factors up to lo6. This indicates that diffusion over a few centimeters may take several years, but if compaction is considered, the rate of fluid movement is increased. In sediments that contain clay minerals, some factors and processes have to be considered that would not be important in pure sandstones. Clays possess the properties of semi-permeable membranes resulting in salt-sieving or reverse osmosis, e.g., when a saline fluid is passed through a layer of clay minerals, some salts remain behind. This complex process is determined by fluid chemistry, porosity and permeability, degree and rate of compaction, burial depth (i.e., overburden pressure), composition of sediments, and possibly temperature (i.e., geothermal gradient). Where studies are undertaken on near-surface sediments, as done by Sharma, the absence of significant overburden and compaction, therefore, will reduce the amount of interstitial
K.H. WOLF AND G.V. CHILINGARIAN
296
fluid movement and the membrane effects will be minimized. Slow compaction, as a result of slow sediment accumulation, would indicate that sedimentary particles were exposed longer to surface fluids and that, at the same time, compaction fluids evolved slower. In the absence of compaction, the sedimentary particles are not exposed to compaction fluids moving from below. Dickey (1972)pointed out that chemical processes associated with compaction result in mineralogical and other chemical changes of importance in both petroleum and ore studies. With increasing depths and geologic time, the pore fluids increasingly deviate from the composition of sea water: sulfate and bicarbonate anions are usually lost, chloride anion is left behind, Ca2+ cation content is increased and Mg2+ is reduced (3 to 5 times as much Ca2+ as Mg2+, whereas in sea water the Mg/Ca ratio is 3/1). Variations in fluid composition with depth, however, are very common, as exemplified by Fig. 3-181 (Schmidt, 1971). The pore fluids of shales or 0
1
I
E - I I-
n w 3000
-
XXJOI 0
I
50
I
IM)
Is0
1
200
TOTAL DISSOLVED SOLIDS, g/l
Fig. 3-181. Variations in concentration of interstitial water with depth, southwest Louisiana. The water in the pores of the shale is much less concentrated than that in the adjacent sands. Water in the sands of the high-pressure zone at Manchester is much less concentrated than that in the equivalent normal-pressure sands at Hackberry. A = water produced from sand, Hackberry field (normal pressure); B = water extracted from shale, Manchester field; C = water produced from sand, Manchester field (abnormally high pressure); D = low-density and high-pressure shale. (After Schmidt, 1971; in: Dickey, 1972,fig. 11, p. 11;courtesy Int. Geol. Congr., Montreal.)
DIAGENESIS OF SANDSTONES AND COMPACTION
so;
29 7
HCO;
Fig. 3-182. Triangular diagram showing the relative amounts of anions in interstitial waters in formations of southwest Louisiana. The waters in the sands contain almost no bicarbonate and sulfate, but these ions predominate in the shale waters. A = Hackberry sand waters (normally pressured); B = Manchester sand waters (abnormally pressured);C = normally-pressured shale water; D = abnormally-pressuredshale water. (After Schmidt, 1971, in Dickey, 1972, fig. 12, p. 12; courtesy Int. Geol. Congr., Montreal.)
muddy sediments are quite different in composition from the solutions in the associated permeable sands, as demonstrated in Fig. 3-182for one particular area: the sandstones contain mainly chloride and are poor in or devoid of sulfate and bicarbonate anions, whereas clayey sediments contain the latter two anions predominantly. The concentration of the solutions in sandstones are related to the degree of compaction of the adjacent clayey deposits, and the water samples taken from the high-pressure zones are often less concentrated in dissolved solids than those from normally pressured zones. Abnormal subsurface pressures are found only in sandstones completely enclosed in shale, with no permeable connection t o the outcrop or into adjacent porous units. According to Dickey (1972),the fluids from high-pressured zones often have salt concentrations much less than is normal for their depth of burial. Dickey suggested that the processes which caused the concentration of the interstitial solutions, were arrested at about the same stage at which the compaction of the clayey sediments was terminated. Dickey stated that on a regional scale, there is a general tendency for the concentration of subsurface fluids to increase with depth, with possible more local deviations and reversals. Von Engelhardt (1960)also commented on the general trend: drilling into sedimentary basins has shown that the salinity of pore fluids commonly increases with depth. This is easy to observe on electrical resistivity curves as exemplified in Fig. 3-183.Another example is given in Fig. 3-184,where at a
298
K.H. WOLF AND G.V. CHILINGARIAN
SP -4
-+
DEmm
RESISTIVITY
W mV
0
C
D
-I
m
i
D
2
D II
<
+
C D 0
m D -4
m II
2
t
2
C
0
D
rn D
-I
rn D
-I P D
< -
Fig. 3-183.Resistivity and SP measurements using the Schlumberger method in a borehole near Stockstadt (Upper Rhine region), Germany. The decrease of the resistivity between 300 and 400 m and the corresponding change in the SP indicate the increase of the salt content of the pore solutions. (After Von Engelhardt, 1960,fig. 59;courtesy Springer, Berlin.) Fig. 3-184.Salt content (96)as related to depth of burial of the formation fluids of the St. Genevieve Sandstone (Mississippian) in the Illinois basin (according to 235 analyses by Meents and coworkers). (After Von Engelhardt, 1960,fig. 60; courtesy Springer, Berlin.)
depth interval of 300-400 m the salt content increases rapidly with depth (Meents et al., 1956). Rieke and Chilingarian (1974),however, found so many exceptions that they are opposed to such a generalization.
DIAGENESIS OF SANDSTONES AND COMPACTION
299
Schmidt (1973)presented data on interstitial fluid composition and geochemistry in a shale-sandstone section with abnormally high pressure zones, and noted significant differences between the total dissolved solids concentrated in waters from normally pressured sandstones (600-180,000 mg/l) and those from overpressured sandstones (16,000-26,000 mg/l). He stated (p. 321): “. . . conversion from predominantly expandable t o non-expandable clays accelerates near the top of the high-pressure zone, which appears correlative with a major temperature gradient change, an increase in shale porosity (decrease in shale density), a lithology change to a massive shale, an increase in shale conductivity, an increase in fluid pressure, and a decrease in the salinity of the interstitial waters.” In this case, decrease in water salinity can be due t o release of water during conversion of montmorillonite to illite and/or compaction effects, which are described in detail by Rieke and Chilingarian (1974). Zimmerle (1963)differentiated between “destructive” and “constructive” diagenesis. In the former case, the complex minerals change to less complex ones or are replaced by other minerals, and this type of diagenesis does not lead to reduction of porosity. Constructive diagenesis refers to overgrowth or authigenic neomorphism of new, commonly simple minerals, usually in connection with chemical mobilization or addition of material. This results in a decrease in porosity and permeability, often related t o mechanical compaction, movements of compaction fluids, and pressure solution. The latter may deliver the chemical components, so that compaction, pressure solution, and cementation can be in some cases interdependent. After describing the various diagenetic features, Zimmerle offered a paragenetic interpretation based on pH of the solutions (Fig. 3-185).Solutions are the “carriers” of the chemical compounds and the dissolution and precipitation of minerals are dependent on many variables. Based on the work by Packham and Crook (1960)and others, however, Zimmerle used pH as the main parameter in Fig. 3-185. Siderite, chamosite and glauconite are of synsedimentary origin under neutral pH and medium values of Eh. Pyrite formed under reducing and neutral t o weak acidic conditions during early diagenesis. It formed prior to neomorphic quartz and feldspar, as indicated in thin sections. Kaolinite, replacing primary clays and clastic feldspar, and neomorphic brookite also formed prior to quartz precipitation under acid pH values. Secondary quartz, which is younger than brookite and older than calcite, indicates a weak acidic to neutral pH, the feldspars (i.e., albite, orthoclase, or microcline) suggest a neutral to weak basic milieu, and calcite requires a basic pH value for its precipitation. A second generation of quartz and kaolinite, apparently related to the introduction of oil into the rock, was followed by barite: It should be pointed out here that movements of petroleum are often considered to be the result of compaction of fine-grained
300
K.H. WOLF AND G.V. CHILINGARIAN PH RANGE.
SEOUENCE
F g $ #--
Syndloqmettc Early Diapcnctic
Late
Diopmcfic
I.
A
2.
m
Fig. 3-185. Schematic paragenesis of the cement and the neomorphic minerals depending on the maximum pH value in the Dogger-fl Sandstone of the Plan-Ost Oil Field. 1 = pH of the formation waters = 5.9-6.6; 2 = range of the “intrastratal solutions” for which evidence is available. (After Zimmerle, 1963, fig. 8, p. 15; courtesy Erdol Kohle.)
sediments. From the paragenesis, Zimmerle deduced that the fluids in the sandstone had a pH range of 5.9-6.6, i.e., weakly acidic, and concluded that both depth of burial and fluid composition were important in determining the sequence of mineral precipitation. During diagenesis a “reversal’,’of porosity occurs, i.e., the originally clay-poor and porous sandstones become calcite- and silica-cemented with elimination of pores, whereas the originally clay-rich, less porous sandstones are influenced only slightly by diagenetic processes and maintain most of the original minor porosity. Hence, in studies of the history of porosities, not only compaction history but also the development of cementation is of significance. Sharma (1969) presented a pH-paragenetic diagram indicating changes with depth of burial (Fig. 3-186),similar to the one offered by Zimmerle (1963). Potter (1968) showed an influence of intrastratal solution on garnet and concluded that for”ancient indurated or strongly compacted sediments the use of heavy minerals for stratigraphic-correlation and provenance studies
301
DIAGENESIS OF SANDSTONES AND COMPACTION
Fig. 3-186.Diagram showing pH relations during all stages of paragenesis in Halfway Formation as observed in cores. Order of dissolution and precipitation of minerals is also presented. (After Sharma, 1969,fig. 9;courtesy Minemliurn Deposita.)
may be of limited value, because these sediments have been corroded and removed. Inasmuch as the survival of minerals depends on a number of interrelated factors, e.g., pH of solution, nature of minerals involved, permeability and porosity of beds, temperature, pressure, rate of flow of fluids, and total geologic time, it is evident that the degree and rate of compaction and the characteristics of the compaction fluids would be influential in determining the corrosion of heavy minerals. Inasmuch as bacteria are extremely important in diagenetic alterations of sediments, it should also be pointed out here that compaction, which modifies the characteristics of the sediments, also alters the number of bacteria and, possibly, the type of bacterial assemblage, bacterial migration, and bioTABLE 3-L Distribution of bacteria in sediments according to particle size (after ZoBell, 1943,in Strakhov, vol. 2, 1969,table 41,p. 424) Sediment
Average diameter of sedimentary particle (p)
Content of N,
Water content
(%I
Number of bacteria per g of dry residue (in thousands)
Sand Silt Clay Colloids
50-1000 5-50 1-5 <1
0.09 0.19 0.37 1.0
33 56 82 98
22 78 390 1500
(%)
MOISTURE CONTENT,'/t
Fig. 3-187. Change in moisture content in sediment cores from the Bering Sea. (After
A.P. Lisitzyn, in Strakhov, 1969, fig. 181, p. 425.)
DIAGENESIS OF SANDSTONES AND COMPACTION
0
I
LOGARITHM OF NUMBER OF BACTERIA 3
5
7
303
9
E u
z I f 60 k
80
Fig. 3-188. Vertical distribution of bacteria (number of bacteria per g of wet mud) in oceanic sediments. (After ZoBell, 1946, in Strakhov, 1969, fig. 182, p. 425.)
genic processes. The factors affecting bacterial growth include: (1)grain size of the sediments (Table 3-L); (2) moisture content of the sediments (Fig. 3-187); (3) content of organic material (Table 3-L); (4) depth of burial (Fig. 3-188); (5) degree of compaction. Inasmuch as the above parameters are interrelated, compaction influences diagenesis directly by modifying the bacterial activity. The literature on the interaction of interstitial fluids and volcanic detritus is becoming increasingly voluminous, in particular on the processes related to zeolite genesis (e.g., Hay, 1966; Utada, 1968; Sheppard and Gude, 1968; Mariner and Surdam, 1970; Utada and Minato, 1971; Ijima and Utada, 1971; Sheppard, 1971; Surdam and Parker, 1972). Specific references to compaction fluids within pyroclastic units and their influence on diagenesis and burial metamorphism, however, are less frequent, probably because the results of compaction-fluid reactions are difficult to nearly impossible to distinguish at the present time from the effects of reactions of any other type of intrastratal solutions. Nevertheless, occasionally early researchers have considered compaction in pyroclastics, and recent work on the compaction of volcanic debris has been treated in Chapter 6 of this book. As early as 1954, Coombs observed in the Triassic succession of New Zealand fairly normal sedimentary diagenetic phenomena, such as overgrowths on feldspars, cementation by quartz and chloritic minerals, and zeolite replacement of volcanic glass. He stated that in addition to connate waters trapped in the sedimentary pile, large quantities of fluid were stored
K.H. WOLF AND G.V. CHILINGARIAN
304
up in the volcanic glass and in the zeolites of early origin. As the temperature during increasing burial rose, perhaps up to 150-300°C at the base of the pile, this stored-up water facilitated a special type of metamorphism with some low-grade hydrothermal effects. Dickinson (1962a,b) discussed diagenetic to epigenetic alterations in mineralogy and chemistry of pyroclastic material and concluded that the metasomatic processes causing zeolitization, phyllosilicatization, and albitization were not initiated by fluids from an igneous source. Instead, the nearest
IN I T l A L
FINAL Water flows through membrane
(A1
semi-permeable membmne transmits water but not salt m re less equal concentrations Concentrated
equal concentrations
concentrated
-
Water flows upward through shale membrane
(C
by hydrostatic head
Fig. 3-189. Initial and final equilibrium states 01 normal (A) and reverse osmosis (B) in a U-tube. C. Reverse osmosis in two sandstones separated by a thin shale acting as a semi-permeable membrane. (After Pettijohn et al., 1972, fig. 10-9, p. 413; courtesy Springer, New York.)
DIAGENESIS OF SANDSTONES AND COMPACTION
01
00
I
80
I
1
1
60
40
20
305
OIL SATURATION,%
Fig. 3-190. Capillary pressure curves (displacement of water by oil): I = coarse-grained sandstone; II = clayey fine-grained sandstone. (After Von Engelhardt, 1960, fig. 34; courtesy Springer, Berlin.)
source of reactive fluids is formed by the connate pore waters and adsorbed aqueous films that must have been expressed from associated and underlying marine mudstones by compaction. Hay and Ijiama (1968) briefly considered compaction by remarking that the tuffs of Hawaii show no compaction features and that the net increase in bulk density observed by them from younger to older tuffs, seems to be a reflection of a net increase in constituents precipitated by inter-pore fluids. Pettijohn et al. (1972, p. 414) discussed a special process, i.e., reverse osmosis, that must also be operative during compaction of sediments. This process that affects the composition of subsurface waters by “salt filtering through a semi-permeable membrane” depends on the content of clayey sediments (see De Sitter, 1947, Bredehoeft et al., 1963, and Rieke and Chilingarian, 1974). The differential pressure at depth within the sedimentary pile supplies the energy. Figure 3-189,a illustrates that in simple osmosis the difference in concentration of salt in fluids on the two sides of the membrane results in a pressure differential (= osmotic pressure), which causes the movement of the solvent from the low-concentration to the highconcentration side. In “reverse osmosis”, the same concentration of salt exists on both sides of the membrane, but a pressure differential is formed that induces a flow of the solvent from the high- to the low-pressure side and results in a higher salt concentration on the high-pressure side (Fig. 3-189,b). According to Pettijohn et al. (1972), applying these principles to a sedimentary environment, the pressure differential with varying depth of overburden
K.H. WOLF AND G.V. CHILINGARIAN
306
will lead to a situation as shown in Fig. 3-189,c. Here, a sandstone below a shale acting as a semi-permeable membrane will have a higher concentration of salts than a sandstone above that shale. It is difficult t o prove at the present stage of knowledge, however, to what degree “salt filtration through a semi-permeable membrane” process is actually occurring in sedimentary basins. A number of important fields of study related t o diagenesis and fluid movements should be given more attention in the future by sedimentary petrologists. These include: (1) capillary pressures and their relationships to pore geometry (see Chilingarian et al., 1972); (2) relative permeabilities of sediments to oil, gas, and solutions of different compositions at varying temperatures and pressures (see Langnes et al., 1972); and (3) relative degrees of adsorption of fluids of different compositions on grain surfaces. Figure 3-190 shows capillary pressure curves for a coarse-grained sandstone and for a clayey, fine-grained sandstone, applied to conditions required for water t o replace oil. Figure 3-191 indicates the relative permeabilities to oil and to water, where water is the wetting agent and oil is the non-wetting phase. The effects of interface phenomena on diagenesis will have to be considered thoroughly in future research, because these phenomena affect the microscopic and sub-microscopic processes during diagenesis in coarsegrained sediments. Several examples have been discussed in this chapter which demonstrate that the presence of a film of oil on sand grains suppress
0
20
40
80
60
WATER SATURATION,
100
o ?‘
Fig. 3-191. Relative permeabilities of a rock to oil (kro) and to water (kr,,,). Water is the wetting phase, whereas oil is the non-adsorbed phase. (After Von Engelhardt, 1960, fig. 36; courtesy Springer. Berlin.)
DIAGENESIS OF SANDSTONES AND COMPACTION
307
-
#
20-
c
UJ
d w
s I-
n W
m
IY
0
m
n l0-
y.'
' .
0
.
INTERNAL SURFACE AREA.S/E..IO'
. 0.mcm a
Fig. 3-192. Amount of adsorbed water (Sh, on surfaces of the pores) and internal surface of 7 5 samples of the Bentheimer Sandstone (Valendis; Scheerhorn Oil Field), area (S/€) Germany. (After Von Engelhardt, 1960, fig. 37; courtesy Springer, Berlin.) Fig. 3-193.a. Curvature of the wetting phase (ring-like) at the point of contact between two spheres (schematic). (After Von Engelhardt, 1960, fig. 38; courtesy Springer, Berlin.) b. Calculated profiles of the drop at the contact of two spheres considering different relationships between the capillary pressure ( P c ) and interfacial tension (0).(After Von Engelhardt, 1960, fig. 39; courtesy Springer, Berlin.)
the chemical interaction of the grains with solutions. The type of distribution of adsorbed water in pore spaces has been rarely observed directly, so that most conclusions are based on indirect considerations. It has been theoretically determined that there is a linear relationship between the surface area and the amount of adsorbed water, which can be substantiated by measurements on natural rocks, e.g., see Fig. 3-192. Such
K.H. WOLF AND G.V. CHILINGARIAN
308
data is of value in diffusion studies of chemical elements during diagenesis. Figure 3-193shows the relationship between capillary pressure ( P c ) ,surface tension ( o ) , and the curvature of adsorbed water rings at the contacts between two grains. More precise work on the relationships of capillary water, its movement patterns, pressures required to move it, shapes of the water films, diffusion, and related factors t o precipitation of minerals in the interstitial spaces is needed. It has already been suggested, for instance, that cement occurring in the textural relationship to its host grains as does the adsorbed water in Fig. 3-193,is the result of precipitation from such adsorbed solutions. The amount of fluid available from one such adsorbed ring, however, would be insufficient to account for the same amount of cement; therefore, it must be assumed that additional volumes of chemical elements were supplied by diffusion or migration of solutions. Magara (1974) concluded that fluid movements from shales into associated sands are the result of both compaction and osmosis, as schematically illustrated in Fig. 3-194.During compaction, the water from the clayey sediments (now shales) will be squeezed into the more permeable sands, as the clay-rich deposits will decrease in porosity accordingly. Maximum reduction in porosity or fluid expulsion will occur in the clayey sediments directly above and below the sands, whereas the porosity near the center of the shale will remain relatively high, but subsequently will be affected (reduced) with increasing compaction. The fluid pressure as a result of compaction can be A I ALE POROSJM
8
FLUID PRESSURE IN WALE
C WATER SALINITY IN SHALE
SHALE
SHALE
SHALE
-.#gfifiw.s* ............. SHALE
owm
.:::::.z.k.:.:.:.: +d .....:L :: Di"" ...............
1 ................
.:.:.:.>>w:p$:::::: ::::::::*.A+:.:.:.:.:.
T
LHYDROSTATIC PRESSURE Fig. 3-194. Schematic diagram showing shale-porosity, fluid pressure, and pore-water salinity distributions in interbedded sand-hale sequence. (After Magara, 1974, fig. 6, p. 288; courtesy Am. Assoc. Pet. Geologists,)
DIAGENESIS OF SANDSTONES AND COMPACTION
309
calculated from the shale-porosity distribution, as done by Magara (1968, 1969). Figure 3-194,B shows schematically the fluid pressure plotted corresponding to the porosity distribution. As expected, the water in the clayey deposit will move from a zone of higher, excessive pressure to a lower-pressure zone. Compaction-water movements are indicated by arrows in Fig. 3-194,B. As a result of ion-filtration, the ions are concentrated in the clayey unit, as schematically illustrated in Fig. 3-194,C. Salinity is the reciprocal of porosity, in that it increases as the porosity decreases. Thus, the salinity in the shales increases toward the sands. The process of osmosis induces water to move from a fresher to a more concentrated locality, as indicated in Fig. 3-194,C, but the osmotic pressure difference in this case is not very pronounced in contrast to that due to compaction. The flow of solution due to the combination of compaction and osmosis will continue until the clay-rich units reach equilibrium, and no fluids can be expelled from them by compaction. Salinity also may reach equilibrium. If, on the other hand, a “freshening mechanism”, such as dehydration of montmorillonite, changes the salinity later on, the osmotic fluid may be changed. Magara (1969, p. 289) stated, however, that the most important in this combined mechanism of flow is that the salinity contrast resulting from ion-filtration starts to appear during early compaction. Consequently, the resultant osmotic pressure difference seems to support fluid migration from the clayey sediments at the early stages of water expulsion. In their thorough discussions on the effect of compaction on salinity, Rieke and Chilingarian (1974) reached different conclusions, e.g.: (1)salinity of solutions in undercompacted shales (higher porosity) should be higher than those in well-compacted ones, providing all other variables are kept constant (p. 25); (2) compaction fluids increase in salinity upon upward migration in a thick shale sequence (p. 274); and (3) ion-filtration does not become significant until overburden pressure reaches about 10,000 psi (p. 238). REGIONAL FEATURES OF COMPACTION
Regional studies of diagenesis must consider all mechanical and chemical effects of compaction, in addition to all other detailed petrologic variations both vertically and horizontally. Numerous case histories are available from the literature of which some are presented below. Various authors showed that many of the diagenetic features are directly or indirectly related to
310
K.H. WOLF AND G.V. CHILINGARIAN
depth of burial or pressure, as well as to temperature, geologic age, fluid composition, variations with time, and others, although not always will one find references made to compaction and compaction fluids. From the published literature and from theoretical considerations, the authors have prepared Tables 3-LI to 3-LIII listing the variables to be considered in regional Compaction studies that may be carried out in the future. Only some of the parameters were considered in the earlier investigations. As in most studies, not all factors can be given the same emphasis for economic reasons and because of numerous other limitations, but a check list, such as the one given here, will assist in choosing those of greatest significance. In presenting the summaries of case histories referring to compaction, it was found best to give them in the order based on the data of publication without losing coherence and continuity. Much of the information was obtained by petroleum geologists during investigations of basin evolution, migration of fluids related to oil accumulation, and related problems. It should be pointed out again that eventually the techniques and concepts developed TABLE 3-LI Factors to be considered in regional studies of compaction (1) Regional facies distribution of coarse elastics, pyroclastics, shales, coal and other organic deposits, and evaporites (very little quantitative information on compressibilities is available on mixtures of lithologies, e.g., (a) sand plus various proportions of clay; (b) sand plus various proportions of clay + silt; and (c) pebbles plus various proportions of sand, silt and clay) ( 2 ) Stratigraphy of all sedimentary deposits (3) Paleotopography, e.g., regional dip variations (4) Paleoenvironments of deposition, which, for example, will control the type of interstitial fluids and diagenesis (and most primary features as listed below) (5) Detailed compositional studies (i.e., granulometric and mineralogic) (6)Detailed textural-fabric and paragenetic studies (e.g., matrix-cement-grain proportions): (a) primary features; (b) secondary features - (i) compactional features, (ii) all other features (7) Detailed mass-property studies: (a) density-porositydepth of burial interrelationships; (b) porosity-permeability-granulometric composition interrelationships (controlling factors: (i) primary features, such as packing, lithology, sedimentary structures, and paleotopography, (ii) secondary features, such as cementation, compaction, leachingsolution, and neoformation of minerals by various processes) (8) Fluid distribution in sedimentary basins (9) Temperature distribution in sedimentary basins (10) Vertical and horizontal variations of all variables listed above (11) Laboratory compaction studies of sediment types encountered in the basin, using uniaxial, hydrostatic, and triaxial compaction apparatuses (12) Comparative investigations between experimental laboratory and natural observations
DIAGENESIS OF SANDSTONES AND COMPACTION
31 1
TABLE 3-LII Texture-fabric indicators of compaction (1)Deformation of matrix (if present), e.g., bending of micaceous grains (2) Interpenetration of grains (3) Type of grain contacts (4)Secondary changes in cementation (if present) (5) Textures and fabrics that might indicate direction of fluid migration (this type of studies has been done on limestones and dolomites with large pores filled with secondary material, but not in the case of sandstones and conglomerates; research and laboratory experiments are needed in this area) (6)Solution features, e.g., stylolites (7) Fracturing of grains, e.g., glass shards in pyroclastics ( 8 ) Conversion of tuffs to zeolitic sediments and bentonite, accompanied by textural changes
by the petroleum geologists can also be applied to the study of ore genesis within sedimentary and volcanic piles. Study by Fiichtbauer fichtbauer (1961) used diagenetic changes on the surfaces of quartz grains to reconstruct the history of oil genesis. He found that the quartz overgrowths in the sandstones of oil fields are more common in the peripheral areas of the oil fields than in the oil-impregnated sections. This indicates that aqueous fluids are required to cause precipitation of silica, whereas oil prevents or interrupts this form of diagenesis. Lowry (1956) had pointed out that the volume of water adsorbed on the surfaces of grains is not sufficient to promote and/or continue silica diagenesis, so that movements of aqueous TABLE 3-LIII Mineralogic changes due to progressive diagenesis-burial metamorphism (1)Clay-mineral changes ( 2 ) Devitrification (e.g., volcanic glass, chalcedony, organic opal) (3) Recrystallization ( 4 ) Changes in mineralogy of coarser detritus, e.g., glauconite to limonite as a result of oxidation; feldspar to zeolite; overgrowths on feldspars and quartz, and corrosion and dissolution of ferromagnesian minerals (5) Changes in mineralogy and paragenetic relations indicating variations in pH, Eh, solubility; temperature, and pressure with geologic time: (a) pH indicator: quartz vs. calcite; (b) Eh indicator: Fe2+ vs. Fe3+ contents; (c) temperature and pressure indicators: zeolite minerals and certain>ypes of phyllosilicates (6) “Crystallinity” of illite (see p. 418 for discussion)
K.H. WOLF AND G.V. CHILINGARIAN
312
fluids through the rock are required. In the basin studied by Fiichtbauer, the sandstones exhibited a range of 1-90% quartz overgrowths. He listed the variables that are responsible for the degree or extent of neomorphism of quartz overgrowths, namely, (a) overburden pressure; (b) history of subsidence and tectonic pressure; (c) time available for diagenesis or age of the rock; (d) chemistry of pore fluids and its changes during diagenesis; (e) interference by carbonate cementation and clay content; and (f) influence of oil migration. The “energy” for silica neomorphism is probably supplied by compaction: (a) mechanical compaction, i.e., turning and rearrangement of grains, which leads to tighter packing; and (b) chemical compaction, i.e., dissolution at grain contacts and deposition at points of smaller hydrostatic pressures. Fiichtbauer (1961,p. 169)reasoned that when one studies individual sedimentary units or reservoir rocks within an oil field, i.e., specimens that were exposed originally to the same pore fluids and have the same age, the same overburden pressure, and the same subsidence history, then the systematic variations in quartz overgrowths of carbonate-cemented and clay-poor sandstones must be a reflection of oil migration. Inasmuch as the presence of oil prevents diagenesis, one must investigate stratigraphic sections that remained unaffected by oil movements to obtain some idea of the type and degree of secondary changes where interruption was at a minimum. Some of the results obtained by fichtbauer are outlined as follows: (1)As to carbonate cementation, Fig. 3-195illustrates that the early diaOO /
OIO
CARBONATE OUARTZ OVERGROWTH MEDIAN
A SP
RESISTIVITY
Fig. 3-195. Restriction of the secondary quartz overgrowths to two dolomitic sandstone lenses in contrast to the remaining water-filled sandstones, borehole Bokel 8, Dogger-fl sandstone “01 ”. (After Fiichtbauer, 1961, fig. 2, p. 170; courtesy Erdol Kohle.)
DIAGENESIS OF SANDSTONES AND COMPACTION YO
OUARTZ OVERGAOWTH
SP
10mV
YO
RESlSTlVlTY QUARTZ MRGROWTH
31 3
SP
RESISTIVITY
210 m
2x
t ”v’ _ p u 29
CARBONATE
KT Nuclei ramoininp withoil
2K m
N0.7
m
No.15
Fig. 3-196. Filtration effect (see text) and coincidence of the present-day with the previous water boundary (= “jump” in diagenesis in the borehole Vorhop-Knesebeck 15). Boreholes Vorhop-Knesebeck 7 and 15; Dogger-0 Sandstone “Ol ” (in both boreholes with water) and “U” (in borehole 7 with water; in the upper part of borehole 1 5 with oil). (After Fiichtbauer, 1961, fig. 3, p. 171; courtesy Erdol Kohle.)
genetic, concretionary dolomite precipitation in the Lias and Dogger sandstones eliminated porosity and prevented the penetration of aqueous solutions. In contrast, the carbonate-free sandstones show quartz overgrowths of up to 70% SOz, whereas in the former instance it was only up to 5%. Similar relationships are exhibited in Fig. 3-196(left side). In contrast, the data presented in Fig. 3-197indicate much earlier carbonate precipitation. (2) To demonstrate a so-called “filtration effect”, i.e., the effect of upward-moving compaction solutions trapped underneath clay-rich layers, Fiichtbauer (1961,p. 170)presented the following discussion. In Fig. 3-196, low at 01,the sandstone shows a 20% quartz overgrowth, whereas in the upper parts of this sandstone it approaches 40%. A similar effect is shown in Fig. 3-196 (right) in the oil-impregnated part. The relationship has been observed in many other cases where the uppermost 1 to 2 m of a sandstone has a higher percentage of silica overgrowths. The parallel increase in staurolite content in Fig. 3-197suggests that in this particular unit no diagenetic corrosion or dissolution has occurred and that silica was precipitated. The silica must have come from an outside source, probablf from upward moving
K.H. WOLF AND G.V. CHILINGARIAN
314
@
SP
@ O/~OUARTZ OVERGROWTH
V~CAR~ONATE
@ yoSTAUROLITE
Fig. 3-197. Tectonic change (inclination) of a structure after oil invasion. (The depths in the diagram are related to the earth’s surface in the structure section NN; also see text for additional explanation.) Dagger$ Oil Field Wesendorf-South, Germany. Numbers refer to well numbers. (After Fiichtbauer, 1961, fig. 4, p. 171; courtesy ErdOJ Kohle.)
compaction fluids derived from deeper clayey units and, possibly, older sandstones undergoing pressure solution. The upward-moving solutions precipitated the silica beneath the impermeable clayey units. Similar mechanisms can be offered to explain carbonate precipitates underneath shales or mudstones. According to Fuchtbauer (1961, p. 170),it is certain that this “filtration effect” was operative prior to oil migration, or was at least penecontemporaneous, because later the remaining water was present in the sandstones only as adsorbed films. An independent explanation also confirms that this phenomenon is an early diagenetic one, i.e., the more pronounced it is, the more intense compaction was or the more compaction fluids were available. Theoretically, the greatest degree of diagenesis as a result of the filtration effect should be present at the upper sections of a thick unit composed of sediments that accumulated quickly and with little or no interruption. With increasing overburden, i.e., with increasing depth in the basin, the filtration effect decreases. Filtration effect is at a minimum in sediments that transgress over older, already compacted or consolidated sediments, because the fluids from the latter have already been pressed out. The effects of filtrating fluids are best observable in deposits which became oil impreg-
DIAGENESIS OF SANDSTONES AND COMPACTION
31 5
nated soon after the formation of early diagenetic features. It should be pointed out, however, that in older sandstones that became water-filled after the filtrating fluids led to quartz diagenesis, the youngest diagenetic precipitation may also have formed quartz overgrowths and thus obliterated all previous diagenetic features. As to the diagenetic differences related to oil migration, Fuchtbauer (pp. 170 and 172) discussed (1)differences within one particular field and (2) differences between oil fields within one large petroleum province. In the case of differences within one oil field, the boundary of the diagenetic effects coincides with the present oil-water boundary. An example of this is shown in Fig. 3-196,where the oil-containing unit has 15-3076 quartz overgrowths, whereas the water-containing part has 40-60% of quartz overgrowths, because diagenesis remained uninterrupted in the latter instance. In between is a transitional unit with specks of oil preserved in it. In addition, the boundary separating the units having higher degrees of diagenesis from those of lower degree does not coincide with the present oil-water contact. Three explanations are possible: (a) oil migration occurred in several steps; (b) the “oil cap” or oil zone was expanded through compaction as a result of reduction in thickness of the oil-bearing unit; (c) the geologic structure was tectonically altered after oil has moved in. Case (a) can be distinguished from the others only very seldom. Early oil migration is probable when a very sharp diagenetic boundary is present within the oil-impregnated unit, but is absent at the present time at the oil-water contact zone. Case (b) is probably applicable in the case presented in Fig. 3-198,where there is an increase in quartz diagenesis towards the bottom. The extension or spreading of the oil cap as a result of compaction is applicable to fields with early oil invasion and impregnation according to the following calculation. Underneath an area of 1m20f the Eldingen Sandstone there were at the time of oil movement about 17.5 m3 of pore space (assuming 50 m thickness and 35% porosity). Since then, the porosity was reduced from 35% to 27%. At the same time, the thickness was reduced at the expense of the pore space by about 5.5 m, from 50 to 44.5 m. Consequently, beneath each 1m 2 of area, the pore space was reduced from 17.5 m 3 to 12 m3. The oil zone must have expanded almost by 50%, therefore. Wherever the sudden change in degree of diagenesis in individual boreholes in an oil field is found at different depths, one has to assume that the structure underwent modification through tilting or disturbances after the oil had migrated into the reservoir rock (case c). An example of this is to be found in Fig. 3-197.The map shows the location of the boreholes 2, 4,5,and 8;borehole 3a is situated far beyond the map. The change in degree of diagenesis in the boreholes 4, 5, and 8 is indicated by diagonally lined areas on the right-hand side. The areas on the left (under the SP curves) indicate presentday position of the bound-
K.H. WOLF AND G.V. CHILINGARIAN
316
t
I
0 2 0 4 0 8 0 1 0 1 0 0
OUARTZ OVERGROWTH,%
Fig. 3-198. Spreading of an oil top through compaction of the sandstone. Borehole Eidingen 55, Liasiu-2 Sandstone. (After Fuchtbauer, 1961, fig. 5, p. 171; courtesy Erdol KOhle. )
ary between oil and water. To reconstruct the form of the structure at the time of silica diagenesis, one has to move the boreholes vertically so that the points at which there is a change in degree of diagenesis form a horizontal line. The distances from the upper surface of the formation to this surface then enables one to reconstruct the former structure. These distances (highs) in the Wesendorf boreholes are as follows: No. 8 = 9.5 m, No. 5 = 8.5 m, No. 4 = 7.1 m, No. 2 < 0 m, and No. 3a < 0 m. One can observe that although the former structure was similar, it was distinctly flatter than the present-day one. Whereas today the borehole No.8 lies 15.7 m above the borehole No. 4, the original structural height difference was only 2.4 m. Because of the high diagenetic quartz content in boreholes 2 and 3a, the sandstones were probably never filled with oil here. In examining the differences between different oil fields within a particular oil province, Fiichtbauer (1972,p. 172) concluded that if two oil fields in an area showdifferences in degree of diagenesis, the oil migration or invasion is older in one case than in the other. An example of this is given in Fig. 3-199,where the Vorhop structure is older than the Hankensbuttel as reflected by the differences in degree of silica diagenesis. Pressure solutionquartz cementation due to the overburden pressure was interrupted in one case by the invasion of petroleum, whereas in the other case the aqueous solutions, which were kept saturated with silica as a result of the compaction processes, caused more extensive silicification. Two examples of the effects of the silica diagenesis on porosity and permeability of the sandstones are presented in Figs. 3-200 and 3-201. In the upper diagrams, the porosity was plotted versus the degree of quartz diagenesis. The porosity magnitude depends on the history of the reservoir rock. The porosity does not show a definite relationship to degree of diagenesis: up t o about 50% of quartz overgrowths, the porosity changes little or not at
DIAGENESIS OF SANDSTONES AND COMPACTION JEDIAN
RESISTIVITY
317
RESISTIVITY
n
&’ 4/20
I I
01
Fig. 3-199.Oil invasions of different ages (older to the left). Boreholes Vorhop 25 (left) and Hankensbuttel M 1 (right). Dogger-fl Sandstone ‘‘01” and “U1”.(After Fuchtbauer, 1961, fig. 6, p. 172;courtesy Erdol Kohle.) At the lower left side, the circle indicates a zone totally impregnated with oil; resistivity curves of left portion of the graph =laterolog 10/100/1,000; resistivity curves a t the left side of figure = ESkl Normal.
all, and only when quartz overgrowth exceeds that value does porosity decrease distinctly (see especially Fig. 3-201).The sediments became compacted down to a porosity of 27-29% without quartz overgrowth development, even if they were oil impregnated. It seems then that quartz cannot dissolve at pressure points, because the silica cannot dissolve in oil. If silica dissolves in the adsorbed water on the quartz grains, it would precipitate in the immediate vicinity. Pressure-solution phenomenon is also unimportant during compaction down to 27-29% porosity, in the case of water-saturated sandstones, because the specimens that have more than 27-295’6 porosity indicate that quartz diagenesis has no effect on the porosity. One would expect an influence, however, in case of pressure-solution occurrence and where the centers of the grains moved closer to each other resulting in pore space reduction. In Figs. 3-200and 3-201,the bottom diagrams illustrate the relationships between permeability and degree of silica diagenesis. Here also, one can observe a-definite intenelationship only when the amount of quartz overgrowths exceeds 50%.
K.H. WOLF AND G.V. CHILINGARIAN
318
0
-
50
I00 0
QUARTZ OVERGROWTH,%
Fig. 3-200
-
50
IW
Fig.3 -201
Fig. 3-200.Dependency of porosity (above) and permeability (below) on the amount of secondary quartz growth. Median diameter of the sample = 0.10-0.16 mm. Eldingen Oil Field, Lias-a-2 Sandstone (-1450-1620 m below the surface). (After Fuchtbauer, 1961, fig. 7,p. 172;courtesy Erdol Kohle.) Fig. 3-201.Dependency of porosity (above) and permeability (below) on the amount of secondary quartz growth. Median diameter of the sample = 0.15-0.24 mm. The low carbonate content was here added to the pore space (see text). Wesendorf-South Oil Field, Dogger-/3 Sandstone (-1480 to 1540 m below the surface). (After Fuchtbauer, 1961, fig. 8,p. 172;courtesy Erdol Kohle.)
Study by Philipp and others Philipp et al. (1963)discussed several new methods in the interpretation of the history of oil migration in their study of the Gifhorn Basin, Germany, containing Middle Jurassic sediments. The results of their investigation can be summarized as follows: (1)A detailed analysis of the structural history, i.e., the trough subsidence and uplift, in space and time yielded information on the age of the reservoir rock development (Fig. 3-202);The thickness of the subsequently-eroded sediments was extrapolated from t6e regional isopach maps (hatched lines, Fig. 3-202). (2)The use of porosity of the shales as a maximum depth indicator is shown in Figs. 3-203and 3-204.A master diagram was employed first considering samples which have never been buried deeper than the present-day
3Im+9MIl m f t
Fig. 3-202. Relationship among age, depth of burial, and accumulation of sediments in the Gifhorn Basin, Germany, based on quartz diagenesis. Location: Liiben to the left and Meerdorf to the right. (After Philipp et al., 1963, fig. 2; courtesy 6th World Pet. Congr., Franfurt/Main.) (See Fig. 3-205.)
K.H. WOLF AND G.V. CHILINGARIAN
320 KAOLIN/CHLORITE RATIO
1111111
INTERSECTION VELOCITY OF CLAYS, m/scc (Liorsic ondDoqger)
‘ “ 1
-
0
POROSITY, %
QUARTZ CONTENT,%
POROSITY, Y.
Fig. 3-203.Variation in diagenetic alterations with increasing depth of burial. Kaolinite/ chlorite ratio has been calculated quantitatively using X-ray diffractograms, whereas the quartz content was estimated from the diffractograms. (After Philipp et al., 1963,fig. 3; courtesy 6th World Pet. Congr., Frankfurt/Main.) Fig. 3-204.Relationship among depth of burial, clay porosity (short vertical dashes), porosity of Tertiary shales, and intersection velocity obtained from sonic logs. The maximum depth of burial (shown by arrows) o r the amount of uplift in Calberlah and Dannenbiittel, Germany, has been determined from the clay porosity and intersection velocity, respectively. The porosity of Tertiary shales is shown by a solid curve. (After Philipp et al., 1963,fig. 4;courtesy 6th World Pet. Congr., Frankfurt/Main.)
depth (Fig. 3-204).X-ray analyses of the shales may support the findings on burial depth, because the diagenetic changes are controlled by the maximum depth of burial as was the porosity, i.e., the kaolinite/chlorite ratios decrease with increasing depth, whereas the quartz content increases (Fig. 3-203). (3)Sonic-log readings were also used as maximum depth indicators (Fig. 3-204),i.e., the interval velocities were employed instead of the porosity values. This method is less time-consuming than porosity determination of the shales and does not require core sampling; however, the exact petrology must be known for proper interpretation of sonic logs. (4) The methods mentioned above are helpful in evaluating the subsidence history of a basin, whereas quartz diagenesis (Fig. 3-205)allows dating of the oil invasion into the reservoir rocks. Although gradual precipitation of quartz cement in clean quartz sandstones is produced mainly by pressure solution without an outside silica supply, some silica precipitated as a result of solubility reduction caused by an increase in salinity of the pore fluids with increasing geologic age and depth of burial. The main controlling variables
DIAGENESIS OF SANDSTONES AND COMPACTION I
M
I
20
LO
30 O/o
50
321 60
quartz grains with overgrowths
W!
11
; I
.25m
m
E
r.ndrtonrr oil-fillrd o w8trr-fillrd A water-fillrd but previously oil-fillrd
Fig. 3-205.Relationship between maximum depth of burial and degree of quartz diagenesis. (After Philipp et al., 1963,fig. 5;courtesy 6th World Pet. Congr., Frankfurt/Main.) The dots are plotted at the maximum depth of burial before oil accumulation; the circles and triangles represent the maximum depth of burial ever reached. Interrupted circles are based o n previous geological interpretations (cf. text). Abbreviations in Figs. 3-205, 3-207 and 3-208:Bk = Bokel; Bo = Bodenteich; Br = Broitzem 4; Ca = Calberlah; D a = Dannenbuttel; Es = Essenrode; G N = Gifhorn-N; GrO = Gross-Oesingen 2 ; Ha = Hankensbuttel-Mitte (in Fig. 3-207 also -N and -0); Ha-Oil = Hankensbuttel-S; Har = Hardesse; HW Hankensbuttel-West; Hz = pit near Bad Harzbuq; II = Ilkerbruch; Lii = Luben; LW = Luben-West; Me = Meerdorf; Oh = Ohrdorf 2; 0 s = Orrel-S 1001; Rii-hi = Ruhme, structural high; Ru-lo = Ruhme, structural low; Ru-HzO = Ruhme, water-filled reservoirs; RU, cse = Ruhme, coarse-grained samples; RU, fin = Ruhme, fine-grained samples; S t = Steinkamp; Th = Thurau 1; Vk = Vorhop-Knesebeck; VN = Vorhop-Nord; Vo = Vorhop; VoH = Vorhop H 1; We = Wesendorf; WN = Wesendorf-Nord; WS = Wesendorf-South; WiS = Wittingen-South 1. The well number appears after the abbreviations, whereas the number of samples is shown below them.
are pressure due t o the overburden and length of burial time. As long as silicification is not too extensive, its degree can be measured very easily under a microscope in transmitted light; however, as more grains merge due to quartz cementation, fewer quartz facets are observable. Philipp et al. (1963,p. 461) defined quartz diagenesis as the percentage of: (1)quartz grains with more than 50% of their surfaces being covered with euhedrd '
K.H. WOLF AND G.V. CHILINGARIAN
322
overgrowths, plus (2)half of the quartz grains with minor silica precipitation. In Fig. 3-205,the open circles represent water-saturated sandstones which have never been oil-saturated during the geologic history. Their degree of quartz cementation has been plotted against the maximum burial depth, which is known in many cases. The distribution of these circles indicates that there is a relationship between quartz diagenesis and depth of burial. The full circles or dots represent oil fields which apparently were formed as a result of early oil migration. To adjust the data, these full circles were plotted versus the maximum depth of burial prior to oil accumulation, as indicated by the quartz diagenesis which has been interrupted by the oil invasion - an assumption proved by the information presented in Fig. 3-205. Based on these well-established test cases, which allow an extrapolation to other similar geologic situations, one can reconstruct more uncertain cases by means of studying quartz diagenesis. The solid triangles inside in Fig. 3-205have a low degree of quartz precipitation, suggesting that the present water-filled sandstones have been oil-filled at an earlier stage. Other geological evidence has confirmed this in some instances. Where the geological evidence for maximum depth of burial was uncertain (vertical lines give a probable range), the solid triangles represent the most likely depth. In stratigraphic traps formed by regression-transgression (e.g., Luben; Fig. 3-195),the quartz precipitation was controlled by the maximum depth of burial prior to the time of transgression. If the oil migration occurred during renewed subsidence after the transgression, but before the rocks reached the previous maximum depth of burial, quartz diagenesis cannot supply information on the time of oil invasion. The broken circles represent the oil pools for which the depth at the time of oil invasion was estimated by using geologic evidence only. The study of quartz diagenesis, however, indicated later oil invasion (full circles or dots) as pointed out by the arrows associated with the broken circles and supported by more detailed geological interpretations. In regard to the scattering of the points, Philipp et al. (1963,p. 463) attributed this t o the different length of burial time. For example, at the Wesendorf-S oil field the maximum depth of burial (1800 m) in the watersaturated wells (WSin Fig. 3-205)is similar to the maximum depth (1500m) before movement of the oil in the oil-saturated WS-wells owing to the late origin of the pool. The water-saturated samples, however, contain 58% quartz grains with overgrowths, whereas in the case of the oil-saturated samples only 34% of the grains have quartz overgrowths. This can be explained by the fact that the water-filled sandstones were submerged to a depth of more than 1200 m for 85 lo6 years until the oil accumulated, whereas the oil-filled ones were buried only for 30 lo6 years. Hence, the different burial time gave rise to different degree of silicification at the same burial depth. Figure 3-205 is based mainly on the wells with the water-
-
-
DIAGENESIS OF SANDSTONES AND COMPACTION
323
.:...D ogger p- sandstones Meerdorf L
I-LL v
20 carbonatemok:-
Fig. 3-206. Influence of carbonate and clay contents on quartz diagenesis. The samples of the upper diagram are poor in carbonate content; the samples of the lower diagram are poor in clay. (After Philipp et al., 1963, fig. 6; courtesy 6th World Pet. Congr., Frankfurtlhlain.)
saturated and the oil-saturated cores, corresponding t o the highest and lowest degrees of diagenesis, respectively. It should be realized that the presence of early carbonate cementation, as well as clay neoformation, could hinder quartz precipitation (Fig. 3-206). (5)Philipp et al. (1963,p. 464) also investigated the intrastratal solution of heavy minerals. Kyanite, staurolite and garnet have been destroyed and, in several instances, a change of kyanite into mica and of staurolite into quartz has been noticed. Figure 3-207 shows that the content of unstable heavy minerals decreases as the degree of quartz diagenesis increases (black symbols). Assuming a uniform chemical milieu, the chemistry of intrastratal solution depends on the maximum burial depth, which, as discussed previously, also affects quartz precipitation. Near the upper and lower sandstone-shale contacts, however, the amount of quartz cement is higher as a result of greater precipitation from compaction fluids, so that here the heavy minerals were not able to be corroded and/or removed (symbols in brackets in Fig. 3-207).It seems that when fluids moving through thick clayey sequences increase in salinity and the salts from these supersaturated compaction fluids may precipitate out immediately upon entering the sandstone. The white or open symbols in the lower part of the diagram represent samples close to the surface, where intrastratal dissolution of the heavy
I
3 60%
I
I
%unst&
-
DSt20 1 DW*rSb 1
OVB
b0,h
D Hz 0I
D
E
4S 1I'0
h.m.
bJ1 25 20 1
matab havY 30% minrmlr
Fig. 3-207. Intrastratal solution of heavy minerals and quartz diagenesis, The percentages of unstable heavy minerals (kyanite + staurolite + garnet) are means of the fractions <0.06, 0.06-0.09, and BO.09 mm. (Signs, see text; for abbreviations, see Fig. 3-205.) (After Philipp et al., 1963. fig. 7; courtesy 6th World Pet. Congr., Frankfurt/Main.)
minerals obviously took place without occurrence of quartz diagenesis. The coarser-grained unstable heavy minerals are the original grains, whereas the fine ones have been modified by dissolution. (6) A very important question that should be answered is: can porosity be used as an indicator of the depth of burial? The data plotted in Fig. 3-208is based on an average of several porosity measurements. Grain size does not appear to have any effect on porosity. In the lower left diagram of Fig. 3-208 the porosity is plotted versus the present-day depth. The lack of distinct correlation is probably due to later uplift and oil invasion (see dis-
DIAGENESIS OF SANDSTONES AND COMPACTION
325
r-POROSITY, O/o
1000
-
E
i In W
0
Zoo0 -
3000 -
k. 0
zoo0
xMo 15
20
25
M
15
20
25
M
E Y
0
9
'
POROSITY, /o ' Oil-filled sondetone8 0 Water-filled sandstone8 \ Present depth - Maaimum depth of burial / Maximum depth of burial before oil accumulalion
Fig. 3-208. Relationship between porosity and depth of burial for Aalenian sandstones. The dots are plotted at the maximum depth of burial before oil accumulation, whereas the circles represent the maximum depth of burial ever reached. For the means presented, only samples containing <4% carbonate, <5% clay (<20 p ) , and 6 2 % by volume of pyrite from wells with a small scattering range of the porosity values have been used. Abbreviations are explained in Fig. 3-205. Lower left-hand diagram: porosities (from the upper graph) are plotted against the present depth. Lower right-hand diagram: porosities of water-saturated sandstones only (from the upper graph), plotted versus the maximum depth of burial. (After Philipp et al., 1963, fig. 8; courtesy 6th World Pet. Congr., Frankfurt/Main.)
cussion under 4 above). The lower right-hand line in Fig. 3-208,which is based only on the "water-filledsandstones, illustrates a definite relationship between the porosity and depth of burial and can thus be used to determine
326
K.H. WOLF AND G.V. CHILINGARIAN
the maximum depth of burial of these water-filled sandstones. The possible influence of oil invasion also has been considered. The porosities of the oil-saturated sandstones (full dots) have been plotted at the maximum depth of burial prior to oil invasion, by making a preliminary assumption, yet to be tested, that the porosity decrease with depth has been interrupted by the oil impregnation. The sandstones filled with oil at shallower depth of about 5 0 0 m and having a high porosity have been compacted from about 35% (water-filled Eocene sandstones have 35% porosity, see 0 s in Fig. 3-208) to 27-31% porosity during further subsidence, in spite of the presence of the oil (Vk, We, Lii, Ha in Fig. 3-208). The compaction is mainly a result of mechanical rearrangement of the grains. As Philipp et al. (1963, p. 465) mentioned, mechanical rearrangement of the grains in laboratory experiments resulted in as little as 27.7% porosity. Although no reservoir rock measurements were made to show the influence of subsidence on the porosity of oil-filled sandstones between a depth of 1000 and 2000 my measurements were available from greater depths (Hur and Me in Fig. 3-208). Based on studies of quartz diagenesis, oil accumulation occurred at a depth of at least 1700 m where the porosity was about 25%. The measured present-day porosity, however, is only about 19% (Fig. 3-208), which may be the result of precipitation from connate waters. From the above data, Philipp et al. concluded that the porosity of sandstones cannot be used at the present state of available information to determine the burial depth during oil invasion. On the other hand, porosities of water-filled sandstones may enable determination of the maximum depth of burial. (7) The variation of properties of hydrocarbons (both gas and oil) with depth have also been discussed by Philipp et al. (p. 466), as they are partly related t o the porosity of the overlying shales and the compaction history in general. Philipp et al. (p. 471) considered the history of basin subsidence and migration of hydrocarbons in individual oil fields based on geological criteria (Fig. 3-195) and quartz diagenesis (Fig. 3-198). They pointed out that the structures presented in their paper appear t o be more gentle than they actually were at various earlier geologic stages because of subsequent compaction. They also mentioned that rates of sedimentation in the central parts of the trough were four times faster than in the peripheral portions. This must have resulted in differential compaction and preferentially directed fluid movements. Also, these investigators pointed out that in certain cases, compaction of the source rocks was interrupted by uplift and erosion, but resumed upon renewed sediment accumulation. Thus, compaction, compaction fluid movements, and related diagenesis can have a complex history and can even be cyclical. It may be of interest that Habicht (1963) in the discussion on the paper by Philipp et al. offered another explanation that is of importance in the study of compaction. He proposed that the principal cause for oil and
DIAGENESIS OF SANDSTONES AND COMPACTION
327
gas migration was not so much the expulsion of oil and gas from progressively compacting sediments, but the increase of temperature in the subsurface. Habicht presented a diagram showing the relationship between quartz precipitation and the time of oil accumulation. The change of depth of burial of the source rock during geological time was also presented. Jankowsky (1963)also employed quartz cementation as an indicator for the relative age of oil versus water. He stated (p. 458)that the formation of petroleum deposits depends on: (a) oil mobilization, which takes place during compaction and is related t o the increase of temperature and pressure in sedimentary basins, and (b) oil migration, first occurring more or less vertically with the compaction fluids and then more or less horizontally into a reservoir rock to form an oil accumulation in a closed structure. If subsequent tectonic modifications occur, e.g., tilting as shown in Fig. 3-209,then remobilization of the fluids can take place. The quartz-diagenesis studies were applied in the latter case. Figure 3-209indicates that the oil trapped by faulting moved from SE to NW within the same stratigraphic or lithologic unit as a result of the tilting. The study of quartz diagenesis of the Dogger Sandstone helped in the reconstruction of these events. In borehole A, the Dogger Sandstone is filled with water today but exhibits little quartz cementation, which indicates pre-Albian presence of oil that prevented more extensive quartz precipitation. There is an extensive quartz diagenesis in borehole B, especially in the deeper sandstone units, suggesting continuous presence of water and presence of small amounts or absence of oil. The Dogger Sandstone in borehole C,which is filled with oil today, exhibits intermediate I.
2'
A
0
C
I
I
I I
I
A BASE OF
'ALB
LOW
n.~i**
of
Hlom
ausm n~~(.n..ii
I
C
I"lOrn*dI.t,
Fig. 3-209. Structural sections to explain the quartz diagenesis. 1 = present-day structure; 2 = pre-Albian structure; 3"=quartz diagenesis in the aquifer. (After Jankowsky, 1963, fig. 6; courtesy 2.Dtsch. Geol. Ges.)
K.H. WOLF AND G.V. CHILINGARIAN
328
CLAY CONTENT,
Oh
Fig. 3-210.Porosity as controlled by clay content. Numbers next to the symbols indicate the sampling depths in cm below the ocean floor. Samples without a number represent surface samples. The curve averages the range of clay contents of the uppermost meter of sediments. I n the diagram in the lower right-hand corner are values obtained from the harbor bay indicating the deviations of porosity from the above-mentioned curve. In Figs. 3-210, 3-211, and 3-212,only samples were utilized that were obtained with a large push-cylinder. (After Fiichtbauer and Reineck, 1963, fig. 4, p. 299;courtesy Sedimentology.) I = sand flat (bank); 2 = mixed (mud + sand) flat; 3 = mud flat; 4 = “surf flat” (= flat with surf and/or breakers); 6 = “flat sea”; 6 = harbor bay or harbor inlet; A = deviation from the upper curve (= average porosity of the uppermost meter) in % porosity; B = depth (in m ) below the upper surface of the sediment.
degree of quartz diagenesis. Together with the information gained from borehole A, this suggests that the oil moved into this sandstone during postAlbian time. Fiichtbauer and Reineck (1963) undertook a study to determine whether or not: (a) sediments of the same grain size, but formed in different depositional environments, have the same porosity, and (b) their condensation ratio* is different. Although the results are related to the material presented in the section on density and compressibility, they are given here because they show some regional or environmental variation and are, therefore, of interest in large-scale compaction investigations. Figure 3-210 shows the increase of porosity with increasing clay content. In pure sands, most porosity values are around 40%. Clays from a young harbor (= Hafenbucht), where
* The condensation ratio = 4w-4 x kv
- $d and &
100
(3-11)
where 6 = natural porosity, and @d = porosities representing the loosest and closest packing arrangements of grains under water, respectively.
DIAGENESIS OF SANDSTONES AND COMPACTION
329
TABLE 3-LIV Summary of average values of densities of water-saturated sediments (after Fuchtbauer and Reineck, 1963, table 11, p. 300) Type of sediments
Density (g/cm3)
Number of samples used
Marsh Dry beach Wet beach Mud-watt Mixed-watt Sand-watt Mud from harbor bay Subaqueous sand bank Shelf
1.427 1.907 1.919 1.538 1.747 1.872 1.332 1.953 1.963
1 15 49 6 7 56 40 7 20
rates of sedimentation are up to 50 cmfyear, have a porosity of about 83% at the surface (determined with picnometer). The clay content of the sediments depends on the rate of sedimentation and other factors of the sedimentary milieu. The clays of the shelf zone (= Flachsee) and Wadden mud (= Schlickwatt) are denser with a maximum of about 70% porosity. Increasing age, coarser grain size, and different sedimentary processes may be responsible for this. The weight of water-saturated sands also depends on the clay content, as presented in Table 3-LIV. As shown in Fig. 3-211, in which porosities were plotted versus median grain size, the three different depositional environments show some variations. These results, however, should not be used for generalization. Noticeable is the small and irregular dependency of porosity on grain size. This, however, is not true when one uses the same sediment samples and performs artificial sedimentation by free fall in the laboratory. In the latter case, there is a distinct decrease in the porosity with an increase in grain size. This cannot be explained by sorting, but may be the result of greater sphericity of the coarse grains that, in turn, results in tighter packing. Similar relationships occur in ancient sandstones, one example of which is presented in Fig. 3-211. These sandstones were buried to a depth of about 1100 m where cementation was prevented by oil impregnation. Fiichtbauer and Reineck concluded that porosity is independent of the grain size in newly-deposited, uncompacted sediments, but interdependence between these two variables increases with increasing burial pressure until there is a normal grain size-versus-porosityrelationship. The increase of porosity with decreasing grain size in recent sands may be compensated by a closer packing arrangement of the finer grains, expressed by a higher condensation ratio. As shown in Fig. 3-212, the condensation ratio increases distinctly with decreasing grain size, which applies to sediments of all depositional environments.
K.H. WOLF AND G.V. CHILINGARIAN
330 I
.
.
.
.
I
.
. . .
I
.
.
.
50%
\ &\
.
h ’
8
ae
i
t
v)
0
a
0
n
\
-30
O!l
012
013
03 MEDIAN DIAMETER,rnrn
OIZ
d3
Fig. 3-211.Relationship between porosity and median diameter in sands. The spread of the values for natural sediments is represented in the area which includes 68% of the values (B). The same area was also duplicated in the right-hand figure, so that a comparison is possible with the artificially accumulated sediments in the laboratory. The diagram in the left-hand corner with a different ordinate axis shows porosity distribution in a Lower Cretaceous sandstone (= “Unterkreidesandstein”). A. Subaerially loosely packed sands as a result of vertical accumulation (dropping down without further shaking). B. Recent sands. C. Subaqueously loosely packed sands as a result of vertical accumulation. D. Subaqueously moved sands with tighter packing. 1 = surf beach o r shoreline beach; 2 = sand flat; 3 = flat sea. (After Fiichtbauer and Reineck, 1963, fig. 5, p. 301;courtesy Sedimentoiogy. )
One should notice, however, that the organically-reworked sediments (by boring and burrowing) have a higher condensation ratio. The results presented in Fig. 3-212are in agreement with those obtained by civil engineers, i.e., that the high energy required to obtain maximum values of condensation of coarse sands is usually not available in natural sedimentary environments, in contrast to the lower energy required for finer sediments. Only during compaction will the coarse sediments be “condensed”. As shown in Fig. 3-210, the porosity of the upper layers of quickly-accumulated clays is 83%, whereas that of slowly deposited clays is about 70%. The decrease of porosity from about 83 to 76% (a difference of about 7%) in the upper 4 m of
DIAGENESIS OF SANDSTONES AND COMPACTION
lo 90-
2
331
I
80 -
70-
'z 0
5
m
60-
50-
W
n
z
8 LO30200 O \0" 0
10-
I
0.2
0.1
0.3
MEDIAN DIAMETER, mrn
Fig. 3-212. Relationship between "condensation" and median diameter (see text). 1 = sand flat; 2 = surf flat (= flat with surf and/or breakers); 3 = shallow sea; 4 = shallow sea (well agitated); 5 = shallow sea (partly agitated); 6 = shallow sea (strongly or well agitated). POROSITY,% IS
20
2s
"00
35
30
//RECENT /
/
OL/GOCENE D
500
000
/
E
/
/
, ,
//
EOCENE
/
2
5 3
m
0 lL
r a n 2500
,'
3000
I
/
/
/
PURE SANDSTONES
/
"/#MAT
Fig. 3-213. Decrease- of porosity in pure sandstones with increasing maximum burial depth. (After Fuchtbauer and Reineck, 1963, fig. 7, p. 304; courtesy Sedimentology.)
K.H. WOLF AND G.V. CHILINGARIAN
332
sediments is shown in the right-hand corner of Fig. 3-210.The porosity in the upper 1 0 0 m is very variable and may depend on the mineralogical composition as well as transportation-deposition mechanisms. The type of clays present plays a major role in determining porosity of clay deposits, e.g., illite versus montmorillonite. The same arguments apply to sands and sandstones with a clay matrix. Figure 3-213 presents one example where pure sandstones decrease in porosity from about 40% at the surface to about 14% at a depth of 3300 m. Usually, straight-line relationships between porosity and depth of burial occur only where the sediments are matrix- and cementfree.
Study b y Adams Adams (1964)studied a sandstone formation which had undergone extensive diagenetic alterations on a regional scale, that influenced its reservoir properties. This formation of Pennsylvanian age consists of two basic sandstone types (Fig. 3-214): (A) nearshore, clean, relatively well-sorted, nonglauconitic, non-calcareous sandstone deposited under high-energy conditions, and (B) sandstone representing seaward facies of type-A sandstone, which usually is argillaceous, poorly-sorted, chloritic or glauconitic, and calcareous, and was deposited by a medium of lower energy than that required by type A. The diagenetic features include effects due to pressure solution
-+TYPE
'6' SANDSTONE
Fig. 3-214. Major depositional facies relations of typical Lower Morrowan sandstone (Pennsylvanian) Anadarko Basin, northwestern Oklahoma. (After Adams, 1964, fig. 3, p. 1570; courtesy Am. Assoc. Pet. Geologists.)
DIAGENESIS OF SANDSTONES AND COMPACTION
(A) W S T D N E TYPE A FABRIC CLASS I
(8) SANDSTONE TYPE A FABRIC CLASS II
(C) SANDSTONE TYPE A FABRIC CLASSI AI
+FAWN cussmi I
n
333
IE) SANDSTONE TYPE A FABRIC CLASSII[ B I
IF) SANDSTONE TYPE B FABRIC CLASS JE B 1-2
(D1 dmmm oB W
” E
TYPE A
SANDSTONE TYPE B FABRIC CLASS B 1-3
m
SANDSTONE TYPE B FABRIC CLASS a B 3
Fig. 3-215. Thin-section illustrations of major fabric classifications. (After Adams, 1964, fig. 5, pp. 1572-1573; courtesy Am. Assoc. Pet. Geologists.) For description of the classes, see Table 3-LV. C is porous.
and presence of several cements, i.e., quartz, calcite, dolomite, and clay. Replacements, decementation, and corrosion are common. Adams discussed: (1)pressure solution; (2) quartz overgrowth; (3) stylolites; (4) sutured, concavo-convex, and other types of contacts versus floating grains; (5) influence of clay on pressure solution; (6) cementation related to compaction; (7) distribution of calcite versus dolomite cement; (8) clay matrix and cement distribution; (9) etched quartz grains; (10) porosity (original and secondary) distribution prior and subsequent t o mechanical and chemical diagenesis. He found a regular and predictable distribution of many of these features from sandstone facies of type A t o type B. Adams offered a petrofabric sandstone classification scheme in which subdivisions are based on degree of pressure solution with various modifiers, including grain outlines and contacts, amount and type of cement(s), and porosity (Table 3-LV) (see modifications made by Pettijohn et al., 1972). Sketches of each of these major petrofabric types are presented in Fig. 3-215. According to Adams, “diagenetic facies”
DOWNDIP DEPOSITIONALLY
STEP 3 Well No. 2
VALUATION: Wildcot well no. I hos encountered good 90%poy. No clue to optimum dirrction for field development. ECOMMENDATION: Study rocks, clossify ond integmta with rock geometry.
EVALUATION: Well no.2 encountered Sondsto$e,yB," we11 developed and Sandstone A tqht. No Clue to optimum dirtction for odditimal development: RECOMMENDATION: Study rocks,classify and integrate with rock geometry.
-
STEP 4 Well No.2
STEP 2 Well No. I
4
13,. SANDSTONE TYPE B FABRIC CLASS&! BI GAS fP/fl, MCFD SANDSTONE TYPE FABRIC CLASS
m
Sample and petrographic anolyses slmw pay to Be 8econdory. Gbomrtry of such a yrervoir is unpredickblr. Sandstone 8 is fine groined with FobricClorsI ond is pmbobly ot updip edge of porous, high energy sondstone. ECOMMENDATION: Drill dopqdip depositionally from Sondstone 8 .
VALUATION:
EVALUATION: Samplr ond pstrogrophic mol ws show pay to be in high ensrgy ronds!me with pood prediction possibilities. RECOMMENDATION: Drill along Cposilionol strike of Sondrtone "8':well no. 2
DIAGRAMMATIC SKETCHES SHOWING APPLICATION TO HYDROCARBON EXPLORATION (Wells No.1 ond 2 Are Appror
1
I Mile Aport
I
MECHANICAL LOG POROSITY
= SANDSTONE TYPE "A"
4I
= SANDSTONE TYPE "8" * SPONTANEOUS POTENTIAL LOG
Fig. 3-216. Sketch of an actual case within study area (subsequent to original study in 1961) depicting importance of detailed rock analyses to hydrocarbon exploration. (After Adams, 1964, fig. 6, p . 1574; courtesy Am. Assoc. Pet. Geologists.)
DIAGENESIS OF SANDSTONES AND COMPACTION
335
TABLE 3-LV Criteria of fabric classification and applicable descriptive terms (modified from Gilbert, 1949,p. 13,and Thomson, 1959,p. 100, in Adams, 1964,table I, p. 1579)
Fabric Class I High degree of pressure solution, flattened grains, no porosity, no cement, sutured contacts Fabric Class I I Moderate degree of pressure solution, approximately equidimensional grains, little or no porosity, little or no cement, primarily concavo--convex contacts with some flat, sutured contacts (A) Cemented by: (1)quartz overgrowths; (2)quartz overgrowths with minor amounts of calcite; (3)quartz overgrowths with some clay Fabric Class 111 Minor degree of pressure solution, mostly original grain outlines, often some porosity, long or tangential contacts (A) Poorly cemented (but normally tightly packed): (1)quartz overgrowths; (2) calcite; (3)dolomite; (4)clay (B) Well cemented: (1)quartz overgrowths; (2)calcite; (3)dolomite; (4)clay Fabric Class IV No pressure solution, original grain outlines, good porosity and/or cementation with no porosity, tangential contacts and floating grains (A) Cemented but porous: (1)calcite; (2) dolomite; (3)clay (B) Cemented, no porosity: (1)calcite; (2) dolomite; (3) clay
could be established based on his classification, and Fig. 3-216represents an actual example of his approach. Although he applied it to oil and gas exploration, eventually such diagenetic investigations will probably be used in solving ore genesis problems, e.g., of uranium ores of the Colorado-Plateau type and copper ores in the red-bed sandstones. Study b y Horn
Horn (1965) undertook a regional study of the reservoir properties of a Jurassic sandstone, i.e., porosity and permeability that depend on lithology and facies types as well as on the characteristics of the pore system (content of matrix and cement, and degree of compaction, for example). The following diagenetic-authigenic minerals were found in pores of the sandstones: chamosite, siderite, pyrite, brookite, quartz, feldspar, kaolinite, barite, ankerite and calcite. Calcite precipitation caused the greatest reduction in porosity from an initial value of 40% to 1.5-7%. The sandstones cemented by kaolinite have 6--11% porosity. The other authigenic minerals did not influ-
336
K.H. WOLF AND G.V. CHILINGARIAN
ence porosity greatly except on a microscale perhaps. Siderite occurs in thin zones associated with siderite concretions. Ankerite, which is confined to one particular sandstone unit, locally reduced the porosity to 5%. Pyrite, brookite, feldspar, and barite have minor influence on porosity, whereas quartz was very important in reducing porosity. In discussing the process of silicification, Horn (p. 251) stated that precipitation of SiOz can be prevented or interrupted by (1)filling of the pore spaces by non-aqueous fluids such as oil, or (2) precipitation of a mineral on the quartz grains so that the latter cannot come in contact with the interstitial fluids and, consequently, no silica can be chemically deposited. Presence of large amounts of matrix in a sandstone also would prevent the entry of fluids. Heath and Dymond (1973) have made similar observations on recent sediments. They stated (p. 181) that on the basis of dissolved silica, the pore waters from North Pacific deep-sea sediments fall into two groups. Pore water of oxidized deposits contain 22-35 ppm, whereas those of reduced sediments contain 10-19 ppm of silica. The above difference in silica concentration remains unexplained. Heath and Dymond (p. 184), however, offered a conceivable possibility, namely, that the fine-grained silicate particles in the oxidized sediments were rapidly coated with ferric hydroxides, which inhibited the attainment of equilibrium; whereas in the reduced accumulations reversible reactions, i.e., decomposition of the silicate minerals, are much more likely to take place (see their equations 1and 2). Horn (1965, p. 252) found that rims of radially-arranged chamosite prevented formation of quartz cement. Chamosite is present only on primary clastic grains and not on any diagenetic precipitates, which indicates its very early precipitation. As Fig. 3-217 indicates, there is development of two types of chamosite, i.e., types A and B. Variety A, which occurs as welldeveloped chamosite rims at least 3-5 p thick, prevented secondary silicification so that the average porosity of sandstone is 22.8%. The type-B chamosite occurs as thin rims (less than 3 p ) , so that precipitation of silica was possible and the average porosity of the sandstone was reduced to 16.2% (Fig. 3-217). The variation in the percentages of chamosite of types A and B occurs both horizontally and vertically within the sandstone units. By examining the regional distribution of chamosite zones, one can explain also the regional variation of porosity. The chamosite formation followed shortly after sedimentation of the clasts. Thus, one can expect some relationship of chamosite occurrence to the sandstone facies, because during the very early diagenetic stages, when the chamosite was formed, there was probably no complete renewal of the intrastratal pore fluids. This assumption is confirmed by: (a) the fact that the rims of chamosite are preferentially present in clay-poor, homogeneous sandstones, but are absent in sandstones with clay laminae, and (b) the association of the chamosite with the chamosite
331
DIAGENESIS OF SANDSTONES AND COMPACTION DiSTRlBUTlON POROSITY
0EPTH.m SEDIMENT
OF
-m &0 a
20 10
15 20 25 POROSITY
10
0
1
W 1
10
100
1000
10
15
20
25
B L
m 5
B
B2 3 + 6
' 7
Fig. 3-217.Dependency of porosity and permeability (averaged values of the measurements normal and parallel to the bedding) on the chamosite content in oil-free, upper coarse-grained "Haupt" Sandstone of the borehole Plon-Ost 105. According to thinsection studies, type A = well-developed chamosite rims, silicification is absent ( 4 ) ;type B = thin chamosite rims, strong silicification (5);transitional between types A and B (6); no thin-section observation (7); fine-grained sandstone (1); mudstone (2); calcitecemented sandstone (3); loss of cores (KV). (After Horn, 1965, fig. 4,p. 251 ; courtesy Erdol Kohle.)
oolites. Figures 3-218A7B give the regional distribution patterns of chamosite; the top diagrams refer to the upper, coarser sandstones (11) and the lower diagrams to the lower, finer sandstone (I) units.'In unit I, there is an elongated SW-NE zone of clay-free sandstone containing conglomerate horizons; an increase in grain size, in general, reflects more turbulent depositional environment. Especially here one finds chamosite oolites and authigenic films of chamosite on the grains. Degree of silicification was low (Fig. 3-218,E) in unit I and high porosity was preserved (Fig. 3-218,F). Clayey sandstones with normal laminations and cross- and flaser-bedding, accumulated outside this elongated zone. Inasmuch as here the rims of chamosite were only poorly developed or absent, a reaction between the fluids and the quartz grains was possible and silicification with consequent porosity reduction took place (Figs. 3-218, D,E, and F). Similar observations have been made in the upper sandstone of unit 11, but the zones of stronger chamo-
K.H. WOLF AND G.V. CHILINGARIAN
338
il
LOCATION
- 1
ow---
MAP
uI.
0 2
TYPES OF SEDIMENTS AND STRUCTURES
I POROSITY, %
TYPE OF CEMENT
1
.A&?
12.8
3 5. 3 6.
20.8
17.3
El?
B 8.
Fig. 3-218.Comparison between sedimentary facies development, cementing material and porosity in two units of the Dogger-/? “Haupt” Sandstone of the Plon-East Oil Field. (In the sketch map to the left, the depth lines of the “Haupt” Sandstone are 100 m apart.) 1 = homogeneous, clay-free sandstone; 2 = clayey sandstone; 3 = unoriented flakes (= plates) of chamosite; 4 = oolites of chamosite; 5 = horizontal thin bedding; 6 = flaser bedding; 7 = cross-bedding; 8 = conglomerate horizons; 9 = authigenic chamosite (bank); 1 0 = extensive secondary silicification. A, B, C = upper coarser-grained unit of ‘‘Hauptsand”; D, E, F = lower finer-grained unit of “Hauptsand”. (After Horn, 1965, fig. 5, p. 252; courtesy Erdol Kohle.)
site development are more pronounced (Figs. 3-218A,B, and C). Very porous sandstones with an average porosity of 21.9-23.5% correspond to type-A chamosite precipitation (Fig. 3-218,C, lower right), whereas the maximum porosity of 18.3%in the upper part of Fig. 3-218,C is present in sandstones having the type-B chamosite genesis that is accompanied by weak or reduced silicification. In his discussion on the influence of oil invasion on diagenetic silica precipitation, Horn (p. 253) stated that inasmuch as the porosity in the sandstones discussed above were determined by the synsedimentary chamosite genesis,
339
DIAGENESIS O F SANDSTONES AND COMPACTION
later differences in the supply of silica were probably less important. On the other hand, in clay-poor sandstones, in which the grain surfaces are not covered by films of chamosite or chlorite, it seems possible that quartz diagenesis was influenced by early oil and gas invasions. These latter conditions apply in another oil field investigated by Horn. As Fig. 3-219,Aillustrates, the fine-grained and coarse-grained sandstones are interbedded with the clayey sandstones, which contain clay laminae. Locally, these sandstones are replaced by sandy claystones (Fig. 3-219,A).In contrast to the earlier example mentioned (Fig. 3-218),Fig. 3-219,Cshows that the lines of equal porosity are parallel to the contour lines of the sandstones, despite the presence of a distinct facies differentiation. This suggests that the porosity distribution is not controlled by primary sedimentary features. The more intense cementation or lithification in the north is the result of more intense silicification (Fig. 3-219,B).Inasmuch as there is no relationship between the LOCATION
MAP
I P E OF SEDIMENT
I , -
TYPE
OF
CEMENT
POROSITY, o /‘
OOO
_.-
21.1
:EL
@L
-
2.
Fig. 3-219. Relationship among sedimentary facies development, type of cement and porosity in the upper coarse-grained “Haupt” Sandstone of the Preetz Oil Field. (In the diagram to the extreme left, the depth lines of the “Haupt” Sandstone (“Hauptsand”) are 100 m apart.) a = strong silicification; b = weak silicification; I = boreholes examined; 2 = facies boundaries; 3 = claystone; 4 = fine-grained sandstone; 5 = coarse-grained sandstone; 6 = claystone with illite-muscovite. (After Horn, 1966, fig. 6, p. 253; courtesy Erdol
Kohle.)
K.H. WOLF AND G.V. CHILINGARIAN
340
lithification and the sedimentary facies, it appears that differences in the cement distribution are the result of diagenetic processes. The absence of quartz precipitation in the south may have been the result of oil invasion that prevented the movement of aqueous solutions. The northern area was situated near the oil pool’s margin, filled with water, so that silica precipitation was not hindered and the sand was cemented more intensely, as exhibited by the presence of 5-6% more cement (Fig. 3-220).The E-W boundary between boreholes 6 and 9, exhibiting striking differences in diagenesis, does not coincide with present-day marginal water boundary, which extends in a E
AVERAGE VALUE OF POROSITY, 10
W J
> W -1
I
2400-
”
“
”
’
PREETZ
15
”
’
”
*
”
O‘/
20
25
12 a
e
0
m
N W
4
2 500-
z 0
2600
A
W
m
a
-
v)
I-
a
3
I
2 700 O A
W
I
I-
LL
r-
2700
P LO N
- o ST
z 3
E
m
2800
m W
0”W
n
2900
W
I I LL
0
3000
rT >
$
3100
8 rT
s
3200
W
f
U
0
I I-
a W n
3300
b. o c.
15
20
AVERAGE VALUE OF POROSITY,
25 O/O
Fig. 3-220.Relationship between porosity and depth of burial of the “Haupt” Sandstone (“Hauptsand”) in the oil fields Plon-East and Preetz. a = oil-free Gamma Sandstone, Preetz 4; b = upper coarse-grained part of “Haupt” Sandstone; c = lower finer-grained part of “Haupt” Sandstone. (After Horn, 1965, fig. 7 , p. 254; courtesy ErdOZ Kohle.)
DIAGENESIS OF SANDSTONES AND COMPACTION
341
N-S direction at about 2630 m depth parallel to the contour lines of the sandstone (Figs. 3-219and 3-220).Oil migrated later into the northern part of the sandstones after the subsurface fluids deposited silica cement. The sandstones in borehole 9 now contain secondary water. Remains of oil and the relatively high porosity indicate that the rocks contained oil earlier. Horn (p. 254) compared the silicification of the sandstones in the PlonOst and Preetz oil fields and found that in the former sandstones the porosity is dependent on the chamosite content, whereas in the latter the silica cementation is controlled by earlier oil invasion. A comparison between these two rock types is desirable because these sandstones are of the same age, occur at the same depth in subsurface, and originally contained similar pore fluids, but exhibit differences in facies and diagenesis. In both cases, there is a linear decrease of the average porosity with increasing depth (Fig. 3-220).The effect of increasing compaction here is related to the pressure solution of the quartz grains at the grain contacts with subsequent or simultaneous precipitation of silica from the saturated solutions at other localities. This mechanism has been described in detail by many investigators in the literature. In the Preetz oil field, silica precipitation took place more or less at the site of pressure solution. In the Plon-Ost oil field, pressure solution also occurred at grain contacts, where the rims of chamosite were missing and are 'cly*>z,?nt only as linings of the pore walls. On the other hand, in the chamosii.e-rich sandstones, the silica precipitation from saturated solutions was hindered. It is possible that the silica was supplied from more remote localities, where pressure solutions due to the overburden pressure was possible and the fluids moved into the clayey sandstones to take part in the quartz diagenesis. The silica may also have been derived from feldspar decomposition and from clays during mineral transformations. Hence, there is a considerable variation in the average values of porosity at the same depth in the sandstones of the Plon-Ost oil field and only a small decrease of the average porosity values of 1-2% per 100 m of depth, in contrast to the gradient of 5% per 100 m reported from the sandstones in the Preetz oil field (gently versus steeply sloping average curve, Fig. 3-220).To what extent the intense decrease in porosity in the Preetz oil field is the result of overburden pressure and whether other factors are involved, has not been determined because a depth interval of 200 m which was studied is not sufficient. A periodic or cyclic oil invasion, for example, may have effects similar to those of overburden pressure, when silicification can occur in lower horizons but is prevented by the oil at higher horizons, until the lower units also receive oil and the silica precipitation is interrupted. As discussed above, high porosities can be maintained in quartzose sandstones when pressure solution is prevented and/or when silica-saturated solu-
342
K.H. WOLF AND G.V. CHILINGARIAN
tions cannot react with the quartz grains because of the presence of rims of minerals around the clastic constituents. As reviewed above, silicification can be prevented also by invasion of oil. Although the chamosite rims can preserve the high porosity to a greater depth, inasmuch as these rims prevent or diminish cementation, the rate of oil migration will be less, because of greater adsorption of oil to the surfaces of the grains. The presence of rims decreases the size of the pores and at the same time increases the surface area per unit of rock volume. The presence of chamosite rims, therefore, is not only important in controlling presence and degree of pressure solution and, consequently, porosity and permeability, but also has to be considered in evaluating the subsurface hydrodynamic flow. The degree of compaction may vary from layer to layer and may depend on the amount of chamosite rims and matrix formed during early diagenesis. (See Table 3-LVI.) TABLE 3-LVI Dependency of the porosity on the chamosite content in sandstones of the same facies (“Haupt” Sandstone Plon-East 102; length of sample = 20 cm; permeability measurements parallel and normal to bedding; (after Horn, 1965, table 1, p. 253) ~
Depth (m)
3239.9 3241.7 3243.6
Chamosite content
4-
-
Secondary silicification
-
+
+
-
-
+
+
-
-
+
~~
~
~~
~~
Porosity (%)
Permeability (md)
Il*
I**
II
1
(%)
22.3 5.9 25.6 9.8 20.8 9.8
18.1 7.5 25.2 9.0 19.5 9.0
125 -
4.3 0.1 163 2.5
1.2 2.1 2.5 0.9
22 1.4
1.3 3.3
189 2.4 83
-
Amount dissolved in acid
* 11 = parallel to bedding; ** 1= perpendicular to bedding. Diagenetic features and chemistry o f pore fluids (Selley, 1966) Selley (1966), in his discussion on diagenetic variations on a regional scale (both vertically and horizontally), suggested that the diagenetic features may be related not only t o pressure solution or burial pressure but also to the chemistry of the pore fluids, for example. In Fig. 3-221, he presented three different units that originated under different environments. Pressure solution and cementation of quartz is confined to those units that accumulated under oxidizing conditions, in contrast to the lacustrine-marine sediments
DIAGENESIS OF SANDSTONES AND COMPACTION
343
Fig. 3-221. Diagram illustrating the relationship among mode of diagenesis, lithofacies, and presumed depositional environment. (After Selley, 1966, fig. 4; courtesy Proc. Geol. Asso c. )
which contain large amounts of chlorite, carbonate, and quartz cements. Cementation does not seem to have been accompanied by pressure solution. The effects of compaction and depth o f burial (Fuchtbauer, 1967) Fuchtbauer (1967a) presented the results of an investigation in which he directly considered effects of compaction and depth of burial. In the section on quartz diagenesis and mechanical compaction, he showed (Fig. 3-222,A) that the porosity of the “Dogger beta” quartz sandstone decreases with increasing depth. The curve nearly coincides with the dashed curve presented by Proshlyakov (1960, in Maxwell, 1964). Examination of the grain surfaces revealed the cause for this porosity decrease, i.e., the percentage of quartz grains, which show secondary overgrowths on crystal faces, increases with depth (Fig. 3-222,B). As the silica is dissolved as a result of pressure solution at points of contact, the grains slip into denser packing (Fig. 3-223,A). The latter figure also shows that the amount of pressure solution required is small. Fiichtbauer (p. 355) pointed out that the rock volume to be dissolved is in general smaller with increasing steepness of the contact faces of the grains. In Fig. 3-223,B, changes in volumes are based on the most unfavorable case of horizontal contact planes. Only 1.5% of the material must be dissolved to result in compaction (up t o 50%), as graphically shown by curve A in Fig. 3-224. The amount of quartz which has to be dissolved evidently increases with compaction, because contacts become longer and the number of horizontal contacts increases. Consequently, as pointed out by Kchtbauer (p. 355), the “mechanical slipping (rearrangement) of grains dominates during early compaction, whereas chemical compaction (pressure solution) is predominant during later stages”. In determining the dissolution of quartz with depth, as shown in Fig. 3-222,B, Fuchtbauer employed curves A and B in Fig. 3-224 and the porosity curve of quartz sandstone in Fig. 3-222,A. The general shape of the curves is probably correct, but the maximum amounts of dissolved quartz may be unrealistic. The dissolution of silica is minor in the upper 1000 m of sediments, but increases roughly
K.H. WOLF and S.V.CHILINGARIAN
344
OUARTZ GRAINS WITH OVERGROWTHS, %
POROSITY, %
0
5
10
DISSOLVED OUARTZ REPRECIPITATED, %
Fig. 3-222. Porosity and Si02 migrations in relation t o depth of burial. The maximum depth of burial is the greatest depth t o which a sandstone was ever buried. (After Fuchtbauer, 1967a, fig. 1,p. 354; courtesy 7th World Pet. Congr.) A. The values for calcareous sandstones were taken from many measurements of the “Bausteinschichten” (= Chattian Molasse; Fiichtbauer, 1964, fig. 22). The curve for the quartz sandstones is for the “Dogger beta” (after Fuchtbauer and Reineck, 1963). H I = average value for silicified sandstones; H2 = sandstones with diagenetic chamosite seams that hindered quartz diagenesis (after Horn, 1965). (Dashed curve is after Proshlyakov, 1960, in Maxwell, 1964.) B. Lower left = maximum amount of quartz dissolved by pressure solution (see text). Upper right = quartz diagenesis, i.e., percentage of quartz grains with secondary overgrowth (after Philipp et al., 1963). Fuchtbauer (1961) defined it as: [100(0.5 b + c ) / ( a + b + c ) ] , where a = number of grains with very few crystal faces, b = grains with some crystal faces, c = grains with crystal faces covering the larger part of the surface. The counting was done with microscope using grains of 0.12 to 0.15 mm fraction on a dry glass slide. These values, however, do not depend on the grain size.
A
B
Fig. 3-223. Mechanical and chemical compaction. A. Incline(- contact planes - mechanical compaction is predominant. The rock volume to be dissolved (black) is even smaller if rotational movement occurs. B. Horizontal contact plane - exclusively chemical compaction. Much more material must be dissolved in this case in order to achieve the same amount of compaction. Its amount is independent of grain size as deduced from geometric considerations. (After Fuchtbauer, 1967a, fig. 2, p. 365; courtesy 7th World Pet. Congr.)
DIAGENESIS OF SANDSTONES AND COMPACTION
COMPACTION, %
345
--c
Fig. 3-224. Theoretical relationship among degree of compaction, porosity (curves B and C), and the amount of rock volume dissolved (curve A ) for the model presented in Fig. 3-223,B. Curve A = dissolved quartz(%), ordinate to the right; curve B = prosity (quartz removed), ordinate to the left; curve C = porosity (quartz reprecipitated), ordinate to the left. (After Fuchtbauer, 1967a, fig. 3, p. 356; courtesy 7th World Pet. Congr.)
linearly with depth; this, in turn, correlates with the linear increase of silica overgrowths below a depth of several hundred meters (Fig. 3-222,B).The upper 1000-1500 m of sediments are, therefore, the domain of mechanical compaction. It seems that both sorting and roundness of grains influenced compaction, roundness being particularly important. The finer-grained sands are more porous than the coarser-grained ones having a comparable clay content, which does not agree with the relationship postulated by Weyl (1959) for the effective range of pressure solution. The appreciable difference in porosity is indicated by the short line marked 0.08 mm in Fig. 3-222,A for some “Dogger beta” rocks. Fine-grained “Bentheimer Sandstein” has a porosity of 27% at a depth of 1100 m, whereas the coarsegrained sandstones have a porosity of 22% (see Von Engelhardt, 1960, figs. 11 and 49). The differences in sorting may be significant, but do not seem to explain completely the higher porosity of the finer-grained sandstones. As Von Engelhardt (p. 21) stated, the higher porosity values of finer sandstones can be explained by the larger number of grain contacts/unit volume of sediment resulting in a higher resistance against compaction. As this has not been found to apply to all petrographically similar fine-grained sediments, other factors must also be active. A more important factor may be the roundness, as it is commonly better developed in comer sandstones (exclud-
K.H. WOLF AND G.V. CHILINGARIAN
346 4
3 2
In
I
B
cn
w z
3
3
2
0 2
0
a
I
3
I
C
2 I
GRAIN SIZE, m m
Fig. 3-225. Rounding (Russell-Taylor) in relation to grain size. Ordinate (roundness): 1 = angular; 2 = subangulir; 3 = subrounded; 4 = rounded. This number was multiplied by the number of grains falling into the respective roundness class. Then the sum of these products was divided by the number of all grains. Each point corresponds to one grainsize fraction. The median diameter is shown by an open circle for each sample. A = “Bentheimer” sandstone from Scheerhorn (Hecht et al., 1962), oil-saturated, 4 samples; B = “Dogger beta” sandstone from Hankenbiittel, oil-saturated, 2 samples; C = “Dogger beta” sandstone from Luben-West, water-filled, 2 samples. (After Fuchtbauer, 1967a, fig. 4, p. 356; courtesy 7th World Pet. Congr.)
ing “textural inversion” of Folk, 1968), and, therefore, the coarser grains “slip” better. From Fig. 3-225, two general relationships become clear: (a) roundness increases with grain size, as has been reported from numerous other studies; however, it reaches a maximum in the coarse sand range (curve A ) ; (b) identical grain fractions are better rounded in the fine-grained sandstones than in the coarse-grained sediments. This can be explained by the differences in transportation mechanisms, namely, rolling versus suspension. The median rounding (= rounding of the median-sized grains, shown by circles in Fig. 3-225) can, therefore, be larger in coarser sandstones (curves A and C) or equal to that of fine sandstones (curve B ) . It follows then, that should rounding affect porosity, the curves A and C should show a relationship between grain size and porosity. But curve B (center of Fig. 3-225) does not show such an interdependency, as demonstrated in Table 3-LVII. According to Fiichtbauer (p. 357), there is an important influence of the degree of rounding on compaction during the mechanical stage of diagenesis resulting in a reduction of porosity. Inasmuch as the specimens contain on the average only 11%quartz grains with overgrowths, the amount of pressure solution must have been comparatively small.
347
DIAGENESIS OF SANDSTONES AND COMPACTION POROSITY ( d o t s )
FILLED
Q U A R T Z OVERGROWTHS
CARBONATE G + S
- 1510 m
,
,
,
f.’L/A5,
I
*
:
,
,
, ,:
j
,
,
I
Md a18 0.20
0.1
1
0.08
0.10
TERTIARY OILFILLED
+ m
Fig. 226. Relationship between the L-gree of quartz diagenesis (centei,, pore filling (left), carbonate content (right), and distance from the sandstoneshale contact for three sandstones of the “Dogger beta” Formation. (After Fuchtbauer, 1967a, fig. 5, p. 357; courtesy 7th World Pet. Congr.) Top = Wesendorf-South 2 (after Fuchtbauer, 1961); center = Hankensbuttel; bottom = Bodenteich; left = median diameter in mm, porosity, and SP curve; right = G + S = garnet + staurolite in the heavy-mineral fraction (after Dmng, 1965).
348
K.H. WOLF AND G.V. CHILINGARIAN
The second stage of compaction is the chemical one, characterized by increased pressure solution combined with precipitation of silica on quartz grains. This can lead to complete silicification of the sandstone as long as no other cementing mineral occurs. Figure 3-226, which shows the source of silica, consists of three parts: (a) top: a section through a water-filled rock; (b) center: a section through an oil-filled rock; and (c) bottom: a section through a formerly oil-filled, but now water-filled sandstone. Each one of these cases is considered separately. (a) The amount of secondary quartz overgrowths is high in the water-filled section, except for portions with carbonate cements. The latter preserved an early stage of diagenesis, which is evidenced by the presence of a number of unstable heavy minerals (e.g., garnet and staurolite). (b) Less quartz precipitation occurred in oil-impregnated sandstones. The minor amounts of quartz present apparently have been preserved by the oil. Using the upper right-hand side (hatched area) of Fig. 3-222, one can deduce that the depth at the time of oil migration-and impregnation was about 1000 m. The importance of this quartz diagenesis to the structural history of one sedimentary trough and the reconstruction of the history of oil migration have been discussed by Philipp et al. (1963a,b). The sections that are close to the sand-shale interface, have much higher amounts of quartz overgrowths (Fig. 3-226). This increase in quartz content on approaching the sand-shale boundary is present in all sections of oil-filled sandstones studied by Fiichtbauer and can be explained only by assuming continuation of quartz precipitation after oil impregnation, possibly at a reduced rate. It seems that silica migrated into the connate water of the oil-saturated sandstones from adjacent shales and siltstones. Inasmuch as the existence of large volumes of connate water is improbable in this case, diffusion must have played a role. The latter process always should be considered together with fluid movements caused by compaction. That silica migration from the shales actually occurred is demonstrable. Chlorite content increases with increasing depth at the expense of kaolinite and with liberation of silica. A mineral balance is presented in Table 3-LVIII (Fuchtbauer, 1967a). Transformation of kaolinite to chlorite upon burial from 1000 to 3000 m gives rise to 5.5% increase in SiOz content, which can be partly explained by the change of kaolinite to chlorite (4.3-1.1 = 3.2%) (see Table 3-LVIII). Additional silica might have been available from kaolinization of feldspar and from pressure solution of quartz at clay mineral contacts. (c) In the bottom part of Fig. 3-226, the sandstone is not filled with oil at the present time. The quartz diagenesis is so similar to that of the oil-filled rock (b), however, that a former oil impregnation is assumed. Fuchtbauer (p. 358) stated that one may find surprisingly high porosities in very deeply buried coarse-grained quartz sandstones. Here, pressure solution is more pro-
El k
TABLE 3-LVII Relationships among grain size, rounding and porosity (%) (after Fuchtbauer, 1967a, table I, p. 356)
Coarse-grained Fine-grained
z
8in
Profile A (Fig. 3-225)
Profile B (Fig. 3-225)
Profile C (Fig. 3-225)
median
rounding1 porosity
median
rounding1 porosity
median
rounding1 porosity
0.42 0.12
3.35 2.4
0.23 0.12
2.08 2.10
0.22 0.08
2.53 2.17
22 27
“he “median rounding” from Fig. 3-225 is listed.
27 26
26 30
: z U
2 0 z
2!
$
U
d 0
5
$
3
0
z
K.H. WOLF AND G.V. CHILINGARIAN
350
nounced in the fine-grained than in the coarse-grained sandstones, which confirms Weyl’s (1959)postulation. F’iichtbauer found that carbonate cementation (calcitic or dolomitic sandstones) is not uncommon. The “minus-cement porosities” (Heald, 1956) have been determined by adding the volume of the cement to the pore volume. This gives the total porosity if no cement were present at the particular stage of burial. The “minus-cement porosities” generally range from 30 to 35%, which correlates with a depth of 600-1100 m according to Fig. 3-222,A. Those carbonate-cemented sandstones with 40% minuscement porosity must have been cemented soon after deposition and prior to distinct compaction. This is also confirmed by the good preservation of labile heavy minerals in such layers (Drong, 1965), as a result of an absence of intraformational corrosion and leaching (Pettijohn et al., 1972). Such cemented layers are most common near the top and bottom of sandstone bodies adjacent to shales or mudstones. The anions are retained in the sandstones by filtration at the sand-shale contact (Fothergill, 1955). The degree of carbonate cementation in the center of sandstones are generally related to grain size. The carbonate-cemented sandstones are most common in the fine-grained layers that are overlain by, or intercalated with, coarse-grained sandstones (Fiichtbauer, 1967a,p. 358). Other authigenic minerals were also mentioned by Fuchtbauer. Authigenic kaolinite is often confined to the purest sandstones (e.g., left part of Fig. 3-227),where its growth was not hindered or masked by detrital clay minerals, whereas early diagenetic dolomite or calcite cementation may have hindered kaolinite precipitation (right part of Fig. 3-227).The pore fluids may also have a controlling influence, e.g., oil impregnation may hinder kaolinite crystallization. As shown in Fig. 3-227,the kaolinite is poorly developed in the shales, is of primary detrital origin, and can be easily TABLE 3-LVIII Silica balance during the late diagenesis of Jurassic shales (after Fuchtbauer, 1967a,table 11, p. 357) Depth (m)
Quartz (%)
1000 14.5 3000 20.0 Si02 content, % increase + +5.5 -
Feldspar (%) (estimated)
Kaolinite (%) (ca. 45% Chlorite (%) (ca. Si02) 28% Si02)
5 5
63 53.6
0
-4.3[=(63-53.6) 0.451
17.5 21.4 X
+1.1[=(21.4--17.5) 0.283
X
DIAGENESIS OF SANDSTONES AND COMPACTION
DOGGER BETA KAOLINITE KAOLINITE - FIRECLAY FIRECLAY Q DIT0,IN DOLOMITIC SANDSTONE + DIT0,IN SANDSTONE NEXT TO SHALE 0
”k10
.
,
20
c
; ”. ’ L 0
30
P E R C E N T
,
40 50 C L A Y
,. .-.;.-. 60
70
351
11
1
80
c 2 O p
of the clay minerals in the fraction <0.02 mm) and the total clay content (ordinate) in sandstones (left portion of graph) and shales (right) of the “Dogger beta” Formation from Hankensbuttel and Luben-West, Germany. The remainder of clay is illite. (After Fiichtbauer, 1967a, fig. 6,p. 358; courtesy 7th World Pet. Congr.) Fig. 3-228. Relationship between the porosity and permeability of the “Dogger beta” Formation from Gross-Hamburg, Meckelfeld, Vorhop, and Wittingen-South, Germany. These measurements are taken from W. Tunn as are those of Figs. 3-229, 3-230A, and 3-230B. (After Fiichtbauer, 1967a, fig. 7, p. 359; courtesy 7th World Pet. Congr.)
distinguished from the well-formed authigenic kaolinite of the cleaner sandstone$. Similar to the authigenic quartz, the kaolinite in these sandstones is concentrated close to the sand-shale boundaries, whereas diagenetic chlorite appears to be more common in the shales associated.with greywackes.
352
K.H. WOLF AND G.V. CHILINGARIAN
Von Engelhardt (1960,p. 83) has offered and discussed a formula in his book on the relationship between porosity and permeability for sands with a restricted size range. Fiichtbauer (p. 359)stated that the shape of the “point cluster” in Fig. 3-228can be explained as follows. The vertical variation results from grain-size differences, and according to Fuchtbauer (p. 359)the “fact that decrease in porosity and permeability does follow the drawn theoretical curves indicates that the specific surface is not higher in the denser samples than in the more porous ones. This is possible only because the clay content of the sandstones of the ‘Dogger beta’ is generally low. Porosity decreases only through denser packing because of greater depth of burial as well as by carbonate, silicate, anhydrite or pyrite cementation. Both processes, however, do not appreciably change the specific surface.” The points above the curve are indicative of samples cemented by pyrite, whereas those below the lower curve represent argillaceous samples. The data presented in Fig. 3-229is different in that the samples are from a narrow depth range and are uncemented. Porosities, therefore, deviate slightly to low values at the same permeability. There is a decrease in porosity with increasing permeability*, resulting from increasing grain size and roundness, which enhances mechanical compaction. According to Fuchtbauer, “the clay content (less than 20 p in size) is responsible for the position of the fine-grained samples (median size is less than 0.12 mm) in Fig. 3-229.It increases with decreasing porosity from 5 to 30%, thus increasing the internal surface area and, therefore, decreasing the permeability. The open circles in the lower part of the diagram represent relatively coarse-grained sandstone with clay seams from the uppermost part of the sandstone body.” The dashed line corresponds to the median grain size of 0.12 mm and the highest median found (0.42 mm) is used for calculating the data for the upper curve. The break between samples with median values higher and lower than 0.12 mm evidently lies in the fact that the clay content is different. As shown on the left-hand side of Fig. 3-222,there is a decrease in porosity with increasing depth for three groups of sandstones having different medians. These arenites are rich in dolomite and calcite grains and have calcite and clay matrices. Porosity decreases with increasing calcite cement, but is independent of the dolomite content. The decrease in porosity with depth is much more distinct than in many other sandstones, and the grain-
* k = 93/[5( 1+)2S,ZI
(3-12)
where k is the permeability; Q, is the porosity; and SO is the specific surface area, which depends on grain-size distribution and shape of the grains. In Fig. 3-228, the two solidline curves, which envelope most of the points, were calculated and plotted for two constant SO values. Accurate formulae relating porosity, permeability, and specific surface area have been presented by Langnes et al. (1972).
DIAGENESIS OF SANDSTONES AND COMPACTION
353
Fig. 3-229. Relationship between the porosity and permeability of the “Bentheimer”
Sandstone, Scheerhorn, Germany. (After Fuchtbauer, 1967a, fig. 8, p. 360; courtesy 7th World Pet. Congr.)
size effect is different from that in the quartz arenites in that the coarsest ones are also the most porous. The carbonates (average content of 15%)in coarser calcareous sandstones occur mainly as detritial carbonate grains (Fig. 3-222,A),and in the form of recrystallized calcilutite matrix which reduces the porosity in the finer sandstones. Figure 3-222 also illustrates that the average porosity of oil-saturated sandstones is higher than that of oil-free sandstones. Figure 3-230shows a clear relationship between porosity and permeability. If this relationship were due to an increased cementation, the points should lie on the smooth curve of equal specific surface area or even beyond its lower left part. The downward-increasing deviation of the data points from the curve, on the other hand, indicates that the specific surface area increases with decreasing porosity as a result of the sediments becoming finer grained and more argillaceous. The curves in Fig. 3-231resemble the curves (Bausteinschichten Sandstone) on the left-hand side of Fig. 3-222.There is a
354
K.H. WOLF AND G.V. CHILINGARIAN
Fig. 3-23OA. Relationship between porosity and permeability of the “Bausteinschichten” Sandstone (= calcareous - calcite + dolomite -, clayey arenites of the Tertiary Molase). Solid-line curve is for constant specific surface area. (After Fiichtbauer, 1967a, fig. 9, p. 360; courtesy 7th World Pet. Congr.)
different reason, however, for the dependency on the grain size. In Fig. 3-222,A, the Bausteinschichten sediments show an increase in calcareous matrix content with decreasing grain size, so that the net result is a lowering of the porosity, whereas in the case of sandstones presented in Fig. 3-231 it is the particularly high clay content that lowers porosity. As shown in Fig. 3-232, the clay content is lower than 10%only in coarse-grained sandstones (Md larger than 0.5 mm). In sandstones with less than 30%clay, more than half of the clay is authigenic kaolinite (Fig. 3-232; see also Table 3-LIX). In studying compaction, one has to distinguish four types of clay occurrences described by Fiichtbauer, i.e., (a) layers, completely separated from the sand; (b) seams, relatively well separated from the sand; (c) uniformly distributed clays, authigenic (late diagenetic) in origin; (d) uniformly distributed, detrital matrix and early diagenetic clay minerals; and (e) combina.
DIAGENESIS OF SANDSTONES AND COMPACTION
1
I
I
000
1
500
UPPER ,CARBONIFEROUS
1
. ..
355
I
I
1
Fig. 3-230B. Relationship between porosity and permeability of the Upper Carboniferous sandstones from the Emsland region, Germany. (After Fuchtbauer, 1967a, fig. 10, p. 361; courtesy 7th World Pet. Congr.)
tions of the above. The influences of diagenetic processes, e.g., compaction and migration of fluids, will be different in each case, as discussed by Fuchtbauer. In cases where organic material is present, mainly H20 (with dissolved humic acids) and COz are released during the first stage of coalification, causing an acid environment in which feldspars are dissolved and replaced by
U
0 I
l-
a W
0
POROSITY, 7-
Fig. 3-231. Relationship among porosity, maximum depth of burial, and median grain size in water- and gas-saturated sandstones of the Upper Carboniferous sandstones in western Lower Saxony. (After Fuchtbauer, 1967a, fig. 11, p. 362; courtesy 7th World Pet. Congr.) Each one of the horizontal lines (a t o i) corresponds to one drilling well; they were obtained from curves showing the relationship between porosity and median grain size for each borehole. Maximum burial probably corresponds to the present depth. In sections a, b, and g, however, greater depth was formerly assumed for geologic reasons. In the lower right, the 0.5-mm curve is compared with the “Dogger beta” quartz sandstones from Fig. 3-222. .
TABLE 3-LIX Mean clay contents (<20 p ) in relation t o the median diameter (after Fuchtbauer, 1967a, table 111, p. 361) Sandstone types
Median diameter (mm) 0.1
Platform: Dogger beta Valendis Molasse : Tertiary Upper Carboniferous
0.13
0.16
0.2
Mean clay content (5%) 2.5 12.5
1.8 2.5
12.5
10 20.5
22
1.3 1.6 8 19
1.0 1.0 6.5
17
DIAGENESIS OF SANDSTONES AND COMPACTION
-z
-90
I-
; a
-80
w
Y
g -70 0
%
W
2 -60 W
357
UPPER CARBONIFEROUS KAOLlNlTE KAOLINITE-FIRECLAY FIRECLAY a-METAHALLOYSITE CHLORITE + MICA 6zPERCENT FELDSPAR IN THE SAND FRACTION '20 p
0
m
(L
2 -50
1 %
& -40 W
a J
t -30
-I
L1
5 -20 W z z tUi-10
c
I
I
Fig. 3-232.Relationship among clay (<20 p ) content, median diameter, porosity, and the relative amounts of kaolinite, mica, and chlorite in clay fraction of Upper Carboniferous sandstones. (After Fuchtbauer, 1967a, fig. 12, p. 364; courtesy of 7th World Petroleum Congress.) The numbers inside the diagram represent feldspar content in percent. The ordinate shows the peak heights of the 001 mica line, the 002 kaolinite line, and the 004 chlorite line in percent of the s u m of these peak heights. Quantitative ratios, therefore, cannot be read from the graph. All values are for one well from about 2,500 m depth, except for the points in the lower graph.
+
-2.4% 2.5-4.9'l.
DIAGENESIS OF‘ SANDSTONES AND COMPACTION
359
kaolinite. This occurs particularly in coarser sandstones that have little detrital clay and are more permeable to fluids prior to kaolinite authigenesis. The numbers in Fig. 3-232 indicate that the feldspar content is lowest in the sandstones containing more kaolinite. Some kaolinite, however, may have formed independently of the feldspar breakdown. According to Fuchtbauer (p. 363), the chlorite in shales is largely of secondary origin, whereas in sandstones it is mostly detrital. In the latter case, some chlorite may be the product of alteration of a detrital matrix. During the second stage of coalification, coal seams only release chemically ineffective methane. This causes an increase in pH, probably leading to a slightly alkaline environment. During this stage, mica (sericite) instead of kaolinite replaces the feldspar. In two publications, Fuchtbauer (1967a,b) described the origin and regional variation of the sediments discussed above. Some of the information is summarized below, because similar approaches are a prerequisite for regional analyses of compaction and movements of compaction fluids. In Fig. 3-233 three fluvial dispersal patterns are illustrated: A. transport direction into the sedimentary environment of K-feldspar-rich coarser sand with tourmaline, apatite and rutile heavy minerals; B. dispersal is characterized by albiteoligoclase finer sand with a garnet heavy mineral fraction; C. dispersal pattern characterized by plagioclase finer sands without garnet. The rivers passed into the brackish-lagoonal parts of a saline basin. The milieu of the northern basin was brackish t o marine through geologic time. In the south there are mainly sandstones, whereas in the north there are predominantly siltstones with numerous interbedded, unsorted sandstone lenses. The sandstones and siltstones contain muscovite and chlorite in the ratio of 6 : 1 to 1 0 : 1. The higher chlorite content in the north may be due to the nature of the depositional environment. The following cements were present: quartz, anhydrite, dolomite to ankerite, calcite, feldspar, halite, clay minerals (muscovite, chlorite, mixed-layer illite-montmorillonite, vermiculite, glauconite), analcime, and barite. The distributions of some of the cements are shown in Fig. 3-233. Distribution of feldspar, quartz and anhydrite cements in the Middle “Buntsandstein” (Lower Triassic) sandstone. Each symbol corresponds to the mean percentage (by rock volume) of secondary cement in one well or surface section. The sections were obtained from boreholes, except for the islands of Helgoland, A3, A8, and the section immediately north of the Main River. The means are based o n thin-section estimations using samples of 0.1-0.9 mm median diameter combined with samples >0.2 mm median diameter, in order to avoid grain size influences. Isopachs of the total Bunter except rock salt are taken from Trusheim (1963) and Sorgenfrei and Buch (1964). A, B, and C are the main dispersal systems (transportation directions). Analc. = analcime; Nu,K = areas of secondary albite and potassium feldspar, respectively; GI = glauconite; M L = mixed layer illite/montmorillonite; V = vermiculite; b = baritocelestite, barite and celestite; H = halite (see Fiichtbauer, 1967a). (After Fiichtbauer, 1967b, fig. 1 ; courtesy Sed. Geol.)
K.H. WOLF AND G.V. CHILINGARIAN
360
TABLE 3-LX Contact strength of sand grains cemented by different minerals (after Fuchtbauer, 1967a, table IV, p. 366) Cementing material
Number of samples Mean (3) Standard deviation (s) Student-t-test
Albite Quartz Anhydrite Open
7 19 13 6
1.20 1.59 2.15 2.26
50.075 50.145 *0.21 2 0.19
*0.11 *O.l 50.19 k0.34
Fig. 3-233.According to Fiichtbauer (p. 366), the following cement paragenesis is the most common: (1)analcime, vermiculite, mixed-layer clays, and illite; (2)feldspar and chlorite; (3)quartz; (4)calcite; (5) anhydrite; ( 6 ) dolomite; (7) barite; and (8) halite. The above sequence of precipitation of the cements was determined from studying the fabric and measurements of “contact strength”. Based on Taylor’s (1950) method, the numbers of tangential ( a ) , long ( b ) , concavo-convex (c), and sutured contacts ( d ) of sand grains were counted separately for areas with albite, quartz, and anhydrite cement. The values of contact strength* are higher with tighter grain contacts (Table 3-LX). Inasmuch as contacts become closer with increasing degree of diagenesis and compaction, the relative age of cementing minerals can be deduced by comparing contact strength of sandstones that contain different cements. A very conspicuous relationship exists between grain size and type of mineral cement: quartz is usually present in finer sandstones, whereas anhydrite cement usually occurs in coarser ones. The supersaturation for quartz in the adjacent fine-grained and coarse-grained sandstones was equal, but because the quartz grains acted as nuclei for precipitation and fine sandstones had a larger specific surface area (i.e., cm2/cm3 of bulk material present), more SiO &m was precipitated in the fine-grained sandstones than in the coarse-grained ones. Inasmuch as the porosity was equal in both fine-grained and coarse-grained sandstones, the porosity w a s reduced much more in the fine-grained sandstones. The already existing difference in permeability was further accentuated, so that the anhydrite solutions of later origin moved preferentially through the coarse-grained sandstones and cemented them. Thus, the grain size may control the differential cementation
____
* Contact strength = ( l a + 2b + 3c + 4d)/(a + b + c + d )
(3-13)
where a is the number of tangential contacts, b is the number of long contacts, c is the number of concavo-convex contacts, and d is the number of sutured contacts.
DIAGENESIS OF SANDSTONES AND COMPACTION
361
that, in turn, could control the rate and degree of compaction. The content of albite cement is higher in the fine-grained sandstones, because they contain more detrital plagioclase than the coarser sandstones and, thus, more nuclei are available for the precipitation of albite (Fig. 3-234). In fourteen out of eighteen occurrences, dolomite cement is present in fine-grained sandstones. Chronologically, NaCl appears as one of the last cements formed and occurs predominantly in the fine-grained sandstones, because the coarser sandstones have been cemented in the meantime by anhydrite. The contents of all other cements are independent of grain size. The distributions of the individual cements show clear regional differences. Vertical differences in cementation in individual profiles are of a more local character, but in general it seems that the deeper parts of the sandstones are more extensively cemented. The distribution of secondary feldspar is illustrated in Fig. 3-233,a. South of the Na-K line, only K-feldspar overgrowths are present; albite is more predominant in the north. It may be no coincidence that the Na-K line coincides with the boundary between the fhviatile basin in the north and the brackish-saline basin in the south. The 30
I
1
I
I
MEDIAN DIAMETER, rnrn
Fig. 3-234. Relation between feldspar content and median diameter of sandstones, based on X-ray estimations by Mrs. Goldschmidt; samples crushed to <35 p. Open circles = albite near A3 (see Fig. 3-233); open circles with dots inside = albite near A8 and east of Hannover; solid circles = potassium feldspar in both areas. (After Fuchtbauer, 1967b, fig. 2; courtesy Sed. Geol.) More or less thick rims of authigenic albite are bordering detrital grains of both albite and potassium feldspar north of the Na-K boundary line in Fig. 233a. Small rims of authigenic potassium feldspar are restricted t o the south of this boundary, which corresponds to the boundary between brackish-marine environment to the north and fluviatile environment to the south.
362
K.H. WOLF AND G.V. CHILINGARIAN
depositional milieu has apparently determined the secondary feldspar genesis, i.e., the earliest cement formed. In the terrestrial fluvial milieu, in which mica and K-feldspar decomposition through weathering occurs, K is enriched in the pore solutions as a result of weathering and the K is adsorbed on the fine-grained fluvial sediments. During compaction, K is remobilized and moves into the sandstones to be precipitated as K-feldspar. Under the conditions existing in the northern basin, where evaporation took place, the pore fluids became enriched in Na content (the concentration was even higher than that of K in the southern basin), leading t o extensive albite genesis. Analcime distribution, which is also shown in Fig. 3-233,b, is similar to that of albite. It is most abundant in the coarser-grained deposits and is confined to the weakly-evaporitic basin sections. Although quartz cement is quite widespread in its occurrence (Fig. 3-233,b), it is more enriched in the fluvial deposits; this becomes more obvious on examining Fig. 3-233,c. The incoming fresh water displaced the influence of the saline milieu. Also, the pore solutions became poorer in the sulfate content, with the resulting decrease in the amount of anhydrite cement. The precipitation of the silica was able to continue. The origin of the silica can be determined through thin-section examinations. In clay-rich sandstones, as well as in siltstones, where two adjacent quartz grains are separated by a clay film, stylolitic contacts are common and indicate dissolution of silica. In contrast t o quartz-quartz contacts, the quartz-clay contacts permitted removal of the silica in solution without interruption, so that pressure solution is more intense in clayrich sandstones (cf. papers in Spec. Publ., 7, S.E.P.M., 1959). The dissolved silica then moved into more pure sandstones t o be precipitated on the cleaner quartz surfaces. Various considerations by Fuchtbauer (p. 367) indicated that the main silicification began when the sandstone had a porosity of 20-30% and was at a depth of about 1OOOm. Chemical compaction becomes apparent below a depth of about 1000 m. Anhydrite is the second most common cement and shows a distribution complementary t o that of quartz (Fig. 3-233,c). It is most common in the central part of the northern depocenter and is absent in the fluvial basin. Fiichtbauer estimated that the anhydrite was precipitated when the sediments were at a depth of 1000 m below the surface. He attributed cementation to large volumes of sulfate in pore fluids, which moved during compaction from the fine-grained sediments (clays) into the sandstones. The anhydrite was precipitated when the NaCl content of the pore fluids reached a certain limit necessary t o exceed the solubility of CaSO,, as a result of the temperature increase during burial and the pressure decrease upon expulsion of the fluids from the fine-grained units. These units were under pressure above the normal hydrostatic pressure of the surrounding sandstones. There are two possible sources for the calcium sulfate: (1)enrichment through
DIAGENESIS OF SANDSTONES AND COMPACTION
363
evaporation in the lagoonal basins, and (2)compaction fluids from the Zechstein evaporite deposits. The original gypsum recrystallized to anhydrite at a critical depth of 500-1000 my with liberation of CaS04-saturated water which could have migrated upward into the sandstones. Baritocelestite, barite, and celestite ( b in Fig. 3-233,c) show a distribution similar t o that of anhydrite. As shown by Fuchtbauer, many of the diagenetic alterations are directly related to mechanical and chemical compaction. Some of the chemical constituents of the cements are of local origin, whereas others appear to have been derived from more remote sedimentary rocks undergoing compaction. Blanche (1973) observed that particularly in fluvial and sabkha facies, the clay matrix is composed predominantly of illite together with kaolinite, chlorite, and traces of montmorillonite. These minerals are derived from the decomposition of feldspars and micas. The permeability in these rocks decreases with decreasing grain size, whereas the porosity remains unchanged, sorting being its most important controlling factor. Following the reasoning by Fuchtbauer (1967,a,b) and Taylor (1950), Blanche expected the textural changes with increasing compaction as outlined by these two researchers. The maximum values of porosity and permeability, however, appear to be little affected by depth of burial down to 4000-4250 m. Below this depth a rapid decrease was observed, i.e., the porosity gradient was 1.3% for each 305 m of burial, which is in agreement with values published for other sandstone formations. Blanche also found that the depositional nature of the sands set physical limits on the reservoir potential: optimum values tended to occur in eolian dune facies, whereas the lowest reservoir parameters were present in the fluvial and sabkha facies. Powers (1967) treated the release mechanisms of fluids from clay-rich sedimentary units (see also Burst, 1969, for another example). Although the present chapter is confined to coarser sediments, the properties and behaviors of claystones, mudstones, shales, and bentonites (= pyroclastic deposits altered usually to ‘montmorillonite-rich accumulations) cannot be completely ignored. At least a brief treatment is in order here, because the stages at which certain amounts of fluids are released from clayey deposits determine how, when, and in what quantities these fluids are made available to the sandstones during compaction. The fluids from the finer sediments may or may not be related to compaction, they nevertheless influence diagenesis in general, including mechanical and chemical compaction, as well as cementation, decementation, and oil and ore-fluid migration. Together with studies such as done by Hitchon (1968) on the total volume of sandstones, shales, and limestones within certain sedimentary basins, the data on the release mechanisms of fluids from the clayey and other sediments and the information on the compaction history of the sedimentary pile will enable the
K.H. WOLF AND G.V. CHILINGARIAN
364
investigator to calculate the amounts of water, hydrocarbons, and, possibly, ore fluids that may be available per cubic meter of rock. At least the total volume of fluids could be estimated. Some of these studies have already proved of practical value in estimating the total amounts of petroleum and ores present in certain regions. Powers (1967)offered a compaction history (Fig. 3-235),which is based on laboratory and field data and the current knowledge of clay mineralogy. His conclusions can be summarized as follows: (1)When montmorillonite clays are buried to a depth of about 3000 ft, most of the water is expelled, except for the last few bound layers between the basal clay surfaces. This bound water may comprise nearly 50% of the volume of the clayey rock, and, apparently, cannot be squeezed out by further increase of burial pressure. (2)At a depth of about 3000 ft, the effective porosity and permeability are essentially zero for the mudstone. The clays and fixed water are "wrapped-around" the sand and silt particles. (3)A change from montmorillonite to illite begins at a depth of about 6000 f t and continues at an increasing rate to a depth of about 900010,000 ft, where no montmorillonite is left. This alteration mechanism is accompanied by desorption of the last layers of bound water from the clays and their release as "free water". (4) The fluids are released suddenly from the montmorillonite in the deep
WT11.mCH
rmJ
IO*I"O.ILLU(I.
DM
ILLIIC llU1.
111.0
"L..
YD I l 0 L I " l T l
Fig. 3-235. Compaction history of different clay minerals when deposited in marine environment and its probable relation to release of hydrocarbons from mud rocks. (After Powers, 1967, fig. 3, p. 1246; courtesy Am. Assoc. Pet. Geologists.)
DIAGENESIS OF SANDSTONES AND COMPACTION
365
subsurface, maybe during the clay-mineral transformation(s) and origin of mudstone f issility . (5)This sudden release of fluids would be absent from illite and kaolinite deposits because of the absence of a mineralogic change. Normal fluid movements by compaction would have occurred earlier. (6)During the montmorillonite-to-illite change, desorption of interlayer hydrocarbon layers and, possibly, trace elements takes place. (7)Below the “no-montmorillonite level” (Fig. 3-235),small amounts of water continue to be lost from the shale as the montmorillonite fraction of the mixed-layer clays collapses to form illite. This occurs at about 14,000 f t in the Gulf Coast area. Magara (1973)studied (1)relationship between differential shale compaction and depth, (2)relationships among porosity, permeability, density, and depth, (3)relationship between clay composition and depth, (4) relationship between the rate of sedimentation and compaction, (5) amount of fluid expelled during compaction, and (6)porosity distribution in shales as an indicator of permeability of associated sandstones and coarse limestones. According to Magara, “shale porosity distribution in incompletely compacted shale zones also may be affected by the permeability and the extent of adjacent sandstone or carbonate rock bodies. A sharp decrease of porosity in shales close to such rock bodies would suggest that relatively large volumes of fluids have been expelled from the shales into the adjacent sandstones or carbonates. If this expelled fluid volume is large, the possibility of hydrocarbon accumulation in such sandstone or carbonate rocks is considered to be favorable.” This has been confirmed by Magara’s studies of Mesozoic oil and gas pools in western Canada.
Compaction and stratigraphic correlation (Conybeare, 1967) Conybeare (1967)considered the results of compaction on stratigraphic correlation and consequent sedimentological-environmental interpretations. Inasmuch as differential compaction of a column of deposits in a sedimentary basin depends, among other variables, on the proportions of clay and sand, Conybeare theoretically considered (pp. 334-337) stratigraphic crosssections composed of different sand/mud ratios (Figs. 3-236 and 3-237). With increasing depth of burial, the total thickness of a section is reduced and the attitude of the sandstone beds is modified. The latter changes from an original attitude to one with a compactional slope, which may influence the fluid migration. In his model, Conybeare assumed that although the mud/sand ratios differ in sediment columns A and A’ (Fig. 3-236),which contain 1150 f t of solid sediments each, the rate of sedimentation was the same. The same assumption applies to cross-section A’--A3. For practical
K.H. WOLF AND G.V. CHILINGARIAN
366
0 I
lo00
II
!i ZOO0 1
aow
SAME LAYERS BURIED UNMR 2500' CLAY SOLIDS (UPPER 3500') Qf COMPACTIONAL SLOPE FROM 1 W 2
SHOWING CHANGE
4000
uoo CLAY :1000'SOLIDS
CLAY :500'SOLIDS
SAND :150'
SAND :650'
CLAY.lOO0'SOLIDS
SAND :1000'
Fig. 3-236.Schematic representation showing the effect of compaction o n sections A-A and A 2 - A 3 when they are initially buried under 3500 f t of clay muds. Sedimentation rates, in terms of total effective solids (100 f t of sand is effectively 100 f t of solids), are the same in columns A and A l , and in A 2 and A 3 . Columns A , A l , A 2 and A 3 have different sand/mud ratios and show variable degrees of compaction depending on the amount and depth of burial of mud layers. Changes in slope from depositional (1)to compactional (2), shown in hatchures between lines 1 and 2, are referred to as the "bellows effect". (After Conybeare, 1967, fig. 5, p. 338; courtesy Bull. Can. Pet. Geol.)
purposes the sandstone framework was considered constant, i.e., uncompactable, which is probably not true in nature (see Chilingarian et al., 1973).If the volume of sand is assumed to be constant, the degree of change of compactional slope will depend on the mud/sand ratio and depth of burial. Figure 3-237shows the thicknesses of sand and mud in columns A , B and B1 of a cross-section through a sedimentary basin. The sedimentation rate at B and B1 is twice that at A , and column B1 consists only of clay. As to the stratigraphic correlation, the sands at a depth interval of 3000-3250 f t in section B are equivalent to the sands at a depth interval of 1750-1900 f t in section A . The compactional slope between A and B at a depth range of 1750-3250 f t is 90 ft/mile or about 1degree. Conybeare calculated (see his table 2, p. 342) that an initial deposit of 837 f t of clayey mud contains 500 f t of solids, so that the amounts of clay solids in Fig. 3-237given as 1625, 2250 and 4550 ft for the sections A, B, and B1,respectively, were originally 2678, 3766,and 7533 f t prior to compaction. But as shown in Fig. 3-237, the actual thicknesses of sections A, B,and B1 are 3000,5250,and 5800 ft, respectively. Subtracting the sand thicknesses of 650 and 2300 f t in sections
367
DIAGENESIS OF SANDSTONES AND COMPACTION A
1000
2000
3000
B‘
B
F
4000 CLAY
. 1625’ SOLIDS
SAND : csd 5000
CLAY : 2250’ S 6000
SAND : 2300’
O
L
I
D
-
A
CLAY : 4550’ SOLIDS
Fig. 3-237. Schematic representation showing a section through part of a sedimentary basin where columns B and B’ each contain twice as much effective solids as column A , but have different ratios of sand to mud. The “bellows effect” illustrated in Fig. 3-236 results in expulsion of fluids from muds into sands; updip migration of these fluids is shown by arrows. (After Conybeare, 1967, fig. 6, p. 339; courtesy Bull. Can. Pet. Geol.)
A and B, the actual thicknesses of the clay deposits in sections A, B, and B1 are 2350, 2950 and 5800 ft, respectively. Subtracting 2300 from 2678, 2950 from 3766, and 5800 from 7533 results in 328, 816, and 1733, which represent the number of feet of water lost during compaction of the three sections. As Conybeare (1967, p. 337) stated: “Although the sedimentation rate in sections B and B1 is the same, section B1 has lost an additional 917 f t of water in comparison with that of section B, and B has lost 488 f t of water more than A ” (Fig. 3-237). Arrows in Fig. 3-237 indicate the directions of possible initial, updip fluid movements within the sand beds, depending on sedimentation and compaction rates. Conybeare called the squeezing of water from the clays into the sands and the consequent changes in the angles of compaction slopes the “bellows effect”. Updip fluid movements commonly occur toward a landmass, but need not necessarily be so. Penecontemporaneous or subsequent tectonism can alter or even reverse the compactional slope. Inasmuch as it has been shown by Chilingarian et al. (1973) that sands are just as compressible as clays, the above-described assumptions and calculations of Conybeare (1967) should be carefully reexamined. Future research
368
K.H. WOLF AND G.V. CHILINGARIAN
work hopefully will resolve this problem and determine under what conditions sands can be considered uncompactable in contrast to conditions when they are as compressible as muds. Baldwin (1971), as well as Conybeare (1967), pointed out that when muds and claystones intertonguing with sands are compacted, the differential compaction (based on the percentage of clay plus silt within the units) causes progressive distortion of the geometry of the initial depositional surfaces. Both of the above-mentioned researchers discussed “decompaction”, which is an analytical technique useful in projecting or extrapolating back the stratigraphy to an assumed earlier condition, i.e., restoration of the original depositional stratigraphy prior to compaction. Compaction is commonly expressed as a change in porosity (= ratio of pore volume to bulk volume) or as void ratio (= ratio of pore volume to volume of solid grains). “Grain proportion” (= ratio of volume of solid grains to bulk volume; Robertson, 1966, 1967) is the complement of porosity. Inasmuch as the volume of solid grains remains constant during compaction, grain proportion offers a simpler frame of reference than porosity. Baldwin proposed that decompaction can be accomplished by multiplying the present thickness of a compacted unit by a “decompaction number”, thus restoring the stratigraphic contacts to their precompaction position. The decompaction number, D, is expressed as: (3-14) where he =-earlier thickness, hp = present thickness, (pp = present porosity, earlier porosity, G, = present grain proportion, and Ge = earlier grain proportion. The grain proportion and porosity are expressed as decimal fractions in eq. 3-14. An approximate value of Gp can be obtained by measuring the dry bulk density of a sample in the laboratory and equating this to grain proportion in Fig. 3-238A, where the solid grain density is 2.66 g/cm.* Table 3-LXI presents the obtained data for North Sea deep-sea sediments (Hamilton, 1969a, tables 1and 2), and includes porosity and solid grain densities. At zero depth of burial (= precompaction stage), the value of Ge can be assumed to be 0.22 or 2276, which is equivalent to an initial porosity of 78% which corresponds to the weighted average of the data presented in Table 3-LXI. It is consistent with the data from Fig. 3-239, with a possible * 5 % @e =
* This is an average value used in soil mechanics (Hamilton, 1969b, p. 26) and is also the weighted average of data in Table 3-LXI.Most mud-forming minerals have an average solid grain density of 2.6-2.7 g/cm3 (Hedberg, 1936, p. 279).
369
DIAGENESIS O F SANDSTONES AND COMPACTION
0.5
1.0 1.1
BULK DIYUlY
a.0
a.s
1 DWTN Of DURIAL (fern0
I,./..)
VOID 14110
Fig. 3-238. Depth of burial graphs for clay and shale. A. Graph for converting bulk density t o grain proportion (and porosity). B. Depth of burial versus grain proportion (and porosity). For each curve, multiply depth of burial values by number shown for that curve. C. Graph for converting void ratio t o grain proportion (and porosity). (After Baldwin, 1971, fig. 1,p. 294; courtesy J . Sed. Petrol.) TABLE 3-LXI Porosity and solid grain density of some deep-sea sediments (after Hamilton, 1969a, tables 12) Sediment type
Continental Terrace Sand-silt-clay Clayey silt Silty clay
Number of samples
Solid-grain density Wee)
(shelf and slope) 17 2.71 40 2.713 17 2.69
Porosity average
(%I1
____-
standard error of the meana
67.5 75.0 76.0
1.66 0.87 0.74
Abyssal Plain (turbidites) Clayey silt 15 Silty clay4 35 Clay 2
2.61 2.55 2.67
78.6 85.5 85.8
1.53 0.49
Abyssal Hills (pelagic) 3 Clayey silt Silty clay 32 Clay 6
2.58 2.71 2.76
76.4 79.4 77.5
-
0.77 1.35
' Salt-free porosity; about 1%greater than without correction for dried salt (Hamilton, 1969b, pp. 25-27). Standard deviation = standard error X (number of Five samples (Hamilton, 196913, table D-1). Hamilton (1969b, table D-2).
*
370
K.H. WOLF AND G.V. CHILINGARIAN
variation in this initial porosity value, because the actual value depends on several variables. These include depositional process and rate and environment of accumulation, which determine the fabric, mineralogy, particle-size variation, organic content, and geochemistry of the deposit (Hamilton, 1969a,b; Meade, 1966). Figure 3-238Bpresents data for inferring the values of G, used in decompacting procedures to a condition of intermediate depth of burial. Where the value of G, has been obtained, the data can also provide the maximum amount of overburden and, as pointed out by Johnson (1950), the discrepancy between G, and the present overburden may indicate the amount of erosion at a surface of unconformity. The data in Fig. 3-238Bpertains to the approximate relation of grain proportion and porosity to depth of burial for clayey and muddy sediments, assuming that abnormal pore pressures were absent. This figure is flanked by conversion charts of bulk density (Fig. 3-238A)and void ratio (Fig. 3-238C).The average curve was obtained from the smoothed-out curve B in Fig. 3-239. 0
20
-x C
L
40
Z
2 c I
P z4
60
I
0
80
1oc DEPTH OF BURIAL ( f eat )
Fig. 3-239. Published porosity-depth data. See table 2 in Baldwin, 1971, p. 295. (After Baldwin, 1971, fig. 2, p. 295; courtesy J . Sed. Petrol.)
DIAGENESIS OF SANDSTONES AND COMPACTION
371
Using the above data, Baldwin (1971,pp. 296-299) proceeded to present examples of restoring compaction structures in interbedded shales (or argillites) and sandstones to their original precompaction geometric configuration. Although the above information is directly related to compaction of the clayey constituents in a stratigraphic unit, it has been presented here with the view that decompaction of the fine-grained units will also restore the interbedded sandstones to their primary, depositional, pre-compaction position, as demonstrated by one example, presented by Baldwin (see his fig. 4,p. 299).The stratigraphic and structural relationships between channel sandstone bodies, with their surrounding fine-grained accumulations, become much clearer after decompaction. As observed today (i.e., after compaction), the original outlines of the channel sandstones and the complex interbedding of the various lithologies can be visualized only with difficulty. Perrier and Quiblier (1974,p. 508) pointed out, however, that the decompaction number of Baldwin only relates to layers of infinitesimal thickness. Tmrnit (1968)has done extensive studies on stylolites (see pp. 133 and 139) and has discussed the various stages of pressure solution during the development of a subsiding sedimentary basin. For example, he presented a conceptual model (Fig. 3-240)on how the regional development of stylolites controls the movement of compaction fluids, an understanding of which is very important in the study of the origin of hydrocarbons and certain types of ore deposits. During the subsidence of the sediments in the center of the basin, stylolites are being formed in the carbonate rocks and, possibly, in quartzites. Carbonates require about 40 m and the quartzites about 1000 m
Fig. 3-240.Comparison between the unbedded or rarely-bedded massive limestones on highs (e.g., geanticlines) and the development of thinly bedded limestones in the basin (e.g., geosyncline). At the beginning, the dissolved material is able t o move vertically in the upward direction with the compaction fluids. As soon as cementation reduces the porosity, a strong vertical flow is eliminated. Consequently, the removal of any dissolved material in pore solutions occurs along the widespread pressure-solution surfaces in the direction of the structural highs. This material is then either precipitated as pore-filling cement or escapes into the ocean surface water. Solid arrow = flow direction of the compaction fluids; large open arrow = direction of basin subsidence. (After Trurnit, 1968, fig. 3,p. 382;courtesy Sed. Geol.)
372
K.H. WOLF AND G.V. CHILINGARIAN
of overburden for pressure solution to commence (see p. 378 in Trurnit, 1968). At the same time, the dissolved materials in the compaction fluids are carried upward into higher horizons where the pore space may become cemented and, consequently, the vertical fluid movements become increasingly impeded. The widespread occurrence of the pressure-solution surfaces, which ascend from the basin onto the geanticlinal ridges, then allows horizontal and diagonal fluid movements. The solutions, therefore, move towards the geanticlines and precipitation of cement takes place there, or the fluids are subaqueously expelled into the ocean water causing increased concentration of sea water in the chemical elements. Trurnit even suggested that microbrecciation of the carbonate sediments may be the result of expulsion of the compaction fluids at the flanks of the geanticlinal ridges. Migration of compaction fluids (Magara, 1968)
Magara (1968) studied migration of compaction fluids in marine Miocene mudstones and associated volcanic and pyroclastic rocks. His approach may serve as an example that can be applied in the future elsewhere. He found that the overlying and underlying mudstones were the source beds and that the water and hydrocarbons moved into the volcanic masses as a result of differential compaction. Magara (p. 2466) stated that it is important in petroleum exploration to determine the directions of movements and amounts of compaction fluids. One might add that this is also applicable to metalliferous ore exploration in cases where the deposits were formed by fluids of compaction and were controlled by stratigraphy. According to Magara (p. 2467),many investigators prepared lithofacies, biofacies, isopach, and subsurface structural maps to determine the migration and accumulation of hydrocarbons in a basin; however, no method that allows the determination of the direction of flow and the amount of the compaction fluids has been proposed. He, therefore, offered two methods to accomplish this, based on the horizontal and vertical porosity distribution, and porosity differences before and after compaction. A similar approach was used by Jobin (1962) in the study of Colorado-Plateau-type uranium deposits in sandstones. His porosity, permeability, and transmissivity data, however, were not related to compaction and, therefore, were not used to determine the migration directions and volumes of compaction fluids. Magara (1968)used the data available from the published literature (Fig. 3-241)on the empirically-determined relations between depth of burial and the degree of compaction of shales and mudstones. The curves show a marked decrease in porosity at shallow depths due to mechanical compaction. Under normal conditions, as compaction occurs and porosity decreases, fluids are expelled from the sediments. If the escape of the compaction
DIAGENESIS OF SANDSTONES AND COMPACTION
37 3
DEPTH, m
Fig. 3-241. Comparison of depth-porosity relations in several regions: Oklahoma (Athy, 1930); Venezuela (Hedberg, 1936); Gulf Coast '(Dickinson, 1951); Japan (Hosoi, 1963). (After Magara, 1968, fig. 5, p. 2473; courtesy Am. Assoc. Pet. Geologists.)
fluids is prevented or lowered, however, compaction may not be great and a high-porosity, high-fluid-pressure condition would result". Hence, more fluids would be expelled from rocks that can undergo compaction than those that do not. Magara stated (p. 2468) that, as a general rule, water is expelled from argillaceous beds into permeable carrier beds and then moves from localities of greater fluid expulsion to localities of smaller degrees of fluid expulsion. For this reason, it is possible to determine directions of compaction fluid movements from horizontal porosity distribution of the mudstone units above and below the carrier beds. Sonic and gamma-gamma logs (formation density logs) were used to determine porosity. Figure 3-242shows one example of porosity distribution of the mudstone overlying a reservoir composed of andesite agglomerate, and indicates differential regional compaction that resulted in controlling the directions of movement of the compaction fluids as shown by the arrows. In the present case, the arrows suggested the locations where oil or gas has accumulated. In other studies, maybe ore bodies could be located. Magara determined which beds acted as barriers and which ones underwent compaction. To calculate the volume of water expelled from the source rocks from a given time up to the Recent, one has to know the original thickness of the source beds. Magara calculated these thicknesses, after determining the normal porosity trend shown in Fig. 3-243.In Fig. 3-244,he reconstructed the formation thicknesses at the end
* For some data on high-pressure zones, see pp. 72-74.
i
Y,WI -WIT* lA
Y~NAMI-&UCHI I
.I
I
-
Fig. 3-242. Differential compaction map o f mudstone beds overlying agglomerates in the Nagaakaregion. (After Magara, 1968,fig. 8, p. 2477; courtesy Am. Assoc. Pet. Geologists.)
DIAGENESIS OF SANDSTONES AND COMPACTION
10
SHlUNJl SK-21 MUDSTONE ,POROSITY, 1
. 30
44
@
50
/
'
375
70% Uoourna Group
Hairune
I
Hahiyarn Famalia
-
Shiiya FOrmcllior
klKbju-m
FOrMtiW
N am lMi Farnatim
Shiinji Tuff
mx Fig. 3-243. Mudstone porosity values in Shiunji SK-21borehole. (After Magara, 1968, fig. 13, p. 2483; courtesy Am. Assoc. Pet. Geologists.)
376
K.H. WOLF AND G.V. CHILINGARIAN
of deposition of the Shiiya Formation, again by using his mathematical technique. The present porosity distribution after burial is also plotted in Fig. 3-244.The Teradomari mudstone just above a depth of 1340 m apparently has acted as a barrier to water movement, so that the water squeezed from the mudstones below this level would have moved downward and the water expelled above it moved upward to the surface and escaped into the sea. In Fig. 3-245,Magara illustrated how he used a simplified method of calculating the volume of compaction fluids. Assuming that the volume of solids remains more-or-lessconstant after compaction, the following formula can be used: V(1-
6 )= V‘(1-8)
(3-15)
where V = volume of sediments before burial, V‘ = volume of sediments after burial, 3 = average porosity before burial, fractional, and $’ = average porosity after burial, fractional. By determining first the after-burial, present-day vertical porosity distribution and then by assuming a normal porosity trend, the average porosity below the barrier before burial, 5,can be determined (Fig. 3-245).Also if the average porosity below the barrier after burial, T‘, and the after-burial, present-day volume of the sediment below the barrier, V‘, is known, the mudstone volume before burial, V, can be estimated (see formula 3-15): v = V”(1 --$)/(l -&)I (3-16)
It follows then, that the volume of compaction fluids, which moved downward, Vd, can be calculated:
v, = a v = (6- 6’) V’(.I -$)/(l - 5)
(3-17)
In applying his method over an entire region, Magara subdivided the area into many blocks and determined the total volumes of compaction fluids which moved downward within these blocks. Then, he was able to determine the directions of movements. From the present writers’ point of view, it seems that in the future such regional studies should take into consideration any data available on different compressibilities of various types of mudstones, sandstones, pyroclastics, etc. Fons and Holt (1966),for example, pointed out that montmorillonite shales (most probably a bentonite formed by the alteration of a tuff) are more resistant to compaction than illite and kaolinite. This, in turn, makes the montmorillonite also more resistant to internal fluid migration, which is related to the degree of water adsorption (see also Chilingar and Knight, 1960).As to the compactability of the different types of pyroclastic deposits, very little is known, but progress is being made at present (see Chapter 6 ) . The composition (i.e., vitric, lithic or crystal constituents), degree of zeolitization, clay neomorphism and other diage-
DIAGENESIS OF SANDSTONES AND COMPACTION SHIUNJI HUDSTONE
S1(-21
WROSllV,
0
377
70%
-Shiiya Farnoti0
-
Fa-0
Nmatanl M i
._--- F
Shiunji Tuff
mv. Fig. 3-244. Reconstruction of formation thicknesses at the end of the Shiiyan stage and the mudstone-porosity plot of Shiunji SK-21 borehole. (After Magara, 1968, fig. 14, p. 2484; courtesy Am. Assoc. Pet. Geologists.)
netic alterations, the rate of deposition, geologic time, and a host of other factors will determine the textural and structural characteristics of the tuffs, which in turn will influence their compactability. For an example on the calculation of the volumes of various lithologies (i.e., sandstones, shales, carbonates, and evaporites in this particular instance) within a sedimentary basin of western Canada and the total volume of fluids present, the reader may wish to consult the publication by Hitchon (1968). It should be remembered that Magara (1968) assumed a normal pressure gradient with depth. Some basins, or localities and units within particular basins, however, have abnormal pressures. Among others, Von Engelhardt (1960, p. 42) has pointed out that pore solutions are expelled. With a
378 MUDSTONE POROSITY
:.’
K.H. WOLF AND G.V. CHILINGARIAN
A S U M E D NORMAL POROSITY TREND
I IP W
D I-
t
W
w cn
a
i
RESERVOIR
Fig. 3-245. Diagram showing method of calculating volume of compaction c s r e n t . D ’ = thickness of mudstone after burial; V’ = Kolume of mudstone after burial; @ = average porosity below the barrier before burial; @’ = average porosity below the barrier after burial; and A@ = mean mudstone porosity decrease during compaction = (After Magara, 1968, fig. 15, p. 2485; courtesy Am. Assoc. Pet. Geologists.)
e@’.
decrease in pore space, the rate of porosity decrease with depth depends on pressure differences and the ability of clayey sediments to let duids pass (= permeability). Inasmuch as the claystone permeability is very small, the rate of compaction may not only depend on the rate of subsidence and increase
DEPTH, m
Fig. 3-246. Above-normal pressures in Tertiary sandstones in Louisiana. Solid circles = measured pressures; open circles = estimated pressures. (After Dickinson, 1953 ;courtesy Am. Assoc. Pet. Geologists.).
DIAGENESIS OF SANDSTONES AND COMPACTION
379
in the overburden pressure, but also on the rate of escape of the pore solutions. In clayey horizons that have not been completely compacted and where fluids are still moving upwards, the fluids in certain horizons must be under pressure which is higher than the normal hydrostatic pressure there. In the presence of interbedded porous sanstones the pressure will be particularly higher here. In the Tertiary clay-rich sediments of the Ventura Basin in California, at a depth of 2700 m the sandstone exhibits a pressure gradient of 0.23 atm/m (Watts, 1948). This pressure corresponds approximately to the weight of overlying sediments. In the Tertiary basin of the Gulf Coast of Louisiana abnormally high pressures have been also measured (Dickinson, 1953), as shown in Fig. 3-246. The lowest measured normal hydrostatic pressure was 0.107 atm/m, whereas the highest abnormal pressure was 0.21 atm/m. The highest pressures were present in sandstones within or beneath larger clay-rich units. A whole chapter was devoted by Rieke and Chilingarian (1974) t o the overpressured formations and their origin. Study on grain-contact types (Phipps, 1969) Phipps (1969) presented the results of a valuable study which showed: (a) the relative importance of grain contact types with burial; (b) the use of a “grain contact index”; (c) that compaction studies can assist in determining the time of cementation by chemical precipitation; (d) that differential compaction may be a function of differential chemical cementation from layer t o layer; (e) that the complex compaction history is a result of both cementation and decementation. A large proportion of Phipps’ data is given below. He studied an Eocene section of the Maracaibo Basin in Venezuela, comprised of a thick sequence of interbedded sandstones and shales. The former are quartzose, generally with more than 90% quartz grains. Chert may comprise up to 10% whereas the amount of interstitial clay is minor and seldom exceeds 5%by volume of sands. Micas, feldspar and pyrite are the most common accessories. According t o Taylor (1950), pore-space reduction and compaction can occur through: (1)solid flow of material under pressure involving grain fracture and distortion, and (2) solution and recrystallization-redeposition of material under pressure. Phipps stated that the sandstones studied by him display abundant evidence for mechanical processes of compaction, which include boehm lamellae, tension cracks in distorted quartz grains, displacement of twin lamellae in feldspars, cleavage along microfractures in feldspars, distortion and bending of micas between quartz grains, and “mushrooming” * of
* “Mushrooming” refers to the relationship between two grains where one became differentially dissolved and began to partly “wrap-around” the second grain.
380
K.H. WOLF AND G.V. CHILINGARIAN
quartz grains. Pressure solution and redeposition of silica cement or quartz overgrowths do not appear to be common and the reduction of porosity apparently took place mainly in response to mechanical processes of compaction. Although long contacts are common, secondary quartz overgrowths are rare. The relative ease with which the sandstones can be mechanically disaggregated indicates that cementation by secondary silica is not significant. The average total number of contacts per grain in a thin section is useful in assessing the degree of compaction (see Taylor, 1950, and section on textures in this chapter, pp. 156-163. This number is a function of grain shape, angularity, and packing, which are likely to be only of second-order importance as compared to sorting in sands composed of uniformly-sized particles. Inasmuch as a maximum value for the number of contactslgrain will be reached at some particular burial depth, the average number of grain contacts decreases rapidly in usefulness with increasing compaction, even for well-sorted sands. Below this depth, increased compaction will give rise to an increase in the number of concavo-convex and sutured contacts. Phipps found that the sorting has no influence on the type of grain contact. He has established an accurate technique for determining the degrees of compaction by employing the “grain contact index”. It gives a semi-quantitative value for the degree of compaction as based on type of grain contact. Phipps (p. 486) measured this index as follows: “The percentages of floating grains, and tangential, long, concavo-convex, and sutured contacts are multiplied by zero, one, two, three and four, respectively. These weighted figures, are then summed and divided by four to obtain the index. On this basis, maximum and minimum values are 100 and zero. Contacts per grain in thin section are derived by counting along regular traverses, the type of grain contact being noted at the same time. The number needed to get repeatable results is a function of the sandstone texture, and is best arrived at by making a few trial counts.” The findings of Phipps, who also examined in detail calcite- and sideritecemented sandstones can be summarized as follows: (1)Calcite-cemented sandstones occur as sporadic (3 inches to 1ft) bands within the main sandstone body. The sand grains are enclosed by the calcite crystals and are identical to those from non-cemented sandstones, but do not form the interlocking mosaic characteristic of the latter. In calcite-cemented sandstones, the boehm lamellae and other indications of mechanical compaction are very rare, whereas floating grains and tangential contacts are very common and the grain contact index and the average number of contacts per grain are very low (Table 3-LXII). In all other respects, these sandstones display the same grain size, sorting, and depositional characteristics as nonEemented sandstones immediately above and below them. Table 3-LXII
TABLE 3-LXII Variations of various properties and characteristics of Eocene sandstones from Maracaibo Basin, Venezuela, with depth of burial (after Phipps, 1969, table I, p. 487) Sample Depth No. (a)
Contacts Grain pergrain contact index
Floating Tangential grains contacts (%)
(40)
Long Concavocontacts convex (56) contacts
Sutured d m 2 contacts (mm) (5)
$0
3
Porosity Permeability (%) (md)
12755.5 12833.5 12834.0 12910.0 12976.5 13100.0 13230.0 13291.0
3.95 2.50 0.80
1.70 4.55 4.00 4.83 3.76
61 29 18 30 56 54 59 54
0
2.0 34.0 6.7 0
0 0 0
20.0 84.0 58.0 72.9 22.5 23.8 17.9 18.5
40.7 10.0 8.0 14.8 36.6 45.6 35.2 50.9
25.6 3.0 0 5.4 33.3 21.8 40.0 26.4
v1
Z
8
m
(%)
77 138 139 170 175 214 286 334
E b
13.7 1.0 0
0 7.6 8.8 6.9 5.2
0.259 0.230 0.180 0.137 0.202 0.230 0.185 0.286
1.25 1.29 1.36 1.46 1.20 1.20 1.24 1.20
12.8 11.4 16.5 3.9 15.3 14.1 10.6 13.9
121.0 25.0 122.0 <.01 112.0 55.0 8.1 368.0
Sandstones with carbonate cement; median diameter of framework fraction of sandstone (i.e., discounting clay fraction); So = Trask sorting coefficient of framework fraction.
2:
U m e l
0
z
*z8
U
2
2
w
00 U
K.H. WOLF AND G.V. CHILINGARIAN
382
shows that the uncemented arenites have undergone compaction, because the average numbers of contacts per grain and grain contact indices are twice as large as those of the cemented sandstones. No floating grains occur here and the percentages of concavo-convex and sutured contacts are high. On the contrary, in the cemented arenites at least 80%of the grains have tangential or floating contacts; concavo-convex and sutured contacts are rare. Porosities of calcitecemented sandstones are very low, but the percentage of carbonates plus percent porosity reaches a value of 35.1% (Table 3-LXIII). It can be deduced, therefore, that the cement was deposited prior to the occurrence of the main compaction. Inasmuch as the original depositional porosity was probably around 40%, there has been a porosity reduction of approximately 5%. This is in contrast to the porosity reduction of approximately 27% for uncemented sandstones. Cementation, therefore, was of early diagenetic origin and occurred after only a very slight compaction. (2) The large crystals of cement in siderite-cemented sandstones have been granulated along cleavage planes. These arenites are characterized by moderate to good porosity and permeability. Comparing Tables 3-LXII and 3-LXIII, it becomes clear that the siderite cement was also precipitated prior TABLE 3-LXIII
Variation of porosity, permeability, and carbonate content of sandstones with depth of burial1 (after Phipps, 1969, table 11, p. 489) Sample No.
Depth
Porosity
212 95 99 108 111 112 114 117 133 1382 13g2 166 169 170 172
12,692.5 12,767.5 12,770.5 12,776.5 12,786.5 12,787.0 12,795.5 12,797.5 12,817.5 12,833.5 12,834.0 12,858.0 12,868.5 12,910.0 12,936.0
14.6 3.5 7.3 3.0 2.2 1.9 2.1 2.7 7.7 11.4 16.5 8.5 7.5 3.9 4.7
1
Carbonate
ca.0.5 13.3 12.5 20.5 25.8 27.8 32.8 32.4 17.1 13.7 16.3 16.5 19.1 23.7 23.5
Porosity (%) Permeability (md) plus carbonate ca. 15.1 16.8 19.8 23.5 28.0 29.7 34.9 35.1 24.8 25.1 32.8 25.0 26.6 27.6 28.2
Percent interstitial clay is not accounted for in this table;
<0.1 0.1 <0.1 < 0..1 <0.1 <0.1 <0.1 <0.1 <0.1 25.0 122.0 <0.1 <0.1 <0.1 <0.1
siderite cement.
DIAGENESIS OF SANDSTONES AND COMPACTION
383
to the main compaction. In some cases, where the siderite cement has been almost entirely leached out, probably by subsurface water, leaving only small remnant granules of siderite in the interstitial spaces, compaction was able to take place. The minute granules, which are rounded cleavage rhombs, still display within the thin section optical continuity over areas up t o the size of the original larger crystal of siderite. Similar leaching is absent in the calcitecemented arenites. The greater porosity and permeability present in the siderite beds is apparently the result of leaching. Phipps (p. 488) reasoned that if the siderite cement was precipitated prior to compaction, then the siderite-cemented bands should have “frozen” the original textures and physical characteristics as the calcite-cemented ones did. Also, both types of sandstones should have similar very low values of porosity and permeability. Unexpectedly, however, the most obvious features of the cldexke arenites are their high porosity, high permeability, and granulated texture, which suggest a later, secondary origin for these features. As Phipps explained, if originally siderite-cemented sediment increased its porosity by the observed percentage, then this would necessitate a considerable expansion of the rock framework in order that porosity plus percent of carbonate cement could exceed the primary depositional porosity, i.e., porosity prior to leaching and granulation. As shown in Table 3-LXIII, however, this is not the case. The arenites with the maximum amount of unleached siderite have nearly the same percentage of combined pore space and cement as the calcitic sediments. Porosity and permeability must have been present to allow the passage of fluids that caused leaching. Phipps, therefore, proposed that calcite was present originally and that the secondary porosity was the result of sideritization, with a consequent volume decrease during replacement of the calcium by the iron ions. For similar possible changes, the reader may examine the data by Schmidt (1965). Inasmuch as complete sideritization of calcite leads to a volume decrease of about 20% (i.e., an increase in porosity of about 5% in an arenite with approximately 25% calcite cement), the process of sideritization invoked by Phipps does not account for the entire increase in porosity. The observed porosity is twice the amount that sideritization could account for. In addition, the replacement of calcium by iron was only partial. It seems, therefore, that leaching of siderite has been responsible for the additional increase in porosity. A stage must have been reached during leaching, however, when the cement was unable t o support any longer the overburden pressure, with a consequent occurrence of compaction. Prior to that, the arenite had its maximum post-burial porosity. Progressive leaching accompanying compaction reduced the porosity until the sediment was fully leached and compacted. Figure 3-247 illustrates the complete paragenetic sequence of the processes involved. Phipps assumed that sideritization of the calcite cement
384
K.H. WOLF AND G.V. CHILINGARIAN
zg
TIME (NoScale)
-
Fig. 3-247. Schematic illustration of the sideritization and leaching processes. (After Phipps, 1969, fig. 1; courtesy Geol. Mag.)
occurred at the present depth of burial of the rock, which at the same time is the maximum depth of burial. There is some evidence that this replacement occurred relatively recently. The sandstones with the greatest amount of remaining siderite cement are more porous than the texturally-identical, uncemented sandstones in the immediate surrounding localities. This suggests that the sediments are still compacting and that their higher porosity is of fairly recent origin, i.e., that leaching is probably occurring at the presentday depth of overburden. Sideritization, which had been an early post-burial event, was followed by subsequent leaching that allowed mechanical compaction processes to be initiated. The continued leaching, concomitant with increasing overburden pressure, maintained a porosity always slightly above that of the surrounding non-cemented sandstones, i.e., slightly higher porosity resulted from the time lag between leaching and compaction. The fact that sample 139 with siderite cement still retains a combined cement plus porosity percentage of 32.8 (Table 3-LXIII), which is of the same order of magnitude as the highest figures for the calcite-cemented sandstones, and that its “Grain Contact Index” is actually the lowest measured, is good evidence that compaction as a result of sideritization and leaching has hardly developed. That is, it has not yet suffered any significant compaction in addition to that which would be expected of a calcite-cemented sandstone. As Phipps pointed out (p. 492), “sideritization must have been a very recent event taking place at or near present depths of burial. In contrast, almost all the siderite cement has already been removed from sample 21” (Table
DIAGENESIS OF SANDSTONES AND COMPACTION
385
3-LXIII), “which only retains 0.5% of carbonate cement distributed as very small remnant granules. This sample has now compacted into a sandstone characterized by the same interlocking mosaic of sand grains common to all ihe non-cemented sandstones, and only differs from them in the retention of the siderite granules. Sample 21 represents the final end-product of the processes I have been describing, and it is clear that, if the siderite cement were not present, its history could not be distinguished from that of the non-cemented sandstones. Accordingly, it is also much more difficult to date the sideritization. The Grain Contact Index of sample 21 is 50 and contacts per grain are 4;both figures are in the same order as those for non-cemented sandstones at the same present-day depth of burial. . . Sample 21 illustrates an interesting point about the nature of the permeability of the sideritecemented sandstones. Despite the fact that sample 21 retains a relatively high porosity of 14.6%, its permeability is unmeasurable with normal methods in common with many of the non-cemented sandstones. This implies that some of the permeability of the siderite-cemented sandstone is a function of texture as well as porosity.” In regard to the mechanism of sideritization, Phipps believed that because the calcite-cemented sandstones have been impervious to interstitial fluids, the Fe-bearing solutions probably were not the cause of siderite formation and the nature of the fluids was changing from that of precipitating to that of dissolving the siderite. He proposed, therefore, that ionic diffusion is the most probable mechanism and that the immediate source for the iron was the pyrite present in all of the sandstones studied. Stylolites, which are very common, frequently contain highly ferruginous material, suggesting that the iron may have been mobilized during stylolite formation. This, in turn, also suggests that sideritization occurred at the present-day depth of burial. The publication by Cadigan (1971)might be mentioned here as an excellent example of very detailed qualitative and quantitative study of petrography and petrology of a formation on a regional scale. He investigated the Triassic Moenkopi Formation, which is one of the major uranium-bearing sandstones of the Colorado Plateau. Clastic grain types, heavy minerals, cement varieties, grain-size variation, sorting, etc. were studied on a regional basis and distribution maps prepared from the data. Similar investigations would be most useful in regional compaction studies. Alluvial fan deposits (Bull, 1972)
Bull (1972;see also Bull, 1964, 1973) undertook a study of land subsidence, compaction, and sediment-filled tension cracks in alluvial-fan deposits. During one of the civil engineering projects in California, thousands of cracks were found. They required investigation, because these fractures have
386
K.H. WOLF AND G.V. CHILINGARIAN
a direct bearing on the design and maintenance of structures, such as canals. According to Bull (1972),the land subsidence, which caused the fractures, is the result of four possible causes: (1)tectonic movements, (2)withdrawal of petroleum and gas, (3) artesian-head decline, and (4)compaction due to wetting of certain alluvial-fan deposits. The removal of subsurface water ( 3 above) has increased the applied stresses and, therefore, increased the normal compaction rate of the upper 1000-3000 f t of sediments. A subsidence rate of up to 1.8 ftfyear was the result. The differential settling occurred in some areas as a result of irrigation, which wetted the soil (4 above). Both overburden load and clay content affected the amount of compaction due to this wetting. An irrigation test was made on one of the fans. The surface of the fan first rose due to the expansion of the clay after water was first added, but as water penetrated to greater depths there was a net reduction in the volume of the deposit. The amount of volume shrinkage increased with depth due to increasing overburden pressure. The measurements on the amount of compaction due to wetting at different depths are shown in Fig. 3-248.There is a straight-line increase in degree of compaction to a depth of 100 ft. Thereafter, compaction decreases because of the higher natural moisture content below a depth of 125 ft, so that the addition of water results in a less pronounced change in the strength of the material and, consequently, less compaction. The effect of clay content on compaction due to wetting is shown in Fig. 3-249.For the samples tested it was concluded that: (a) Samples with less than 2% clay content compacted only slightly when wetted. (b) Samples with more than 30% clay content not only resisted compaction but also expanded by swelling. (c) Maximum compaction due to wetting occurred in water-laid sediments at a clay percentage of about 12%. (d) Resistance to compaction increases with increasing amounts of clay content (above 12%) when wetting takes place. The swelling of montmorillonite clays also reduces compaction. (e) The net compaction reached zero at about 30% clay content. Bull also studied the possibility of future near-surface subsidence in California, especially of large alluvial fans, which could damage the California aqueducts. Poland (1972) undertook similar studies. Figures 3-250 and 3-251 show a regional compilation of land subsidence from 1920 to 1972. More than 20 f t of subsidence due to artesian-head decline, sometimes at a rate of 1.8 ft/year, was measured. The compaction effects reached down as far as 3000 f t below the surface. In his most recent research results on the same area in California, Bull (1973)demonstrated that the estimated specific unit compaction (= compac-
DIAGENESIS OF SANDSTONES AND COMPACTION
387
tion during a time period, per unit thickness, per unit applied-stress increase) of the sedimentary pile in the northern subarea is four times that of the southern subarea. This indicates distinct variations in sediment compressibility. Bull demonstrated quantitatively that one-third of this compressibility difference is a result of lower prior total applied stress in the northern in contrast to the southern subarea. The remaining two-thirds of this difference are the consequence of variations in the water-expulsion rates in the clay-rich beds of different depositional environments. The northern area is characterized by flood-plain deposits composed of extensive sand beds with thin clay beds, resulting in a relatively faster dewatering under increasing effective overburden stress. The accumulations in the south, on the other hand, are
14
12
COMPACTION, 9.
0
CLAY CONTENT, 9.
Fig. 3-248. Effect of overburden load o n compaction due t o wetting. Inter-Agency Committee test plot B, after 42 months of operation, Arroyo Ciervo fan, California. (After Bull, 1972, fig. 3, p. 6; courtesy U.S. Geol. Surv.)
Fig. 3-249. Effect of clay content on compaction due to wetting under a simulated 50 ft overburden load. Squares = water-laid sediments; circles = mudflow deposits; triangles = deposits intermediate between mudflows and water-laid sediments; initial point on curve (ordinate) = Ottawa sand devoid of clay. (After Bull, 1972, fig. 2, p. 7 ; courtesy U.S. Geol. Surv.)
5311W P I
01
5
u
DIAGENESIS OF SANDSTONES AND COMPACTION
I ...I
I c . ~ surr., ”.,u sI*rso.&-l.Y
kmm
*4h,
389
1 c.ntr.i
Fig. 3-251. Land subsidence, 1926-197 2, LOBBanos-Kettleman City, California. (After Dr. J.F.Poland, personal communication.)
390
K.H. WOLF AND G.V. CHILINGARIAN
Experiments o f Pryor (1973) Pryor (1973) pointed out, as has been done by many others before him, that sandstones are the result of long and commonly complex histories of geologic evolution involving combination of: (1)sedimentary depositional mechanisms; (2)burial; (3) chemical and physical diagenesis (including compaction); and (4) structural deformation. In order to understand the complete history from the beginning to the end, a detailed knowledge of the primary (initial) grain accumulation characteristics is required. Little precise data has been collected in this area until recently by occasional investigations, in contrast to the more extensive information available on the secondary mechanical and chemical alterations, even though unresolved problems still exist in the latter field, too. Pryor (1973, p. 185) reported on some experiments with artificial sediments and stated that “the investigators jarred the sediments in their containers until no further compaction was observable and maximum packing was attained. Although natural sands do not receive this treatment, the winnowing processes of the beach and dune environments result not only in better sorting of the sands, but also result in a closer packing geometry. The ‘dumping’ processes of the fluvial environment yield a much looser, propped-open packing style than the beach-dune winnowing processes. Although no geometric data are available on the relative packing styles of sands from different depositional environments, packing styles may vary significantly and may be the cause of the textural and fabric differences shown in Table 3-LXIV. Von Engelhardt and Pitter (1951), among others, have shown that packing differences exert a strong control on variabilities in permeability and porosity. ” TABLE 3-LXIV Variation in carbonate content and composition, porosity, and grain density of Eocene sandstones from Maracaibo, Venezuela (after Phipps, 1969, table 111, p. 491) Sample Depth No. (ft)
Porosity
(%I
Carbonate Carbonate composition (%) (W wt) CaC03 MgC03 FeC03
Grain density2 (glcc)
117 139l 170
2.7 16.5 3.9
32.4 16.3 23.7
2.669 2.917 2.695
12,797.5 12,834.0 12,910.0
79.8 17.8 89.8
4.8 6.8 6.8
15.4 75.4 3.4
Siderite cement;2 grain density = total density of sandstone, including sand and cement, but not pore space, i.e., solids density not bulk density. Note markedly higher grain density of siderite-cemented sandstone.
DIAGENESIS O F SANDSTONES AND COMPACTION
100
400
boo
800
39 1
1000
Time (Seconds)
Fig. 3 - 2 5 2 . Effects of repacking by elutriation o n permeability and porosity of selected recent sands from different depositional environments. (After Pryor, 1973, fig. 25, p . 188; courtesy A m . Assoc. Pet. Geologists.)
To test the effect of depositional environment on packing, Pryor used duplicate samples collected from different sedimentary environments and allowed the sands t o be repacked through elutriation. Figure 3-252 demonstrates the change in permeability and porosity through time of elutriation as repacking takes place. In all cases, an initial rapid reduction in permeability during the first hundred seconds of the experiments was observed. This was followed by a levelling-off and stabilization of the rate of packing. Porosities also exhibit a marked reduction from the undisturbed to repacked grain arrangements. The data also illustrated that river-bar sediments change most extensively in packing in contrast t o beach and dune samples which change the least. Table 3-LXV summarizes some of the data collected by Pryor on river-bar, beach, and dune samples. He also concluded (p. 1 8 5 ) : “(1)Permeability increases with increasing grain size, and ( 2 ) porosity increases with increase in sorting; both relations agree with the experiments of Fraser (1935) and the determinations of Krunibein and Monk (1942).
TABLE 3-LXV
0 W
lu
Relations of permeability (k)and porosity (@)of sands to textural parameters and depositional environments (after Pryor, 1973, table 2, p. 188) River Bar
Grain size increases Sorting increases
Beach
Dune
Fraser (1935)
k
9
k
9
k
@
k
@
increase increase
increase increase
increase decrease
decrease increase
increase decrease .
decrease increase
increase decrease
decrease increase
x x
DIAGENESIS OF SANDSTONES AND COMPACTION
39 3
The relations of porosity versus mean size and permeability versus sorting separate the data into two different groups: (1)the river bars, where porosity increases with increasing sorting and permeability increases with increasing sorting (the reverse of Fraser’s, finding), and (2) beaches and dunes, where porosity increases with decreasing mean grain size and permeability increases with decreasing sorting (both in accord with the findings of Fraser)
...
9,
Pryor also described the permeability characteristics of bedding units of the beach, dune, and river-bar sediments; Fig. 3-253 depicts the latter. In river-bar sediments, there is a high variability in textural characteristics, permeability, and porosity, as well as many truncations and intersections as a result of cut-and-fill sedimentation. The permeability increases downward in each dipping laminae due to textural changes. The directions in which permeability increases are generally parallel to the length of the river point-bar deposits. Sand bodies of the river-bar and beach environments, the sediments of which exhibit well-organized and different patterns of permeability variations, are compared in Fig. 3-254. Pryor (pp. 188-189) stated that “permeabilities in river-bar sands are generally higher and more variable than in beaches and dunes and decrease systematically downstream and bankward. In beach sand bodies, permeability values have a small variability, but are
Fig. 3-253. Diagram of river-bar laminae packet with permeability characteristics. (After Pryor, 1973, fig. 15, p. 181; courtesy Am. Assoc. Pet. Geologists.)
394
K.H. WOLF AND G.V. CHILINGARIAN
RIVER BAR
HIGH VARIABILITY
BEACH
LOW VARlABlLlM
Fig. 3-254. Generalized relation of permeability trends in river bars and beaches. (After Pryor, 1973, fig. 26, p. 189; courtesy Am. Assoc. Pet. Geologists.)
relatively low on the beach faces, high on beach crests, and variable on the beach berms. Permeability-variation trends are parallel with the length of river-bar sand bodies and perpendicular to beach sand bodies. Dune sand bodies show no distinctive trend of permeability variation. There are greater variations of grain characteristics and permeability-porosity within bedding and lamination subunits than between them; this is especially pronounced in river-bar sand bodies. Packets of laminae and beds in beach-dune sand bodies are more uniform in character than those in river-bar deposits. Clay laminae and other low-permeability units commonly are present between bedding and laminae packets in river-bar sand bodies. In all deposits the boundary conditions between sediment packets are important factors in determining the effective reservoir characteristics of the sand bodies because sediment packets surrounded by other units of lower permeability will have effective permeabilities influenced by, and largely determined by, the lower permeabilities of the bounding units and hence will not demonstrate their ultimate through-flow capabilities. The ideal reactions between permeability-porosity and textural characteristics that have been shown by various authors for artificially packed particles are demonstrated only weakly by beach and dune sands and are not demonstrated by river-bar sands. The different styles of natural grain packing between river-bar sands and beach-dune sands are probably the cause of these deviations from the ideal models.”
DIAGENESIS OF SANDSTONES AND COMPACTION LATE DIAGENESIS-EPIGENESIS-BURIAL
395
METAMORPHISM
In the sequence late diagenesis-epigenesis-burial metamorphism*, the epigenesis (the term catagenesis is preferred by some researchers) and metamorphism probably are outside the realm of compaction of sedimentary rocks. Inasmuch as the various mechanical and chemical processes related to compaction grade into these stages, however, the writers decided to present a brief discussion on the burial processes and products beyond compaction. That there is no definite boundary between diagenetic compaction and epigenesis-metamorphism, can be easily seen by the progressive stages from orthoquartzite to a genuine metaquartzite, as well as by the transition from a clayey arenite through a chlorite-rich greywacke into a greywacke-schist. With these progressive changes in mineralogy, one would expect to find also changes from an uncompacted to compacted and, finally, t o recrystallized I
40,
TEMPER ATU R E ,a C
Fig. 3-255.Ranges of metamorphism. Conditions represented by points in the blank field on left are unlikely to be met in nature. A-B = range of measured geothermal gradients; C-D = range of geothermal conditions considered t o be common; E = melting range of granite in equilibrium with water vapor. 1 = diagenesis, weathering; 2 = dynamic metamorphism; 3 = thermal metamorphism; 4 = dynamothermal metamorphism; 5 = onset of melting; 6 = possible transient conditions in a geosyncline. (After Bayly, 1968,fig. 18-2; copyright @ 1968 Prentice-Hall, Englewood Cliffs, N.J.)
* It is important to note here that some investigators include epigenesis (or catagenesis, as preferred by some) in late diagenesis, whereas others consider it part of metamorphism. Some refer to the transitional stage between diagenesis and metamorphism as metagenesis (see Larsen and Chilingar, 1967). Some of the data presented here have been updated in the third edition of the book by Winkler (1967)published in 1974.
396
K.H. WOLF AND G.V. CHILINGARIAN
textures. Another example that illustrates the transition from diagenesis through epigenesis to metamorphism is supplied by the data on organic matter (“organic diagenesis-metamorphism”), which is discussed here also. The types of metamorphism as based on depth of burial or pressure and temperature are presented in Fig. 3-255.The processes related to diagenetic compaction and burial metamorphism belong to the shaded area depicting possible transient conditions in geosynclines. Two similar diagrams are given in Figs. 3-256and 3-257(Winkler, 1965)with supplemental data, especially important mineral associations. Figure 3-258presents data on the temperature variation in the Cenozoic sedimentary basins in the Gulf Coast of Texas and Louisiana. This data alone indicates that the late diagenesis and epigenesis may occur at temperatures near the boiling point of water and above, and that based on the temperature gradient there are no definite demarcations among diagenesis, epigenesis, burial metamorphism, and regional metamorphism. Fyfe (1973)presented the results of his study on dehydration reactions and stated that “the vapor pressures of hydrous materials indicate that thermal gradients in wet sediments may be influenced by endothermic dehydration reactions. Stress and salinity may influence the temperature of dehydration. The large solubility of stressed grains must lead to rapid elimination of porosity and to the condition that fluid pressures approach lithostatic pressures during sinking of wet sediments”. Examples, or case histories, of the various types of low-grade metamorphism are offered below, and the information on organic diagenesis is considered first. Landes (1967)called the incipient alterations of organic material “eometamorphism” in his investigations on the limits placed on the distribution of hydrocarbons by the intolerance of petroleum in the earth’s crust to TEMPERATURE, ‘C
Fig. 3-256.Schematic pressure-temperature diagram for different types of metamorphism. The ELT region below the lowest possible geothermal gradient of about lO*C/km is not realized in nature. The indicated depths corresponding to given pressures are maximum values and may be less than shown. (After Winkler, 1967, fig. 1 ; courtesy Springer, New York.)
L 66
NOIJ,L3VdpJO3 a N V SBNOJ,SaNVS 6 0 SIS3N3DVIa
m
398
W
(D
m
K.H. WOLF AND G.V. CHILINGARIAN
x
3:
3 G *z Fig. 3-258B. Temperatures in South Louisiana at a depth of 3048 m. Contour interval (CI) =, 5OC. Hot belt with temperatures between 100 and 113'C is near present shoreline. Rest of the area has temperature near 95 C. (After Jam et al., 1969, fig. 5, p. 2148; courtesy Am. Assoc. Pet. Geologists.)
U 0
. + Q
5
DIAGENESIS OF SANDSTONES AND COMPACTION
399
position between the other sedimentary rocks and petroleum in vulnerability to chemical and physical change brought about by increased heat and pressure . . . rank of the coal provides an excellent indication of the character of any hydrocarbons which may be stratigraphically nearby” (Hilt’s Law). Figure 3-259 represents an attempt to show the results of incipient metamorphism, as recorded by carbon ratios in coal and hydrocarbons, whereas Table 3-LXVI presents the relation between coal reflectance and hydrocarbon occurrence. The average relation between subsurface temperature and the density of hydrocarbons are demonstrated in Fig. 3-260. The amount of liquid hydrocarbons decreases whereas the percentage of gas and condensate increases with depth. Eometamorphism limits the occurrence of hydrocarbons in space both laterally in the basins bordered by mobile rims and at depth in deeper basins, due to increasing heat and pressure. According to Landes (p. 828), “it is estimated, based on past experience and therefore subject to modification in the future, that the commercial oil floor ranges from a depth
COAL
HYDROCARBONS Heavy oil and gos fields
?to Sub- Bituminous
OD0
000
Oil and gos fields
High volotile Bituminous c High volatile Bituminous 8
High volatile Bituminous A
Light oi! and gos fields Increasing percentage of gas fields Oil phase-out zone Gas oredominant
Gas only.
8o
I
Medium volotile Bitummous Low volotlle Bituminous
?---
decreoses downward
Semi- Anthracite Anthracite
00
Volume
Melo-Anthracile lo Graphite
lnodequote porosity
Graphite from enlropped gas
?
Fig. 3-259. Carbon ratios and occurrence of oil and gas. (After Landes, 1967, fig. 7, p. 836; courtesy Am. Assoc. Pet. Geologists.)
400
K.H. WOLF AND G.V. CHILINGARIAN
TABLE 3-LXVI Coal reflectance and hydrocarbon occurrence (after Ammosov and Syn-i, 1961, in Landes, 1967, fig. 8, p. 837) Stage
Vitrinite reflectance
~~
Coal type
Oil prospects
Lignitic Sub-bituminous Bituminous Bituminous Semi-anthracite Anthracite
Good Good Fair Gas only Nil Nil
~
1 5 8 11
60
74 83 92 125 170
17 22
of about 14,000 ft to 27,500 ft where the gradient is 1"/100ft. The commercial gas floor is an 'assay floor'; it is the level at which the volume of gas obtainable from reservoirs with decreased porosity does not yield a profit. It is concluded that commercial deposits of oil and gas do not extend to the basement rock in deeper basins." Staplin (1969) described progressive organic metamorphism, the next step beyond diagenesis as he pointed out, which affects particulate organic matter. He stated (pp. 56, 57): "The progressive changes include darkening of the organic matter, decrease in light transmissibility and electrical resistivity, increase in index of refraction, reflectance and lustre, and loss of fine structural detail. Increase in carbon ratio, spectrometric changes and other evidence of chemical transformation accompany the physical transformations .
..
200 LL 250 a 300
-1
!n -
w
0
a '
z
350 w 400+
I
2owo sIty and Gas Volume
Decrease Downward
1
1
Fig. 3-260. Earth temperatures and occurrence of oil and gas. (After Landes, 1967, fig. 9, p. 839; courtesy Am. Assoc. Pet. Geologists.)
DIAGENESIS OF SANDSTONES AND COMPACTION
401
The degree of organic metamorphism is determined by observing the organic debris, especially plant spores and non-woody cuticle and amorphous sapropelic debris. Of the organic materials in the rock, these seem to be the components most sensitive to heat, and their alteration in a clayey matrix progresses at roughly the same rate. Phytoplankton, less sensitive and more variable in its response, is less useful as a ‘thermometer’. The progressive severity of organic thermal metamorphism is subdivided into five degrees (thermal-alteration index, Table 3-LXVII), which are easy to determine for plotting on maps. As shown on this Table, sediments in areas with unaltered to moderately altered organic matter (thermal index 1-3) contain free and distillable hydrocarbons where the facies are suitable. In strongly to severely altered areas (thermal index 4 and 5) only dry gas is evident on the basis of cuttings, gas and hydrocarbon analyses, and experience gained through drilling shows that petroleum prospects in such areas, except for dry gas, are minimal.” Staplin studied (1)the distribution of gas in the Devonian carbonates of northeast British Columbia and equivalent areas in the Northwest Territories, and (2) the occurrence of oil and wet gas in the same carbonates in northwestern Alberta (Fig. 3-261). He found that the color and preservation of TABLE 3-LXVII Interpretative summary: type of organic matter, degree of alteration, and expected hydrocarbons (after Staplin, 1969,table 1, p. 57)
T y p e o f organic matter, matured facies (a) Sapropelic, amorphous (b) Plant cuticle, charcoal (c) Mixture of a and b Organic metamorphism (a) Blackening,:increase in index of refraction
(b) Pyrobitumen (c) Graphite, mineralization Thermal-alteration index (1)None (2) Slight (3) Moderate (4) Strong (5)Severe
Organic matter fresh, yellow brownish yellow brown black black, with additional evidence of rock metamorphism
Associated hydrocarbons wet gas and oil dry gas wet gas and oil wet or dry, depending on thermal alteration index
wet or dry wet or dry wet or dry dry gas dry gas to barren
K.H. WOLF AND G.V. CHILINGARIAN
402 WEST I CRETACEOUS MISSISSIPPIAN
z I WINTERBURN
EAST 2
3
4
5
6
7
8
D(I 11 1 f, 1
Fig. 3-261. Color of organic matter in eight wells along an e a s t w e s t section (northeastern British Columbia and northwestern Alberta). Color of organic matter: A = yellow; B = brown; C = dark brown and black. (After Staplin, 1969, fig. 2; courtesy Bull. Can. Pet. Geol.)
the organic matter, distribution of sedimentary hydrocarbons, presence of pyrobitumens, mineralization (e.g., galena, chalcopyrite), geothermal gradients and other variables all support metamorphism as the reason for the occurrence of dry gas in northeastern British Columbia. Staplin (p. 61)concluded: “Mineralization, high density and strong compaction, high-temperature gradients, clay organization, and deep depths of burial are evidence for thermal alteration, but the correlation between the appearance and properties of organic matter and the degree of metamorphism is the most rapid method of determining the effect of thermal activity in an area.” Baker and Claypool (1970) and Baker (1972)mentioned that incipient (low-grade) metamorphism leaves few obvious mineralogic or textural effects. Consequently, mildly-metamorphosed sedimentary rocks (other investigators may consider them epigenetic or late diagenetic in origin) are very difficult t o study. These authors also suggested that organic compounds are more susceptible to alteration and, therefore, may be more sensitive indicators of changes at temperatures and pressures existing between diagenesis and metamorphism than minerals. Thus, together with textural investigations, as discussed elsewhere in this chapter, it might be possible to develop a system of metamorphic facies classification for the realm of incipient metamorphism, based on the molecular and compositional character of organic substances. In Fig. 3-262,the hydrocarbon content is plotted against organic carbon content on a log-log graph paper. There is a positive correlation between these two variables for the unmetamorphosed samples, whereas no systematic relation is apparent for the metamorphosed samples. The hydrocarbon content of metamorphosed specimens is generally less than that of unmetamorphosed samples, and the difference is particularly evident for metamorphosed samples containing more than 0.5%organic carbon. There is
DIAGENESIS OF SANDSTONES AND COMPACTION ,J A x
403
UNMETAMORPHOSEO ”METAMORPHOSED” Martinsburg a equivalents Milligen a other Idaho samples
1.0 10.0 ORGANIC CARBON WT. %
100
Fig. 3-262. Relationship between hydrocarbon concentration and organic carbon content in metamorphosed and unmetamorphosed rocks. (After Baker and Claypool, 1970, fig. 1, p. 461; courtesy Am. Assoc. Pet. Geologists.)
a net loss of hydrocarbons during incipient metamorphism. Other data indicate that there is a relative increase of saturated compared to aromatic hydrocarbons as a result of incipient metamorphism. The relationship between saturated/aromatic and hydrocarbon/organic-carbon ratios in sedimentary rocks is presented in Fig. 3-263.Baker and Claypool (1970, p. 460) stated: “The unmetamorphosed samples fall into a linear trend with saturated/aromatic values less than 2 and a wide range of hydrocarbon/organiccarbon values. The metamorphosed rocks mostly fall in an area of low hydrocarbon/organic-carbon ratios and saturated/aromatic values greater than 2.” It is assumed that the metamorphosed samples were derived from material that originally would have fallen into the field defined by the unmetamorphosed organic matter. The metamorphic trend lines in Fig. 3-263illustrate the effects of incipient metamorphism on extractable hydrocarbons in sedimentary rocks: the absolute amount of hydrocarbons decreases, the ratio of hydrocarbons-to-organic carbon decreases, and the ratio of saturated-toaromatic hydrocarbon increases. The results of these investigations, therefore, indicate that the amount, composition, and molecular structure of the extractable hydrocarbons and the 3Ccontent of kerogenic carbon, according t o Baker and Claypool, can be used for the recognition and classification
HYDROCARBON/ORGANIC CARBON RATIO x
lo-’
Fig. 3-263. Effect of incipient metamorphism on extractable hydrocarbons in sedimentary rocks. Solid circles and dashed area = unmetamorphosed; x and stippled area = metamorphosed; triangles = Martinsburg Shale (Ordovician) and equivalents; x = Milligen Shale (Mississippian) and other Idaho samples (Paleozoic mudrocks); open circles = Triassic Lockatong Formation of New Jersey and Pennsylvania. (After Baker and Claypool, 1970, fig. 9 , p. 462;courtesy Am. Assoc. Pet. Geologists.)
+20 r +18. y+16
-
3 +14.
t +I2 +lo -
J
INCREASING METAMORPHISM MlLLlGEN
FM.
+8u ‘6’-
9
+2 -
+4
Fig. 3-264. 6C13 values of kerogenic carbon versus relative degree of metamorphism. (After Baker and Claypool, 1970, fig. 6, p. 465; courtesy Am. Assoc. Pet. Geologists.)
DIAGENESIS OF SANDSTONES AND COMPACTION
405
of incipient metamorphic facies. The general trend of the data obtained from the unmetamorphosed samples is controlled by the characteristics of the primary organic matter in the original sediments. Baker and Claypool (1970) concluded that (1) 613C values of the sediment extracts cannot be used as reliable indicators for either recognizing or classifying incipient metamorphism, and (2)the carbon isotopic composition of the total organic carbon of some metamorphosed clayey rocks may reflect the degree of metamorphism. Each of the groups of metamorphosed specimens used by Baker and Claypool (Fig. 3-264)show some qualitative progressive variation in 613C of kerogen carbon related t o the degree of metamorphic alteration, i.e., proximity to an igneous source or geographic location relative to the direction of increasing metamorphic grade, despite considerable spread of the data. In analyzing the data, the reader should be aware of various assumptions made by Baker and Claypool as well as of difficulties involved in the interpretation of the data. Kisch (1969) presented data on the relationship of some characteristic burial-metamorphic mineral assemblages to the rank of more or less co-equal coals (see also Frey and Niggli, 1971). He discussed: (1)alterations of clay minerals, and (2)appearance of zeolite-facies mineral assemblages. No details are presented here aside from the diagrams (Figs. 3-265-3-269),as they illustrate sufficiently well the principle of parallel mineral and coal modifica-
50
VOLATILE MATTER,d.o.f. Ye
Fig. 3-265. Derivation o f the vertical coal-rank scale from Karweil’s (1965, fig. 1; after Huck and Karweil, 1955) calculated volatile matter versus relative depth curve. (After Kisch, 1969, fig. 1, p. 410; courtesy Pergamon Press, Oxford.)
+=
MUSCOVITE AND ILLITE IN SANDSTONES
KAOLINITE
.
Fig. 3-266. Effect of late diagenesis on the distribution of clay minerals with depth in some deep boreholes. (After Kisch, 1969, fig. 2, p. 412; courtesy Pergamon Press, Oxford.)
?
cd
8
P
E z
*
h
i.
SIDERITE
t i
CULORITE IN CLAVSTONES
NVIIIV3NITIH3 ' A ' 3 CINV d?OM * H X
a
.i
j i
-
.a
CHLORITE AND MOTITE-CHLORITE IN SANDST.
!; - MUSCOVITE AND ILLITE \W CLAVJTONES
i
...
4
~HLORITE
.:
t
'
Aj
i
1
j
\
1
,
/ ...............
--
MONTMORILLONITE(l7~)
-
IIXED-LAVED ILLITE-MONTMORILLONlTE
..........
--
ILLITE(IOA)
M.-L. ILLITE-MONTMOR
Ll o o 0
90P
lb
0
UJ
DIAGENESIS O F SANDSTONES AND COMPACTION
407
tion with transition from the late diagenesis-epigenesis into low-grade metamorphism. Figure 3-265shows the vertical coal-rank scale (Karweil, 1956), as determined by calculations assuming a decrease of 2.2% volatile matter per 100 m increase in depth in the fat coal range and an average geothermal gradient of 40"C/km. In this figure, the volatile matter content is plotted versus the relative depth of burial and corresponding coal-rank parameter (based on volatile matter content). These relations hold for the Carboniferous Coal Measures of western Europe. Similar diagrams, however, should be prepared for areas with different coalification gradients. Figure 3-266demonstrates the diminution and, finally, disappearance during late diagenesis of the kaolinite and montmorillonite and the appearance of illite-muscovite and chlorite with increase in pressure and temperature. Figure 3-267presents schematically a correlation between the distribution of clay and carbonate minerals and the coal rank from Germany, Australia, and U.S.A. Kaolinite is PPRAMETERS
(vlrnlrr)
QUEENSLAND
MONSTE~LAND I M)PEWOLE. Wt5TWALIA
Fig. 3-267. Schematic distribution of clay and carbonate minerals with respect to the rank of associated coal (vitrite) in four late-Paleozoic areas. Coalification gradient is based on Karweil's (1956) curve (Fig. 3-265). Carbon contents for vitrites of stated volatilematter yield after Teichmiiller (1963, figs. 1, 2) and after Kotter (1960) (in brackets). Data schematized after ( I ) Kisch, 1968; (2) Scherp, 1963, Stadler, 1963, Esch, 1966; (3) Quinn and Glass, 1958; ( 4 ) Eckhardt, 1964, Teichmuller and Teichmiiller, 1966a, b. (After Kisch, 1969, fig. 3, p. 413; courtesy Pergamon Press, Oxford.)
K.H. WOLF AND G.V. CHILINGARIAN
408
I A M I C l r n
Fig. 3-268. Schematic distribution of the zeolites (analcime, heulandite-clinoptilolite, and laumontite) with respect to the rank of associated coal (vitrite). Coalification gradient as in Fig. 3-267. (After Kisch, 1969, fig. 7 , p. 419; courtesy Pergamon Press, Oxford.)
replaced by illite (muscovite), chlorite, and/or pyrophyllite with increasing rank of coals, but in different localities different changes in the clay mineralogy correspond to the varying coal ranks. As pointed out by Kisch, these discrepancies could be the result of differences in the primary mineralogy of the deposits and the availability of cations from labile detritus and circulating pore fluids. Figure 3-269indicates that during late diagenesis the kaolinite and montmorillonite disappear earlier in feldspathic sediments than in kaolinite-quartz-rich rocks. On the other hand, in the areas considered by Kisch (p. 415) “the burial-metamorphic disappearance of kaolinite in both these ‘rock families’ is associated with somewhat higher coal ranks (lean coal to meta-anthracite) than the coking coal to anthracite indicated by the correlation of mineral zones and coal types after Kossovskaya, Logvinenko and Shutov (1957)”(Fig. 3-269). Kisch (1969,p. 415) mentioned one case where the sedimentary rocks are
Fig. 3-269. Correlation of the distribution of some primary (stippled) and newly-formed (in black) silicate minerals in various terrigenous and volcanic sedimentary rock types during late diagenesis; compiled and schematized after Kossovskaya and Shutov (1961,1963).Broken bars indicate extension of some mineral zones (after Coombs, 1961). Correlation of mineral zones in terrigenous rocks with coal type after Kossovskaya et al. (1957).N o t drawn to scale. (After Kisch, 1969, fig. 5, p. 416; courtesy Pergamon Press, Oxford.)
410
K.H. WOLF AND G.V. CHILINGARIAN
rich in unstable detritus of igneous origin, which supplied alkalies and Mg and Fe ions. Here, the kaolinite disappears during the lean coal rank stages. Another example illustrates the absence of kaolinite and predominance of sericite-chlorite cements in sandstones and argillites associated with lean coals and anthracites (Fig. 3-269). The case of little or no availability of cations is represented by the kaolinite-coal tonsteins (Kisch, 1969, p. 415). He pointed out that in this case no noticeable replacement of kaolinite by illite and chlorite takes place as coal rank increases, i.e., the kaolinite persists up to anthracite rank, but its b-axis disorder progressively decreases up to a fat coal rank. Replacement by pyrophyllite may take place at the metaanthracite stage. Figure 3-268 illustrates that in the case of feldspathic-lithic and tuffaceous rocks, another group of mineral modifications occurs during burial metamorphism, i.e., the appearance of a succession of diagnostic zeolites and other Ca-aluminosilicates. With increasing depth of burial, these minerals tend to be progressively less hydrous and more dense. The coal-rank
Fig. 3-270. Precision of crystallite size measurements as a function of peak width. (After Griffin, 1967, fig. 1, p. 1008; courtesy J. Sed. Petrol.)
DIAGENESIS OF SANDSTONES AND COMPACTION
41 1
parameters and depth of burial have been presented along with the critical zeolite assemblages in Fig. 3-268. Griffin (1967) described an X-ray diffraction method to identify humic materials ranging in rank from lignite to meta-anthracite (or even graphite) as a result of diagenetic and metamorphic processes. Figures 3-270-3-273 present some of the results obtained by Griffin. Bishop (1972) studied sandstones containing pumpellyite and coal from relatively shallow depth. The inferred depths of burial suggest a pressure of not over 0.5-1 kbar and a temperature of not over 50--130°C. He presented
4
m
-01
Lignite
I roo
7
500
-02
400
Subbituminous A Subbturninous A Hlgh-Volotile 811 C
Hlgh-VolotlieBlt C High-Volotile Bd B Hlgh-VolOtlk Bit B
High-Volatile Bit A Medum-Volatk Bit
-03
Pr.
-04
200
-05 -06 -07 -08 -1 00 9
-20
-30 -40
-ao -60 -70
::1
-100
z I
' $
::
?ir
Low-Volatile Bit Sernionthrocite
- 200
Anthracite
-n o
- so0 a00
Mela-Anthrocite Grophite
Fig, 3-271.X-ray diffraction patterns of HCl, H F residues of the humic coal sequence (lignite to meta-anthracite). A pattern of standard graphite is presented for comparison. (After Griffin, 1967, fig. 2, p. 1008; courtesy J . Sed. Petrol.) Fig. 3-272.Variation in crystallite thickness and oxygen and hydrogen contents with coal rank. (After Griffin, 1967,fig. 3, p. 1010; courtesy J. Sed. Petrol.)
412
K.H. WOLF AND G.V. CHILINGARIAN
Fig. 3-273. C-axis dimensions of crystallites as a function of repeat distance (= d spacing) within the crystallites. (After Griffin, 1967, fig. 4, p. 1010; courtesy J . Sed. Petrol.)
an empirical straight-line curve of coal rank versus depth of burial for one particular area. From the rank number of the coal specimens associated with the pumpelIyite-containing sediments, Bishop was able t o determine depth of burial. In investigations such as those discussed above on organic matter, one has to consider the possible changes in composition that are the result of evolution and not due to secondary diagenetic or metamorphic modifications. Jackson (1973)observed that “humic” matter, isolated from bitumen extracted from suites of sedimentary rocks varying in lithology and depositional environment and ranging in age from Archean to Miocene, exhibited a systematic change in the molecular structure through geologic time. He attributed this change to major events in the evolutionary history of organisms. Jackson also found that despite diagenetic alterations, the primary differences between various humic deposits probably tend to persist under mild conditions long after burial. He did not comment on the possible effects of higher grades of diagenesis and metamorphism. A progressive change of the composition of organic matter with increasing burial depth and temperature was documented by Tissot et al. (1974). Murata et al. (1969,1972) discussed the isotopic changes of diagenetic
DIAGENESIS OF SANDSTONES AND COMPACTION
413
carbonates in terms of chemical processes that operate in deeply-buried marine sediments. This may also be a good approach in studying certain types of terrigenous sediments (shales, siltstones, sandstones, and conglomerates) if they contain carbonates in the form of fossils, cement, and/or concretions, as well as fracture fillings. If they can be shown to undergo changes with depth of burial, the information obtained is indicative also of the changes which occurred in the host rock. The alterations include changes in carbon isotopes, which are temperature-dependent, and in degree of dolomitization, sideritization, and pyritization, all of which can be influenced by composition of compaction fluids expelled from shales into coarser sediments. Murata et al. (1969)discussed certain isotope reactions between methane, generated by burial metamorphism of organic matter, and the carbonates that could account for the presence of abnormally heavy carbon in the carbonates. The effect of such reactions would be greatest in a deposit in which the ratio of organic shale to carbonate is large and in which the reacting fluids moved only along certain zones rather than by diffusion (Murata et al., 1969). Work on isotopes allowed them to conclude that the overall chronological succession of dolomite types, during the history of a sediment rich in organic material, seems to be: (1)lightcarbon dolomite, (2)heavy-carbon, low-iron dolomite, and (3) heavy-carbon, iron-rich dolomite. Types 1 and 2 belong to the long anaerobic phase of the sediment history, in the presence of organic matter in the sediments; whereas type 3 marks the first stage of the post-orogenic aerobic phase as a result of uplift of the formation. The latter allowed invasion of oxygen-rich surface waters, which released iron through oxidation of diagenetic pyrite. Thus, the minerals formed by fluids of compaction could be differentiated from those formed by post-tectonic fluids. Another example of the application of oxygen isotope study to problems of burial metamorphism is provided by Eslinger and Savin (1973). Anyone interested in the importance of compaction fluids in controlling chemical reactions, may wonder to what extent they have been influential in the origin of zeolites. Many publications on zeolite genesis mention the importance of such factors as (1)alkalinity of lake waters (e.g., Hay, 1966; Coombs and Whetten, 1967;and many others), (2) salinity of fluids (surface and subsurface solutions), (3) pH, (4) ionic composition, and ( 5 ) SiOz content. Little reference has been made to compaction fluids, however, most likely because they are very difficult to differentiate from other fluids that took part in the chemical reactions. Hay (1966)has shown that zeolites can form very early during diagenesis from volcanic glass reacting with alkaline lake water. Does that mean that should the latter be unavailable in the sedimentary environment, no zeolites can form? Or is there a possibility that at least later, during burial, saline compaction fluids, derived from another stratigraphic section, can move into
414
K.H. WOLF AND G.V. CHILINGARIAN
a volcanic glass-rich unit and give rise to zeolites? The answer to the latter question seems to be affirmative as demonstrated by late diagenetic and burial metamorphic zeolites. The composition of the fluids may be critical. Muffler and White (1969) reported on active metamorphism of Upper Cenozoic sediments of the Colorado River delta in the Salton Sea geothermal field in California, where a continuous transition from sediments through indurated sedimentary rocks to low-grade metamorphosed rocks of the greenschist facies occurs without the formation of zeolites. The authors suggested that the high aC02/aH20ratio prevented the formation of zeolites. For other examples of burial metamorphism, the reader may refer to the work of Crook (1963) and Coombs (1961). Packham and Crook (1960) presented a useful discussion on the problems related t o the transitional stages from halmyrolysis and diagenesis to metamorphism, and supplied diagenetic depth sequences. They pointed out one interesting difference between diagenesis and metamorphism: the original mineral composition of the rock is important in diagenesis, whereas the bulk composition is more significant in metamorphism. Coombs et al. (1959) assigned the zeolite facies to the lowest metamorphic stages, whereas Packham and Crook (1960) considered them as high-rank diagenetic facies because of the presence of recognizable essentially sedimentary textures. These facies owe their existence to the chemically-reactive volcanics. Thick, deeply-buried sediments lacking volcanics do not contain these authigenic minerals. There is a tendency of the zeolite facies to change with increasing age, as most of them are of Cenozoic or Mesozoic age and absent in Paleozoic and Precambrian rocks. Especially in eugeosynclinal sequences, pyroclastic sandstones may contain authigenic zeolites and other silicates that definitely show a progression of mineral facies as a function of pressure and temperature and, therefore, can serve as an index of depth of burial. It must be pointed out that certain types of zeolites can form under very low surface temperatures and pressures, i.e., those originating in saline lakes. For those readers who wish to gain information on the higher grades of progressive regional metamorphism, which influenced the chemical composition of a meta-arkose, Schwarcz’s (1966) paper can serve as a good starting point (see also Den Tex, 1965; Angel, 1965; Hietanen, 1967; Nelson, 1969; Ernst, 1971). It seems noteworthy t o mention Schwarcz’s suggestion that some of the reactions were isochemical in nature because (1)arkoses are composed of “inert” constituents (i.e., quartz, plagioclase, and alkali feldspar), which are stable over a wide range of temperatures and pressures, and (2) there is a lack of water of hydration to provide an aqueous fluid phase that, in turn, may have inhibited migration of chemical species and internal equilibrium during metamorphism. In contrast, the interbedded shales changed to pelitic schists which, according to Schwarcz, underwent many
DIAGENESIS OF SANDSTONES AND COMPACTION
415
stepwise reconstructive mineral transformations, with evolution of a waterrich fluid phase with increasing temperature. In this case, although compaction and compaction fluids had no obvious influence, the clayey material released fluids during metamorphism. Fluid release during diagenetic clay mineral transformations at much lower temperatures and pressures has been described previously in this chapter.
Zones of secondary alterations In terrigenous rocks of geosynclinal and cratonic (= platform) sedimentary sequences, Kossovskaya and Shutov (1958)established zones of secondary alterations on the basis of vertical change of various parameters, e.g., mineral composition and textures. Kisch (1969)used their results in his comparative investigations on coal rank. Three major zones (Fig. 3-269)were recognized, ranging vertically from the least to the most altered rocks, i.e., epigenetic, metagenetic, and metamorphic, with four subdivisions of the former two. The information applies to one locality with Mesozoic and Paleozoic rock sections, and the zoning reflects all stages of alteration in response to (1) slow basin subsidence, (2)effects of interstitial fluids, (3)increase in pressure and temperature, and (4)stresses in the geosynclinal area. The four zones are briefly described here.
Zone of unaltered clay cement. This zone is found on platforms and in the upper horizons of the marginal parts of geosynclines. It is up to 15002000 m in thickness and is composed of loose or poorly-indurated sandstones and water-saturated claystones, with original mineralogy and texturp being preserved. Gravitational compaction and consolidation increases with age in older rocks, as the bulk specific gravity increases from 1.4to 2.10 and porosity decreases from 40 to 20%. There is a variation in the mineralogical composition of the clay present as cement in the sandstones and clay in the claystones. This zone is characterized by the presence of carbonate, sulphate and, rarely, chloride waters. Some diagenetic reactions have taken place here, e.g., corrosion of ferromagnesians and feldspars. Zone of altered clay cement. This zone is present in the lowest parts of the platform regions, whereas in geosynclines the zone includes the fringe areas of folded strata. This zone is quite thick (up to 6000 m and more) and consists of shales and argillites having a porosity of 4-5% and less in the lower part of the zone. The recrystallization of the clayey matter and the neomorphism (newly-formed minerals) are characteristic of this zone. Montmorillonite disappears and kaolinite changes to illite (= hydromica), and the clay cement in the sandstone occurs as chlorite-illite aggregates. The excess
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K.H. WOLF AND G.V. CHILINGARIAN
of Si02, resulting from the clay mineral transformation, forms quartz (and chalcedony?) micropatches. The paragenesis usually is colored Fe-chlorite, colorless chlorite, illite, and quartz. The mineralogy of the cement minerals depends on the original clay composition and the type of clastic grains, e.g., acid plagioclases are accompanied by illite cement and ferromagnesians predetermine the presence of chlorite. Quartzites with a primary kaolinite cement (chemically precipitated?) may remain unaltered except for some recrystallization of the kaolinite (for a similar change of illite crystallinity, see Ludwig, 1972), whereas in polymict sandstones the kaolinite changes to illite. The change of original clastic textures and fabrics is accompanied by quartz precipitation in the intergranular spaces. At a depth above 2000-. 2500 m, the subsurface fluids are highly mineralized and the solubilities of quartz and feldspars increase as a result of higher temperatures. Pressure solution is common here resulting in a supersaturation of the interstitial fluids with silica and consequent precipitation of quartz cement and formation of overgrowths on plagioclase. The transition from stage l to stage 2 can be recognized by the textural relationships of the neomorphic minerals, several examples of which were presented by Kossovskaya and Shutov (19 58). Zone of quartz-like structures and illite-chlorite cement. This zone is present in geosynclines only and did not develop in platform regions. This zone contains a complex of features determined by stress and depth of subsidence and is characterized by the presence of quartzites, schistose argillites, and clay-slates. Nearly complete change to quartzite and obliteration of primary clastic textures occur here. Grains show serrated contacts and composite blastesis. Silicification increases as a result of increasing temperature and pressure and even lithic fragments are recrystallized. Remnants of earlierformed cements are preserved as rims of chlorite and illite and are included in the blastic patches as separate entities. Where protective films or matrix is present to reduce the influence of interstitial fluids, however, the above alterations are absent. In this zone, the clay is altered to a chlorite plus illite assemblage and the detrital biotite undergoes a similar change. Zone of spine-like aggregates and muscovite-chlorite cement. This zone is developed in the central fields of folded regions and gradually changes to the upper zones of regional metamorphism. This zone is characterized by slates, phyllites and quartzites with complex spine-like structures, which are the result of interdigitation of chlorite and muscovite lamellae that penetrate the quartz and feldspar grains. The segregation of the two platy minerals is particularly evident in slates with large lepidoblasts of chlorite and muscovite set in a groundmass of scaly chlorite, muscovite and quartz. Illite is replacing
DIAGENESIS OF SANDSTONES AND COMPACTION
417
muscovite in various degrees depending on the stage of alteration, which is controlled by the loss of water and an increase in K content. The latter is derived from the decomposition of acid plagioclase and potash feldspars, as shown by corroded and muscovite-replaced orthoclase and microcline. The lower part of this zone consists of minerals that correspond t o the stable equilibrium of stress minerals at low-temperature regional metamorphism, i.e., chlorite, muscovite, quartz, albite, and epidote. Zones 1and 2 described above correspond to the so-called epigenetic stage and the dominant factors are vertical pressure and the influence of subsurface fluids. Zones 3 and 4 are stages of early metamorphism or “metagenesis” controlled by stress, vertical overburden pressure, and temperature. There is a gradation into regional metamorphism. (For a discussion on the various terminologies from diagenesis t o metamorphism, see Dunoyer de Segonzac, 1968, and Teodorovich, 1961.) Like the zones of regional metamorphism, zones l to 4 may cut across the stratigraphic boundaries. Sherwood and Huang (1969)studied highfy indurated carbonate rocks (limestones, dolomitic limestones, and dolomites) and igneous and metamorphic rocks (slates, schists, gneisses, diabases, and granites). They noted fundamental differences in the character of pores between the carbonate rocks and the igneous and metamorphic rocks (which need not necessarily be of general application in other areas): “Pores in the individual crystalline rocks are concentrated predominantly at two t o four different sizes, whereas those in the carbonate rocks generally are of one size. Also the mean pore diameters in the carbonate rocks are generally smaller; the pores are of a lesser size range than those in the igneous and metamorphic rocks, but the porosity of the carbonate rocks generally is higher and shows a greater variation. The low porosity and small pores measured in the dolomitic rocks from Virginia appear to indicate that most dolomitization preceded the strong compactive forces of Appalachian folding.’’ If studies such as these are performed on all major rock types in a sedimentary basin or in a region where mineralization in specific rock units has occurred, the results may suggest reasons for preferential or differential fluid movements and/or diffusion, as has been demonstrated in some cases already. As to diffusion processes per se, Elliott (1973)reviewed the theoretical data of the diffusion flow laws and included in his discussions pressure solution phenomena in sedimentary and metamorphic rocks. He found that microscopic observations could allow a clear distinction between deformation dominated by dislocation processes and deformation accompanied by mass transfer by diffusion. He concluded, however, that no microscopic criteria exist at the present time to recognize the path along which diffusion has occurred. As to detailed investigations of clay-mineral changes within sandstone
418
K.H. WOLF AND G.V. CHILINGARIAN Al,Si,O,,
A I z S ~ ~ O , o ~ O nH,O tI~~ idealized montmorillonite ( w i t h No, M Q )
(OH),
+si K AI,(Si,AI)
O,(OH),
ideolized i l l i t e
DEHYDRATION‘
Fig. 3-274. Possible dehydration among the aluminous clay minerals. Glauconite is chemically intermediate between the three molecules indicated. (After Bayly, 1968, fig. 13-10; courtesy Prentice-Hall, Englewood Cliffs, N.J.) Glauconite is of very early diagenetic origin, whereas other minerals may form as a product of late diagenesis and/or burial metamorphism within the sedimentary rock.
units during burial, much remains to be done, e.g., studies on the partition of chemical elements among coexisting phases. Figure 3-274 (Bayly, 1968) shows the possible diagenetic and epigenetic trends during burial in a sedimentary basin, excluding glauconite genesis of very early diagenetic origin; however, more complicated models with numerous intermediate stages can be set up, which is beyond the purpose of this chapter (e.g., Velde and Bystrom-Brusevitz, 1972). Another approach in which the length/thickness ratio of biotite is used, is discussed by Jones and Galway (1972)and Etheridge (1973).Although their approach was applied to higher-grade metamorphism, possibly one can use similar techniques to study minerals formed as a result of deep burial. Referring t o Kubler (1967,1968),Ludwig (1972)outlined a procedure by which the “crystallinity” of illite (i.e., the width of the 10 A illite peak at half height) is used in conjunction with index minerals t o determine the grade of regional metamorphism (Fig. 3-275).The X-ray pattern of illite clay minerals changes as a result of diagenesis and metamorphism. Thus, by using standards from illite-containing rocks of specific grade of secondary alterations, one can use the X-ray patterns to define the degree of secondary modifications. Four illite standards (i.e., Morris, Fithian, Marble Head, and OECD illites) were used after treatment with ethylene glycol (Fig. 3-275).The limit between anchimetamorphism sensu Kubler (= very low-grade metamorphism sensu Winkler) and epimetamorphism (= low-grade) is located at 4 mm along
DIAGENESIS OF SANDSTONES AND COMPACTION
419
-
“INCIPIENT TO WEAK METAMORPHISM:
r------------
’ A\
DIAGENESIS
I
OECD
I
‘\‘c SOmm
I
2
3
1
I
4
5
6
Fig. 3-275.Weaver’s “sharpness ratio” (= ratio of height of 10 8, peak to height of 10.5 8, peak; ordinate) versus Kubler’s “crystallinity” (i.e., width of the 10 illite peak at half height; abscissa: 0-20 mm, compared to quartz = 1-6 /A) for four illite standards in the grain fraction of 2-6.3 p. TS = Schwarzschiefer (black slate standard sample from the Staatliche Geologische Kommission, Berlin-Ost); TI3 = Tonschiefer (clay slate, source as for 2’s);1 = Morris illite; 2 = OECD illite; 3 = Fithian illite; 4 = Marble Head illite. Arrows indicate the oriented specimens treated with ethylene glycol. The limit between “anchimetamorphism” sensu Kubler (“very low stage of metamorphism” sensu Winkler) and “epimetamorphism” (“low stage”) is drawn at 4 mm. This corresponds to the value of 1.25, if the width at half height of the 4.26 8, quartz peak (3.2 mm) is taken as unit ( = 1.00). (Modified after Ludwig, 1972,fig. 1;courtesy Neues Jahrb. Geol. Paluontol.)
the abscissa. This corresponds to the value of 1.25, if the width at half height of the 4.26 8 quartz peak (3.2 mm) is taken as unit value (= 1.00). The above technique may also be useful in the study of progressive changes from syngenesis through diagenesis to burial and higher grades of metamorphism by using illite and other minerals, as demonstrated by Frey (1970), Frey and Niggli (1971), and Frey et al. (1973). The reader may wish t o refer to the publication by Velde and Bystrom-Bmsevitz (1972), who have experimentally studied the illite-montmorillonite evolution as a result of simulated burial metamorphism. The older studies of metamorphism of sedimentary piles have not as heavily leaned on the knowledge of the earlier diagenetic history of the sedimentary rocks as would appear to be desirable, so that, in a way, the metamor-
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K.H. WOLF AND G.V. CHILINGARIAN
phism has been frequently considered somewhat “out-of-context”. Although the general composition of the pre-metamorphic rocks was always considered, more precise data is required in the future, and compaction must find its deserved place in these investigations. One example, provided by Huang and Wyllie (1973), is presented here. Although the petrologic and geochemical details fall outside the scope of this chapter, the information by Huang and Wyllie demonstrates, for example, the importance of interstitial water during metamorphism. The secondary processes and results thereof are different for sediments containing water as contrasted to those lacking fluids. Huang and Wyllie pointed out that plate tectonics, with its concept of consuming plate margins or sinking lithosphere slabs, can throw light on the styles of metamorphism occurring along the continental margins. Many of the intermediate and acid calc-alkaline magmas of the island arc and continental margins are derived from the slabs by progressive partial fusion of the siliceous sediments and oceanic crust of the slabs. They proposed experiments to solve some of the remaining problems: one should evaluate proposed natural processes by comparing the results of melting relationships of the theoretically inferred source rock materials (i.e., oceanic crust with sediments), on one hand, with those of the observed products of magmatic activity, on the other. The pressure range used would simulate that from the surface to great depth. The starting materials employed in their experiments were similar to many granitic plutonic rocks and may represent the lowmelting portion of some subducted oceanic sediments with a composition capable of metamorphism to muscovite-bearing quartzo-feldspathic rocks. Inasmuch as the compaction history of fine-grained and coarse-grained sediments determine the amount of fluids retained in a stratigraphic section, and because the sedimentary-volcanic pile will undergo burial and higher-grade metamorphism (related to plate consumption or not), the H20content will determine the paths of the secondary processes (compare figs. 1 and 2, in Huang and Wyllie, 1973). Consequently, the more complete and precise the available data is on the compaction history of sedimentary and volcanic deposits, combined with the information obtained from laboratory experiments, the more factual will be the theoretical reconstructions of metamorphism based on observed mineralogy, textures, and geochemistry. In their section of anatexis of metamorphic sediments in subduction zones, Huang and Wyllie pointed out that the following information must be available: (1)compositions and mineralogy of the rocks including water content and composition; (2) phase relationships of the materials present; and (3) temperature distribution in and around the subducted slab as a function of pressure, or depth. They proposed a model for magma generation with a zone of partial fusion, which gradually melts the lithosphere as the temperature increases. The first liquids originate from sediments or their
DIAGENESIS OF SANDSTONES AND COMPACTION
421
TEMPERATURE "C Fig. 3-276.A model for anatexis of subducted oceanic sediments in pressure (depth)temperature projection showing two episodes of melting. The curves b-u and b'-a' are the solidus curves from figs. 1 and 2 in the original paper. The dashed curves indicate published estimates of temperature distribution along the surface, and 1 km below the surface of a subducted lithosphere slab, according to Oxburgh and Turcotte (1970).The dotted curve indicates another estimate along the surface of a subducted lithosphere slab published by Toksoz et al. (1971). (After Huang and Wyllie, 1973, fig. 4; courte-7 Springer, Berlin.)
metamorphosed equivalents, whereas the oceanic crust will melt at greater depths. The paths for the surface of the slab and at a depth of 1 km are given in Fig. 3-276by the dashed line that bounds the stippled area, indicating when the sediments melt. Inasmuch as the solidus curve for material that contains excess water differs from the one for dry sediments (see figs. 1,2,and 3,in Huang and Wyllie, 1973), there are at least two solidus curves in Fig. 3-276 (i.e., a-b and ~ ' 4 ' ) . According to Garrels and Mackenzie (1971),the metamorphism of many ocean sediments could form metamorphic rocks composed dominantly of mica, quartz, and two feldspars, so that it is possible to use the phase diagrams for predicting the melting behavior of subducting sediments. The line a-b presents the conditions in the case of sediments with interstitial aqueous pore fluids, whereas line a'-b' is for sediments with all free water removed, possibly as a result of compaction. When subduction of sediments occurs with a lithosphere slab dipping at 45", then the temperature at the surface of the slab increases along the dotted line a-a' in Fig. 3-276.In the case where pore water was trapped in the sediments and was not removed by mechanical and/or chemical diagenesis, melting begins at point a, around 60
422
K.H. WOLF AND G.V. CHILINGARIAN
km below the surface. The first liquid would be saturated with water. Inasmuch as all of the pore fluids dissolve in a narrow temperature range, progressive fusion would result in a magma that is HzO-undersaturated (see fig. 3 in Huang and Wyllie). Although Huang and Wyllie have presented more information, the above discussion seems to be sufficient for the present purpose of pointing out that there are several gaps in the petrogenetic knowledge related to the boundaries (or transitional zones) between sediments and sedimentary rocks and sedimentary rocks and their metamorphic equivalents, which in turn change to igneous-like rock types. To close this gap in the genetic data, numerous models have to be established that include a complete paragenesis from early to late diagenesis, through low- and high-grade metamorphism, into the realm of origin of igneous rocks by remelting or fusion (= anatexis). A comprehensive study of compactional diagenesis would be required to determine the reasons for the amount of water retained in sedimentary-volcanic piles, as this quantity, in turn, will determine the fusion and mechanisms of igneous petrogenesis. STRUCTURES IN SEDIMENTARY ROCKS AS A RESULT OF COMPACTION
Enough data is available already on the origin of structures formed by the various mechanical and chemical compaction processes that could form the basis for a separate chapter. This is, however, beyond the scope of the present book, so that only a brief, alas very incomplete, list of references is given in tabular form without discussion, and the reader is referred to the corresponding references (see Table 3-LXVIII).
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TABLE LXVIII Structures and related features formed or influenced by compaction Features formed by andfor influenced by compaction
References to pertinent literature
(1) Stylolites
Lerbekmo and Platt (1962);Golding and Conolly (1962); Trurnit (1968a,b, 1967); Schidlowski and Trurnit (1966); Park and Schot (1968) Franks (1969) Clifton (1965); Judson and Barks (1961) Gocht (1973); Fursich (1973) Rhoads (1970)
(2) Cone-in-cone (3) Polish on pebbles (4) Deformation of fossils by compaction (5) Reworking of sediments by organisms as a function of compaction; rheologic properties of sediments (6) Cleavage as a function of compaction (7) Concretions and their relation to compaction (8) Miscellaneous post-depositional structures influenced by compaction and rheology, e.g., load structures (9) “Quickstone” genesis (10) Mud lumps (11 ) Sandstone dikes (12) Faults formed by compaction (13) Sandstone geometry
(14) Surface topography due to compaction (15) Relationship between compaction and sedimentation, e.g., cyclicity (16) Dip changes due to compaction (17) Classification of sedimentary structures that consider compaction and rheology
Dunnet and Moore (1969); Alterman (1973); Powell (1972a,b) Oertel and Curtis (1972); Rukhin (1958); Fursich (1973) Dzulynski and Walton (1965); Swarbrick (1968); Elliston and Carey (1963); Wobber (1967) McNeill(l966); Boswell(l963) Dickey (1972) Steinitz (1970); Bull (1972); Shelton (1962); Conybeare and Crook (1968); Peterson (1968) Carver (1968); Phillips (1972); Powell (1972a,b); Bruce (1973) Pettijohn et a]. (1972); Baldwin (1971); Brown (1969); Rittenhouse (1961); Conybeare (1967); Bloom (1964); Kaye and Barghoorn (1964) Hrabar and Potter (1969); Dolly and Bush (1972); Halbouty (1972); Wobber (1967); Rittenhouse (1961); Brown (1969) Duff et al. (1967); Halbouty (1972) Rittenhouse (1972); Conybeare (1967); Borradaile (1973) Elliott (1965)
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REFERENCES Aalto, K.R., 1972. Diagenesis of orthoquartzites near Bogota, Colombia. J. Sediment. Petrol., 42: 330-340. Adams, W.L., 1964. Diagenetic aspects of Lower Morrowan, Pennsylvanian, sandstones,, northwest Oklahoma. Bull. A m . Assoc. Pet. Geol., 48: 1568-1580. Akal, T.,1972. The relationship between the physical properties of underwater sediments that affect bottom reflection. Mar. Geol., 13: 251-266. Allen, J.R.L., 1962. Petrology, origin and deposition of the highest Lower Red Sandstone of Shropshire, England. J. Sediment. Petrol., 32: 657-697. Allen, J.R.L., 1969. Notes towards a theory of concentration of solids in natural sands. Geol. Mag., 106: 309-321. Allen, J.R.L., 1970. The systematic packing of prolate spheroids with reference to concentration and dilatency. Geol. Mijnbouw, 49 (3):211-220. Alterman, I.B., 1973. Rotation and dewatering during slaty cleavage formation: some new evidence and interpretations. Geology (3outder), 1 : 33-36. Amstutz, G.C., 1959. Syngenese und Epigenese in Petrographie und Lagerstattenkunde. Schweiz. Mineral. Petrogr. Mitt., 39;2-84. Angel, F., 1965. Retrograde Metamorphose und Diaphthorese. Neues Jahrb. Mineral. Abh., 102: 123-176. Archie, G.E., 1950.Introduction to the petrographics of reservoir rocks. Bull. Am. Assoc. Pet. Geol., 34: 943-961. Athy, L.F., 1930a. Density, porosity and compaction of sedimentary rocks. Bull. Am. Assoc. Pet. Geol., 14: 1-24. Athy, L.F., 1930b. Compaction and oil migration. Bull. A m . Assoc. Pet. Geol., 14: 25-35. Athy, L.F., 1934. Compaction and its effect on local structure. In: W.E. Wrather et al. (Editors), Problems of Petroleum Geology - a Sequel to Structure of Typical American Oil Fields. A m . Assoc. Pet. Geol., Sydney Powers Memorial Volume: 811-823. Atwater, G.I. and Miller, E.E., 1966. The effect of decrease in porosity with depth on future development of oil and gas reserves in south Louisiana. A m . Assoc. Pet. Geol., Progr. Annu. Meet., New Orleans, p. 48. Baker, C., 1972. Aquathermal pressuring-role of temperature in development of abnormal-pressure zones. Bull. Am. Assoc. Pet. Geol., 56: 2068-2071. Baker, D.R., 1972. Organic geochemistry and geological interpretations. J. Geol. Educ., 20: 221-234. Baker, D.R. and Claypool, G.E., 1970. Effects of incipient metamorphism on organic matter in mudrock. Bull. A m . Assoc. Pet. Geol., 54: 456-468. Baldwin, B., 1971. Ways of deciphering compacted sediments. J. Sediment. Petrol., 41: 293-3 01. Bayly, B., 1968.Introduction to Petrology. Prentice-Hall, Englewood Cliffs, N.J., 371 pp. Beard, D.C. and Weyl, P.K., 1973. Influence of texture o n porosity and permeability of unconsolidated sand. Bull. A m . Assoc. Pet. Geol., 57: 349-369. Beaudry, D.A., 1950. Pore-Space Reduction in Some Deeply Buried Sandstones. Thesis, Univ. Cincinnatti, unpublished. Berg, R.R. and Davies, D.K., 1968. Origin of Lower Cretaceous muddy sandstone at Bell Creek Field, Montana. Bull. Am. Assoc. Pet. Geol., 52: 1888-1898. Berner, R.A., 1971. Principles of Chemical Sedimentology. McGraw-Hill, New York, N.Y., 240 pp.
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Berner, R.A., 1972. Chemical kinetic models of early diagenesis. J. Geol. Educ., 20: 267272. Billings, G.K., Hitchon, B. and Shawe, D.R., 1969.Geochemistry and origin of formation waters in the Western Canada sedimentary basin, 2. Alkali metals. Chem. Geol., 4: 211-223. Bishop, D.G., 1972. Authigenic pumpellyite and other metamorphic effects in the Kyeburn Formation, central Otago. N . Z . J. Geol. Geophys., 15:243-250. Blair, G.W.S., 1969. Elementary Rheology. Academic Press, New York, N.Y., 158 pp. Blanche, J.B., 1973. The Rotliegendes Sandstone Formation of the United Kingdom sector of the southern North Sea Basin. Znst. Min. Metall., Trans., Sect. B : 85-89. Blatt, H., 1966. Diagenesis of sandstones: processes and problems. In: Recently Developed Geologic Principles and Sedimentation o f the Permo-Pennsylvanian o f the Rocky Mountains-Wyo. Geol. Assoc., 12th Annu. Conf., pp. 63-65. Blatt, H., 1967. Original characteristics of clastic quartz grains. J. Sediment. Petrol., 37: 401-424. Blatt, H. and Christie, J.M., 1963, Undulatory extinction in quartz of igneous and metamorphic rocks and its significance in provenance studies of sedimentary rocks. J. Sediment. Petrol., 33: 559-579. Blatt, H., Middleton, G.V. and Murray, R., 1972.Origin o f Sedimentary Rocks. PrenticeHall, Englewood Cliffs, N.J., 634 pp. Bloom, A.L., 1964.Peat accumulation and compaction in a Connecticut coastal marsh. J. Sediment. Petrol., 34: 599-603. Blount, C.W. and Dickson, F.W., 1969. The solubility of anhydrite (CaS04) in NaClHzO from 100' t o 450'C and 1 to 1,000bars. Geochim. Cosmochim. Acta, 33:227245. Boehm, P.D. and Quinn, J.G., 1973. Solubilization of hydrocarbons by the dissolved organic matter in seawater. Geochim. Cosmochim. Acta, 37: 2459-2478. Borg, L.Y. and Maxwell, J.C., 1956. Interpretation of fabrics of experimentally deformed sands. A m . J. Sci., 264: 71-81. Borradaile, G., 1973. Curves for the determination of compaction using deformed crossbedding. J. Sediment. Petrol., 43: 1160. Boswell, P.G.H., 1948. The thivotropy of certain sedimentary rocks. Sci. h o g . , 143: 412-422. Boswell, P.G.H., 1952. The determination of the thixotropic limits of sediments. Liverp. Munch. Geol. J., 1: 1-22. Boswell, P.G.H., 1961.Muddy Sediments. W.H. Heffner, London, 140 pp. Bredehoeft, J.D. and Hanshaw, B.B., 1968. On the maintenance of anomalous fluid pressure, 1.Thick sedimentary sequences. Bull. Geol. SOC.A m . , 79: 1097-1106. Bredehoeft, J.D., Blyth, C.R., White, W.A. and Maxey, G.B., 1963. Possible mechanism for concentration of brines in subsurface formations. Bull. A m . Assoc. Pet. Geol., 47: 257-269. Brewer, R., 1964. Fabric and Mineral Analysis o f Soils. Wiley, New York, N.Y.,470 pp. Brooks, M.,1966. A study of density variations in New Red sandstones from the English Midlands. Geol. Mag., 103:61-69. Brovkov, G.N., 1964. Main features of the diagenesis of the Aalenian Coal Measures of Dagestan. Int. Geol. Rev., 6:912-919. Brown, L.F., Jr., 1969. Geometry and distribution of fluvial and deltaic sandstones (Pennsylvanian and Permian), north-central Texas. Trans. Gulf Coast Assoc. Geol. SOC., 19: 23-47. Brown, L.F., Jr., McGowen, J.H., Seals, M.J., Waller, T.H. and Ray, J.R., 1967.Role of compaction in development of geometry of superposed elongate sandstone bodies. Bull. A m . Assoc. Pet. Geol., 51: 455-456 (abstr.).
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Rubey, W.W., 1927. The effect of gravitational compaction on the structure of sedimentary rocks. Bull. Am. Assoc. Pet. Geol., 11: 133-136, 625-628. Ruhl, W. and Schmid, C., 1967. Uber das Verhaltnis der vertikalen zur horizontalen absoluten Permeabilitat von Sandsteinen. Geol. Jahrb., 74: 447. Rukhin, L.B., 1968.Grundziige der Lithologie. Akademie-Verlag, Berlin, 806 pp. Runnels, D.D., 1969. Diagenesis, chemical sediments and the mixing of natural waters. J. Sediment. Petrol., 39: 1188-1201. Sarkisyan, S.G., 1971. Application of the scanning electron microscope in the investigation of oil and gas reservoir rocks. J. Sediment. Geol., 41: 289-292. Sasaki, A. and Krouse, H.R., 1969. Sulfur isotopes and the Pine Point lead-zinc mineralization. Econ. Geol., 64: 718-730. Sawabini, C.T., Chilingar, G.V. and Allen, D.R., 1974. Compressibility of unconsolidated, arkosic oil sands. SOC.Pet. Eng. J., 14 (3):132-138. Saxby, J.D., 1969. Metal-organic chemistry of the geochemical cycle. Rev. Pure Appl. Chem., 19: 131-160. Scheidegger, A.E., 1957. The Physics o f Flow through Porous Media. Macmillan, New York, N.Y., 236 pp. Scheidegger, A.E. and O’Keefe, J.A., 1967. On the possibility of the origination of geosynclines by deposition. J. Geophys. Res., 72: 6275-6278. Schidlowski, M. and Trurnit, P., 1966. Druck-Losungserscheinungen an Gerollpyriten aus Schweiz. Mineral. Petrogr. Mitt., 46: 337-351. den Witwatersrand-Konglomeraten. Schmidt, G.W., 1971. Interstitial Water Composition and Geochemistry o f Deep Gulf Coast Shales and Sands. Thesis, Univ. of Tulsa. Schmidt, G.W., 1973.Interstitial water composition and geochemistry of deep Gulf Coast shales and sandstones. Bull. Am. Assoc. Pet. Geol., 57: 321-337. Schmidt, V., 1971. Facies, diagenesis and related reservoir properties in the Gigas Beds (Upper Jurassic), northwestern Germany. In: L.C. Pray and R.C. Murray (Editors), Dolornitization and Limestone Diagenesis-A Symposium-Soc. Econ. Paleontol. Mineral., Spec. Publ., 13: 124-168. Schwarcz, H.P., 1966. Chemical and mineralogical variations in an arkosic quartzite during progressive regional metamorphism. Bull. Geol. SOC.A m . , 77 : 509-632. Selley, R.C., 1966. Petrography of the Torridonian rocks of Raasay and Scalpay, Inverness-shire. Proc. Geol. Assoc., 77 : 293-314. Serruya, C., 1969. Problems of sedimentation in the Lake of Geneva. Verh. Znt. Ver. Limnol., Stuttgart, 17: 2019-2018. Serruya, C., Picard, L. and Chilingarian, G.V., 1967. Possible role of electrical currents and potentials during diagenesis (electrodiagenesis). J. Sediment. Petrol., 37 ( 2 ) :
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compaction of natural sediments and notes on allied phenomena. Geol. Mag., 87: 26-40. Surdam, R.C. and Parker, R.B., 1972. Authigenic aluminosilicate minerals in the tuffaceous rocks of the Green River Formation, Wyoming. Bull. Geol. SOC. A m . , 83: 689-700. Swarbrick, E.E., 1968. Physical diagenesis; intrusive sediment and connate water. Sediment. Geol., 2: 161-175. Swolfs, H.S. and Friedman, M., 1972. Mechanical and chemical effects of pore fluids on rock properties. Bull. A m . Assoc. Pet. Geol., 55: 2090 (abstr.) Talash, A.E. and Crawford, P.B., 1965. Rock properties computed from random pore size distribution. J. Sediment. Petrol., 35: 917-921. Tanner, W.F.,Jr., 1963. Crushed pebble conglomerate of southwestern Montana, J. Geol., 71: 637-640. Taylor, J.M., 1950. Pore-space reduction in sandstones. Bull. A m . Assoc. Pet. Geol., 34: 701-7 16. Taylor, S.R., 1968. Geochemistry of andesites. In: L.H. Ahrens (Editor), Origin and Distribution o f the Elements. Pergamon Press, Toronto, pp. 559-583. Teodorovich, G.I., 1961. Authigenic Minerals in Sedimentary Rocks. Consultants Bureau, New York, N.Y., 120 pp. Teodorovich, G.I. and Chernov, A.A., 1968. Character of changes with depth in productive deposits of Apsheron oil-gas-bearing region. Sou. Geol., 1968 (4): 83-93. Terman, M.J., 1973. Nuclear-explosion petroleum-stimulation projects, United States and U.S.S.R. Bull. A m . Assoc. Pet. Geol., 57: 990-1026. Terzaghi, K. and Peck, R.B., 1968. Soil Mechanics in Engineering Practice. Wiley, New York, 2nd ed., 84 pp. Terzaghi, R.A.D., 1940. Compaction of lime mud as a cause of secondary structure. J. Sediment. Petrol., 10: 78-90. Teslenko, P.F. and Korotkov, B.S., 1967. Effect of arenaceous intercalations in clays on their compaction. Znt. Geol. Rev., 9: 699-701. Thomson, A., 1959. Pressure solution and porosity. In: Silica in S e d i m e n t s S o c . Econ. Paleontol. Mineral., Spec. Publ., 7: 92-110. Tickell, F.G. and Hiatt, W.N., 1938. Effect on angularity of grain on porosity and permeability of unconsolidated sands. Bull. Am. Assoc. Pet. Geol., 22 : 1272-1274. Tissot, B., Califet-Debyser, Y., Deroo, G. and Oudin, J.L., 1971. Origin and evolution of hydrocarbons in Early Toarcian shales, Paris basin, France. Bull. A m . Assoc. Pet. Geol., 55: 2177-2193. Tissot, B., Durand, B. Espitali6, J. and Combaz, A., 1974. Influence of nature and diagenesis of organic matter in formation of petroleum. Bull. A m . Assoc. Pet. Geol., 58: 499506. Toksoz, M.N., Minear, J.W. and Julian, B.R., 1971. Temperature field and geophysical effects of a downgoing slab. J. Geophys. Res., 76: 1113-1138, Towe, K.M., 1962. Clay-mineral diagenesis as a possible source of silica cement on sedimentary rocks. J. Sediment. Petrol., 32: 26-28. Trask, P.D., 1931. Compaction of sediments. Bull. A m . Assoc. Pet. Geol., 15: 271-276. Trask, P.D., 1959. Effect of grain size on strength of mixtures of clay, sand and water. Bull. Geol. SOC.A m . , 70: 569-580. Troger, W.E., 1963. Der geothermische Gradient im pt-Feld der metamorphen Facies. Beitr. Mineral. Petrogr., 9: 1-12. Trurnit, P., 1968a. Pressure-solution phenomena in detrital rocks. Sediment. Geol., 2: 89-1 14.
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Chapter 4 COMPACTION AND DIAGENESIS OF VERY COARSE-GRAINED SEDIMENTS J.R. HAILS
INTRODUCTION
An extensive review of the literature has shown that the compaction of very coarse-grained sediments, with a diameter in the range of granules through pebbles and cobbles t o boulders (2 mm to larger than 256 mm), is a geotechnical process which has not been studied in any great detail. Even authors of basic and advanced soil mechanics and civil engineering texts pay little attention to this particular subject which covers, on average, little more than a few paragraphs. Most information on the subject has been published in soil and engineering laboratory reports, but the quantitative data contained in these technical papers are somewhat limited compared with the data available on sands and finer-grained sediments. One seemingly obvious reason for this disparity is the fact that engineering geologists, besides civil and soil engineers, have focused their attention on soil mechanics in order t o obtain quantitative information on the geotechnical properties of sands, silts, and clays. A second reason may be that, from both the geological and engineering points of view, very coarse-grained sediments do not seem to present, as a rule, so many difficulties as clays in foundation design. Other possible reasons include the difficulty of recovering gravel samples without disturbing their original texture and structure(s) during coring operations, and the impracticability of petrological studies in the laboratory. Nevertheless, studies have been made of the effects of the shape and grain-size distribution of particles and texture of particle aggregates upon the mechanical properties of aggregates in the context of their suitability as foundation and embankment materials. The purpose of this chapter is to summarize briefly the more important general facts pertaining to the geological and engineering properties of very coarse-grained sediments, as they influence compaction. Because of the limited available data, this review is inevitably subjective in nature and covers a fairly wide spectrum of topics instead of dealing with a specific sample from an extensive literature. Perhaps at this stage it is useful to define compaction and consolidation as they are used in this chapter, because there is some disagreement between geologists and civil and soil engineers as to the exact meaning of the terms
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(see, for example, Handy, 1972, and Parker, 1972).In an attempt to resolve this problem and in order to standardize vocabulary, the definitions contained in the Glossary of Geology (Gary et al., 1972)are adopted here. The term compaction may be defined as a reduction in bulk volume or thickness of, or the pore space within, a sediment body in response to the increasing weight of the overlying material that is continually being deposited, or to the pressures resulting from earth movements within the crust. It is expressed as a decrease in porosity brought about by a tighter packing of the sediment particles. Consolidation is any process whereby loosely aggregated earth materials (such as gravels) become firm or coherent rock (conglomerate). Thus, a consolidated rock need not be compacted. In the fields of soil mechanics and civil engineering, consolidation refers t o the reduction of volume of soil under load, owing t o drainage from the pore spaces which results in a closer packing of particles. The reader’s attention is drawn to this definition, with regard to the results of laboratory compaction studies referred to later in this chapter. According to Fraser (1935)porosity and permeability vitally influence the compaction and consolidation of sediments, but it must be borne in mind that both the porosity and permeability of clastic sediments are subject to marked changes in response to metamorphic and diagenetic processes during geologic time. Porosity is defined here as the ratio of the volume of interstices in a rock or soil to its total bulk volume, whereas the permeability or perviousness of a rock, sediment or soil is its capacity for transmitting a fluid without impairment of the stmcture of the medium. Its value depends, in part, upon the size and shape of the pores, the size and shape of their interconnections, and extent of the latter (see Chapter 3,pp. 188-241). In order to study as comprehensively as possible the rate and amount by which very coarse-grained sediments compact, it is necessary to determine and to evaluate, if possible, how and to what extent their physicochemical properties broadly interact with each other during loading by superincumbent sediments, or during mechanical compaction by vibratory or static pressure methods. Soil mechanics parameters, such as apparent cohesion (the cohesion between particles in granular soils due to capillary forces), shear strength and moisture content, are mentioned, because these factors are of considerable significance with respect t o those sediments having a high matrix content. Inasmuch as the textural (grain size, shape and degree of sorting) properties of coarse-grained sediments affect compaction and also, in general, reflect their sedimentological history, whether it has been relatively short or long, there is an indirect relationship between compaction and environment of deposition. Consequently, the way in which glaciers, rivers, waves and currents transport and sort sediments, and contribute to packing density, porosity, permeability and ultimately compaction are also considered briefly.
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GRAVELS
Introduction According to most standard references (e.g., Pettijohn, 1957), gravels consist of a framework of pebbles, cobbles, and boulders with voids or openings between these framework components. The voids are rarely empty and are generally filled with detritus of various sizes, together with precipitated cements. Compared with other types of sediment, gravels are generally restricted in their areal distribution, and most commonly occur locally as shoestring bodies infilling river channels, or as beach and glacial outwash deposits. There appears, however, to be no general agreement on the exact proportion of constituents comprising a gravel. Most textbooks, technical reports, and scientific publications cite a gravel as containing between 10 and 50% gravel-size particles.* For example, according to Pettijohn (1957), “actual analysis shows that the field geologist is prone to call a deposit gravel even if pebbles and like sizes form less than one half of the whole. Some rocks, such as tillite or indurated boulder clay containing less than 10% of gravel-sized fragments, are designated, none the less, as conglomerates”. Therefore, sensu stricto, sandy gravels, pebbly sands, as well as glacial tills, fanglomerates, gravelly soils and other relatively poorly sorted sediments with a wide range of particle size, legitimately fall within the category of very coarse-grained sediments and the ambit of this chapter. Individual gravel particles can approximate, if they are well-rounded and symmetrical, the shape of a sphere, a spheroid (prolate or oblate), or an ellipsoid. These latter solids are simply defined mathematically by a secondorder equation which, in turn, contributes to simpler quantitative analysis. For purposes of comparison these solids may be used widely in geology, because various terms such as disc-, tabular-, rod-, and roller-shaped (commonly referred to as cuboidal, elongate, and flaky in aggregate testing) imply shapes which approximate closely to one or other of them. Considerable attention has been paid to roundness and sphericity of gravels (Wentworth, 1922; Wadell, 1932, 1933, 1934; Zingg, 1935; Krumbein and Griffith, 1938; Cailleux, 1945, 1947; and Berthois, 1951, amongst others). Yet, none of these researchers has considered the effects of shape or grain size on the compaction characteristics of these very coarse-grained sediments.
* Editorial
note: See classification scheme outlined in the Introduction chapter of Vol. I, figs. 1-6 and 1-7, and tables 1-11 and 1-111.
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Clean gravels Gravels with unfilled voids are invariably referred to as open-work gravels. In this chapter, clean gravels are defined by the writer as being entirely free of interstitial material which is smaller than the granule-size fraction (i.e., less than 2 mm), regardless of their sorting characteristics. Such aggregates may be packed densely or loosely, affording the largest amount of void space in the latter case, with any gradation in between. Various configurations of loose and dense packing that might be identified in clean gravels, as envisaged by the writer, are shown in Fig. 4-1,A-N. These are merely diagrammatic models, because the shapes of particles in gravels, conglomerates and other types of very coarse-grained sediment, rarely approximate spheres of uniform size. Obviously, shape varies according to the resistance of individual particles r-------i
C
A
u r-------i
F
G
H
I
J
K
L
M
N
V
P
Fig. 4-1. Diagrammatical representation of particle arrangement and packing density in clean gravels ( A - N ) and in those with a matrix (0-4').Loose packing is illustrated in A, C, E, F, G, and I , and dense packing in B, D, F, G, H, and I.
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to abrasion and weathering, whether chemical or physical, and time. These two processes, in tum, are related respectively to the transportational and sedimentological history of the entire aggregate population. In the writer’s experience, the arrangement of loosely- and densely-packed well-rounded gravels into a limited number of discrete laminae, as exemplified in Fig. 4-1, E-I, is often difficult to see in the field. This can be attributed to the fact that the necessary combination of excellent sorting and uniformity of grain size seldom results from wave or current action, whatever the history of the deposit being studied, even if wave action is highly selective with regard to shape sorting. Inasmuch as clean gravels, as defined here, lack interstitial material, it does not necessarily follow that the grain-to-graincontacts shown in Fig. 4-1will be entirely unaltered, because the effects of pressure-solution phenomena must be considered. It is not the intention to comment at length on pressure-solution phenomena in detrital rocks because, although this subject has been discussed and reviewed in considerable detail by Tmmit (1967, 1968a,b), it has not been explained adequately with regard to compaction sensu stricto. According to Trurnit (1968b), “pressure-solution phenomena have been referred to as pitted pebbles, indentations or impressions in pebbles and sand grains, interlocking grains, microstylolites between pebble and sand grains, stylolites and sutures”. He has attempted to demonstrate “that the geometry of the contact surface depends on the radii of curvature of the partners at. the contact and on the relative pressure solubility of the partners along the direction of stress. The grain size of the partners and the amount of solution residue accumulated also plays a role”. But the actual effect of compaction on the solution of particles in contact with each other in gravels and conglomerates is apparently still obscure. Conjecturally, it might be argued that the effects of solution resulting from compaction could possibly be more pronounced in those deposits composed of either chemically unstable or deeply-weathered particles. If this contention is true, it would seem that clean beach gravels would be least affected, because they rarely contain a high proportion of weathered components as abrasion ultimately eliminates less resistant material either in the surf zone or during transportation from the source area to the point of deposition. Apart from the classic work by Graton and Fraser (1935) on the systematic packing of spheres in relation to porosity and permeability, there has been little research of substance directly related to the arrangement and packing density of pebbles and cobbles in unconsolidated or consolidated gravel populations. Graton and Fraser’s study was focused on the geometry of assemblages of ideal spheres of uniform size and the relations thereto of porosity and permeability. Perhaps it is pertinent to mention here that Fraser (1935) stated that “the
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extent to which depositional processes can affect porosity is partially dependent on their ability in sizing". He was of the opinion that glaciers are probably the least effective in this respect with the consequence that glacial deposits are generally characterized by a wide variety of shapes and particle sizes, which can often be intermixed t o give a mass with relatively low porosity. In contrast, beach and fluviatile deposits display a higher porosity because of better sorting and size grading. The original or primary porosity of unconsolidated clean gravels, then, is largely controlled by the uniformity of particle size, shape, arrangement or packing of individual particles, as well as the process and manner of deposition. Although the initial porosity is determined at the time of deposition, it may be decreased or increased subsequently as a result of compaction and other diagenetic processes. Sometimes in situ weathering, in the case of chemically unstable components, may also alter porosity. The general consensus among engineers and geologists is that the porosity of all types of gravels, as initially deposited, cannot be greatly reduced by compaction until pressures are sufficiently great to cause crushing. Because of the lack of reliable field observations, the pressures required to cause crushing are generally unknown. Thus, in the absence of great pressure, compaction after deposition is a relatively unimportant process and, as Fraser (1935) pointed out, in the great majority of cases post-depositional reductions of porosity result from cementation by introduced mineral material. In the light of these comments, one may well inquire whether or not there is a relationship, in the case of gravels, between packing density" and compaction, bearing in mind the definition of compaction used earlier in this chapter. If no relationship exists, it can be claimed that the packing density of clean gravels reflects their textural characteristics and environment of deposition in toto, whether it is fluviatile, fluvio-glacial, glacial, or marine. Consequently, there is likely to be some variation in this packing density. At least some corroborative evidence for this variation, and data contributing towards a better understanding of the packing density of gravels, has been obtained from studies of shoreline equilibrium in general and beach stability in particular. Carr (1971) used quartz granulites, quartz-jasper conglomerates (conspicuous fragments of jasper and quartzite in an essentially purple-colored matrix of quartzite material, which had been subject t o beach processes on previous occasions), and basalts to study longshore transport and the sorting of pebbles on Chesil Beach, England (Fig. 4-2).He found that the packing
* Packing density is defined as a ~
measure of the extent t o which the grains of a sedimentary rock occupy t h e gross volume of the rock in contrast t o the spaces between the grains, equal to t he cumulated grain-intercept length along a traverse in a thin section (Kahn, 1956). (See Chapter 3 . )
COMPACTION OF VERY COARSE-GRAINED SEDIMENTS
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Fig. 4-2. Variation in mean particle size, transverse section, Chesil Beach, England. Grading is best displayed at high-water mark (HWM).Standard deviation is greatest immediately behind the beach crest where there is coarser, loosely-packed material, a legacy of storm wave action. Extreme backslope (top photograph) shows grading pertaining to an earlier beach feature. (See text for details of a-axes or long diameters; Photo by A.P. Carr.)
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density of the indigenous pebbles varied appreciably according to the wave conditions. Carr also noted, within the various grades of material used in his experiments and from the wave energies encountered, that there is a relationship between pebble size and extent of longshore movement. The best correlations were found with the short axis ( c ) , particularly in the case of the quartz granulite, although workers elsewhere have found better relationships with shape rather than linear parameters. Five hundred particles were measured for each of the sites illustrated in Fig. 4-2. In the vicinity of the backslope (top photograph) the length of the long A-axis had a mean value of 33.8 f 6.7 mm (minimum = 14.7 mm, maximum = 65.8 mm), whereas immediately behind the beach crest there was a coarser admixture apparently thrown over and stranded in storms with a mean value of 45.6 8.0 mm (minimum = 25.8 mm, maximum = 73.2 mm). The best-sorted material was on the beach face at high-water mark (HWM),where the mean value of the A-axis is 39.6 5.3 mm (minimum = 22.4 mm, maximum = 65.8 mm). Particle size and shape vary systematically across gravel beaches, as exemplified in Figs. 4-3A,B, which show the arrangement of granules and pebbles at high-water mark and mid-tide level, respectively, on Slapton Beach, South Devon, England. The individual particles are arranged in such a way that the smaller pebbles and granules are more densely packed than the larger particles, which have been left as lag deposits by the backwash. It might be argued, however, that the larger particles have been rejected in a lateral plane as the beach prograded. The granules at about mid-tide level are fairly uniform in grain size and are generally more angular and more densely packed than the pebbles at high-water mark. Another interesting point is that there is a tendency for the smaller pebbles and granules shown in Fig. 4-3A to occupy voids between the larger particles, probably as a result of filling-in from above. Fraser (1935) examined the degree of packing produced by wave action by comparing the porosities of coarse-grained beach sands and gravels both in the field and in the laboratory. In a beach deposit, composed mainly of pebbles, he found that there is a tendency for medium-sized pebbles to be heaped on the landward side of a boulder and for the larger pebbles and smaller material to be on the seaward side. The medium-sized material, however, tended to collect on both the seaward and landward sides of the boulder, with a larger amount on the landward side, when there were only a few large pebbles present. These observations partly reflect, as Fraser explained, the uneven dissipation of wave energy accompanying swash-backwash action across the beach face. If, as some authorities would argue, the shape of clean gravels reflects their environment of deposition, then there is an interrelationship between the shape, packing density, and mode of deposition. Perhaps more attention should be given to theoretical packing densities of
*
*
Fig. 4-3. Variations in particle size and shape across Slapton Beach, South Devon, England. Smaller pebbles and granules (B) are more densely packed than the larger particles left as lag deposits by backwash near HWM (A). Uniformity in grain size and greater angularity of granules at approximately mid-tide level are particularly well displayed. The coin has a diameter of 25.40 mm.
4 54
J.R. HAILS
well-rounded pebbles or cobbles. According t o an intriguing study of the packing structure of spheres reported by Gardner (1966), there may be some form of irregular packing that would be denser than about 74% grain volume (porosity = 26%), which corresponds to the hexagonal close-packing density. The latter is widely held to be the densest possible packing, although no one has yet proved that denser packing is impossible. Gardner stated that experiments made in studying random packing, by pouring a large number of steel balls into spherical flasks, show that stable random packings have a density that varies from about 0.59 to 0.63 grain-volume fraction. Gardner concluded that if there is a packing denser than 0.74 grain-volume fraction, it will have to be carefully constructed on a pattern that no one has yet considered. Gravels with interstitial material and a high matrix content
So far, then, the thesis has been advanced that clean gravels, unlike sands, are seldom altered by compaction. To what degree they are altered by cementation and recrystallization has yet to be determined. Obviously, the amount of potential cementation should increase as the amount of interstitial material increases assuming, of course, that cementation is not entirely induced by ground-water circulation and percolation. Inasmuch as little is known about the relationship between compaction and other diagenetic processes in gravels and conglomerates, inferences can only be made, with some reservation, from studies of the ultimate bearing capacity of fine-grained soils and pore-space reduction by solution and cementation in sandstones. In the case of sandstones, the interplay of chemical and physical processes is still not understood fully, and it is reasonable to comment here that the role of diagenesis in sedimentary geology has been somewhat neglected compared with most other processes. (See Chapter 3 for present-day status of the data.) Recrystallization leads to low porosity and correspondingly a marked diminution in the pore space volume, unless it is assumed that the pores were filled before crystallization and that prior to the commencement of recrystallization, the sediment had been compressed until all its pores were closed (see, for example, Fraser, 1935). Porosity is controlled by the composition of the individual gravel particles and the interstitial matrix. The quantity of soluble material, such as silica (quartz), calcite, dolomite, and magnesite, that may be precipitated either contemporaneously during deposition or during diagenesis, depends not only upon the physical properties and chemical composition of a gravel population itself, but also upon the composition and weatherability of the overburden, the rate and degree of compaction and consolidation, and, indirectly, the rate of burial. Diagenesis occurs more readily in the more soluble coarse-grained sediments, such as limestone gravels or those with a matrix high in calcium carbonate content, than in the
COMPACTION OF VERY COARSE-GRAINED SEDIMENTS
455
sediments composed of less soluble material, such as silica. As a result, porosity often increases in the sediments composed of more soluble minerals if the rate of solution exceeds that of precipitation. Although it has been mentioned already that the original porosity of clastic sediments influences both compaction and consolidation, these processes, in turn, influence the development of secondary porosity so as to alter the original porosity. The porosity of sandy gravels, like very coarsegrained sandstones and pebbly sandstones, may decline because of additional cementation in response to pressure. On the other hand, the porosity of such deposits may increase as a result of fracturing or by dissolution during diagenesis and epigenesis. Bearing in mind the limited available evidence from field investigations at present, it is difficult to establish what effect fracturing, cementation and compaction will have on porosity in other types of very coarse-grained sediments. It seems reasonable to suggest that cementation and compaction will play a more important role in fluviatile or glacial gravels with a heterogeneous matrix, whether or not it is comprised of soluble or insoluble material. Gravels with isolated particles in a predominantly fine-grained matrix (cf. lower part of section in Fig. 4-1,O) should compact more readily than clean gravels under static loads, or by ramming, rolling and vibration. In the situation illustrated in Fig. 4-1,0, it is generally more difficult to compact the upper part of the horizon showing particle-to-particle contacts and void spaces than the lower part, where the fine fraction is uniformly distributed throughout the large and small spaces between the subrounded and subangular gravel particles (see also Fig. 4-19).Thus, the density of the interstitial material is fairly uniform throughout the two horizons as well. In fact, compaction tests of gravelly soils, in particular, have shown that if the gravel content is small, in the order of 10% or less, the density of the fine fraction is not affected by the presence of the gravel. The greater the variability of the gravel matrix in terms of mineralogical composition and proportion to the gravel, the more difficult it is to specify and control the moisture content requirements for artificial compaction (e.g., stabilization of building foundations), not ignoring of course density needs as well. Moisture content is one of the primary factors affecting the compaction of sediments and its importance increases with increasing clay content.
Gravelly soils and other very coarse-grained sediments Significant soil mechanics and applied geological work on the compaction characteristics of gravelly soils has been undertaken in the laboratory, often under controlled conditions, and documented in various research and soils
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engineering reports and technical memoranda (e.g., Maddison, 1944;Shergold, 1953;Lewis, 1954;Holtz and Lowitz, 1959;Merriman and Houston, 1963; Gordon et al., 1964; Pike, 1972; and Dean and Southgate, 1972, amongst others). For example, Shergold (1953)described a method developed by the Road Research Laboratory, Crowthorne, England, for expressing the degree of angularity of a gravel in terms of its percentage of voids when compacted in a standard way. This proportion was found to vary from about 33% in a well-rounded beach gravel to about 45% in a very angular crushed rock. Since the publication of this paper there has been little information published on the subject. The results published to date provide an overall picture of the parameters 140
-
135 -
130 -
r a
3 > t
125
-
B
120 -
115 -
110-
0
Fig. 4-4. Relation between dry density and moisture content obtained with different types of compaction equipment (plant) when fully compacting loose gravel-sand-clay layers, 22.5 cm thick. (After Lewis, 1954, fig. 8a, p. 20.)
COMPACTION OF VERY COARSE-GRAINED SEDIMENTS
457
influencing the compaction of gravelly soils and very coarse-grained sediments. The general application of results is often difficult, however, owing to the number of interrelated factors influencing compaction, which are effective concurrently. Interest has been directed in road research towards the moisture content at which soil should be compacted, the degree of compaction that should be obtained, and the performance of various equipment in compacting soil (Figs. 4-4 and 4-5). So far, tests have shown that angular gravel is more stable and, therefore, is better suited for use as a base-course in highway construction than well-rounded gravel. This factor is partly exemplified in Figs. 4-1,K-N, which show that angular particles tend to interlock and, GRAVEL -SAND-CLAY Limit of rolling
0 4-cwt vibrating smooth-wheeled roller 2k-ton vibmting m d - w h e e i e a mlkr ziton vibrating-plate compactor l k - t o n vibmting-@ate cmpactor 0 4-cwt vibrating-plate compactor H 40-h.p. tmck-laying tractor X 80-h.p track-laying tractor 2-cwt power-rammerfMean resat,
A A
-
‘ation line voids)
I air
\
MOISTURE
COMENYper cent
16
18
Fig. 4-5. Relation between dry density and moisture content obtained with different types of compaction equipment (plant) when fully compacting loose gravel-sand-clay layers, 22.5 cm thick. (After Lewis, 1954, fig. 8b, p. 21.)
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therefore, provide more strength and higher resistance to deformation than well-rounded gravels. An important problem that needs to be resolved is the degree to which the shearing resistance and mechanical behavior of gravelly soils and other types of very coarse-grained sediments are controlled by the shape of the gravel components and their density. Huang et al. (1963) conducted a laboratory study in order to examine the effect of discernible geometric characteristics of coarse aggregate particles on the compaction characteristics of soil-aggregate mixtures. Gravel and crushed stone were used in their experiment, the shape, angularity and surface texture of which were determined by the “particle index test”. According to Huang et al., “the test is based on the concept that the void conditions in a uniform-sized coarse aggregate when rodded in a standard rhombohedron mold show the combined features of shape, angularity, and surface texture of the aggregate. The result of this test is expressed as the particle index of the aggregate, for which a mass of single-sized, highly polished aluminium spheres is taken as zero. Typical index values of the aggregates that have been tested in the aggregate laboratory at the University of Illinois range from about 4 for a gravel composed of rather spherical particles with rounded corners and smooth surface to about 20 for a crushed limestone of flaky particles with angular comers and edges and a very rough surface”. The gravel, obtained from Wisconsian glacial outwash deposits, was composed of a heterogeneous assortment of limestone, dolomite, basic igneous rocks (unspecified), quartzite, sandstone and some chert. In contrast, the crushed stone was derived from Pennsylvanian fine-grained limestone. The soilaggregate materials were artificially prepared according to a mathematical expression to yield nine different and fully controlled gradations. Though both the gravel and crushed stone contained varying amounts of particles of different shapes, the latter was characteristically more angular and rougher in surface texture than the gravel material. Two main conclusions have been drawn from the results of the testing. First, the geometric characteristics of the coarse aggregates, as indicated by the “particle index”, have a definite bearing on the resultant void characteristics of a compacted soil-aggregate mixture. There appears to be an almost linear relationship between void content in a compacted sample and the particle index of the coarse aggregate. The percentage of voids in a compacted sample increased with increasing values of particle index. Second, for the soil-aggregate mixture containing coarse aggregates of given geometric characteristics, there is an optimum gradation index at which maximum dry density is achieved under a given compactive effort. For a mixture with a 19.05 mm maximum aggregate size, this gradation index is equal to 0.50. The effect of particle shape, porosity, and moisture content on the shear strength of coarse-grained road construction aggregates (maximum size = 25
COMPACTION OF VERY COARSE-GRAINED SEDIMENTS
459
mm) has been evaluated by Eerola and Ylosjoki (1970) by means of a large-diameter (25 cm) vacuum triaxial apparatus (Geonor k-900). They tested the samples of: (1) rounded granitic gravel; (2) normal macadam* crushed from granitic rock and containing 25% schists, and (3) elongated macadam crushed from amphibolic rock. The samples were sorted so as to acquire the same gradation curve for all the samples. The results of the study disclosed a difference of 13" between the friction angles of the rounded gravel and the elongated macadam with the same gradation.** This difference corresponds to an increase of 31% in the friction angle of rounded gravel. Eerola and Ylosjoki inferred from their tests that the water content produces an effect chiefly upon the friction angle of the material and the soil structure originating during the compaction work. One may well inquire whether or not the mineralogical differences between the gravels have influenced the friction angles as well. Research has shown that moisture has a small influence on the friction angle of many minerals. It was found that the rounded gravel was the easiest to compact. With a compaction effort of 1.72 kp/cm2, the porosity of this gravel attained the value of 18%at the optimum water content. With the same compaction effort the porosity was 26%in the case of elongated macadam which was difficult to compact. The minimum compaction effort (0.085 kp/cm2) used in the tests resulted in porosities of 25.8% and 35.3%for the same materials, respectively. The porosity of.norma1 macadam falls in between the values referred to above, inasmuch as the minimum compaction effort produced a porosity of 32.0%. According to Merriman and Houston (1963), shear strength, permeability, and consolidation (i.e. , compaction in geological terminology) are three primary soil properties t o be evaluated when gravelly soils are used for construction purposes. They stated that if the gravel content of the soil is less than the critical gravel content (ie., when gravel particles come into contact and interfere with each other), these properties are primarily governed by those of the fine fraction of the soil-gravel mixture. When the gravel content is greater than the critical gravel content, however, the density of the fine fraction is usually decreased because of increased difficulty in compacting the gravel. Because some of the gravel particles are bearing against, and interlocking with, other gravel particles, the degree of consolidation (compaction) is decreased and the shear strength is increased.
* Made in the manner and with the materials advocated by J.L. McAdam; that is, with successive layers of broken stone of nearly uniform size, each subjected to pressure before the next is laid. ** Samples o f the three materials were sorted so as to acquire the same gradation curve for all samples; this curve falls within the ideal gradation range of the material to be used for road construction.
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Maddison (1944)studied the effect of stone (crushed or natural angular rock particles that will pass a 76-mm sieve) content on the compaction of soil mortar. By using brick earth and a stone aggregate, he found that coarse material impedes the compaction of the soil mortar, which is composed of grains less than 2 mm in size. A slight reduction in density of the compacted soil mortar was noted for increasing stone contents of up to 25%. Density decreased considerably with large stone contents, and the moisture content, corresponding to the maximum density obtainable under standard compaction procedures, increased. The degree of compaction obtained, however, was not affected to any marked extent by the sizes of stone used. Laboratory compaction studies of alluvial deposits, composed of large amounts of subrounded to subangular gravel in a matrix of sandy clay of low plasticity, for the impervious core of the Oroville Dam, California, indicated that maximum compaction was sensitive to small changes in moisture content and that, in the range of rock contents studied, total sample density increased with increase in gravel content (Gordon et al., 1964). It should be pointed out that this trend is partly in conflict with that reported by Maddison (1944). If the moisture content of a gravelly soil can be controlled accurately during compaction, cohesion can be optimized by taking advantage of capillary forces. Inadequate compaction leaves a large void content, which may lead to differential settlements over a period of time. Holtz and Lowitz (1957) conducted tests, including individual Proctor tests, on sandy, silty and clayey soils to which a subangular-to-subrounded river gravel (4.76 to 76.20 mm in size) was added in amounts to provide 20, 35, 50 and 65% gravel content in the total mixtures. In addition, large-scale tests were performed by a large automatic compaction machine on a silty soil containing 80%gravel, on each type of soil without gravel, and on the gravel material itself. They found that the size and gradation of the gravel has .an effect upon the compaction of the fine-grained fraction and concluded that the finer and less well-graded the gravel, the greater the particle interference. Also, because the particles are smaller, the finer gravel does not transmit the compaction effort to the fine-grained fraction as effectively as the coarser material. With increasing plasticity of the fine-grained fraction, particle interference occurs at higher gravel contents. The maximum density of gravel-clay mixtures is at 65-70s gravel content. After the gravel exceeds the optimal amount, there are insufficient fines to fill the voids between the gravel particles and the density decreases rapidly with increasing gravel content. Thus, it is evident that the way in which very coarse-grained sediments are deposited will not only influence particle-to-particle contacts, but also the total and effective porosity (the percentage of void space) and the permeabil-
COMPACTION OF VERY COARSE-GRAINED SEDIMENTS
461
ity, which is a function of the type of void interconnections. As the gravel content decreases, eventually the gravelly sediments will respond to compactive forces very much like fine-grained soils. The degree of compaction will then depend upon the composition and moisture content of the matrix or soil and the load applied. Studies at various laboratories have shown that some clay soils compact best at a relatively high water content, whereas granular soils may compact t o a much higher density than clays at a relatively low water content. It 'follows from these findings that the matrix is obviously more significant in determining degree of compaction than the gravel content in some gravelly soils, including glacial tills. According to Meade (1968),who has studied the compaction characteristics of alluvial sediments with diverse particle sizes ranging from fine clay to gravel in the San Joaquin and Santa Clara Valleys of California, simple correlations between void ratio (ratio of pore volume to solid grain volume) and particle sorting are, as would be expected, obscured by the more prominent relations of void ratio to overburden load and other historical factors (compaction by earthquake vibrations, etc.). Obviously, unless a laboratory method of compaction simulates compaction in the field, the results obtained may not necessarily be valid. As Meade pointed out, two factors complicated his studies of compaction based on void ratios determined in the laboratory from core samples obtained over a period of three years. Firstly, the void ratios reflect on the one hand the combined effect of natural compaction, owing to the slow long-term increase in effective stress (i.e., pressure = effective overburden load per unit of area) accompanying gradual burial, and, on the other hand, the man-made compaction resulting from the rapid short-term increase in effective stress resulting from decline in artesian head. Perhaps it should be noted here that the historic compaction and land subsidence in the two valleys are clearly related to the depletion of artesian pressure and resulting increase in grain-to-grain stress in the sediments, which is caused by pumping the confined ground water from the sediments faster than it is being replenished. Secondly, data are not available for differentiating the increments of void ratio decrease owing to the natural and man-made increases in effective stress partly because, at the time of coring, the man-made increment was transient (see Chilingarian i d Wolf, 1975, Chapters 2 and 8). Bearing in mind the limitations of laboratory and field compaction simulation studies, it is apposite to comment on other very coarse-grained sediments, such as glacial tills and debris flow deposits. Little is known about their compactability. It seems probable that these types of sediments, like conglomerates and breccias, will respond differently to pressure exerted by the overlying deposits according to the overall composition of their matrix and the degree of weathering sustained by individual mineral and rock frag-
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ments. There will also be notable differences in their packing densities, reflecting their environment of deposition. Culley, (1971) investigated the effect of closed-system freeze-thaw cycling on resilient and residual strains, residual moduli, and volume changes of calcareous, oxidized till specimens (with pebbles 1.27 cm and larger in size removed), that had been compacted to various densities and water contents and subjected t o repetitive loading in a constant triaxial stress system with Ames dials of 0.025-mm and 0.0025-mm calibration. This study, which was undertaken to assess the benefits t o be gained from increasing compaction specifications for subgrades, showed that if compaction of the till was increased to a higher density at a lower water content, strains, modulus, and freeze-thaw effects would be improved to such an extent that a significant reduction in pavement structure would be possible. It also showed that the extent to which a reduction could be made was dependent on the ability of the subgrade to maintain the compaction density and water content throughout its design life. Obviously, there is scope for investigating in more detail the compressive strain behavior of very coarse-grained sediments, including gravelly soils like tills, under loading. This aspect of compaction has not been evaluated as widely as that associated with soil-freezing phenomena such as frost heaving. Factors affecting the vibratory compaction of highway embankment fills have been discussed by Forssblad (1971). These fills range from granular soils with large stones and boulders to clay. In the broadest context, rock fill, granular soils, and gravel are non-cohesive sediments which have high permeability. According to Forssblad, such materials possess high load-bearing capacity in the compacted state and are not susceptible to frost action in Scandinavia. Although up to 10% of fines smaller than 0.06 mm in size can be tolerated normally, the maximum quantity of fines varies according to particle size and other properties of the material with grains smaller than 0.06 mm in diameter. Rockfill and other material comprising large stones must be compacted in thick layers, because the maximum diameter of the stones should normally not exceed two-thirds of the layer thickness. Forssblad reported that medium to heavy vibratory compaction is effective in compacting gravel in thick layers. Inasmuch as gravels contain large particles, they are not affected to any appreciable extent by capillary forces. If the gravel contains on average more than 10% of fines, however, the soil will become elastic and springy when the water content is high, and in this state the material will be more difficult to compact. In the case of uniformlygraded gravel and sand, experiments show that the low shear strength of the graded mixture makes it very difficult t o achieve a high degree of compaction near the surface of the fill.
COMPACTION OF VERY COARSE-GRAINED SEDIMENTS
463
Laterite and silcrete gravels Laterite and silcrete gravels also merit a brief mention in the context of this chapter. A comprehensive review of the genetic and environmental relationships that appear to exist between laterite and silcrete has been presented by Stephens (1971). Laterite in Australia, for example, is thought to be a “product of relative accumulation formed in the oxidation-reduction zone of groundwater, seasonally rising and falling in a deeply weathered mantle associated with old landsurfaces of low relief, but with minor occurrences of absolute accumulation in areas of greater relief” (Stephens, 1971). According to Stephens, it is now concluded that “silcrete is predominantly an absolute accumulation by precipitation and crystallization induced by evapotranspiration following the lateral movement of discharge from groundwater carrying in solution the silica released by the weathering of silicate minerals during the formation of the laterite profiles and other dedicated soils in areas more or less distant”. Some insight into the compaction characteristics of lateritic gravels has been gained from laboratory tests of the physical properties and mechanical strength of aggregates used primarily for road and airfield construction in countries in Africa and Asia. Simply, lateritic profiles are developed in gravels by the capillary movement of iron-rich water from formations underlying such deposits. In general, such gravels are an admixture of quartz and lateritic gravel particles in a matrix of grains having finer sizes. Quartz gravels are derived from weathered quartzite, granite, and pegmatites, whereas lateritic gravels and similar gravelly materials are the weathered products of practically all rock types in Ghana, for example (Gidigasu, 1970). Besides varying in shape from angular to rounded, lateritic aggregates can be flaky or elongated depending upon the effect of hydrological factors. Opinions differ as to their resistance to shear and with regard to their possible use as highway base course materials.* Despite differing points of view, available evidence indicates that they would not withstand pressures emanating from overburden or thick overlying deposits. Irwin (1959) evaluated the strength of some lateritic gravels by applying the “ten per cent fines” aggregates crushing test, which determines the load in tons required to produce 10% fines, from a sample of a given single-sized aggregate. The main conclusions of Irwin’s study were as follows: (1)for the 12.70 mm-9.52 mm size fraction, the “ten per cent” values ranged from 1.6
* A stable base material contains suitable proportions of gravel, silt, and clay in the form of a soil or low-grade aggregate with or without an admixture of cement, bitumen, or other stabilizing agent. Besides the quality of the base material, the stability of the base is dependent on its thickness and degree of compaction.
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t o 17.2 tons for the tropical gravels and aggregates; (2)in general, the “ten per cent” value increased as the fractional size of the gravel within the range decreased from 12.70 mm to 4.76 mm; (3)immersion in water for 24 h prior t o testing, reduced the “ten per cent” values of all the gravels and aggregates. Gidigasu’s (1970)study on the effect of pretreatment on the compaction characteristics of lateritic gravels shows that: (1)it is necessary t o define the maturity of the gravels in terms of strength of the laterite particles, and (2)a relationship exists between specific gravity, water absorption, and the mechanical strength of lateritic gravels. There is a very significant correlation between the silica t o iron ratio and the mechanical strength of lateritic gravels, although no such correlation seems t o exist in the case of silica to silica sesquioxide ratio. Thus, the main conclusion is that iron plays a relatively more significant role in the development of mechanical strength in lateritic gravels than alumina. It follows from this that geotechnical characteristics of lateritic gravels are a function of weathering and degree of laterization. Doleritic or basaltic gravels are commonly avoided in road construction, because of their susceptibility to rapid chemical weathering and decomposition under certain climatic conditions. In addition, some types of basalt and dolerite are believed to be unstable because they contain relatively large proportions of clay and limonite produced hydrothermally in the interstices of the rock of the geological formation. De Graft-Johnson et al. (1972)have studied the physical, chemical, mechanical, and weathering characteristics of coarse lateritic gravels in Ghana, West Africa, and reported that such aggregates are suitable for road construction. Yet, laboratory studies by Bhatia and Hammond (1970)showed that soils having weak lateritic gravel fractions tend t o break down during compaction. De Graft-Johnson and his co-workers concluded from their study that the lateritic gravels, which have better mechanical strength and compaction properties, are derived from rocks having high concentrations of iron and aluminium; this indicates the dominant role of iron oxide in laterites. In western Nigeria, several distinct types of gravel useful in roadmaking have been recognized, including hardpan laterite (ground-water laterite), concretionary gravel composed of fenuginous material, and quartz and quartzite gravels. The latter two are derived from coarse-grained acidic igneous and metamorphic rocks as angular rock-scree or talus material (Ackroyd, 1971). Grading and plasticity tests indicate that many of these gravels are potentially superior to the conventional laterites as road base material, even if they are probably inadequate as combined base and surfacing material. Silcrete is widely distributed in Australia and southern and Saharan Africa, in river terraces, as depression deposits, in stream beds of low gradient and ephemeral flow, and beneath rocks such as basalts in these regions.
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Sands and gravels, usually waterworn, most commonly make up the granular portion of rock material cemented by the matrix, with the result that frequently the silcrete strongly resembles quartzite and conglomerate (Stephens, 1971). The conglomeratic nature of silcrete is quite common and often results from abraded fragments of silcrete itself being cemented together by a later-formed matrix. Once formed, silcrete is extremely resistant to weathering and has a hardness of about 7. CRUSHING STRENGTH OF COARSE-GRAINED SANDS AND GRAVELS
Owing to the limited geologic data available at present on the compaction and diagenesis of very coarse-grained sediments, the reader is inevitably left with important unresolved problems and, of course, many theoretical considerations. Whether or not some of the factors influencing and contributing to the compaction of sands also apply to the compaction of gravels, conglomerates, and breccias is still a subjective and somewhat speculative matter that awaits future research. Obviously, from what has been said in this chapter, further work is necessary on determining the relationship between the postdepositional compactional history of sediments and depth of burial. So far, little evidence has come to light of the magnitude of pressure which would exceed the crushing strength of the average gravel and conglomerate. Fraser’s (1935) work indicated that the initial porosity of coarse-grained sands is not considerably reduced by compaction until the pressure is sufficiently high to crush the grains and cause more or less complete collapse of the arching across the voids. I t is doubtful that similar circumstances would apply to gravel particles. Supporting evidence for Fraser’s conclusions has come from more recent research. One-dimensional, high-pressure compression tests on soils by Roberts and De Souza (1958), who obtained compression indices [defined as the gradient obtained from the curve of compression data plotted as void ratio (e) versus pressure (p) on a log scale, i.e., C, = -de/d(logp)] from 0.5 to 0.7 for quartz (Ottawa) sand in the pressure range from 1000 t o 10,000 psi. Their results indicate that at the high pressures encountered in deep sedimentary deposits, sand may be more compressible than clay owing to the shattering of individual grains. This implies that deeply-buried sands, with low void ratios, almost certainly d o not have the same grain-size characteristics that they possessed when first deposited. Thus the medium grain size will tend to be smaller, the angularity more pronounced, and the gradation wider. Roberts and de Souza state that these findings should be of considerable importance in the interpretation of the origin and depositional environment of sands. So far, it seems that this type of post-depositional change has been
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given little consideration by sedimentologists. Athy (1930) reported that porosity measurements on cores of poorlycemented sands from deep borings have shown that many sands at about 1219 m are more porous and less dense than sands which appear to be similar but which are taken from shallower depths. This disparity reflects the wide variation in the mineralogical and granular composition of sands and in the amount of cementing material which they contain. Yet, Levorsen (1967) cited an instance where clean sands drilled below 4572 m afford no evidence of crushing, whereas muddy or dirty sands would have been made impermeable by pressure at far shallower depths. (For further details see Chapter 3.) CONGLOMERATES AND BRECCIAS
Conglomerates are indurated or consolidated gravels. Common generic sources of conglomerates are accumulations of alluvial fans, braided streams, flood plains, and beach gravels. They vary appreciably with respect to composition, size, and shape of the constituent particles, although most of them are composed of rounded fragments or pebbles larger than 2 mm in diameter. The pebbles or larger rock fragments are normally in contact with one another and are cemented together by siliceous, ferruginous, or calcareous material. It is commonly believed that the sandy matrix of many ancient conglomerates was introduced after deposition, because the large interstices in the well-sorted “open-work” gravels would facilitate the infiltration of sand (Krumbein and Sloss, 1963). There are several types of breccia, although all of them are not necessarily important for compaction studies.* Several classes of conglomerates have been identified and classified by Pettijohn (1957) into two main groups, namely, orthoconglomerates and paraconglomerates. The former have an intact framework of pebbles and coarse sands and are characterized by a mineral cement. They are interpreted as having been deposited in highly turbulent fluviatile and marine (nearshore or surf zone) environments. The paraconglomerates are sensu stricto mudstones with occasional pebbles and/or cobbles. Examples include tillites,
* The reader’s attention is drawn here t o the unfortunate and somewhat misleading terms “intraformational” conglomerate and “intraformational” breccia, because sensu stricto “intra” means “within”. “Within a formation o r unit” would pertain to a solution breccia which has not undergone any transportation and sorting. On the other hand, the second variety o f intraformational conglomerate is formed by erosion of earlier-formed rock and redeposition within the same environment - hence “intra”. In summary, the former is a genuine “intraformational” sediment per se, whereas the other is an intraenvironmental sediment and can be composed of “sorted” material (Editors, personal communication, 1973).
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pseudotillites, pebbly mudstones, and some structureless clay or shale bodies in which pebbles or cobbles are randomly distributed (Krumbein and Sloss,
1963).
From the foregoing brief description of their properties, it is reasonable to invoke that, as in the case of clean gravels and those with a high matrix content, there will be some variation in the compaction properties of conglomerates. How these properties vary is not well known because of limited evidence from laboratory studies and the very sparse amount of field data. Theoretically, a well-sorted quartz-pebble conglomerate composed of 90% or more of quartz pebbles in a matrix of well-sorted, clean quartz sand, which was probably introduced after deposition, could compact differently from a poorly-sorted granite-pebble conglomerate with a high matrix content consisting of quartz, feldspar, and kaolinite. The degree of compaction will probably be greater in the latter conglomerates and, certainly, the effect of diagenesis is likely to be more pronounced in view of their varying chemical composition. Certain characteristics of consolidated and compacted gravels have been observed in the Budleigh Salterton Pebble Beds, South Devon, England, a Triassic formation composed of well-rounded metaquartzite cobbles and boulders (some of which are fossiliferous) up to 0.45 m in diameter, set in a coarse t o fine gravel and silty sand matrix (Henson, 1971). Orthoclase feldspars in the sand matrix and porphyries have been subject to postdepositional in situ kaolinization and the silty matrix has been recrystallized so that original grain sizes are indeterminate. It is, therefore, difficult to distinguish between primary cementing material and that introduced subsequent to consolidation, if, indeed, this distinction is really necessary. Nevertheless, the clays, formed from the weathered orthoclase, and the recrystallized silty matrix have influenced, no doubt, both the compaction and consolidation of these somewhat unique pebble beds. Many of the cobbles and pebbles are oblate spheroids, which display imbrication. They show evidence of cracking and fracturing in that entire clasts in cliff exposures disintegrate easily in the hand on being removed (Figs. 4 6 and 4-7).Henson concluded that the fracturing must be postdepositional and is related to stress release after loading by the overburden, because fractured cobbles that formerly underwent transport and abrasion after breakage are uncommon. His comment is of interest because cobble-to-cobble contacts are not particularly common and, consequently, there is an almost total absence of surface pitting resulting from pressure solution at the points of contact. From cursory examination it appears that there are two distinct populations of pebbles in these beds. One population is comprised of relatively unweathered material derived at the time of deposition from neighboring rock outcrops, whereas the other is composed of deeply-weathered pebbles
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Fig. 4-6. Budleigh Salterton Pebble Beds, South Devon, England. Thick cross-stratified sand unit, possibly of aeolian origin, overlain by cross-stratified gravelly sands with gravel pavements and cobble and pebble beds. (Photograph reproduced by kind permission of M.R. Henson.)
which, at the time, probably mantled a pediment surface. "he latter appear to have been more susceptible to subsequent fracturing under increasing overburden pressure referred to by Henson. Thus, the absence of less resistant rock types and the.*well-rounded form of abrasion-resistant quartzarenites indicate that the Budleigh Salterton Pebble Beds are a product of multi-cyclic sedimentation.
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Fig. 4-7. Budleigh Salterton Pebble Beds showing large-scale cross-stratification formed by the development of braid bars. Scale measures 50 mm along the lower side. (Photograph reproduced by kind permission of M.R. Henson.)
SUMMARY AND CONCLUSIONS
(1)A prominent recurring theme throughout the technical reports and scientific papers that have been reviewed by the writer is the important influence of texture (particle size, shape, and sorting) and chemical composition (proportions of various minerals) of gravels and matrix on compaction. There is a lack of quantitative information pertaining to field investigations on the effects of pressure exerted by overlying sediments. Although a great deal of knowledge has been gained from laboratory investigations, many conclusions obtained must be considered as tentative until they are confirmed by field studies. It is extremely difficult to simulate in the laboratory the conditions that exist in the field. Natural compaction rates are rarely rapid enough to be measured in the field. Rapid compaction in the laboratory represents an acceleration of the natural processes that normally occur with the further accumulation of overlying sediments. (2) If the contention advanced in this chapter is correct, i.e., that little compaction occurs in clean gravels after deposition, unless great pressure, the magnitude of which is not generally known, is applied, the packing density of gravels primarily reflects their environment of deposition. The need to
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examine more fully the effects of diagenesis on the compaction of very coarse-grained sediments is readily apparent. Those gravels with a high matrix content can be cemented, with apparent evidence of particle interlocking, through the solution of matrix material whether it is silica (quartz), calcite, dolomite, siderite, pyrite or anhydrite. Unless such gravel aggregates, however, are composed of components that are widely varied in chemical composition, and are highly unstable, it is reasonable to suggest that the mineralogical composition of individual particles will undergo little or no change during the diagenetic transformation of gravels into conglomerates. Virtually no data have been published on the relationship between compaction and the fabric of tills, and outwash and flood-plain gravels, despite numerous geologic and geomorphic studies of fabric patterns. (3) Brief theoretical consideration has been given to the possible differential compaction of conglomerates dependent upon variations in their chemical, mineralogical and granulometric composition. (4) Permeability is obviously reduced by compaction and cementation, and increased by solution. Void or pore-space reduction after deposition and subsequent burial is influenced by the composition of the sediment and interstitial fluids, conditions of ground-water circulation, and the magnitude of pressure exerted by superincumbent deposits. (5)The effect of particle shape and moisture content on the shear strength of very coarse-grained sediments has been reviewed in the light of the evidence obtained from studies of road construction aggregates. The packing density and the shear strength of graded aggregates are interdependent, with an increase in dry density usually producing an increase in shear strength for a given material. It is necessary to ensure during engineering compaction operations that the moisture content of a gravelly soil is sufficient to optimize cohesion. (6) The ratio of silica to iron has a very significant correlation with the mechanical strength of lateritic gravels, the geotechnical characteristics of which are a function of weathering, degree of laterization, and iron concentration. (7) Finally, the writer is aware that the engineering compaction of gravels and gravelly soils has been reviewed only briefly. It is considered that the fields of earth and rock fill dam construction and concrete technology are more relevant in a review specifically written for engineers rather than for geologists. ACKNOWLEDGEMENTS
The writer wishes to thank Dr. A.P. Carr and Dr. T. Leslie Youd for their comments and criticisms of an earlier draft of this chapter.
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REFERENCES Ackroyd, L.W., 1971. The engineering classification of some western Nigerian soils and their qualities in road building. Trans. Road Res. Lab. Rep. (Crowthorne, England), T R R L Overseas Bull., 10: 43 pp. Athy, L.F., 1930. Density, porosity and compaction of sedimentary rocks. Bull. A m . Assoc. Pet. Geologists, 14: 1-24. Berthois, L., 1951. Faconnement et granulometrie des galets au Delec pr6s Brest (Finist6re). Rev. Ge‘omorphol. Dyn., 2: 259-275. Bertram, G.E., 1963. Rockfill compaction of vibratory rollers. Proc. 2nd Panam. Conf. Soil Mech. Found. Eng., Brazil, 1: 441-455. Bhatia, H.S. and Hammond, A.A., 1970. Durability and strength properties of lateritic aggregates of Ghana. Build. Road Res. Inst. Kumasi, Ghana, Proj. Rep., 9: 1 5 pp. Bluck, B.J., 1967. Sedimentation of beach gravels: examples from South Wales. J. Sed. Petrol., 37: 128-156. Cailleux, A., 1945. Distinction des galets marins et fluviatiles. Bull. SOC. Gkol. Fr., 15: 37 5-404. Cailleux, A., 1947. L’indice d’emousse: definition et premiere application. C.R. Somm. Se‘ances SOC.Ge‘ol. Fr., pp. 250-252. Carr, A.P., 1971. Experiments on longshore transport and sorting of pebbles: Chesil Beach, England. J. Sed. Petrol.; 41 : 1084-1104. Chilingarian, G.V. and Wolf, K., 1975. Compaction of Coarse-Grained Sediments, I. Elsevier, Amsterdam, 552 pp. Clare, K.E., 1969. Road making gravels and soils in Central Africa. Road Res. Lab., Min. Dansp., Overseas Bull., 12: 61 pp. Culley, R. W., 1971. Effect of freeze-thaw cycling on stress-strain characteristics and volume change of a till subjected to repetitive loading. Can. Geotech. J., 8: 359-371. De Graft-Johnson, J.W.S., Bhatia, H.S. and Gidigasu, M.D., 1972. The engineering characteristics of a lateritic residual clay of Ghana for earth dam construction. Symp. Earth Rockfill Dams, New Delhi, India, 1: 94-107. Dean, R.C. and Southgate, H.F., 1972. Degradation of limestone aggregates during construction. K y . Dep. Highways, Res. R e p . , 324: 44 pp. Eerola, M. and Ylosjoki, M., 1970. The effect of particle shape on the friction angle of coarse-grained aggregates. In: First International Congress of the International Association o f Engineering Geology. Comite Franqais de Geologie de l’Ingenieur, Paris, pp. 4 4 5-4 56. Forssblad, L., 1971. Classifying soils according to compaction characteristics. World Constr., 24: 23-25. Fraser, H.J. 1935. Experimental study of the porosity and permeability of clastic sediments. J. Geol., 43: 910-1010. Gardner, M.,1966. Packing spheres. In: New Mathematical Diversions from Scientific American. Simon and Schuster, New York, N.Y., pp. 82-90. Gary, M., McAfee, R., Jr. and Wolf, C.L. (Editors), 1972. Glossary of Geology. American Geological Institute, Washington, D.C., 858 pp. Gidigasu, M.D., 1970. The influence of pretreatment on the compaction characteristics of lateritic soils. Build. Road Res. Inst. Kumasi, Ghana, Proj. Rep., SM.6: 1 2 pp. Graton, L.C. and Fraser, H.J., 1935. Systematic packing of spheres - with particular relation t o porosity and permeability. J. Geol., 43: 785-909. Gordon, B.B., Hammond, W.D. and Miller, R.K., 1964. Effect of rock content on compaction characteristics of clayey gravel. A m . SOC. Test. Muter., Spec. Tech. Publ., 377: 31-46.
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Handy, R.L., 1972. Compaction versus consolidation. Geotimes, 17 (8): 10-11. Henson, M.R., 1971. The Permo-Triassic Rocks of South Devon. Thesis, Exeter Univ., Eketer, Devon, 329 pp. Holtz, W.G. and Lowitz, C.A., 1959. Compaction characteristics of gravelly soils. U.S. Bur. Reclam., Earth Lab. Rep., E.M. 509: 26 pp. Huang, E.Y., Squier, L.R. and Triffo, R.P., 1963. Effect of geometric characteristics of coarse aggregates on compaction characteristics of soil-aggregate mixtures. High way Res. Rec., 22: 38-47. Irwin, M.J., 1959. A laboratory investigation of the resistance to crushing of some tropical gravels and aggregates. D.S.I.R. Road Res. Lab., RN/3489/MJI: 9 pp. (unpublished). Kahn, J.S., 1956. The analysis and distribution of the properties of packing in sand-size sediments, 1, 2. J. Geol., 64: 385-395; 578-606. Krebs, R.D. and Walker, R.D., 1971. Highway Materials. McGraw-Hill, New York, N.Y., 428 pp. Krumbein, W.C. and Griffith, J.S., 1938. Beach environment in Little Sister Bay, Wisconsin. Geol. Soc. Am. Bull., 49: 629-652. Krumbein, W.C. and Sloss, L.L., 1963. Strafigraphy and Sedimentation. Freeman, San Francisco, Calif., 2nd ed., 660 pp. Levorsen, A.I., 1967. Geology o f Petroleum. Freeman, San Francisco, Calif., 2nd ed., 724 PP. Lewis, W.A., 1954. Further studies in the compaction of soil and the performance of compaction plant. Road Res. Lab. Tech. Pap., 33: 46 pp. Maddison, L., 1944. Laboratory tests on the effect of stone content on the compaction of soil mortar. Roads Road Constr., 22: 37-40. Meade, R.H., 1968. Compaction of sediments underlying areas of land subsidence in Central California. US. Geol. Surv., Prof. Pap., 497-D: 38 pp. Merriman, J. and Houston, W.N., 1963. Research on compaction control testing for gravelly soils. U S . Bur. Reclam., Soils Eng. Rep., EM-662: 1 4 pp. Parker, C.A., 1972. Compaction consolidated. Geotimes, 17 (11): 7-8. Pettijohn, F.J., 1957. Sedimentary Rocks. Harper and Row, New York, N.Y., 718 pp. Pike, D.C., 1972. Compactability of graded aggregates 1. Standard laboratory tests. Transp. Road Res. Lab. Rep. (Crowthorne, England), LR 447: 34 pp. Road Research Laboratory, 1952. Soil Mechanics for Road Engineers. H.M.S.O., London, 541 pp. Roberts, J.E. and De Souza, J.M., 1958. The compressibility of sands. Proc. A m . SOC. Test. Mater., 58: 1269-1277. Shergold, F.A., 1953. The percentage voids in compacted gravel as a measure of its angularity. Concrete Res., 5: 3-10. Stephens, C.G., 1971. Laterite and silcrete in Australia: a study of the genetic relationships of laterite and silcrete and their companion materials, and their collective significance in the formation of the weathered mantle, soils, relief, and drainage of the Australian continent. Geoderma, 5: 5-52. Terzaghi, K. and Peck, R.B., 1967. Soil Mechanics in Engineering Practice. Wiley, New York, N.Y., 729 pp. Trurnit, P., 1967. Morphologie und Entstehung diagenetischer Druck-Losungserscheinungen. Geol. Mitt., 7: 173-204. Trurnit, P., 1968a. Pressure-solution phenomena in detrital rocks. Sed. Geol., 2: 89-114. Trurnit, P., 1968b. Analysis of pressure-solution contacts and classification of pressuresolution phenomena. In: G. Muller and G.M. Friedman (Editors), Recent Deuelopments in Carbonate Sedimentology in Central Europe. Springer, Berlin, pp. 7 5-84.
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Wadell, H., 1932. Volume, shape and roundness of rock particles. J. Geol., 40: 443-451. Wadell, H., 1933. Sphericity and roundness of rock particles. J. Geol., 41: 310-331. Wadell, H., 1934. Shape determination of large sedimental rock fragments. Pan-Am. Geol.,61: 187-221. Wentworth, C.K., 1922. The shapes of beach pebbles. U S . Geoi. Surv., Prof. Pap., 131-C: 75-83. Williams, F.H.P. and Maclean, D.J., 1950. The compaction of soil. Road Res. Lab. Tech. Pap., 17: 46 pp. Zingg, Th., 1935. Beitrag zur Schotteranalyse. Schweiz. Mineral. Petrogr. Mitt., 15: 3 9-1 41.
Chapter 5 ORE GENESIS INFLUENCED BY COMPACTION KARL H.WOLF
INTRODUCTION
A book on the compaction of sediments would not be complete without considering the types of ores in sedimentary and/or volcanic stratigraphic piles, chemical elements of which may have been supplied to the sites of precipitation by “water of compaction” (Noble, 1963). Although this concept has been with us for some time now, its theoretical and practical applications to specific cases has been mostly on a very general, superficial basis. Exceptions occur, as in the field of paleohydrology, possibly in conjunction with paleogeomorphology, which has been applied in some detailed quantitative and/or qualitative investigations more recently (see especially the 1971 Economic Geology Symposium (Vol. 66,5),and the papers by Jobin, 1962, Brown, 1971,and White, 1971). Inasmuch as compaction is only one small topic of paleohydrology, the discussions given in this Chapter are frequently not confined to compaction, but may branch out into related disciplines, such as those covering porosity and permeability. Some of the presented material may not be new, because, as pointed out by Elliston (1966):“. . . the idea that ore deposits originate from sediments is a simple and very old one. It was proposed by Georgius Agricola in his book De R e Metallica in 1556. The basic idea or certain aspects of it have been resurrected many times since and workers such as Gray, Knight, Garlick, and Noble have invoked aspects of it recently.” Although the general concept of compaction related to ore genesis is not new, this theory is undergoing constant refinement as a result of data supplied from geological and geochemical investigations. The concept of compaction is only one selectively chosen aspect of the evolution of thoughts on sedimentary ore genesis and, in a way, it may seem that this concept has been presented here somewhat out of context. As a result of the large amount of data on compaction and compaction fluids made available recently, a synthesis such as this one, is required as part of a contribution to the general evolution of ideas in sedimentary ore petrology, so well outlined by Amstutz (1964),for example. Those particularly interested in the historic developments of genetic hypotheses, may wish to read the paper by King (1968)on the Broken Hill mining area in Australia. (His appendix, a poem entitled “The Blind Men and the Elephant”, makes provocative reading.)
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Another publication with a historic perspective is that by Butler (1935 and 1967)on “Some facts about ore deposits” in which he discussed numerous “genetic myths”, some of which have long been revised, however. These publications on the historic development of genetic concepts in ore petrology demonstrate that at any particular time several theories may exist for a specific ore district. Some of the theories have been forcefully presented by their advocates and remained unchallenged for a comparatively considerable period before a “geological heretic” had the audacity to propose an alternative. Subsequent verbal and written debates have usually resulted in a final concensus of opinion on the various hypotheses with the consequence that either the older ideas were modified or rejected, or remained as a possible alternative to await future repeated evaluations based on fresh data. Every idea in geology will undergo the scrutiny prescribed by the Scientific Method - and so it will be with this chapter. The future will show to what extent the ideas brought together here, as well as the intellectual extrapolations made by the writer himself, are correlative with the geological realities. This chapter is based predominantly on a survey of papers that make references to compaction as an important process in the formation of ores in sediments, and possibly volcanics; however, it also includes some speculative ideas and suggestions in the hope that they will promote additional work. If the latter has been achieved, then one of the purposes of this summary and review has been fulfilled. Most of the information is qualitative and may even lack deeper analysis due to the generalized nature of the data; this is rather frustrating for the reader who wishes a more penetrating, quantitative approach to the problems under consideration. Despite this one limitation, constructive reviews, based on a conceptual rather than a quantitative basis, are necessary for students, teachers, researchers and explorationists to permit them to consult a reference in their daily work. It should also be made clear that only one view-point, i.e., of the possibility of compaction fluids giving rise to ore mineralization, is emphasized. This concept is not necessarily in agreement with other proposed geologic mechanisms and there is, therefore, room for disagreement. Inasmuch as this contribution is a review and was in preparation for some time, certain ideas may have been modified and supplemented by others in the meantime. Nevertheless, these ideas are stimulating and require additives to any serious thinking on the subject. We may do well to remind ourselves of the thoughts expressed by R.G.H. Siu in The Tao of Science: “Ingenious theoretical superstructures live in constant dread of factual termites that continually gnaw at their foundations. They topple at the first inconsistency with observation. Concepts glory only in a relatively short term of office . . . This does not necessarily mean that no statements can be made about reality. It just
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means that science herself deals in temporary hypotheses of perfection at any given moment” (From the preface of Helgeson, 1964;for other quotations, see Wolf, 1970.)As Arnold Toynbee put it: “Thus any presentation, whether particular or general, of scientific truth is always precarious and temporary.” Progress in the various scientific disciplines is increasing rapidly, so that it is becoming necessary to write comprehensive reviews and summaries, in particular of specifically selected, narrow topics, such as the one on the relationships between compaction and ore genesis. These reviews are especially useful if they are done constructively and lead to new conceptualized conclusions. As Carl L. Hubbs expressed it so succinctly; “Science’s only hope of escaping a Tower of Babel calamity is the preparation from time to time of works which summarize and which popularize the endless series of disconnected technical contributions” (in: Copeiru, 1935). Consequently, there need not be any controversy as to the “usefulness” of summaries and reviews in comparison to “original” and “fundamental” research - there is little doubt that both styles of contributions are required. Both approaches have their peculiar challenges, demands, hazards, and rewards. According to A.M. Weinberg: “AS I see it, at least part of the conflict amounts to a philosophic judgement whether science is the search for new knowledge or the organizer of existing knowledge” (in: Reflections o n Big Science, Pergamon Press, 1967). DELINEATING AND DEFINING SOME COMPACTION PROBLEMS AND CONCEPTS
General comments The meaning or definition of compaction has already been given in Chapter 1 “Introduction”; nevertheless, some of the general effects and causes of compaction are to be considered here. As Table 5-1 indicates, there are at least three mechanisms of compaction (A), that done or in combination can be active at different times during the geologic history of the sediments (B). The facies relationships between the different lithologic units are important in controlling the presence of fluid escape during compaction, the direction and rate of flow, the possibility of mixing of different types of subsurface fluids, etc. (C).The mode of compaction is given in a very simplified list (D), and the reader is referred to other chapters in this two-volume book and in its companion volume on the compaction of fine sediments (Elsevier Series Developments in Sedimentology, 16). As to the geologic stage when compaction occurs, one can see that the terminology used in the field of diagenesis can be applied here also, possibly
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TABLE 5-1 Outline of compaction in sediments: mechanisms and controlling factors ~
~~~
~
~~
A . Mechanisms 1. “Burial compaction” due to overburden: (a) intergranular; (b) unit-over-unit 2. “Vibration-to-shock compaction” 3. “Tectonic squeezing compaction” (e.g., Berry, 1973) B. Geologic time B1: 1. Early-stage compaction, e.g., during diagenesis when sediments are still in close proximity of original depositional environment 2. Late-stage, epigenetic (or catagenic) compaction, when sediments are removed from original depositional environment. (a) deeply buried in sedimentary basins (burial metamorphism); ( b ) after uplift and formation of a land-mass B2 : 1.Pre-ore compaction 2. Syn-ore compaction 3. Post-ore compaction C . Facies control 1. Compacting unit completely enclosed by sediments that formed more or less in the same sedimentary environment (may lead to abnormally high formation pressures) 2. Compacting unit is connected with the environment in which it was formed 3. Compacting unit is connected with the subaerial environment, e.g., seaward-dipping coastal plain sediments into which fresh water can penetrate. (Artesian pressure can move fresh water up to many miles offshore) D. Mode o f compaction 1. Pressing out of interstitial fluids, e.g., in clays and mudstones 2. Rearrangement of particles, e.g., in sandstones 3. Removal of material through solution: (a) carbonates (limestones and dolomites): (i) caves; (ii) breccias; (iii) stylolites; (iv) complete removal by solution; (b) evapcrites; (c) quartzose sands 4. Breakage of particles; brecciation E. Mobilization o.f .fluids 1. Compaction 2. Tectonic compression (e.g., Berry, 1973) 3. Change in hydrostatic head 4. Changes in thermal gradient 5. Diffusion 6. Capillary action 7. Metamorphic processes: (a) Recrystallization: decrease in pore space as a result of increase in mineral volume and a consequent increase in fluid pressure; (b) actual release of fluids from minerals 8. Diagenetic neomorphism (e.g., clay mineral transformation with concomitant release of water) (related to 7,b)
F. Source rocks for metals 1. Fine-grained, clay-rich ( T silt) sediments
479
ORE GENESIS INFLUENCED BY COMPACTION TABLE 5-1 (continued)
2. Coarse-grained, silt-, sand-, and pebble-sized sediments (e.g., see Helgeson, 1967) 3. Carbonate rocks 4. Evaporites 5. Volcanics 6. Plutonic rocks (or conglomerates containing clasts thereof) 7. Ores (reworked to form second-cycle ore concentrations) Note. All the above can be sources in the intra-basinal as well as in the extra-basinal environment. In the latter case, the metals were supplied to the area of precipitation by surface andfor subsurface waters.
G. Reservoir or host rocks for metal deposits in sedimentary volcanic piles Same as F, except evaporites, which do not act as host rocks for ore deposits
with some modification. Compaction and compaction fluids, therefore, can be of diagenetic, epigenetic or catagenic, and burial-metamorphic origin. When investigating parageneses of ores, one could classify the compaction features as of pre-, syn-, and/or post-ore origin (cf. B1 versus Bz in Table 5-1). In regard to the various mechanisms that cause mobilization and/or remobilization of fluids, one has to remember, of course, that in addition to compaction there are many other concepts that have to be applied if one is TABLE 5-11 Flow diagram depicting the general factors involved in ore genesis by basinal fluids (EXpanded after Beales and Jackson, 1968) <
Source -+
Transportation + Concentration +
Terrigenous rocks undergoing chemical weathering
Adsorption of ions on clay minerals and organic matter
COMPACTION Release -+ Transportation + Release of ions into compaction fluids
t
Precipitant + Precipitation
Fluid movePrecipitation is ment: rate, direc- controlled by tion and distan- different geoce of migration chemical and is controlled physical condiby compaction tions, e.g., by history in com- mixing of fluids, bination with change in temperbasin configura- ature, and prestion and other sure factors
K.H. WOLF
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considering the broad spectrum of origins that apply to ore minerals in sediments and volcanic rocks (a partial list is included in section E in Table 5-1). The various lithologies that are known to be sources of metallic metals and can act as reservoir rocks for ores, are shown in sections F and G in Table 5-1. One type of investigation that remains to be done on numerous ore deposits, on the interesting topic of the relationship between ore genesis, compacTABLE 5-111 Variations of some factors that influence compaction Conditions conducive for compaction t o occur
Transitional
Factors less conducive for compaction t o occur
(1)Presence and increasing proportion of clays and mica in coarsegrained sediments
(1)Absence and decreasing proportion of clays and mica in coarsegrained sediments
(2) Mineralogy: e.g., certain clays release water more easily
(2) Mineralogy: e.g., certain clays release water less easily
(3) Thick overburden; rate and magnitude of geologic subsidence distinct
(3) Thin overburden; geologic subsidence small
(4) Long geologic time
(4) Short geologic time
(5) Presence of numerous dia-
stems and unconformities
(5) Diastems and unconformities less abundant or absent due to continued deposition
( 6 ) Presence of many coarse and fine interbeds (better escape routes for fluids)
(6)Presence of few coarse and fine interbeds, which makes fluid escape more difficult
(7) Presence of conducive structure: e.g., depositional dip up into reservoir rock
(7) Lateral migration routes impeded, not yet developed, or sealed, e.g., by a fault
(8) Corrosive type of intraformat i n a l fluids’ composition; these fluids can increase porosity and permeability by removing cement
(8) Corrosive fluids absent, or the fluids’ composition is neutral or otherwise noncorrosive in composition
(9) Hydrostatic formation pressure only
(9) Presence of “abnormally” high formation pressures
(10) Late diagenetic cementation permits compaction over long geologic period
(10) Early diagenetic cementation reduces degree of compaction
(11)Rate of sediment accurnulation slow
(11)Fast rate of sediment accumulation
ORE GENESIS INFLUENCED BY COMPACTION
481
tion and fluids, would be to fully integrate paragenetic textural relations (e.g., Amstutz and Park, 1971).This should be carried out on all scales, from regional studies to micro-analyses. It would be comparable, say, to structural petrofabric investigations from the micro- through the meso- to the megascopic scales. This type of approach, as a supplement to others, will lead to information that can establish: (a) whether the ore is syngenetic-diagenetic or of late-stage epigenetic origin, and (b) to what degree compaction fluids may have been responsible for the precipitation of the ore minerals. Several examples of such studies are already available from the literature and are referred to below. Certain generalities on the relationships between compaction and ore deposition have been known for some time and are exemplified by the flow diagram of Beales and Jackson (1968),reproduced here with some changes and with supplemental data (Table 5-11). As to the details, much on compaction is available from several scientific disciplines, but the results are only now being increasingly applied to the origin of ores and to exploration problems (e.g., hydrologic principles to the White Pine copper deposits in Michigan; see White, 1971,and Brown, 1971,for example). Adsorption and release mechanisms of chemical elements from clays and organic matter have been investigated to a certain degree (e.g., Kashirtseva, 1960;Weiss and Amstutz, 1966; Jackson and Beales, 1967;Saxby, 1969). Eventually, research and exploration geologists should have more information at their disposal on the causes and effects of compaction in general, so that they can make an attempt to apply the data in mineral-genesis studies. Thanks to numerous investigations by the oil companies it is possible to compile many of the variables that will vary regionally as shown in Table 5-111. More quantitative data, however, is required. It is one problem to prove that the ore in the sediments is of syngenetic, diagenetic or epigenetic origin and formed by processes related to paleogeography, stratigraphy, sedimentation, lithology, and sedimentary properties such as porosity and permeability, which are all, relatively speaking, easy to determine. It is another matter, however, to show: (a) that compaction has preconditioned both the aquifers and host rock (i.e., that compaction was responsible in controlling the petrographic characteristics); and (b) that compaction waters have actually supplied the metals precipitated in the host rock. The latter is by far the most difficult to demonstrate and much is based on speculation. Many ore deposits, e.g., the Colorado Plateau-Wyoming type uranium deposits in sandstones may never have been influenced by compaction. Only the primary porosity and permeability may have been the direct controlling parameters, which were originally influenced by compaction only very indirectly. As soon as the sediment was indurated and had a more or less definite porosity and permeability, compaction had little or no
482
K.H.WOLF
influence on these parameters. Yet, the fluids within the fine and coarse sediments must have been pressed out, resulting possibly in some minor compaction folding, supplying chemicals for the precipitation of cement within the coarser sediments, delivering fine organic debris, and so on, so that compaction and its fluids had at least an indirect influence in conditioning the host rock for subsequent ore mineralization. On the other hand, this preconditioning may have been of a nature that was less conducive to ore mineral precipitation. Although studies, like that by Jobin (1962),have shown that porosity, permeability and related parameters are of the firstorder importance, compaction as a process in ore genesis may have been of the fifth- or tenth-order significance (if one lists all the factors and processes which are believed to have been of importance in their order of decreasing influence). One may look at the problem from another point of view and divide the source and host rocks into: (a) those that do not have preserved any evidence of compaction; and (b) those that do contain such evidence. The dilemma in petrology is that the rocks that undergo the most intense compaction, i.e., the fine-grained ones and those that may have acted as source rocks, usually belong to group a, whereas most of the direct visual petrographic evidence for compaction is confined to the coarser deposits that often are the host rocks. The evidence in fine sediments that is available is so small in size and not readily obtainable to be used in either qualitative or quantitative studies, that it cannot be employed at the present time for determining the style and degree of compaction. As briefly mentioned in the chapter on sandstone diagenesis in this book, however, considerable quantitative data is available from investigations by the petroleum industry that permit inferences and extrapolations to be made on the compaction of various types of sediments. Certain techniques in determining the compaction of carbonate rocks (see the section on the Mississippi-Valley type ores) and of sandstones (see Chapter 3 on diagenesis) have been developed and will have to be employed in ore petrology in the future, before a full understanding of the significance of compaction can be established. It seems obvious, however, that once the host rock and the ore have undergone metamorphism, it may become impossible to decipher the early compactional history. Although the writer speaks collectively in this chapter about “compaction fluids”, one must consider at all times the following. Of the four types of ore solutions mentioned by Stanton (1972),namely, sea water, ground water, deep geothermal water, and igneous solutions, the compaction fluids per se are more closely related to the sea-water and ground-water groups. It should be clear, however, that because sea water can become interstitial subsurface fluid, usually with extensive diagenetic chemical changes, and because both geothermal and igneous solutions can become mixed with more “normal” ground water, the subsurface waters in general, including compaction fluids,
ORE GENESIS INFLUENCED BY COMPACTION
483
can be of complex mixed or transitional origin as well as composition. Details are beyond the purpose of this chapter. Whatever the precise origin of the compaction fluids may be, they may be important in the diagenetic development of some base metal sulfide ores of the manto, stratabound and stratiform types (Fig. 5-1). Although in this chapter reference is made mainly to “sediments”, it should be understood that volcanics are to be included. Stanton (1972) mentioned, in support of Fig. 5-1, that “. . . a unit of volcanic material, largely pyroclastic, underlies a unit containing lenses of sedimentary iron sulfide. The pyroclastic unit is presumed, on the basis of observations about modern volcanoes, to contain substantial traces of heavymetal halides (chiefly chlorides) sublimed in vesicles and fractures in the tuffs, breccias, and lavas concerned. Most of the halide material is dissolved into the pore waters and, with increasing burial and compaction, those solutions are expressed and moved upward. As they rise, most of the metallic halide is dispersed and lost. Some however, inevitably encounters overlying iron sulfide lenses, and metals such as copper, lead, and zinc displace some of the iron from the sulfide, thus precipitating as sulfides themselves. The result is a pyritic base-metal deposit, the nonferrous metal fraction of which may be regarded as being the result of diagenetic, low-grade metamorphic processes.” Where does the process of compaction fit into the overall genetic classification scheme and in the investigation of ores? Approaches to the understanding of mineralization must vary in scope, if investigations are to be meaningful and successful. Classification schemes, conceptual models (Wolf, 1973a,b) and multiple genetic theories may have to be adopted to the particular situations to prevent the investigator from being trapped in an “intellectual dead end”. For example, the understanding of the origin of ores in
v
V
v
1 ”
V
I
V
Tuff unit containing traces of base metal halides V
V
V
v
‘ V
1
V
V
-
t
-
Fig. 5-1. Possible mechanism of formation of some pyritic base-metal orebodies during diagenesis and very early regional (burial) metamorphism; heavy-metal halides in the pore waters of a pyroclastic unit may encounter pyritic lenses when forced upward by compaction. (After Stanton, 1972, fig. 18-2, courtesy of McGraw-Hill Book Co, New York, N.Y.)
K.H. WOLF
484
sediments may be delayed or hindered in particular cases if one approaches the problem by comparing the mineralization with those in a classification scheme based on host rock, i.e., an approach employing the “ore-type by ore-type” comparison until one finds the “pigeonhole” in the scheme that best fits the natural occurrence. In other words, if one assumes that ores in black shales and argillites, siltstones and sandstones, and in carbonates are very distinct varieties (i.e., as in the case of lead--zinc ores), one may miss some underlying fundamental genetic similarity. Ores that show certain petrographic similarities usually differ in other aspects, whereas, on the other hand, ores that at first sight appear to be different may have some common genetic relationships. It should not be forgotten that different processes can result in similar end-product, whereas one and the same mechanism under different environmental conditions may give rise to dissimilar deposits. Whatever the difficulties, the process of compaction should be granted full status in any investigation of ores in sedimentary host rocks, and there is no need to establish a special “type” of ore mineralization as a result of compaction. The flow chart in Table 5-IV lists some of the more obvious factors to be considered in the investigation of ores in sediments and volcanics - compaction has to find its place in all instances on all scales of studies. Table 5-V merely presents some of the more “philosophical” aspects of such investigaTABLE 5-IV Summary list of some major parameters that control compaction and factors that, in turn, are controlled by compaction (i.e., there is a reciprocal relationship) COMPACTION
<
Regional factors
Local geologic factors
Orogenic vs. cra-Geometry of rock tonic bodies Geosynclinal vs. Source rocks geanticlinal Reservoir rocks Eu- vs. miogeoRatios of the +- above two synclinal Facies distribution Fractures, etc. for Lithologies passage of fluids Metamorphism Hydrodynamic sysGeothermal pattem terns Fluid chemistry .
t
.4
Ph ysicochemical and biochemical factors
Other factors
Geologic time Evolution of the earth’s physical and chemical character istics
ORE GENESIS INFLUENCED BY COMPACTION TABLE 5-V
-
485
Varieties of investigations involved in the study of compaction Data gathering Large-scale General “Classical” geology 4 Field approaches 4
conceptual model (s)
conclusions
’small-scale ’particular ’“modern” geology ’experimental, laboratory, and theoretical approaches
tions, whereas the reader is referred to the publication by Wolf (1973a) for discussions on the scientific method. Formation of ore deposits by water of compaction The paper by Noble (1963) is one of the few publications wholly devoted to examining the possibility that lateral secretion* by water of compaction can form mineralization of ore-grade in sedimentary and associated volcanic rocks (see also Knight, 1957, and King, 1965). Noble listed numerous examples of ore districts which are mentioned in subsequent sections of this chapter, e.g., sandstone-type vanadium- uranium ores; red-bed copper deposits; Mississippi Valley-type lead-zinc ores; and uranium in ancient conglomerates. Other genetic interpretations have been offered for all these, however, and should the concept of compaction eventually be shown to be less likely applicable to the specific ore districts cited by Noble, it does not nullify the validity of the concept in general, so that the theory of ore genesis by compaction fluids may still be a plausible process for many other ore deposits. As pointed out by Noble, chemical elements are disseminated in common rocks, which lends support to the hypothesis of lateral secretion*. Although these metals are widely dispersed and occur in parts per million or even parts per billion (Krauskopf, 1967, 1971; Boyle, 1969; Hirst, 1971), calculations have demonstrated that the total tonnage may be enormous (Pustowalow, 1959). Calculations have also shown that in most basins the amount of fluid is sufficient to have been able to transport the metals to a favorable site of precipitation, assuming that the chemical composition was conducive. Pb, Zn, and Cu can be transported in solution as components of a carbonate or
* Lateral secretion: a theory of ore genesis formulated in the 18th century that postulated the formation of ore by leaching of the adjacent wall rocks. A contemporary term for such a mineral deposit is the adjective “lithogene”. Certain types of vein deposits are formed by lateral secretion.
486
K.H. WOLF
bicarbonate complex, and U and V as carbonate complexes; other metallic ions are transported by chloride-rich solutions. Saxby (1969) presented a useful summary on the importance of metal-organic components. During early diagenesis water may acquire soluble constituents from the host sediments, especially because a vast aggregate surface area is available in the fine-grained, clayey material. The physico- and biochemical milieu must be favorable because sorption, desorption and ion exchange may either hinder or enhance the processes of release and removal of the chemical elements (see Kashirtseva, 1960; Rashid, 1971, 1972a,b; Rashid and Leonard, 1973). Organic matter, for example, may not allow desorption and release of metals because of a reducing environment. The expulsion of formation fluids from fine and coarse sediments has been discussed by numerous investigators (e.g., Meade, 1964; Ostroff, 1967; Powers, 1967; Magara, 1968; Mangelsdorf et al., 1970), whereas the investigation of metamorphic waters is more or less only in its initial stage (e.g., Barnes, 1970, and Huang and Wyllie, 1973). No details are presented here on the derivation of saline compaction fluids from evaporites or as a result of maturation of subsurface fluids. The reader is referred to studies of Davidson (1965), Dunham (1966, 1970), Jodry (1969), Tooms (1970) and Bush (1970) on the possible role of brines in the formation of ores. Noble stated that the flow of mineralization fluids can be inferred from the distribution of alteration zones and from the relation of the ore deposits to zones of alteration as well as to basin configuration and facies. Lead and zinc ores are associated with dolomitized shelf carbonates, whereas vanadium and uranium deposits, with zones of maximum heavy-mineral alterations. The most perplexing problem in the application of the lateral secretion theory is related to the leaching process of the elements, their transportation and, finally, the precipitation of the elements. When all possibilities are considered, water of compaction seems to offer the most plausible explanation because it can leach individual particles during diagenesis when compaction and consolidation are still incomplete. Leaching of chemical constituents from lithified rocks is more difficult to envisage. The importance of brines derived from evaporites, for example, has been discussed by several researchers (see below). As to fluid movements, it has been suggested that hydrostatic pressure is one of the major sources of pressure to motivate ground-water circulation, which places severe limitations on the effectiveness of the lateral secretion concept, as pointed out by Noble (1963). On the other hand, lateral secretion by compaction fluids may give rise to pressures greater than hydrostatic pressure. Movement of fluids can be in many directions, especially initially, but will decrease with geologic age until more definite “aquifers” have been established. Compaction fluid movements can occur to a depth not reached by meteoric fluids. In addition, large volumes of the water are
ORE GENESIS INFLUENCED BY COMPACTION
487
moved within a very short geologic time, which under favorable conditions can result in large quantities of metal being deposited. As to the causes of precipitation, Noble mentioned that many have been proposed and that all are conjectural. No unique process is envisaged, because several or all may have been operative. Possible causes mentioned by Hey1 et al. (1959), and others, for the precipitation of lead and zinc minerals are : (1) Decrease in temperature. (2) Decrease in pressure in compaction fluids could result in disruption of metal-bearing complexes; in addition there is an indirect effect of accumulating a gaseous precipitant, such as H2S, as a result of evolution from the fluid because of pressure reduction. (3) Neutralization of acidic magmatic solutions, petroleum, and H2S. (4) Interaction of thermal ore solutions with connate or meteoric waters, or mixing of different types of compaction fluids (e.g., Runnels, 1969). (5) Changes in chloride concentration (e.g., Helgeson, 1964; Tooms, 1970; Dunham, 1970). (6) Presence of different types of precipitants (e.g., in the case of uranium it would be organic matter), organisms, clay, H2S, natural gas, sulfides, CaC03, and phosphate (see for example Beales and Jackson, 1966, and Dunham, 1970). Much remains to be done on establishing the characteristics of ores formed by compaction fluids. As Noble pointed out, there should be an absence of features accepted as diagnostic of ores of a magmatic hydrothermal origin (unless subsequent remobilization occurred). One should be cautious, however, of a fortuitous or coincidental association of igneous rocks with the ore deposits, i.e., the mere presence of dikes, for example, may not rule out a diagenetic origin of the ores. The following characteristics of ores formed by compaction fluids were listed by Noble: (1)A likely source was available. (2) Location is in or adjacent to zones that have transmitted large quantities of fluids. (3) Fine-grained sediments that have been possible sites of build-up of anomalous fluid pressure may be nearby. (4)Evidence of passage of large quantities of fluid should be present: (a) extensive alteration; (b) addition and/or removal of material in permeable zones. One should remember here that compaction fluids may have been at different localities, and at different times, an oil- and/or ore-forming fluid, a solvent, a carrier of cementing material, or an agent of various alterations. (5) Primary structural, mineralogical and age (= paragenetic) relationships of the ore (ignoring subsequent secondary geologic modifications) are consistent with deposition accompanying expulsion of the compaction waters.
488
K.H. WOLF
(6)The minerals are of low-temperature and low-pressure types, and commonly constitute a simple assemblage. (7) Fluid inclusions also indicate the above (e.g., Roedder, 1968). (8) The isotope data e.g., 3zS/34Sratio of sulfide minerals from the Pine Point lead-zinc deposits in Canada, are characteristic of non-magmatic sources. Deviations may be present as a result of mixing of components from more than one source. Considering all the above criteria, one must point out that many of the features listed above could be characteristic of deposits formed not necessarily by fluids of compaction, but also by fluids mobilized by other mechanisms, e.g., mere hydrostatic pressure, geothermal gradient, and burial or low-grade metamorphism, so that the problem of recognizing specific ores precipitated hom compaction waters is still complex. Beales and Jackson (1966)utilized comparable ideas and supplemented them with additional concepts, involved in the normal evolution of a sedimentary basin to produce a specific model to explain the formation of the big group of Mississippi Valley-type, lead-zinc ore bodies. In recent years, ideas of basin evolution have become standard in exploration thinking and many new ore bodies have been found partly as a result of having a logical basin-oriented philosophy. Further comments on the specific place of compaction studies in Beales and Jackson’s (1966,1968)work are found later in the text. Relationship between rheological properties of sediments and ore genesis The rheological properties of sediments have been considered only occasionally in ore genesis, although physicists and physical chemists have developed the fundamentals (e.g., Blair, 1969). The theoretical principles have found application in soil mechanics and civil engineering, and sedimentologists have begun to explain some structures as a result of the thixotropic properties of water-containing sands, silts and clays (Weyl and Ormsby, 1960;Boswell, 1963;Elliott, 1965).In ore petrology, Elliston (1963),Hunter (1963)and Wright (1969)have used the concepts of thixotropy of unlithified sediments to explain the origin of the ore mineralization of the Peko mine of Australia. The writer (Wolf, 1973c)has proposed that the rheologic phenomena in igneous petrology should be given more attention in the future althpugh the phenomenon has not been completely ignored in the past. As to its application in sedimentology, more data has to be obtained on the relationship between shock (induced, for example, by earthquakes) and compaction, on one hand, and the various possible responses by the sedimentary deposits, on the other. As Elliston (1963),for example, has pointed out, composition is one of the major factors that control the rheologic behavior of sediments, in particular the grain-size variation together with the inter-
ORE GENESIS INFLUENCED BY COMPACTION
489
granular fluids present. The ratios of sand and silt, as well as the amount of clay and colloidal material, also determine the thixotropy of a deposit. All newly-formed sediments are thixotropic to some degree, except for clean sands. Textural variations, especially dependent on compaction, will be influential also. The characteristics of the colloidal material, such as its state of flocculation, is particularly important, because the colloidal “gelling” components, when present in the intergranular spaces in the sediments, may give the whole deposit “gelatinous” properties. The material behaves like an elastic solid up to a certain yield point at which it will flow. In particular, when this sedimentary mass is exposed to shock waves, a thixotropic transformation (gel to sol or to more-fluid gel) may occur and the sediment may become a fluid and flow as a slurry. Little information is available to estimate the depth to which sediment may remain thixotropic, but some relationship with depth of burial is to be expected because of increasing compaction of the clay matrix, increasing number and areas of contacts between particles accompanied by expulsion of water, and increase in degree of lithification, for example. As a result of the thixotropic properties of sediments, they can become mobile and, therefore, can flow, intrude and form dikes and sills. This flowage can lead to accretion of the colloidal material into globular aggregates where the colloidal particles adhere to each other and “snowball”, as has been described by Elliston (1963) and Wright (1969), to explain such features in the rocks of the Peko mine district. Many similar textures, fabrics and structures have been interpreted as of igneous and/or metamorphic origin, but Elliston believed that the origin of the various features in the host rocks of the Peko ores as well as the origin of the ore mineralization itself is related to the thixotropy of originally unlit,hified sediments. The reader must be referred to the original publications for details, except for the summary of the relationship between diagenesis and ore formation given below (Elliston, 1963, pp. QS-Q9): (1) Unconsolidated to weakly indurated sediments can flow by the sudden yielding of the interstitial colloidal gels if subjected to shock produced by earthquakes, for instance, which would change the sediment’s rheologic or thixotropic properties. The re-slurried material can slump and flow along the surface or intrude the overlying deposits. (2) Under conditions when the colloidal fraction in the sediment is stable in the gel state, it segregates into a fine granular matrix as larger aggregates or clots. The adsorbed metallic ions are segregated with the colloidal material. In time, with increasing compaction, dehydration and temperature, the colloidal clots will crystallize. (3) Ions will be desorbed and removed from the gel of the crystallizing colloid as insoluble metal hydrosols peptized by dissolved silica in the water released during crystallization.
490
K.H. WOLF
(4) Following seepage zones or extrusion channels, the water with the ions are moved by compaction fluids. Under certain circumstances the fluids are enriched in the desorbed metal hydrosols. (5) Concretionary precipitation of colloidal minerals in the extrusion channels may occur as overgrowths, clots, intergrowths and replacements, for example, similar to those described by Elliston. Remobilization to form higher metal concentrations is possible. A particular precipitation paragenesis can result from successive desorption of metals in the source sediments. Elliston concluded that “an orebody could be formed simply by shaking or disturbing natural sediment accumulations”, and the physical-chemical appraisal of colloidal matter by Hunter (1963) indicates that Elliston’s theory of ore genesis is feasible. Thus, although it has been proposed that compaction alone could be responsible for the origin of ores in sediments, in some instances a mere change in the thixotropy of a sedimentary pile, together with compaction possibly, could be responsible for mobilizing and moving ore-forming fluids.
Relationships between the origin of petroleum and metalliferous concentrations Once the hypothesis, that the components of ore deposits within sediments were brought to the site of deposition by subsurface basinal fluids, is accepted, it is only logical to continue to search for geologic and geochemical similarities and differences between the genesis of petroleum and metalliferous concentrations. Pospelov (1969) discussed certain analogies between ore and petroleum occurrences as to their structure, geologic setting or position, and origin. The oil-ore fluid link is a most useful one as Jackson and Beales (1967) have shown in their studies of the Pine Point lead-zinc district. Concepts and techniques developed in the oil exploration industry, and heretofore not applied to the search and development of metalliferous deposits, will then be considered as useful exploration guides. Pospelov defined “fluids” as “carriers endowed with fluidity but differing in their material and phase compositions”. Any fluid, whether: (1)cold, warm or hot; (2) containing hydrocarbons or metallic chemical elements; (3) fresh or saline; (4) connate, igneous hydrothermal, or heated and recirculated surface water will be controlled by similar geologic factors. Inasmuch as the literature from the petroleum industry has many references to the importance of compaction that is part of “basin evolution” and to its contribution to the origin and migration of hydrocarbons, and because it has also been suggested that the sediments in the same or similar basins may have supplied the chemical elements for ore deposits, the oil-ore analogy is not only in order, but most important.
ORE GENESIS INFLUENCED BY COMPACTION
491
Ore deposits
Oil deposits a
h
d
C
J .. .:
.,:: . ,.,.;)' .
..-.
Fig. 5-2. Comparison of oil and ore deposits: types of deposits. (After Pospelov, 1969, fig. 1.)
Bedded: a . Tuymazy, U.S.S.R. ; b. metasomatic lead (phenocryst), Mississippi, U.S.A. (after Buckley e t al.). Stratigraphic: c . Suez Bay; d . lead-zinc, Monek, Canada. Veins: e. ozocerite, Borizova, U.S.S.R., 260 m depth; f. gold (after Azhgirey). Stocks (intricate configuration; scattered phenocryst type) : g . asphaltites in andesite intrusions, Mendoza, Argentina; h . uranium-molyhdenum, associated with subvolcanic intrusions. Stringers (pillar-shaped): i. oil and gas, Kansas, U.S.A. ;j . fluorite, localized in underlying limestone by subsidence, East Green, Ill., U.S.A.
K.H.WOLF
492 Oil deposits
Ore deposits
+
+
C
f
i
Fig. 5-3. Comparison of oil and ore deposits: lithological screens and traps in arches. (After Pospelov, 1969,fig. 2.)
Contiguous structures: a. oil on periphery of ancient volcanic massif, Mexico; b. metasomatic lead ores in Missouri Cobalt Mine, screened by buried body of porphyries. Lithological damming: c. Dan River, Michigan, U.S.A. ; oil accumulation in dolomitized vertical zone of limestone near fissure, capped by clayey sediments; d . metasomatic polymetallic deposit, Sierra, U.S.A. (Butler's projection), in dolomites overlain by shales, with vertical leader veinlets.
ORE GENESIS INFLUENCED BY COMPACTION Oil deposits
a
493 Ore deposits
b
I-
d
Fig. 5-4. Comparison of oil and ore deposits: screening fissures and traps. (After Pospelov, 1969, fig. 3.) u. Luling, U.S.A. (after Brucks); b. Artemovskoye, Azerbaijan; c. metasomatic hematite,
bedded, Cumberland, U.K. (after B. Smith); d. sulfide body, Konsay, Uzbekistan (after F.I. Vol’fson et al.); e. Surakhany, Azerbaijan; f. uranium stock associated with subvolcanic intrusive; g,h. Pirsagat and Kala, Azerbaijan; screened by fissures (projection); i. metasomatic body of magnetite, Krivoy Rog, Ukraine; screened by fissures; j . tubular bodies, Ekaterino-Blagodat’, Zabayakal’ye (projection).
Bends of folds: e. Serafimovka, Bashkiria; f. mercury in sandstones overlain by shales, Sofiyevskiy ‘Dome, Nikitovka; g. stratoisohyetal lines of the roof, Subkhankul, Bashkiria; h. stratoisohyetal lines of the roof, polymetallic metasomatic deposit, Sokol’noye, Rudnyy Altay (after K.F. Yermolayev); i. Santa Fe Springs, U.S.A.; j . siderite, Baykal, Ural (after Z.M.Starostina); k. skarnsulfide, Gaspe Copper, Canada.
K.H. WOLF
494 Oil d e p o s i t s
0
C
Ore d e p o s i t s
b
d
Fig. 5-5. Comparison of oil and ore deposits: fissures as channels and screen for fluids. (After Pospelov, 1969, fig. 4.) a. Neftechala, Azerbaijan; b. Leadville, U.S.A. (after G. Laughlin); c. Kum Dag, Turkmenia, U.S.S.R.; d. secondary lead (black) and zinc (striped) ores developed from a polymetallic ore body associated with fissures (after I. Knyazev).
ORE GENESIS INFLUENCED BY COMPACTION Oil deposlts
a
C
495
O r e deposits d
Fig. 5-6. Comparison of oil and ore deposits: multi-level and pillar-shaped deposits produced during ascending migration of fluids. (After Pospelov, 1969, fig. 5.) a. Lokbatan, Azerbaijan; b. Ventura Avenue, California; c. gold ore veins, saddle-shaped, Bendigo, Austria; d. copper skarn, Nickel Plate, Canada; e. auriferous quartz vein, saddle-
shaped, complicated by metasomatism (after N. Borodayevskiy); f. Nebit Dag, Azerbaijan; g. massive sulfides, Lena Lake, Canada; h. Surakhany, Azerbaijan; i. magnetite, Kunzhunkul, Kazakhstan.
496
K.H.WOLF
Pospelov (1969)has shown that there are many resemblances in structural-lithological-stratigraphic controls of the occurrences of ore and oil, i.e., many similarities occur in the types of deposits (Fig. 5-2),in the lithological screens and traps in “arches” (Fig. 5-3),in screening fissures and traps (Fig. 5-4),in fissures which acted as channels and screens for fluids (Fig. 5-5),in regard to multi-level and pillar-shaped deposits produced during ascending fluid movements (Fig. 5-6), and in migration paths of fluids during formation of the two types of deposits (Fig. 5-7).A cursory examination of these diagrams will show that many of the structural-stratigraphic relationships presented in Figs. 5-2to 5-7 are very similar t o many of the ore deposits for which a sedimentary origin has been advocated. Some have syngeneticdiagenetic and others epigenetic characteristics. Rittenhouse (1972)has offered a rigorously developed stratigraphic-trap classification from a petroleum geologist’s point of view, which should be given consideration by those interested in sedimentary ore genesis. All of Rittenhouse’s traps have counterparts in ore deposits within sediments and volcanics. Silver (1973)discussed a special type of stratigraphic trap which he called “isolani”. These traps “formed in an isolated mass of porous rock surrounded by a less permeable matrix, preferably, but not necessarily, with unoxidized organic-rich shale as a source bed for hydrocarbons”. The isolani may be found throughout a sedimentary basin. The thoughts expressed by Silver on the accumulation of oil and gas as isolated bodies may find future application in explaining the origin of certain ores within sedimentary and pyroclastic rocks because they too often occur as “unexplainable” isolated masses. If one accepts the possibility that there is a genetic relationship between hydrocarbons and ores, or at least that fluid mechanisms in sedimentary basins are similar in both instances, the ideas of Silver should be considered by ore petrologists. Isolani play a special role in the accumulation and preservation of hydrocarbons. Following Hubbert (1953)that the slide/ sandstone interface is a surface of unidirectional flow when the two lithologies are water saturated, due t o the “screening effect” discussed by Hubbert, petroleum will enter an isolani as a result of compaction or any other means. Figure 5-8 presents the similarities and differences between conditions of hydrodynamic versus compaction filling of isolani. The diagram depicts that there are two types of isolani. The basic isolani, on the right, is encased in clay-rich material and formed early in the sedimentary process, whereas on the left t h e second variety of isolani is prevalent in transgressive and regressive sandstones and shales, with the interfingering of the porous sandstones and shales. As is common in such stratigraphic associations, the coarsegrained units grade into finer silty and then into more clay-rich units. Variations in the sedimentary environment may produce thin shale laminae barely discernible. The former traps are called “primary isolani”, whereas the latter
ORE GENESIS INFLUENCED BY COMPACTION Oil deposits a
@ -$ .* .- .
+
+ +
y
+
+ + +
-
Ore deposits b
+ + +
+
... .
497
+
+
+ + + +
t
+
+ ++++\+ +
+
t
+
s
+
Fig. 5-7. Comparison of oil and ore deposits; migration paths o f fluids during formation of fluidogenic deposits. (After Pospelov, 1969, fig. 6.) a. accumulations of oil at structural barrier during migration of oil from source rocks in depressed area (after K.B. Ashirov); b. circulation of ore-bearing solutions in system of ore-yielding, ore-distributing, and ore-accumulating structures (after V.I. Smirnov).
“secondary”, because of the time relationship between the original stage of sedimentation and the formation of the traps. Silver (1973)pointed out that ~
~
HYDRODYNAMIC FLOW
COMPACTION
Fig. 5-8.Some distinguishing features of accumulation in isolani under compaction with and without accompanying hydrodynamic flow. (After Silver, 1973, fig. 1 ; courtesy of Bull. Am. Assoc. ,Pet. Gevlogists.)
498
K.H. WOLF
occurrences of hydrocarbons in isolated porous bodies have been noted by many investigators, but that only few have analysed such accumulations in the context of a basin. This statement also applies to certain ore deposits in sedimentary-volcanic piles. Silver also discussed the significance of depth of burial, and geologic time therefore, in forming some oil pools that range from shallow to deeply buried and early t o late diagenetic. Again, the same arguments may apply to ores in sediments formed by compaction water and/or other basinal fluids. He stated (p. 731), for example, that “greater depth of burial, and hence compaction and diagenetic alteration, increase the chance that siltstones, shaly siltstones and silty shales surrounding more porous sediments will become the confining matrix by establishing barrier permeabilities”. In an excellent publication, Weeks (1961) summarized the most salient factors that control the origin, migration, and occurrence of petroleum, and it is hoped that within the near future a similar comprehensive publication can be prepared for ores in sedimentary and volcanic rocks. Weeks (pp. 5-6) stated that “hydrocarbons occur incidentally associated in minute quantities with certain forms of metallic ores and mineralizations, whose occurrence is subject t o similar basic environmental controls, such as pH, Eh, etc. Similarly, many metals occur disseminated in trace quantities in bituminous, finegrained, basinal sediments such as shales, mark, and dense shaly limestones”. And: “Metals that occur associated with hydrocarbons in the sediments or in the ash of the petroleum include V, Ni, U, Pb, Zn, Cu, Cr, Fe, As, Sb, Mg, Mo, Ag, Al,Cd, and probably others. These are commonly spoken of as trace metals; they are of world-wide occurrence in basinal sediments. They are extracted from the sea water by organisms, including bacteria, which probably use them in their metabolism.” Weeks proceeded to offer a list of some of the principal factors that tend to make a basin highly prospective or, conversely, the absence of which indicate that it is likely to produce little or no oil. Weeks’ summary table has been offered here (Table 5-VI) in an unaltered form, except that asterisks have been added t o indicate those factors that are of significance in the study of chemical sedimentary ores. The crosses point to factors that control directly or indirectly characteristics of compaction. It should be emphasized that there is a fundamental generality to the table such that some of Weeks’ factors may apply equally to the Colorado-Plateau type of uranium occurrences, whereas the same factors plus or minus other points, may be significant in the genesis of some Mississippi-Valley type lead-zinc deposits. Yet, another combination of factors may have led to the formation of the Kupferschiefer, Rhodesian copper and many other deposits. Pirson (1952),in a geologic note on the influence of permeability on ore distribution in sedimentary carbonate units published by Ohle (1951),men-
ORE GENESIS INFLUENCED BY COMPACTION
499
TABLE 5-VI Summary table of important factors that control petroleum occurrence with some supplemental information related to origin of ores in sedimentary-volcanic piles (modified after Weeks, 1961)
+* +* +*
(1)Favorable basin form; this is a highly important factor (2) Optimum degree of mobility (3) Environments, as reflected in the lithofacies and biofacies, favorable for the deposition of both source and reservoir facies + * (4) Favorable physical relationship between source and reservoir facies + * (5) Favorable volumetric and thickness ratios of reservoir/source-rock facies + * (6) High degree of sorting or screening of muds from the sands + * (7) Adequate rate of deposition + * (8) Variability in the rate of deposition laterally + * (9) Presence of lateral lenticularity rather than uniformity of deposition across the basin * (10) Presence of unconformities, including diastems or lesser breaks + * (11) Abundance and timeliness of effective basin: wide distribution of structural, compactional, or stratigraphic traps or, as is so common with good oil occurrence, combination of these + * (12) Presence of deposition sinks flanked by adequate porosity (“deposition basin” or “deposition sink” rather than “depositional basin or sink”, because it means a basin in which deposition was occurring rather than a basin that was produced by deposition) + * (13)Presence of broad structural noses or bulges plunging alongside the basin or out into the deeper parts of the basin * (14)Preservation from erosion of oil-bearing members where they lie near the top of a formation or series that is terminated upward by an erosional unconformity * (15) Lack of fresh-water flushing of pre-unconformity accumulations * (16) High percentage of the original deposition basin (particularly the more favorable or oil-bearing trend parts of the basin) that remains uneroded and otherwise intact in the present structural basin * (17) Favorable evidence of oil and gas in the formation * (18) Favorable results of any exploratory drilling + (19) Absence of regional or local metamorphism + (20) Optimum depth for prospective oil horizons + * (21) Adequacy and timeliness of impermeable cover and of base and updip seals; these are among the most important of all factors * (22) Presence of evaporites, either below as a favorable structure former (e.g., salt domes or salt anticlines) o r above as a cap rock or as an indicator of the probable existence of silled basin conditions + * (23) Aquifer system having good porosity and permeability; conditions of wide open artesian flow of wide basin dimensions, however, are not generally favorable; blanket sands require more closure and closure permanences than do porous bodies with lenticular porosity; also, lenticularity usually indicates more rapid deposition and closer, hence more favorable, lateral and vertical association of source and reservoir rocks
* = Applicable to sedimentary ore investigations, if
appropriate extrapolations are made.
+ = Compaction should be considered here, e.g., different compaction mechanisms,
degree of compaction, differential compaction, and regional variations of compaction.
500
K.H. WOLF
tioned the relationship between the carbonate rocks and the occurrence of hydrocarbons and ores. Pirson visualized the mechanism as follows: (1)Mineralization fluids dolomitized the host rock which resulted in an increase in porosity and permeability. (2) As a result of the presence of a permeability barrier, possibly offered by the non-dolomitized rock, sulphide-bearing mineralization fluid caused precipitation of zinc sulphide in the dolomitized porous country rock. (3) The sulphide-depleted solutions moved out of the rock without leaving traces. Pirson believed that, whatever the exact process of oil formation may be, it is acceptable that the oil-field fluids carrying the hydrocarbons, may also have been the mineralizing fluids. Moving from the source beds through the reservoir rocks, where the oil is strained and screened by permeability barriers such as shales, the solutions may cause dolomitization wherever limy beds are encountered (see also Jodry, 1969). Thus, permeability is increased and ore mineral precipitation could follow. Pirson mentioned also that the oil field fluids were mobilized most probably by compaction from the source rocks. Grutt (1957)proposed that natural gas may have been the reductant for the uranium deposits in the Gas Hills area of Wyoming. According to Harshman (1970),however, most of the ore mineralization here is not related to the paleo-outcrops of the hydrocarbon-containing units, as would be the case if gas were the precipitant, unless secondary migration was possible. Nevertheless, he did not deny an influence elsewhere and stated that an excellent case can be presented for gas being the precipitant for the recently discovered roll-type uranium ores in the coastal plains of Texas. Nininger et al. (1960) also discussed the role of petroleum in the origin of uranium deposits, but offered no conclusive evidence and stated (p. 47) that “the association of crude oil with uranium deposits, apparently is fortuitous, in as much as very few ores contain obvious traces of oil”. H2S produced during maturation of petroleum and by bacteriogenic reduction of sulfate, however, is recognized as a factor in the deposition of uranium. In addition, oil can transport uranium and other elements in minor quantities, so that a more definite and a more fundamental role of hydrocarbons in the origin of metalliferous deposits cannot be ruled out. Perel’man (1967)also discussed the origin of uranium in sedimentary rocks and described several types of “geochemical barriers” which cause a reduction in migrational ability of elements in moving intraformational fluids. These barriers may be syngenetic, diagenetic and/or epigenetic and may be mechanical, physicochemical and/or biological in origin. Compaction processes may have an effect on the properties of these barriers. What is of particular interest at the moment, however, is the advocated relationship between petroleum- and/or bitumen-containing rocks, on one hand, and the
ORE GENESIS INFLUENCED BY COMPACTION
501
metal-precipitating barriers, on the other. Perel’man stated that one type of barrier develops at the boundary between water and oil when oxygenated waters invade the hydrocarbon-containing rocks. These fluids encounter a reducing barrier which is created in the water near its contact with oil or bitumen, conditioned by the H2S which is generated. In turn, the presence of HzS results in pyritization of the rock and fixation of U4+. Evans et al. (1968)outlined the requirements for the formation of the western Canadian Mississippi-Valley type ores of the Pine Point district which are associated with reefal facies containing petroleum and H2S-bearing gas: (1)Favorable source beds, such as clay-rich sediments, containing hydrocarbons, sulfur and metals. (2) Migration to and trapping of the hydrocarbons in porous traps, such as reef limestones. (3) Concentration of metals, in this case especially Pb and Zn, in the brines beneath the gas caps and petroleum pools. (4) The interface between the metal-bearing brines and the H2S-rich gases acting as the locus of sulfide ore mineralization. In instances where the metal-rich solutions move through rocks that do not offer a suitable trap and H2S is absent, no sulfide will result. It seems pertinent then that the hydrocarbons migrate prior to the metal-bearing waters. This in turn implies that hydrocarbons may be generated earlier in the evolution of a basin or region than the metal-rich solutions. There is a possibility that this need not be the case in all basins. The question then arises, what happens to the metals if they do not encounter organic barriers? What other geochemical barriers (see Perel’man, 1967) could provide a means of precipitation, and if none exists would the metals escape with the fluids? Future research may find the answer as to what conditions in the basin evolution will release either hydrocarbons or metals first into the subsurface fluids. Golovin (1970)in his section on “Ore zoning of artesian basins” stated that there is a close genetic, paragenetic and spatial relationship between the occurrence of subsurface waters, hydrocarbons, and many metalliferous and non-metalliferous minerals. Subsequent to a burial of the sediments to a depth of about 1.5-2 km, there is a separation of liquid and bituminous material and the accumulation of oil and gas, possibly as a result of compaction and clay mineral transformations. Under further geologic developments, the basin changes into several zones that differ in their hydrodynamic relationships: (1) the upper crestal or recharge zone with actively descending waters; (2) the limbs with passively descending waters; (3) the central zone with comparatively stagnant fluids; and (4) the terminal zone of discharge with ascending waters. In zone 1 of the artesian basin, gas prevails over oil
502
K.H. WOLF
because the liquid hydrocarbons were removed by descending fluids. Large metalliferous precipitations of Cu, Pb, Zn, S, together with celestite, are absent here. Along the limbs (zone 2), both gas and oil accumulations are present. The Cu and polymetdic (Pb-Zn) ore deposits (e.g., Naukat, Dzhezgazgan) formed here by ascending 02-free solutions. The largest oil and gas occurrences are in zone 3, whereas in zone 4 oil pools prevail over gas accumulations. Major celestite deposits as well as asphaltite and ozocerite are characteristic of zone 4. In the same symposium other researchers described syngenetic uranium (uraninite and pitchblende), which is characteristically associated with solid insoluble organic bitumen in skeletal and oolitic shallow-water marine limestones. High contents of V and Se, together with increased contents of Co, Ni, Mo, and sometimes As and Ti, were observed. The ore-bearing carbonate beds are also oil- and/or water-bearing. Dunham (1970) is one of several researchers who have utilized the information on subsurface fluids in sedimentary basins in an attempt to show that these solutions can be responsible for numerous types of ore concentrations. He mentioned that the members of the United States Geological Survey studying the Mississippi-Valleytype ores, for example, have found that fluidinclusion studies of ore minerals (e.g., Roedder, 1967, 1972) indicated that the composition and temperatures of the mineralizing fluids are very similar to those of adjacent oil pools (see also White et al., 1963). Referring to the works of Skinner (1967), Beales and Jackson (1968), and Jackson and Beales (1967), Dunham (1970) accepted the plausibility that ore genesis and petroleum accumulations can be complementary to some extent and both belong to the grand diagenetic process in sedimentary environments. Dunham pointed out then that if this hypothesis is accepted, certain practical implications follow, namely that basin analyses, as done by the oil industry, will have to be applied in the exploration for ores in sedimentary and volcanic piles. If the Mackenzie Basin is the generator of the Pine Point ores and the Illinois Basin is the source area for the mid-continental mineral deposits of the U.S.A., then the North Sea Basin could be considered to be responsible for the Pennine ore fields in Great Britain. Dozy (1970) in a contribution that is part of a sequence of papers on related problems (cf. Tooms, 1970; Dunham, 1970; Davidson, 1965; Bush, 1970)*, together with Beales and Onasick (1970) made a very strong appeal for compaction in their geological models for the origin of lead-zinc ores of the Mississippi-Valley type. Dozy stated that there is little opposition to the concept that the metallic elements in sedimentary ores have been supplied, either partly or wholly, by
* All these authors presented their ideas in the symposium in which Dozy’s paper of 1970 was published.
N E OKLAHOMA PLATFORM
OZARK U P L I F T
I I
I
I
PERMIAN
ILLI NO1S - IN D I A N A KENTUCKY
G ARCH
I
PERMIAN
PENNSY LVAN 1Ab
PENNSYLVANlAh :ALLEGHANYl
MERAMEC
MISSISSIPPIAN
L S t Wl.
MISSISSIPPIAN
A
DEVONIAN
DEVONIAN
Lst, C h t
ALEXANDRIA
Lst,
HUNTON
SI LURl AN
er:.
3zz
Lst 6 D d
GALENA Do1 OECORAH Do/, Lst Dal, L DLATirOLLE
MAQUOKETA TRENTON
ORDOVICIAN
DEVONIAN
ST. P E T E R
SILURIAN
MAUUOKETA Sh
a,
ORDOVICIAN
S t PEJER S s t __c_
ARBUCKLE (=ELLENBURGER1
PRAIRIE 3 U CHIEN
Oaf BONNETERRE 001
CAMBRIAN
DO1
MADISON
SSt
MAOISON S s t
LS?
LSt
CAMBRIAN
PRECAMBRIAN
PRECAMBRIAN
O C C U R R E N C E OF
I
OiL
Fig. 5-9. Schematic stratigraphic table indicating distribution of oil and ores in the Paleozoic of part of the Central Shield of North America. (After Dozy, 1970, fig. 1; courtesy of Trans. Inst. Min. Metall.)
504
K.H. WOLF
brines, connate and/or formation waters, and that subsequent concentrated precipitation formed ore districts. Hydrocarbons were intimately associated with these formation fluids from their origin through migration to final accumulation. Thus, proper evaluation of the geological conditions of a sedimentary basin is required to comprehend the origin of both ores and hydrocarbons. As Fig. 5-9indicates schematically, both occur roughly in the same rock units. There are even more striking correlations (Figs. 5-10and 5-11), which is not surprising because subsurface fluids were the medium of transportation for both metallic ions and hydrocarbons that accumulated in porous and permeable rock units. As Figs. 5-11 and 5-12demonstrate, oil was preserved in the basin, if trapped on its way up by the structural and stratigraphic barriers, whereas it often escaped on regional highs. The ores, on the other hand, occur mainly on these paleo-highs. As is generally accepted, the oil source rocks are clayey or micritic sediments with a high organic content (e.g., black shales and marly limestones after consolidation). It is also known that the total salt content of the original brines may have been approximately 150,000-200,000 p.p.rn., and that the high mineral content of the formation waters is often the result of a release of adsorbed ions to the subsurface fluids during compaction or diagenesis. Dozy believed that compaction and the release of adsorbed water on clay minerals is one of the main driving forces to move formation fluids. But he also lists osmotic pressures, caused by semi-permeable clay layers, and tectonic stresses as contours in feet on top Trenton and basement Trenton, mainly permeable dolomites Trenton, tight limestones
Fig. 5-10. Facies distribution of Ordovician Trenton Formation and distribution of ores in Upper Mississippi Valley district and hydrocarbons in the Lima-Indiana fields, both on Kankakee Arch. (After Dozy, 1970, fig. 2; courtesy of Trons. Znst. Min. Metall.)
ORE GENESIS INFLUENCED BY COMPACTION
Mississippian
51 lurian
L 1
/ Devonian
Ordovician
Carnbro - Ordoviaan Granite
505
‘%a
ollpools
($&,Ore districts A B
Tri-State N Arkansas
c
5 E MISSOUPI
D
tllinois /Kentucky
Fig. 6-11. Distribution of oil and ore occurrences around the Ozark uplift. (After Dozy, 1970, fig. 3; courtesy of Trans. Inst. Min. Metall.)
causes. Commonly, fluids will pass into porous and permeable contiguous formations, but under certain conditions the water released during compaction and diagenesis cannot move from the fine-grained units into adjacent coarse-grained rocks, and the fluid pressures, therefore, can exceed the hydrostatic levels and approach lithostatic or overburden pressures. The results are overpressured zones, flowing clays, and mud volcanoes, for example. Tectonic unrest may lead to occasional disturbance of such unstable conditions (cf. discussions on thixotropy) and the sediments may settle, accompanied by mobilization of the excess formation water under high pressure. The released flood of water will tend to find an outlet towards the surface and the stream of matured brines may push from the deeper portions of the basin interior, where diagenesis is most advanced and compaction most intense, towards the upper portion of the basin rim (Fig. 5-13).Dozy proposed that precipitation of ore can then take place in the shallower deposits. As he mentioned, however, the brines responsible for the ores of the Illinois-
K.H.WOLF
506
.-.-
Kan.
I
Post - Ordovician Cambro-Ordoviciar Precambrian Ore district
Fig. 5-1 2. Salinity of formation water of St. Peter Sandstone (Ordovician) (isoconcentration lines in p.p.m.). Partly redrawn from Levorsen, 1967. (After Dozy, 1970, fig. 4; courtesy of Trans. fnst. Min. Metall.)
Kentucky district followed a different migration route, namely along the system of the NE-SW faults that occur in the whole area. On their way up these fault-zone conduits, the heavy saline fluids penetrated laterally the most porous and permeable units, thus replacing lighter and colder intraformational fluids, and precipitated the fluorite and other minerals (Figs. 5-14 and 5-15). Dozy does not maintain that water of compaction is the only fluid responsible for the origin of the Mississippi-Valley ore deposits and accepts the possibility that other types of fluids may have been at least partly responsible and became mixed with the water of compaction. Dickey (197 2) discussed the migration of interstitial water in sediments
ORE GENESIS INFLUENCED BY COMPACTION
507
Sediments ot old tectonic u p l i f t
'
Precipitation of minerals
Fig. 5-13.Sketch showing proposed flow of brine from center of a basin towards a high at its edge. (After Dozy, 1970, fig. 6 ; courtesy of Trans. Inst. Min. Metall.)
and their role in concentrating petroleum, gas, and minerals in terms of hydrodynamics, compaction process in sediments, physical and chemical processes associated with compaction, and chemical composition of interstitial waters. In regard to the precipitation of minerals, he mentioned that the interstitial water will be hot in the deeply-buried sediments and often
$$$\
Illinois-Kentucky d i s t r i c t
Fig. 5-14. Structural map of Illinois Basin, contoured in feet on top of Devonian, indicating location of the Illinois--Kentucky district with respect to basin configuration and Rough Creek Fault zone. Redrawn from Ver Wiebe, 1957. (After Dozy, 1970, fig. 7 ; courtesy of Trans. Inst. Min. Metall. )
K.H.WOLF
508 Sediments ot basin
-
j
Invasion o t brine replacing Ilghter
formation tluid
r*ihr$i ’ Precipitation of minerals
Fig. 5-15. Sketch of model for the XHinoisKentucky district indicating “vertical” escape of fluids along fault. (After Dozy, 1970, fig. 8; courtesy of Trans. Znst. Min. Metall.)
saturated with various minerals. As the solutions move up into shallower horizons, they cool and, consequently, precipitate silica, carbonates, feldspar and others, often in a predictable paragenetic sequence (see Chapter 3 on sandstone diagenesis). Many so-called hydrothermal Cu, Zn, Pb, Au, and Ag ore veins were most probably formed by connate fluids, rather than by solutions from igneous sources. According to Dickey, the process has many similarities to the concentration and accumulation of hydrocarbons, except that it occurs at higher temperatures. Diagenesis and metamorphism, as well as the origin of oil and ore minerals, occur in an ambient environment of saline fluids, which enhances the chemical reactions. As a result of maturation, the composition of subsurface fluids changes greatly with depth of burial and geologic time. Compaction is one mechanism that causes the movement of these solutions and, as Dickey pointed out, the flow pattern in the basin determines the eventual location of the hydrocarbons and the ore mineralization. Hitchon (1971) described several common features of oil and ore mineral accumulations in sedimentary rocks. Both are composed of aggregates of originally widely disseminated matter that upon remobilization, transportation, and secondary accumulation become concentrated at sites that are structurally and/or stratigraphically controlled. As described by Hitchon, the metal-bearing fluids within the sedimentary basin are controlled by the various hydrodynamic potential fields. In agreement with other investigators, he recognized the importance of compaction in moving subsurface solutions. Hitchon then referred to the Pine Point ores as an example of a deposit formed by the precipitation from formation waters. The numerous studies undertaken (Jackson and Beales, 1967; Cummings and Robertson, 1969; Fritz, 1969; Jackson and Folinsbee, 1969; Sasaki and Krouse, 1969) pointed to an origin of ores from formation waters. Trace-element studies of the present down-dip intraformational fluids (Billings et al., 1969), together with the examination of regional flow pattern and the chemical variations of formation waters by Hitchon (1964, 1969a,b), support this interpretation.
ORE GENESIS INFLUENCED BY COMPACTION
509
The detailed distribution of the hydraulic head in the Middle Devonian Keg River Formation of northern Alberta is given in one of Hitchon’s diagrams. The flow direction is generally updip towards the Pine Point lead-zinc deposit, which is located at the outflow end of a porous and permeable lime stone reef complex. The recent hydraulic head pattern indicates that the Pine Point ores and the Zama-Rainbow oil occurrences are the result of updip formation water movements along the porous and permeable carbonate rocks. These channels allowed the fluids expelled from the dark clay-rich sediments to move updip into reservoir rocks, where first the oil and then the ores were deposited. The review of Saxby (1969)of the metal-organic reactions in the geochemical cycle supported the theory that the close relation of some oil pools and sedimentary ores indicates an intimate association of organic and inorganic components in subsurface fluids. Cordell (1972)in his critical review of the depths of oil origin and primary migration offered information and raised questions that are directly applicable to ore genesis within sedimentary-volcanic piles (see also McCulloh, 1967). The problem as to whether oil originates and migrates at shallow depths or in deeper horizons, and all the numerous factors that determine this depth, therefore, are just as important to ore studies. If effective barriers exist at shallow depth that prevent the movement of hydrocarbon precursors, the same should be true for metal precursors. Cordell concluded that oil resulted from thermocatalytic transformations and primary migration at depths between a few thousand to about 10,000f t because of: (a) chemical alteration of organic matter at elevated temperature-pressure conditions; and (b) as a result of clay-mineral transformation that releases water and organic matter (and possibly trace elements), thus allowing good source-bed drainage and fluid migration. All this should be given consideration in future investigations of the origin of ores in the sedimentary environment. Paxticularly noteworthy to investigate would be the relationship, if any, between the factors that cause the release of organic material from the source beds into the subsurface fluids, on one hand, and the release of the metallic elements, on the other. According to Cordell, the primary migration of hydrocarbons is apparently related to the stages of clay-mineral dehydration, of which the most important stage is temperature-dependent. The result is a continuous or recurrent migration over a long time and in broad depth intervals. The same arguments apply to the origin of metalliferous precipitates, if one accepts the proposal of some researchers that the elements are present as “disseminations” (adsorbed on fine particles) in the sedimentary basin. Upon release into fluids, migration, and precipitation, the metallic elements would then form ore deposits. In the chapter on sandstone diagenesis (Chapter 3) as related to compaction, the reader will find summarized information on the relationships be-
K.H. WOLF
510
IN SOURCE BEDS
ACTIVITY
Fig. 5-16. Comparative activity of oil origin and primary migration as related to burial depth. Width of two graphs represents postulated amount of activity on a worldwide basis. (After Cordell, 1972, fig. 5; courtesy of Bull. Am. Assoc. Pet. Geologists.)
tween hydrocarbon distribution, the chemical properties of oil and gas, and the mineralogical and diagenetic-metamorphic characteristics of the host rocks, on one hand, and the depth of burial and geologic time, on the other. Some researchers have proposed a “depth rule” (Fig. 5-16) for the origin and accumulation of hydrocarbons, e.g., that oil in a basin becomes lighter and more paraffinic with increasing depth. It is only one additional step t o extrapolate this t o the present discussion and ask whether it will eventually be possible t o establish a “depth rule” for the mechanisms that form ore deposits in sedimentary-volcanic milieus. Dallmus (1958), in his paper on the mechanisms of basin evolution and its relation t o the habitat of oil, mentioned that a basin floor can subside nearly 2,000 m before any deformation occurs. Tectonic pressures, which are added at that depth to the overburden-compaction pressures (Table 5-I), will facilitate further fluid expression from the sediments. This, according to Cordell (1972), explains the more than coincidental fact that the 2,000-m depth is close to the top of the depth range of flush primary migration of hydrocarbons. If this depth relationship applies t o the migration of ore-forming compaction fluids, then one can establish the minimum depth in a sedimentary basin at which ores can form. Future studies, however, may indicate that the mechanism of release of metallic elements into the compaction fluids may be somewhat different than that of hydrocarbons, so that the minimum depth may be different for the metallic deposits. Also, depth of burial may control certain
ORE GENESIS INFLUENCED BY COMPACTION
511
types of fracturing, with the result that particular varieties of ore veins are confined below a specific depth of burial. The water content is highest in sediments, especially fine-grained ones, at very shallow depth, and most of the fluid is expelled to the surface during the initial 150-200 ft of burial. Under these conditions the concentrations of metallic compounds in the fluid may be too low due to dilution effects and the solutions, which are flushed out of the system, generally cannot be responsible for the origin of large ore bodies, although minor amounts of sulphide mineral can form. Also, at that early stage of basin evolution, there may not be suitable barriers to migration to result in accumulation of the solutions in a reservoir rock where precipitation of the minerals could otherwise take place. Three related publications should be mentioned here before concluding this discussion. The reader will find supplemental data on the correlation between the amounts and compositions of hydrocarbons and depth of burial by consulting Tissot et al. (1971). They too showed that thermal degradation of organic matter is a function of depth of burial in sedimentary basins. Connor and Gerrild (1971) described geochemical differentiation of crude oils from closely spaced sandstone units and demonstrated that the trace element content of the oil was correlative with their approximate stratigraphic position. The study also revealed that the oil consists of at least three distinct compositional types, which they interpreted as having been formed through natural fractionation of a single indigenous crude oil parent. Vredenburgh and Cheney (1971) have presented interesting data on the sulfur and carbon isotopes of petroleum that, when combined with the information obtained from other investigations, would be applicable in petrogenetic interpretations of sulfide ores. Of particular interest is Fig. 5-17 in which the sulfur isotope history of the petroleum, including the evolution of HzS, has been illustrated. The three publications just mentioned, especially when combined, offer a geochemical approach that appears to be promising in the study of ores associated with oil evolution. Let it suffice to say that: (a) the variations in types and amounts of hydrocarbons with depth of burial, (b) the variations in trace-element content, and (c) the variations in sulfur and HzS contents are possibly the result of chemical differentiation and fractionation, for example, which in turn may also explain the local and regional variations in ore mineralization. Investigations such as these (together with those on formation waters, e.g., Hodgson et d., 1964; Hodgson and Hitchon, 1965; Hitchon, 1968a,b; Billings et al., 1969; Boehm and Quinn, 1973) will have to be performed in the future on those ore deposits that are believed to have originated from the chemical components carried by crude oil and/or the formation waters involved in the origin, migration and accumulation of hydrocarbons.
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The oil exploration industry has made numerous contributions to the understanding of evaporite mineral and rock genesis, and their relationships to oil origin, migration, and accumulation. It seems appropriate, therefore, to at least mention again here an extension of these investigations into the possible connections between the presence of evaporites in sedimentary basins and the origin of ores, either with or without a relationship to hydrocarbons. The following authors have published on the latter possibility: Seidl (1959);Davidson (1965,1966); Brongersma-Sanders (1965,1967, 1970); Dunham (1966, 1970); Bush (1970);Tooms (1970);Vine and Tourtelot
(1970).
The mere fact that both hydrocarbons and certain metalliferous enrichments occur in sediments and their lithified equivalents, indicates that the geological, geophysical, and geochemical techniques and theories developed by oil explorationists can be applied, and this has been done so already in the search for ore deposits. Certain approaches, however, still remain to be tested. Although some of these techniques are of a very detailed nature and can be employed only if the necessary manpower and financial support are available, they eventually will give a better understanding of the genetic mechanisms in sedimentary ores. For example, the investigations on the diagenesis and paragenesis of quartz and other minerals has resulted in the determination of the relative, and even absolute, time at which oil migrated into host rocks and interrupted the diagenetic processes (for details see the
L
-
I
PRESENT-DAY SULFUR DISTRIBUTION
Fig. 5-17. Schematic sulfur-isotopic history of Paleozoic petroleum, Wind River Basin, Wyoming. (After Vredenburgh and Cheney, 1971, fig. 9; courtesy of Bull. Am. Assoc. Pet. Geologists.)
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discussions in Chapter 3 on sandstone diagenesis). Consequently, similar studies should be performed on ores. A comparison also should be made between the world-wide stratigraphic distribution (i.e., according to geologic age) of hydrocarbons in the geologic record, on one hand, with the distribution pattern of certain varieties of ores in sedimentary basins of all ages. The data on the distribution of hydrocarbons is much more complete at the present time than that of the ores. Are specific geologic periods better endowed with certain ores that could have had a genetic relationship to the origin and accumulation of hydrocarbons, or not? One word of caution is in order in relation to the above discussions. In spite of the plausible genetic connections between the hydrocarbons and ores, there is always the chance of their coincidental associations, as pointed out by Shchurkin (1973).He described one case where he believed that oil migrated along deep regional faults into higher horizons, and that the same routes were followed by hydrothermal solutions. Thus, ore minerals and bituminous material are present in the same veins. EXAMPLES OF STRATABOUND AND STRATIFORM ORE DEPOSITS AND THEIR RELATIONSHIPS TO SEDIMENT COMPACTION
Mississippi Valley-type l e a d z i n c deposits
It is believed that the limestone and dolomite host rocks are not the source rocks of the chemical elements constituting the ore minerals of the Mississippi Valley-type lead-zinc ores, but that, as noted earlier, the latter were possibly derived from the associated clay-rich units. In such a case, compaction in the reservoir rocks may be expected to have different significance from that in the source rocks. Indeed, whereas compaction in the clayey rocks is a fundamental process that mobilizes and moves the fluids into the host rock, compaction in the latter would not be conducive to the formation of ore concentrations as it would tend to reduce porosity and permeability. "he precise influence of compaction in the carbonate rocks would depend on whether compaction (together with other diagenetic processes such as solution and cementation) has occurred during a pre-, syn-, or post-ore stage. What is extremely important is the relationship between the compaction and other diagenetic effects of the clayey units, on one hand, and that of the adjacent reservoir rocks, on the other. For example, even if there is an absence of early diagenetic cementation in the carbonate sediments to allow early compaction, the physical and chemical parameters related to the fluids of Compaction, that come from the clayey sediments somewhat later and
514
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move into the host rock, may be conducive to cement precipitation in the carbonate sediments to eliminate the pore system. Inasmuch as a large amount of fluids had to pass through the rocks to account for the deposited volume of lead and zinc ore, the system must have been open for a relatively long geologic period. If, on the other hand, the early compaction fluids assisted in dolomitization of the original limestone, then there may have been an increase in porosity and permeability, thus actually preparing the host rock for the ore precipitation. One more example may suffice. As is well known, many reef complexes are composed of a core of in situ organic growth which is interbedded or intertongued with transgressive and regressive units of shoreward and seaward sediments (e.g., lagoonal and off-shore, deeper-water limestones, respectively). If early compaction fluids from the basinal clayey deposits are preferentially moving within certain horizons in the reef complex and if cementation occurs due to CaC03 precipitation, then these horizons may become tight, become cap rocks over porous and permeable units in which eventually the ore minerals can be precipitated because a new generation of compaction or other fluids can be trapped there. Here, then, the question arises as to what extent compaction fluids may prepare a host rock for subsequent ore precipitation as a result of dolomitization, decementation, and formation of solution channels. Possibly, the interbedded and cavity-filling evaporite minerals can be removed to cause collapse and brecciation of limestones to form features that may be similar to subaerially formed karst structures. In general, one may say that future studies should be carried out on the recognition of compaction processes that will lead to a decrease in porosity and permeability of the carbonate host rock and of those that may increase these properties. For example, the work by oil company geologists on subsurface fluids (see Chapter 3 on sandstone diagenesis) has shown that the release of compaction fluids may occur in stages, related to clay mineral transformation. Thus, the supply of waters may be periodic and if with it the fluid composition changes, then they will have a different effect on the invaded host rocks. This effect may change with geologic time, e.g., the result may be cementation, decementation, dolomitization, dedolomitization, and possibly subsequent oil and/or ore accumulation. A survey of the literature shows that compaction has been mentioned by numerous investigators who studied Mississippi Valley-type ore deposits, e.g., Amstutz et al. (1964).The information presented below is not complete by any means, but illustrates to what degree the process of compaction has been considered. Most of the published work is qualitative and brief with no details presented by the early investigators. More detailed studies, however, have been made more recently. The information presented here is given chronologically, with some exceptions. A number of investigations referred
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t o here have not mentioned compaction, but have been included here because it seems that in these cases the process of compaction could have been playing a role in ore genesis. In numerous cases, the present writer has made additional comments to supplement the information given by the original author of the particular publication. Ellison (1955) emphasized that lead-inc ores tend to occur in limestones and dolomites and were dependent on porosity, permeability and chemical properties of the carbonate host rocks. He mentioned that in the Southeast Missouri district the algal reefs were favorable passageways for the ore solution because of their permeability. By using paleoenvironmental techniques one could outline these conduits as well as the host and source rocks. Independent of whether the passageways were used by compaction or igneous hydrothermal ore-carrying fluids, environmental and sedimentological techniques would still be useful in exploration work as it is the facies distribution that controlled localization of the ore. Two aspects are considered here. (1) In Figs. 5-18A and 5-18B (Ohle and Brown, 1954) references are made to faults that might be considered by some to support derivation of the ore from deep-seated igneous hydrothermal solutions, which had passed via the faults into the algal reef passageways. This need not necessarily be so, because not all fractures were formed tectonically at a much later stage in the geologic history of the rocks. Particularly in reef complexes, where one often finds an association of various lithologies having different susceptibilities to compaction, differential regional settling, which is possibly induced by earthquakes, may cause joint and fault systems t o originate more or less penecontemporaneously. These systems could be either parallel to the boundary between the shallow-water carbonates and deeper-water basin shales, or could be normal to this transition zone. Of course, both systems could be present, depending on the breaking-up pattern of the reefoid masses. Thus, subsequent diagenetic compaction fluids could have moved into the fractures and from there into the primary cavities of the limestones and/or dolomites. (2) The second aspect involves the proposal of Mountjoy et al. (1972) who have suggested a slide mechanism responsible for moving large blocks of carbonates into off-shore, deeper-water environments. If this idea is accepted, then it may well be that these blocks, often hundreds or even thousands of feet in width and length, after moving into a basin may rest in very close proximity of the clayey source sediments. Compaction fluids from the surrounding sediments may move into these broken-up carbonate rock blocks t o react with them and possibly precipitate ore minerals. Again, only detailed regional and local paleoenvironmental studies, in combination with other approaches, may allow one to recognize this process. Knight (1957), in his paper on the source bed concept, postulated that all
K.H. WOLF
516 r
i
Fig. 5-18. Relation of algal rolls to lead orebodies in southeast Missouri. (After Ellison, fig. 4, 1955, and after Ohle and Brown, 1954; courtesy of Econ. Geol. and Geol. SOC. Am. Bull.) Note the relationships between the trends of the faults and the ore rolls. A. Portion of No. 5 Federal Mine, Flat River, Missouri. B. Portion of No. 9 Federal Mine, Flat River, Missouri.
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the ore bodies he discussed were derived from syngenetically precipitated sulfides along one particular sedimentary unit of the basin, and that the ore was subsequently remobilized in varying degrees as a result of rise in temperature within the host rock. He suggested that the Mississippi Valley-type ores of the (a) Upper Silesia field, Poland, (b) Northern Morocco field, (c) Upper Mississippi Valley, Tri-State area, U.S.A., and (d) Southeast Missouri fields in the U.S.A., owe their origin to the processes related to his sourcebed concept. Knight did not, however, like Noble (1963),for example, emphasize compaction as an important process. Instead, he seems to have believed that the rise of temperature was responsible for mobilization or remobilization of the metals from the source beds. Snyder and Odell (1958)in their study of the breccias in the Southeast Missouri lead district considered three factors involved in the origin of the breccias, namely, character of depositional surface, differential compaction, and the effect of water in the sediments. This study is, of course, significant because the breccias were important ore-controlling structures. Four major breccia zones exist, each embracing over 60,000 tons of rock, with individual ore bodies in the breccia zones ranging from over 6,000f t long and containing several million tons of ore to a few hundred feet long with several thousand tons of ore. During mining operations, dips as steep as 25" were measured; the flanks of the larger depositional ridges exhibited dips of 8"-10", which continued up to several hundred feet downslope and for several thousand feet along the strike. These dips may be partly the result of differential compaction, but it has been established that an original depositional dip of several degrees existed along the flanks of the carbonate ridges. In their section on the role of differential compaction, Snyder and Odell (1958)stated that the conspicuous facies differences in the lower Bonneterre Formation were reflected in differential behavior of the sediments during the early diagenetic history, especially in degree of compaction. The carbonate banks and the widespread algal bioherms growing on the banks were particularly influencing sedimentation. Depressions or basins between the calcarenite banks received large volumes of fine-grained material, e.g., lime mud, calcisiltite and argillaceous muds. The calcarenitic banks and the rigid biohermal framework were only slightly compacted, whereas the finer grained basinal sediments were strongly compacted. As Terzaghi (1940, p. 88) showed, calcareous muds may undergo almost as much compaction as muds composed of clay minerals, if lithification is delayed. Inasmuch as the degree of compaction in part depends on the amounts of fine-grained material in the individual sedimentary units, the intensity of compaction may gradually decrease from the basin to the bank facies where the intertonguing and intermixing of the finer and coarser debris occurs. This gradual change of compaction may have its greatest influence on the attitude of beds along the
K.H. WOLF
518
flanks of the depositional ridges, so that at this location the original dips would be greatly exaggerated. This process, which is accompanied by the squeezing out of connate fluids, continues until compaction is complete. These compaction fluids may have escaped vertically into the basin surface waters; however, relatively impermeable layers of fine sediments may have forced some of the fluids to migrate horizontally, especially during late diagenesis. As to the origin of the breccias, Snyder and Odell (1958) said that the
1
SEA L E V E L
I BANK FACIES
A
DEPOSlT ION
BASIN FACIES
3
t-------l SEA L E V E L
BANK FACIES
R u
I’
GOM PAGTl ON
BnsiN FACIES”)
EA LEVEL
SEA LEVEL
C
1
SEA L E V E L
I
E
I
SLIDING
D SEA L E V E L
I BANK FACIES GO M PAC T I 0 N
DEPOSITION
L
F
.
. ’ . . ‘BnSN. I
SLIDING
I
Fig. 5-19. Stages in formation of breccia bodies. (After Snyder and Odell, 1958, fig. 9; courtesy of Geol. SOC.Am. Bull.)
ORE GENESIS INFLUENCED BY COMPACTION
519
sedimentary depositional surface changed from the horizontal to inclined and back to the horizontal from the shallow-water banks t o the fore-reef slopes and, finally, to the basin, respectively (Fig. 5-19A). With progressive sediment accumulation, the overburden forced the connate water out of the basinal sediments, causing compaction and thinning of the fine-grained units. Within the coarser-grained limestones, large volumes of solutions could be stored in the intergranular voids. Although the overburden pressure causes tighter packing of the granular framework, there may be only very little decrease in volume (see Chapter 3 on compaction of carbonates in Vol. I of Compaction of Coarse-Grained Sediments). As the dips of the beds along the flanks of the carbonate ridges increased, as the process continued, the unconsolidated sediments became unstable and slid downslope (Fig. 5-19C). The initial movement, according t o Snyder and Odell, took place by sliding and flowage of the upper thin veneer of loose, uncompacted sediments. The frictional drag on the underlying beds disturbed successively older units, thus increasing the slide thickness and the volume of the moving material. The oldest beds affected were sufficiently coherent to act as a gouging and cutting tool on their underlying beds. This downward cutting was localized by a fracture zone formed during compaction. As mentioned already, differential compaction occurred between the bank and basin sediments. Fracturing, rather than intergranular compaction, took place in those units which did not easily adjust to the differential settling. The lower boundary of the slides in the upper sediments cut down rapidly through the fractured zone to an older unbroken calcarenite unit. As shown in Fig. 5-20, Snyder and Odell believed that this process localized the slide plane. Following the initial slide, the breccia ridge was buried as normal deposition continued. If the bank and basin facies relationships remained unchanged with continued deposition, the process would be repeated when a sufficient thickness of sediment was deposited and compaction again resulted in oversteepening of dip (Fig. 5-19E). Two slump episodes along a basin margin are common; three are not unusual. A t least one case was recorded with five distinct periods of slumping. In calculating the breccia thicknesses and amount of compaction, it was found that the pile-up of breccia at the foot of the slope above the floor of the basin is nearly equal t o the depth of gouging. Snyder and Odell also believed that the compaction of the carbonate muds, probably aided by recrystallization, was essentially completed under a load of 40 ft of sediment. As the origin of the Mississippi Valley-type ores, Ohle (1959) reviewed five possible modes of genesis; however, he favored a “hypogene-hydrothermal” mechanism related to deep-seated igneous activity. When he considered the two genetic alternatives involving ground water, i.e., one invoking artesian circulation and the other downward percolation or lateral movement of ground water, he concluded that both concepts cannot be rejected and de-
K.H.WOLF
520
INCREASED COMPACTION
FAILURE
Fig. 5-20. Stages in localization of glide plane. (After Snyder and Odell, 1958, fig. 10; courtesy of Geol. SOC.Am. Bull.)
serve further attention. Ohle also mentioned the theory of primary, syngenetic sulphide precipitation which has been proposed for the Rhodesian (South Africa), White Pine (Michigan, U.S.A.), and the Kupferschiefer (Germany) copper deposits, as well as for several others. Ohle rejected this hypothesis for the Mississippi Valley-type deposits because the ores occur as open-space fillings mainly. When one reads his explanation of the process of ground water origin, however, i.e., that the original deposition of the metals occurred during more or less normal sedimentation and that the widely scattered bits of metal were gathered up by leaching from the rock and concentrated by redeposition, then a possible relationship between the two separate theories of “original, syngenetic deposition” and “ground water precipitation” can be considered. The only difference between these two theories lies in the concept that in one case the metals supposedly remained in the source muds, whereas in the other the metals were remobilized by leaching and transported and redeposited in a more “favorable” host rock. The differences can be reduced by considering the following: (1)One can assume that the ore minerals can be precipitated in both black
ORE GENESIS INFLUENCED BY COMPACTION
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muds and in carbonate sediments, as long as the geochemical requirements are met. (2) In case 1, where the ores are found in black shales, a remobilization and concentration of the metals may not have been necessary, if the original metal was precipitated in highly-concentrated quantities to constitute economically mineable ore. On the other hand, if the metals were deposited in a disseminated form through a thicker pile of mud, subsequent leaching, transportation and redeposition may have been necessary in the same lithologic environment to concentrate the chemical elements into a unit large enough to form ore. Inasmuch as in this case there was no transport of the metals from a mud “source rock” t o a carbonate “host rock”, one should not make a theoretical lithologic distinction between the two. Either one of the above cases may apply to deposits like those of the Kupferschiefer or Rhodesian copper ores. (3) In case 2, in which the ore is found in a carbonate host rock, the possibility exists that the metals were remobilized in the associated muds or shales and were transported to and precipitated within the carbonate rocks. As discussed elsewhere in this chapter (pp. 547 to 552), it is geologically possible to form fracture and cavity systems very early, so that syngenetic and/or diagenetic ore veins could form. The above discussions show that the main difference between the “original, syngenetic deposition” and “ground water theory”, as mentioned by Ohle, lies in the presence or absence of remobilization and subsequent reprecipitation of the chemical elements. Studies of the regional stratigraphy, regional and local paleohydrology, and geochemical facies relationships, in conjunction with compaction studies, should be undertaken to determine which theory is the more plausible one. Metals may have been originally deposited with the clays and organic matter which formed part of a mud. This mud was initially very rich in interstitial fluid and offered an internal geochemical milieu that under one set of conditions (a) will not allow the leaching or dissolving of the metals and, therefore, no remobilization can occur (or the process is at a minimum), independent of the degree of compaction and the volume of squeezed-out water of compaction. The metals remain, or are only slightly diagenetically remobilized, in the mud. On the other hand, under another set of geological and geochemical conditions (b) the water of compaction can first dissolve and move the metals and then precipitate them in a favorable environment. The precipitation can take place either: (1)in the same type of lithologic unit that had also served as a source rock, i.e., the precipitation occurs in a mudstone or shale; or (2) in a carbonate rock or a sandstone, into which the compaction solutions or ground water moved. Regional studies should aid in determining whether or not good reservoir rocks were available in the former case. Or could this
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K.H. WOLF
situation be analogous to the origin of oil shales with nonhydratable clays? Here, the hydrocarbons may have formed in a mud but also remained behind in the same mud due to the lack of “oriented” and “adsorbed” water to push the oil out (see Chilingarian and Rieke, 1968). Obviously, some of the original “free” water moved out during compaction, whereas some mechanism allowed the oil to stay in place. Kendall (1960) believed that the numerous features studied by him permitted proposal of another theory of ore genesis in addition to the more traditional one based on tectonic structures and hydrothermal activity. He also gave consideration to a syngenetic precipitation of the metals from sea water and that subsequent migration into the fractured, fine-grained dolomite host rock could explain the existing “epigenetic” features. Extraction of zinc by marine organisms (Wolf et al., 1967), possibly together with preferential adsorption on clays, could be responsible for the mono-metallic zinc deposit. Conversion of the zinc from the ionic state or organo-metal complex (see Saxby, 1969, for example) to zinc sulfide, may have happened during burial. Here, one may remember that Knight (1957) has envisaged that mobilization of the chemical elements, which formed ore upon concentration, was largely due to an increase in temperature. Although temperature may have been important in the process of remobilization of elements under some conditions, however, it has been shown in some experiments that a release of metal ions from clays into fluids can occur without an increase in temperature (Weiss and Amstutz, 1966). Amstutz (1963) mentioned diagenetically-squeezed marcasite nodules which were “healed” by galena, possibly during the syn-ore stage. Here, and in other papers, e.g., Amstutz et al. (1964), Amstutz used Illing’s (1959) diagram to show movement of fluids during diagenesis, i.e., movement of fluids during compaction from the deeper-water muds into the shallow-water reef complexes. The work of Amstutz and his co-workers became so influential that it has been referred to as the “Amstutz school” (Brown, 1970). Similarly, Noble (1963) also has proposed that the Upper Mississippi Valley lead-zinc deposits and the lead ores of the Southeast Missouri district may have originated by precipitation of the metals from compaction fluids that moved from the basin onto the flanks of the carbonate rock-rich highs (Fig. 5-21). Callahan (1964) emphasized that ore localization can be controlled by paleogeographic features, e.g., unconformities, and associated structures, such as karst-cave collapse breccias below unconformities, reef talus, compaction and drape structures, landslide breccias, and zones of facies change occurring above unconformities. In describing the Tri-State district of Missouri-Kansas-Oklahoma, Callahan stated that the mineralized section is bracketed by unconformities and disconformities as a result of periodic up-
ORE GENESIS INFLUENCED BY COMPACTION
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Fig. 5-21. Map showing relation of major ore concentration to the major structural basins. (After Noble, 1963, fig. 4; courtesy of Econ. Geol. )
m Typcol lead occurrences
0.
5
x) miles
Fig. 5-22. Diagrammatic cross-section of the southeast Missouri lead district, The New Jersey Zinc Company. (After Callahan, 1964, fig. 6 ; courtesy of the author.)
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K.H.WOLF
lift and erosion. Thus, the carbonate rocks developed a subsurface drainage system along the unconformities, and breccias were formed by solution and collapse. The domes and depressions along the erosional surfaces created compaction and drape structures in the younger, overlying beds which were conducive in trapping the metal-bearing fluids. Callahan proposed a similar compaction origin for the broad folds of which some are associated with the unconformities in the Upper Mississippi Valley lead-zinc district of Northwest Illinois-Southwest Wisconsin-Northeast Iowa. In the opinion of the present author the best example is the Southeast Missouri lead district (Fig. 5-22) about which Callahan said that this district is confined to a belt on the Ozark dome with a primary lead production (Pb/Zn ratio is 50 to 1)from a 375-400 f t thick dolomite bed of Late Cambrian age (i.e., the Bonneterre Formation). The unit contains sedimentary features significant from the standpoint of ore localization, e.g., compactional arch and drape structures reflecting irregularities in the surface of the underlying Early Cambrian sandstone or biohermal reefs (Fig. 5-22), formed along the paleo-shoreline, sedimentary breccias generated by landslides, and talus accumulations pinching out against Precambrian highs (Fig. 5-22). Although some of the ore occurs in fractures, sedimentary features, such as those mentioned above, control the fundamental pattern of the distribution, shape and size of the ore bodies. According to Callahan (1964),exploration in Southeast Missouri, based on paleogeographical and sedimentological principles rather than on tectonic deformational theories, has been very successful in revealing new ore districts. Here too, at least some of the ore-controlling structures were formed by compaction, but what remained unsaid was the possibility that water of compaction could have been important in preconditioning the conduit beds as well as the host rock through diagenetic processes. The location of the typical lead occurrence suggests that compaction fluids, possibly from all the three formations presented in Fig. 5-22, passed through the reefs. Also, fluids from the on-shore carbonates of the Bonneterre Formation may have passed through the reef and possibly caused dolomitization, thus increasing the porosity and permeability of the host rock. It seems then, that even if compaction fluids did not supply the metal for the ores (a possibility not to be dismissed entirely), they as least may have preconditioned the sediments to become receptive for the metal-bearing fluids. The French school of syngeneticists, composed of A.J. Bernard, J.C. Samama, P. Nicolini, and others, emphasize sedimentation, sedimentary petrology, paleogeography, and low-temperature and low-pressure geochemistry in explaining the origin of many stratabound ores. Fuchs (1964),for example, presented two diagrams (Figs. 5-23 and 5-24) suggesting that intrastratal fluids, e.g., compaction water, may have been responsible for the ore
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Fig. 5-23. Interpretive scheme of the formation of the lenticular barite in the upper part of a structure. (After Fuchs, 1964, fig. 11; courtesy of Sci. Terre.) 1 = breccias and calcarenites; 2 = fine dolomitic facies; 3 = horizon of siliceous induration; 4 = barite mass; 5 = evaporites; 6 = disseminated barite patches; 7 = direction of migration of the fluids during diagenesis.
mineralization which often is zoned. In one case, the theory was applied to barite deposits, but lead-zinc mineralization in carbonate rocks could form under similar geologic settings. Hoagland et al. (1965) in their attempts to explain the origin of certain
Fe-Mn
Barite
\
Sulfide:5 ( z n , PIb )
B A S IN
Evapori tes
Fig. 5-24. Theoretical division of the syngenetic deposits as a function of the paleogeography. (After Fuchs, 1964, fig. 14; courtesy of Sci. Terre.)
K.H.WOLF
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features associated with the Ordovician zinc deposits in East Tennessee resorted to the occurrence of compaction during the various geologic stages. They believed that compaction took place during at least four episodes: (1) during the formation of the host rock in a marine environment; (2) during the subaerial origin of the karst features, when the erosion penetrated into older uncompacted carbonate sediments and may have released large volumes of connate water to the surface (thus also changing the formation pressures, possibly); (3) during dolomitization, which may have been contemporaneous with ore genesis; and (4) during the post-ore geologic history as a result of the overburden pressure of younger rocks. Ford and King (1965) in their treatment on the layered epigenetic galenabarite deposits in Carboniferous limestones of Derbyshire, England, suggested that the metal-bearing solutions (possibly, ground-water brines concentrated by nearby uprising hydrothermal solutions) migrated either laterally or downwards to precipitate the ore in caverns, for example. Subsequent ground-water movements resulted in further cavernization of the host rock, which was filled by the sand supplied from large sink holes on the surface (Fig. 5-25). Ford (1967) considered various alternatives for the genesis of the galena, sphalerite, fluorite and barite, and concluded that “it is tempting to speculate whether the source was ‘emanative’ from a buried magma to the east of Derbyshire, and for which there is no direct evidence, or whether it was the result of the concentration of metals from the Permo-Triassic or underlying Carboniferous of the North Sea by means of K-rich brines. If a source for the latter is sought in or beneath the deep Permo-Triassic basins of the North Sea, a mechanism to provide a laterally upward hydraulic gradient is required . . Ford proposed that the depth of burial was sufficient to have provided compaction pressure to move the fluids and to increase the temperature of the solutions so that they can be called “hydrothermal”. He concluded, however, that the addition of juvenile waters of magmatic origin may have occurred also. As to ores generally deposited in karst caverns within limestones and dolomites, the following considerations of their origin are important: (1)Are the caverns within the shallow-water limestones, of very early geologic origin when the basinal sediments were still undergoing accumulation and compaction? In this case, one can envisage a renewed subsidence of the shallow-water rocks leading to the deposition of a sedimentary cover over them to form a cap rock. Then, any of the compaction fluids moving from the basin into the shallow-water rocks can be trapped and precipitate material into the karst cavities. (2) Or can it be established that the caverns were formed much later in the geologic history of the rocks when all early diagenetic processes, including compaction, were completed in both the shelf and basinal sediments? In this
.”
ORE GENESIS INFLUENCED BY COMPACTION
527 2
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4
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.. . .... . .... .. .....,..,.... . . . .: . . . .......
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6
n
Fig. 5-25. Diagrams to illustrate the evolution of the Golconda ore deposits. (After Ford and King, 1965, fig. 11;courtesy of Econ. Geol.) 1. The dolomitized limestone rests with an undulating boundary on unaltered limestone, both with scattered clay beds. 2. Early solution cavities develop at the base of the dolomitized limestone and above some clay beds. 3. The solution cavities are partly filled with bedded ores composed of precipitated galena and barite, which are interbedded with derived dolomite crystal sand. Barite, or later calcite, lines some of the resulting cavities. 4. Repeated phases of solution lead to intermittent collapse of parts of the bedded ores and of the dolomite roofs. The Tertiary sand cover may have been deposited during this stage. 5. Further solution leads to the collapse of more dolomite and allows the sand cover to subside, with some washed into both solution cavities and collapse caverns. 6. While solution and collapse continue, surface erosion, probably glacial, removes all the sand cover except that preserved in the solution-collapse pits.
528
K.H.WOLF
instance, it seems unlikely that water of compaction from contemporaneous sediments gave rise to the deposition of the ores. (3)In considering the above two possibilities, future work should establish how long it takes for compaction fluids to be pressed out of basinal sediments, and also determine the factors that control the variation in time necessary to reach equilibrium. This, in turn, would determine how much time is actually available for the shallow-water rocks to develop a karst system through sub-aerial exposure, before renewed subsidence can result in a cap rock that can trap the compaction fluids from the basin. The extent to which volcanic-exhalative solutions may contribute to lead-zinc mineralization, is a matter of extrapolation of the existing genetic theories because direct evidence may be absent. King (1965)discussed four types of stratiform lead-zinc ore types in Australia and believed that they are of a volcanic-exhalative origin. These irregular ore bodies are stratigraphically controlled and originated contemporaneously with the enclosing carbonate rocks which are often of the reef-facies type. Schneider (1964)and Schulz (1964)also accepted a submarine, volcanic-exhalative (i.e., extrusive-sedimentary) origin for the Alpine lead-zinc ores. Snelgrove (1966), who described the lead-zinc stratabound deposits of Ireland, pointed out their paleogeomorphologic control reflected in the ore bodies’ characteristics related to stratigraphy, sedimentation, petrography, structure, and volcanism. In this case, the presence of pyroclastics offer at least some evidence of the possible influence of volcanism in the genesis of the ore. A coincidental association, however, should not be ruled out either. Although R.L. Stanton has been an advocate of the volcanic-exhalative ore-forming process for a long time, he also accepted as more likely the theory that basinal compaction fluids may be responsible for many lead-zinc deposits in carbonate rocks (Stanton, 1972,pp. 551-553). This points to the fact that, although many copper deposits are definitely of volcanicexhalative origin (see Tatsumi and Watanabe, 1971, for example, for details on four subvarieties of the Kuroko-type deposits), this hypothesis need not hold true for any other ore mineralization. In other words, the volcanic mechanism does not offer an easy “band-wagon” theory. If genuine igneourhydrothermal solutions reach the intraformational fluids and/or surface waters and become mixed with them, then a geochemical mechanism must be found that could explain the predominance of the bi-metallic (often even mono-metallic) precipitation in the Mississippi Valley-type ore deposits. In the Economic Geology Symposium on the Mississippi Valley-type ores edited by Brown (1967),the present author has found only an occasional direct reference to “water of compaction”, but numerous researchers discussed syngenetic and diagenetic processes (cf., pp. 12, 18,28, 49, 58, 65, 84-85, 196,217, 349, 364, and 381,in Brown, 1967).The following state-
ORE GENESIS INFLUENCED BY COMPACTION
529
ments made in the symposium suggest that in these particular instances the process of compaction may be given its warranted consideration: (1)the brines were derived from nearby sedimentary basins; (2) the metallic epoch involved coincided with the time of formation of the host rock so that the latter was still unconsolidated at the time of ore precipitation; (3) the sulfide was derived from the marine environment; (4) the derivation of the metal ions was from a lateral, possibly volcanicrexhalative source; (5) the lowtemperature hydrothermal deposits were probably formed from circulating ground water heated in a granitic province; (6)the connate water was heated by an intrusion from which fluorine was derived; (7)precipitation was facilitated by the presence of hydrocarbons in the reef limestones; (8)the source of the metals is to be sought in the shales, which may have been leached by migrating solutions; (9)the metal-bearing brines were connate, similar to oil field brines; and (10) precipitation occurred in rocks overlain by shale acting as a cap rock. It is interesting to note here the suggestion in one paper (Helgeson, 1967)that arkoses can supply sufficient lead from K-feldspar. Amstutz and Bubenicek (1967)presented a sequence of diagrams that summarize their (as well as those of other co-workers) concepts on the origin of several ore deposits, which include the influence of diagenesis and com-
Fig. 5-26. Diagenetic crystallization sequence (paragenesis) in oolitic limestone of the southern Illinois fluorspar district. Fluorite is mostly contemporaneous with the clear calcite cement. (After Amstutz and Bubenicek, 1967, fig. 6.)
530
K.H. WOLF
paction. The paragenetic sequence deduced from many rock samples collected from widely separated localities is schematically presented in Fig. 5-26.The diagenesis has been divided into three stages, namely, the depositional, the early burial, and the pre-metamorphic (or late burial?). The stylolite formation was placed within the period of compaction for many reasons discussed by Park (1962)and Amstutz and Park (1967).In regard to the origin of the fluids that may have formed ore deposits within the sedimentary rocks, Amstutz and co-workers (Amstutz and Park, 1967,and as early as 1963) referred to the work of Illing (1959)and stated that if one considers the diagenetic circulation paths of pore or compaction solutions in reef complexes (Fig. 5-27),then a good explanation is found for the high metal content in the porous flanks and on the crests of the biohermal reefs. Inasmuch as there are four basic theories on the genesis of the Mississippi Valleytype ores, Amstutz used Fig. 5-28to present them conceptually. As shown in Fig. 5-29,numerous criteria used to establish the paragenesis of the sulfide minerals, together with gravity-density or geopetal features, provide strong support for putting the Mississippi Valley deposits into group l a or l b . Bernard (1964)described diagenetic paragenesis in French deposits of the same type. For comparison, the four basic theories on the genesis of the Arkansas barite belt deposits are given in Fig. 5-29.The barite nodules are of early and the matrix is of late origin. Numerous new observations support
Fig. 5-27. Diagrammatic cross-section of the Leduc Reef chain, showing the escape paths of connate water squeezed out of the surrounding compacting Ireton and Duvernay shales. (After Illing, 1959, fig. 4, and Amstutz and Bubenicek, 1967, fig. 8.) Some of the heavy metals may have been extracted “chromatographically” from the hydrocarbons in the connate water, while moving through zones having different pH-Eh values and consequently of different bacterial content.
ORE GENESIS INFLUENCED BY COMPACTION syng ene t i c la Supergene o r e minerals deposited contem-
poraneously in and with t h e sedimepts.the o r e matter.Ba.Pb.Zn Fe.Cu.Ni,Co.S,etc , i s of erosional and thus of supergene origin
I b Hypogene o r e minerals deposited contem. poraneously in and w i t h t h e sediments. t h e o r e m a t t e r is of exhalative-volcanic and thus of hypogene-hydrothermal origin
531
epigeneiic i7a Supergene o r e minerals deposited by
supergene solution and replacement I e . by epigenetic lateral solution. migration and secretion caused by groundwater movements, the source of the o r e m a t t e r is the same o r some adjacent sedimentary bed
a b Hypogene o r e minerals deposited by solutions and replacements i n t h e course of upwards percolations of telemagmatic hydrothermal o r regenerative hydrothermal fluids o r emanations. e i t h e r along faults o r f r a c t u r e s o r t h r o u g h pore spaces along grain boundaries. both f r o m unknown sources
Fig. 5-28. The four basic theories o n the genesis of the Mississippi-Valley-type deposits. Diagenetic fabrics point to a syngenetic mode of formation. (After Amstutz and Bubenicek, 1967, fig. 9.)
the genetic theory l a in Fig. 5-29.In Figs. 5-28and 5-29,the four possible processes have been named: (1)syngeneticaupergene; (2) syngenetic-hypogene; (3)epigenetic-supergene; and (4)epigenetic-hypogene. This is a useful approach which has also been employed in studying diagenesis. What is significant here is that the processes of compaction of both the host and source rocks can be of importance in all four of the above-mentioned processes in: (a) modifying the host rock; (b) supplying metals from the source to the host rocks; and (c) supplying compaction fluids that, when mixed with the metal-bearing solutions, may give rise to precipitation. Amstutz and Bubenicek (1967) showed schematically (Fig. 5-30)the major types of sulphide deposits affected by diagenetic changes. This diagram in effect means that most of these deposits can be influenced by compaction and compaction fluids. In the case of A’-H type of deposits, in volcanic flows, compaction may have little effect on the rigid volcanic rocks themselves, but will influence the surrounding sediments. The compaction fluids,
K.H. WOLF
532
S y n g e n e t i c
As
I '1
At=o
$ I b Volcanic-exhalative Ba and/orS mixed
andprecipitated with normal sediments, thus, exogenous source for Bo and/or 5.
Epig e n e tic
Atz-,
Bas04 exogenous to the sediments, introduced from the Cretac Magnet Cove ring dike.
Fig. 5-29. The four basic theories on the genesis of the Arkansas barite belt. (After Amstutz and Bubenicek, 1967, fig. 1 2 . )
however, can migrate into the volcanic rocks. If pyroclastics are involved, then compaction may have several types of effects, and the compaction processes are related to those of clastic deposits with some differences, as discussed in Chapter 6 on the compaction of volcanic sediments. Compaction fluids may pass into and/or out of the volcanic debris before, during and/or after divitrification of the volcanic glass, for example. The H'-L type of deposits may be affected by compaction and water of compaction, which may have been influenced by volcanic exhalative fluids and gases. In the K-type deposits, the influence of compaction and its solutions may be very prominent also. Park and Amstutz (1968) studied a number of features associated with the southern Illinois fluorspar deposits among which there are compaction features, e.g., stylolites which lead to the reduction of up to 37% of the original thickness of the carbonate rock under consideration. The stylolites formed during the cementation, and the dissolved CaCO, may have been the source of the cementing material. Park (1968) stated that the sedimentary-diagenetic origin of the ores in the carbonate host rock can be established on the basis of: (1) the large-scale geometric features of the ore beds; (2) the relationships between the ore and the enclosing carbonate host rock; (3)the
ORE GENESIS INFLUENCED BY COMPACTION ......
intra-,
533
c
pori-,
volcanic
or
apo-
k rypt 0 or
or
purely
tole
- vo L tan1 c ,
sedlrnentary
Fig. 5-30. Schematic diagram of types of mineral deposits (especially sulfide deposits) which are affected by diagenetic processes. To the left (A'-H) are types of deposits which are clearly associated with volcanic rock, whereas on the right (H'-L) are deposits which may or may not be related to the volcanic-exhalative activities. The K-type of deposits are largely contained in late diagenetic compaction fractures in and near organic reefs. Types A'-F' refer to common types of distribution patterns in lavas, whereas H-L are deposits located within sediments, with which they were formed. (After Amstutz and Bubenicek, 1967, fig. 15.)
close relationship between ores and the primary sedimentary structures; (4) the textural relationships between the fabrics of the ore and carbonate host rock minerals; ( 5 ) geochemical data; and (6) the absence of conduits along which metal-bearing solutions could have moved. The ores within the faults and in certain types of vugs, were formed either during the late diagenetic or epigenetic stages. The metals for the ores were derived either from the redistribution of earlier ore through solution or brought in by saline formation waters. One should remember here that joint and fault systems can be formed through differential compaction, for example, so that the filling of these fractures need not rule out the importance of compaction in supplying fluids to form ores. Boyle and Lynch (1968) visualized the origin of lead-zinc ores in carbonate rocks occurring in three stages: (1)marine organisms initially concentrated heavy metals and sulfur from the sea water (cf. Wolf et al., 1967); (2) an additional concentration by bacterial processes followed as metals were precipitated by HzS; and (3) dissolution by brines and redeposition as a final concentration of the sulfides in permeable zones or in structurally prepared sites. In this theory, there is plenty of room to accommodate the processes of compaction in making a contribution to the genesis of the ores, as will be
534 E
K.H. WOLF W
I
Thickness
Fig. 5-31. Locations of the precipitation of lead, zinc, barium and iron as a function of the facies of the Muschelkalk along the edge of the basin. Facies 1 to 5 have been described in Fig. 5-33 below. (After Bernard and Samama, 1968, fig. 34; courtesy of Sci. Terre. )
discussed below in several cases where a syngenetic origin has been invoked. Bernard and Samama (1968)in their study of the regional distribution of facies and ore minerals, as depicted in Figs. 5-31,5-32,and 5-33,indicated that lithology, paleogeography, and diagenesis, among others, controlled ore genesis. The distribution of their ore cements (Fig. 5-32)and the zoning of the Pb, Zn, Ba, and Fe (Figs. 5-31and 5-32),both imply that the ore-forming fluids changed in composition. It may well reflect: (1)changes in the chemistry of compaction fluids as compaction progresses and, possibly, due to the clay mineral transformations; and (2)mixing of fluids. Roedder et al. (1968) discussed the genetic problems of the Mississippi Valley-type Mex-Tex deposits in New Mexico and reached the following conclusions based on fluid inclusion investigation. The ore-forming solutions were moderately to strongly saline brines, probably with a high chloride
535
ORE GENESIS INFLUENCED BY COMPACTION
Cements
(Bo.Sr)SO4
CaCO,.MgCO,
+CoSo,
'
Fe203 ~
MnO
log Ba ~
log S r
Fig. 5-32. Diagram showing lateral variations of the clastic sediments, types of cements, and of some chemical constituents indicating mechanical-chemical-mineralogic, and chemical-elemental differentiation, respectively, in the red-bed type of sedimentary host rock of the ores. (After Bernard and Samama, 1968, fig. 27; courtesy of Sci. Terre.)
content, and had temperatures of approximately 200"C, except for the last stages of precipitation when temperature was lower. Suspended organic matter was present also, probably mobilized from the sediments. They concluded that the mixing of near-surface meteoric water with deeper brines seems precluded by the uniform, above-normal geothermal gradient and by the high salinity. Two possibilities mentioned by Roedder et al. are that: (1) either the earlier fluids were flushed out by new, more saline solutions; or (2) the presence of fluids of high salinity represents progressive dewatering
K.H.WOLF
536 E
W
I
Zone of restricted circulation
Dilution zone
Zone of alternate dilution and Depth saturation
Fig. 5-33.Theoretical division (a) and interpretation (b) of Muschelkalk facies. 1 = reduced facies with traces of crystalline salt; 2 = varicolored facies; 3 = green facies; 4 = black facies; 5 = facies of anhydrite and dolomite; 6 = location of the middle carbonate unit; 7 = synsedimentary fault; 8 = top of the Buntsandstein. The environmental interpretation (b) is based on the classical scheme offered by Sloas (1953) and corresponds to the theoretical facies reconstruction given in (a). (After Bernard and Samama, 1968, fig. 33; courtesy of Sci. Terre.)
by compaction of clay-rich units in the sedimentary basin (oral communication by P.M. Bethke to Roedder et al.). The work of Jackson and Beales (1967)and Beales and Jackson (1968) relating to the great Pine Point lead-zinc ore field in northwestern Canada is noteworthy for its repeated reference to the role of compaction in the ore-forming process. The host rocks 'are the Pine Point and Presqu'ile dolostones of Middle Devonian age that seem to have originated as part of a reef complex marginal to a major cratonic basin (Fig. 5-34). The ore is controlled by the paleoecology of the reef as indicated by Fig. 5-35. This complex formed a barrier along an axis of reduced subsidence, possibly controlled by
ORE GENESIS INFLUENCED BY COMPACTION
537
basement faulting. As burial proceeded, the barrier complex plunged in west-southwest direction in the subsurface toward the geosyncline as a narrow prism of highly-permeable and porous sediments; the porosity is of the intra- and interskeletal type, modified by diagenesis and solution of syngenetic evaporites. This conduit facilitated the escape of the basinal compaction fluids from the west through the barrier. The metals are believed to have been brought in from the basin in a chloride brine. Polysulphide complexing permitted remobilization and reprecipitation in the ore zone. Gas and oil also may have invaded the carbonate host rocks (see also Hitchon, 1971),as evidenced by cavity-filling bitumens associated with some ore bodies. If the subsurface rocks were contaminated with anaerobic bacteria, H $3 formed and accumulated upon mixing of organic matter and sulphate solutions. Thus, sulphides of lead and zinc precipitated to form the ore.
Fig. 5-34. Location map illustrating the position of Pine Point and the paleogeography of western Canada during deposition of the Middle Devonian Presqu'ile Barrier Reef Complex. (After Jackson and Folinsbee, 1969, fig. 1, modified from Grayston e t al., 1964; courtesy o f Econ. Geol. )
K.H. WOLF
FORE REEF SHALE AND LIMESTONE
L AGOONAL Gdnekic subdivisions
of Pine Point and Presbu’ile Formations
Fig. 5-35. The ore bodies at Pine Point occur in the Presqu’ile Dolostone and its interface with the Pine Point Formation, and are controlled in part by the paleoecology of the barrier complex. (After Jackson and Folinsbee, 1969, fig. 2; courtesy of Eeon. Geol.)
Jackson and Folinsbee (1969)in their geologic summary of the Pine Point deposits, stated that the evaporite solution breccias and lattice work stmctures in the Presqu’ile dolostones can be examined in open pits and it is clear that brecciation and dissolution played a role in the porosity development of the dolostone. Although the ore precipitation was not confined to one particular facies of the reef complex, most of the ores are confined to the Presqu’ile. Initial sedimentary trends and solution-prone strata, particularly when affected by faulting, fracturing and karsting, seem to have been the fundamental controls in the origin of porosity trends and breccia zones which at least to some degree controlled ore localization. In the removal of the evaporite by solution, some type of compaction, most likely “unit-overunit” instead of “intergranular” overburden compaction (see Table 5-I), was involved. Billings et al. (1969)supported the hypothesis that the Pine Point ores were formed by expelled formation waters of a sedimentary basin of deep origin that mixed with shallow, H2S-rich fluids. As Fig. 5-36indicates, the fluids may have come from at least two sources, namely, from the Elk Point evaporite basin as well as from the Mackenzie shale basin to the north. To what degree mixing of these two fluids may have been important in ore genesis, remains to be shown by future studies. Also significant would be the
539
ORE GENESIS INFLUENCED BY COMPACTION
MACKENZIE BASIN
B.c. CAFGARY
‘
REGINA
,\
WINNIPEG
--_- - - - 1 _ _ _ _ ~ _ _ _ - - - - - - - -_ - ALTA
0 I
’
I
L
I
500 Km ,
I I
SASK.
I
MAN.-*-
-1
U N I T E D STATES Carto. Sect..
Sch. of Geol., I S U
Fig. 5-36. Devonian paleogeographic map showing present location of Pine Point ore and approximate location of formation water samples. (Modified from Jackson and Beales, 1967, and Billings et al., 1969, fig. 1; courtesy of Bull. Can. Pet. Geol., and Econ. Geol.)
determination of the time when karsting of the host rock took place versus the time when the ore-forming solutions migrated into the reservoir unit. The direction of movement proposed by Billings et al. differs from the earlier suggestion made by Jackson and Beales (1967). The brines leached the metals (Zn, Cu and, probably, Mn and Fe) from the clay-rich basinal sediments as the adsorbed ions on clay minerals and organic matter can be released relatively easy. The material balance calculations made by Billings et al. indicated that sufficient volumes of leachable metals could have been present in the shales to account for the ores. Davidson (1965) discussed the observations made by Lebedev (1967a) that are relevant to the work of Beales and Jackson and other researchers (see below) on the Pine Point ore deposits. Lebedev reported on hot saline solutions tapped by drill holes in the petroleum province of the Cheleken peninsula (Caspian Sea). These solutions present some of the best examples of “telethermal” ore-forming fluids which could supply 300 tons of Pb per year. The boreholes are cutting an anticlinal structure in Pliocene sediments which overly 2800 m of Miocene red beds, the latter containing an extensive evaporite suite. The geothermal gradient is high and the deep formation
540
K.H.WOLF
fluids reach the surface at temperatures up to 80°C. Yet, no young igneous rocks are known within several hundred kilometers. The hot metal-bearing solutions (essentially Na-Ca chlorides) come mainly from sandy horizons in the red beds with the following composition (in percentage equivalents) : 75% Na+, 22% Ca2+, 3% Mg2+, 99.9% C1-, 0.05% and 0.05% HC03-. The Pb content averages 10 to 15.0 mg/l and Zn 0.19 to 5.35 mg/l. The high content of Br and I (i.e., 450 and 24 mgfl, respectively) and the local geologic relationships strongly indicate that the brines were derived from the adjacent evaporites. Within two years, the drill hole casings were coated by Ca- and Mg-carbonates and native Pb, the latter up to 1.5 tons in weight in one well and up to 2.5 tons in another. The lead precipitate contains 0.60%As, 0.49%Cu, 0.08% Ag, and 0.02% Sb. The metal-bearing thermal solutions of Cheleken, as Davidson (1965) mentioned, are very similar to the brines of the Salton Sea and the Red Sea deeps. The latter two appear to have derived their composition through diagenesis or solution of evaporites and leaching of the country rocks, so well described in several publications (e.g., White et al., 1963; White, 1968; Degens and Ross, 1969; and White, 1971). It should be noted that, although in most instances where the geothermal areas have been suggested as possible models of ore-forming systems, little or no references have been made to compaction. Hence, future research must attempt to reconcile certain differences found from one thermal area to the next, i.e., similarities as well as differences must be established between, for example, the Salton Sea, Red Sea, and the Cheleken area, on one hand, and the characteristics of other localities, on the other. A preliminary examination may show, for instance, that the Salton Sea and Red Sea hydrothermal fluids are mobilized by hot igneous rock near the upper surface of the earth (e.g., White, 1971), whereas the Cheleken thermal solutions are heated by the high geothermal gradient, not necessarily the result of shallow igneous activity. It may be possible t o establish different genetic models (or submodels), therefore, of which the Cheleken model may be more analogous to the one proposed by Jackson and Beales (1967) than the model applicable to the other two localities. Compaction may definitely be of importance in one case, but may be of subordinate significance (if not entirely absent) in the other. The Pine Point ores may have been formed very early in the geologic history of the sedimentary basin, so that it is appropriate to propose that compaction fluids were the metal-bearing solutions. These fluids may also have been the medium of transportation for hydrocarbons. They were neither derived from an igneous source, nor were subsurface waters heated by hot igneous rocks. In the case of the Pine-Point-type of ores (as well as the Cheleken ores), therefore, it is appropriate to draw an analogy between oil and ore evolution in the sedimentary basin, whereas this is less applicable
ORE GENESIS INFLUENCED BY COMPACTION
541
to the Salton Sea model. (See publications listed in the earlier section on the relationships between the origin of hydrocarbons and ores, pp. 490-513.) Fritz (1969) demonstrated that the secondary changes present in the carbonates in which the Pine Point lead-zinc ores occur, are the result of a TEMPERATUl RANGE
PAR-
25
REEF -DOLOMITE
- - - - - - - - --
M
FLUIDS
75 I
--
WHITE VEIN DOLOMITE
IDlOMORffllC D0LOMITE
I SULFIDE MINERALS
CALCITE
intapmm with s%d
nimsmk
I
CALCITE w=%tr3.iwid,
CALCITE
boe cwldr m v
w
--- -- -- - --_ CALCITE
recent freshwater deposits
Fig. 5-37. The cycle of mineralization at Pine Point and factors influencing the deposition of the carbonate gangue. (After Fritz, 1969, fig. 4; courtesy of Econ. Geol.)
542
K.H.WOLF
cycle of mineralization, as schematically shown in Fig. 5-37. His geochemical efforts concentrated on the calcite and dolomite, and he concluded that although the origin of the mineralizing fluids is not known, they are very likely comparable to formation waters of oil fields. As shown in Fig. 5-37, Fritz believed that at least four types of fluids have been involved, namely, sea water, pore water, hydrothermal fluids, and percolating surface waters. Theoretically, there may be gradations between these fluids in that trapped sea water may become pore water, which in turn when deeply buried (or heated and remobilized by heat given off from deep-seated magmas) can change into the so-called hydrothermal fluid. Thus, the process of compaction may be influential in the movements of both pore and hydrothermal solutions. The mode or style of compaction involved (see Table 5-1) will change with increasing age and evolution of a sedimentary basin. Particularly noteworthy is the information provided by Fritz (1969) and Fritz and Jackson (1972) on the origin of the dolomites. They suggested that a genetic relationship exists between the dolomitization of the original sedimentary limestone and the mineralization in several Mississippi Valley-type ore districts. Their isotope data suggested four periods of dolomitization which were the result of: (1)capillary concentration and refluxion; (2) cannibalization; and (3) replacement and neoformation (?). As an extension of an earlier work (i.e., Fritz, 1969), Fritz and Jackson (1972) stated that brines (modified sea water or modified fresh water) and hydrothermal fluids were the respective solutions responsible for the dolomitization. From all the geological and geochemical data available at present it seems that formation waters (some of which no doubt were compaction fluids) rather than igneous hydrothermal solutions were involved. Further details on the composition, origin, and the direction of flow of the fluids responsible for the precipitation of the Pine Point ores, as well as on the significance of aquifers and fault zones as conduits, have been presented by Roedder (1968), Evans et al. (1968), Sasaki and Krouse (1969), Billings et al. (1969), Hitchon (1971), and Kesler et al. (1972). The existence of several types or sub-types of Mississippi Valley-type ores is supported by the following findings. Sangster (1970) described two types of lead-zinc deposits in carbonate rocks in Canada: (a) the Pine Point ore field, which he mentioned as a representative of the Mississippi Valley-type ores; and (b) the Remac-type ore of the Kootenay arc in southeastern British Columbia, which is now a successful lead-zinc producer from Middle Cambrian rocks. The mineralization of these two types of deposits is quite different. The host rock of the Remac type is a well-banded, fine-grained dark dolomite containing a well-banded, fine-grained ore. The ore is very similar to the so-called volcanic-exhalative type such as that of the Bathurst area, New Brunswick, Canada. In contrast, the Mississippi Valley-type ore occurs
ORE GENESIS INFLUENCED BY COMPACTION
543
as open-space fillings. Whereas the latter was formed in shallow-water carbonates, adjacent t o a rapid shale pinch-out of the carbonate sequence, the Remac deposit formed well away from the basinal margins, possibly in an environment of somewhat deeper, quieter water. The dolomite host rock is situated near the center of a shallow basin. Rapid facies changes, such as those present in the Mississippi Valley-type ore districts, are absent in the Remac ore host sediments. The differences in genesis of these two types of lead-zinc deposits, described above, are envisaged by Sangster (1970) as follows. In contrast to the Mississippi Valley-ore host rock of a well-aerated, shallow-water, reef environment, the Remac-type carbonates appear to have originated in a deeperwater, more restricted and less turbulent, euxinic milieu which gave rise to the dark-layered, graphite-bearing carbonates and local ribbon cherts. Under these geochemically reducing conditions, sulphides precipitated simultaneously with the carbonate host sediments more or less syngenetically andfor formed during the diagenetic stage. The HzS which was generated by bacteria would not be continually flushed out in such a quiet-water environment but instead would accumulate t o form a local HzS-saturated microenvironment on the sea floor. Any metal-bearing solutions finding their way into this environment would immediately give rise t o the precipitation of sulphides as a layer on the sea floor, or these solutions may become interstitial fluids within the sediments from which the ore mineral could precipitate subsequently. Hence, the Remac-type ores differ in several important aspects from the Mississippi Valley-type ores, that will also determine to which extent compaction contributed t o the origin of ore in each case. The mode of origin (i.e., syngenetic-diagenetic vs. epigenetic) and the paragenetic relationships, as well as the stratigraphic, textural and structural relationships between the host rock and the ores, exhibit the most obvious differences. The Remac-type mineralization would be more related to that of Mount Isa or Sullivan-type ores. Inasmuch as many ore classification schemes are based on the type of host rock, the following should be considered. As Sangster (1970) pointed out, the association of the Remac ores with the carbonates may be fortuitous because deep-water carbonates, though not uncommon, are not as widespread as those of shallow-water origin. Also, Remac-type ores are rare. The Remac milieu may merely represent a clay-mineral-starved environment in which carbonates accumulated relatively uncontaminated, whereas the more common sedimentary situation elsewhere allows the predominant deposition of clays. Also, in contrast to shales, the carbonates may be less susceptible to compaction (cf. Chapter 3 on compaction of limestones in Vol. I of this book) because of early lithification. This would allow metal-bearing connate solutions to remain in situ longer than in normal clay-rich sediments, thus permitting more time for the precipitation of me-
544
K.H.WOLF
tallic sulphides. It is important to recall here the slide mechanism proposed by Mountjoy et al. (1972), and the origin of carbonate turbidity currents described in the literature, that can move shallow-water limestones into deeper basinal environments. Should metal-bearing compaction fluids from the surrounding muds penetrate into the porous limestone, the resulting ore deposit may exhibit the Remac-type mineralization. Violo (1969) performed experiments on the “metasomatism” of limestones by galena under temperatures ranging from 150°C to 300”C, which caused a recrystallization of the limestone and a cloudy diffusion of galena. Other experiments were undertaken on: (a) the precipitation of various combinations of galena and sphalerite in the presence and absence of calcite at lower temperatures (i.e., 50°C) with variations in pressure, and (b) the precipitation of galena in sandy sediments (see also Helgeson, 1967). Although his results are of wider interest in ore genesis, only some remarks on Violo’s work that may be related to compaction under natural conditions are presented here. For example, Violo found that pressure has a very important influence on sulphide texture, even at low temperature (50°C). Even moderate hydrostatic pressure (100-160 atm) in the presence of water can cause recrystallization and structural modification, in contrast to their absence if conditions remain at 50°C and 1atm. This suggests that compactional pressures prior to complete expulsion of the intraformational fluids may cause secondary “diagenetic” modifications of both the metalliferous and host rock minerals. Violo also found that galena is more sensitive to pressure variations than sphalerite in the presence of a solvent and at low temperature. Textural variations, therefore, could be expected in ores that have undergone compaction. In his experiments on the remobilization of galena at variable pressures and at constant temperature (50-60” C), Violo found that rhythmical, chemical precipitates of galena and calcite formed at pH of 3-5 and 8-9, respectively. When subjected to a pressure of 180 atm at constant temperature over 240-280h, the laminae became folded and undulated. These structures may resemble those formed by diagenetic compaction, as pointed out by Violo himself. On the other hand, in laminated sands-plus-galena deposits, no such folding occurred but, instead, the galena underwent strong diffusion from the laminae into the whole sandy host sediment. The above differences in diagenesis in limy versus sandy sediments suggests that chemical and mechanical compactional diagenesis under natural geological conditions may result in varying textural and structural effects. Zimmermann (1970) described geopetal features, rhythmic banding, and compaction structures of syngenetic--diagenetic origin from the Meggen
ORE GENESIS INFLUENCED BY COMPACTION
545
barite-pyrite-sphalerite deposits and compared them with the Arkansas barite accumulations. In 1969, Zimmermann described sag structures and geopetal and compaction features in and around ore layers of the Mississippi Valley-type lead-zinc district in Wisconsin which suggested a syngenetic origin for the barite and associated minerals. On the other hand, Brown (1970)reviewed the main genetic problems and concluded that the geopetal features, paragenesis, stylolites, etc., proposed by the Amstutz school as evidence of a syngenetic-diagenetic origin are not sufficient. Brown mentioned that it has not been emphasized sufficiently that the mixing of unlike solutions, for example, igneous hydrothermal and compaction, has been a factor of considerable importance in the epigenetic precipitation of ore minerals and that their textures and fabrics could be similar to those found in ores of diagenetic origin. Future research will have to reconcile these differences in opinions. Amstutz and Park (1971)discussed the shape and size, or fabric, of ore minerals, the compositional gradients, and the time relations (= paragenesis) of ores of diagenetic origin, and pointed out that the metamorphism of ores and their surrounding host rocks can only be correctly interpreted when the original diagenetic features are known from unmetamorphosed examples. They described two diagenetic situations and fabrics, i.e., one common in carbonate and the other in sandstone host rocks, both of which are quite universal in occurrence. Similar observations have been made in at least fifty important Cu-Pb-Zn-Co-Ni and pyrite-marcasite deposits. Figure 5-38 is an example of the approach used by Amstutz et al. (1964)and Park (1968), illustrating the paragenetic sequence of the pyrite-sphalerite-galena-fluorite in oolitic carbonate rocks from the Cave-In-Rock fluorite district, Illinois. This sequence occurs in numerous other geographical types of ore deposits with perfect agreement in general, but with some interesting deviations that enables one to establish sub-types. The preliminary work of Amstutz et al. (1964)suggests delicate paragenetic criteria for distinguishing between different types of sedimentary and diagenetic ore-forming environments. Lateral and vertical zonation may have been caused by variations in intrastratal fluid chemistry during diagenesis and not only by differences in fractional crystallization. In considering the influence of compaction on ore genesis, it is significant to note that Fig. 5-38 shows pre-, shallow-, and deep-burial stages; at the bottom, the porosity values assigned to the shale and limestone correspond to these burial stages, thus reflecting degree of compaction and cementation. Inasmuch as the compaction fluids could have been responsible for the precipitation of the ore minerals and there appears to have been a similar paragenesis in the deposits at different localities, it is suggested that the development and evolution of interstitial and compaction fluids are remarkably similar in different sedimentary basins.
K.H.WOLF
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Hill and Wedow (1971) studied collapse breccias in the East Tennessee zinc district as indicated by compaction and porosity features of the Lower Ordovician Mascot Dolomite and Kingsport Formation. They also determined whether or not the latter was an aquifer in the early Middle Ordovician time and whether or not it was similar t o the Tertiary limestone aquifer system existing in Florida. The present physical characteristics of the carbonate rocks are due t o deep burial under thousands of feet of Middle Ordovician and Later Paleozoic rocks. The bulk density, porosity and permeability of the Lower Ordovician rocks, therefore, were different from those existing
ORE GENESIS INFLUENCED BY COMPACTION
547
today as a response to later compaction. The investigation showed that quantitative estimates can be made of the relative compaction or reduction in thickness undergone before and after the collapse of brecciated zones that occurred in a Lower Middle Ordovician karst terrain. The karst and brecciation features are important because they served as loci for ore mineralization. In determining compaction, two sets of measurements were taken by Hill and Wedow (1971). (1)One set of observations permitted the estimation of the total compaction since the rock was still in a soft, unlithified state up to the time of maximum depth of burial beneath thousands to tens of thousands of feet of overlying Paleozoic sedimentary rocks. (2) The other set of mea-
0.
1
2
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Fig. 5-39, Sketches E rutlines of compacted desiccation crack fillings in representative slabbed drill-core samples of the G-430/434 bed of the Mascot Dolomite, Mascot-Jefferson City district, Tennessee. Stippled pattern = dark-colored dolomite; no pattern = light-colored dolomite. Horizontal lines through vertical sections indicate positions of the semi-circular cross sections shown at top of diagram. Vertical and horizontal scales are equal. (After Hill and Wedow, 1971, fig. 2; courtesy of Econ. Geol. )
548
K.H. WOLF
surements allowed determination of the amount of compaction and thinning of the Lower Ordovician strata that occurred after collapse brecciation up to the time of maximum depth of burial. These measurements were then compared with the results of laboratory experiments on the relationship between compaction and absolute pressure or equivalent absolute depth of burial, thus determining the approximate depth of burial of the Ordovician rocks. Some of the details are considered here. Estimates of the total reduction in thickness were based on measurements of highly contorted desiccation-crack fillings (Fig. 5-39)in fine-grained dolomite beds. That they were mud cracks or desiccation cracks is shown by the linear outline in cross sections in the plane of the bedding (Fig. 5-40) and by the polygonal mud-crack pattern on bedding surfaces. According to Hill and Wedow (1971),the original cracks probably formed in an unlithified carbonate mud with approximately 40% or more water. With further accumulation
Fig. 5-40. Comparison of cross-section of compacted mud-crack filling with model of crack and filling before compaction. Dense stippled pattern = dark-colored dolomite; open stippled pattern = dark-colored dolomite mud; no pattern = light-colored dolomite or dolomitic mud. Weighted average compaction indicated by model is 38.3%. (After Hill and Wedow, 1971, fig. 3; courtesy of Econ. Geot.)
ORE GENESIS INFLUENCED BY COMPACTION
549
of sediments, the muds were progressively compacted, the interstitial water was pressed out, and the porosity was reduced. The cracks were contorted as shown in Fig. 5-40.The original walls of the cracks can be assumed to have been essentially vertical. Through calculations, these two investigators have determined the most probable dimensions of the triangle representing the cracks. After contortion of the crack, the natural surface represents in crosssection the side of the triangle which retains its original length. Hill and Wedow (1971) described a technique for measuring the bounding surfaces and, thus, the original depths of the cracks. The length of that line representing the surface minus the thickness represented by the particular sample (corrected for dip of bedding, if necessary) is the reduction in thickness caused by compaction. The results showed a 43.1% relative compaction value with a standard deviation of 11.8%, with some regional variations that may reflect thinning of the overburden and, therefore, lower degree of compaction. Estimation of post-breccia compaction can be made from the thickness of certain marker beds that are thicker in the detached blocks of the collapse breccias than in the surrounding undisturbed rock (Fig. 5-41),as the rock was less compacted at the time of brecciation than it has been at any time since. It was assumed by Hill and Wedow that the stresses of the post-breccia loading during compaction was taken up largely by the breccia matrix rather than by the fragments and blocks themselves, so that the thicknesses of the markers at the time of brecciation were preserved and are, therefore, greater in the breccia fragments than in the surrounding rock. Also, the less compacted rock of the breccia fragments should exhibit a lower bulk density and a greater intergranular porosity in contrast to the more compacted rocks of the unbrecciated wall-rock. It is assumed, of course, that, first, no significant amount of cement has been introduced to reduce porosity and, secondly, that no minerals having high specific gravity were precipitated after brecciation. Figure 542 demonstrates that the bulk density of the rock increases as the thickness decreases due to compaction. The plots for the wall-rock and breccia sample groups, although slightly overlapping, are characteristically different. A similar diagram showing the correlation between thickness and porosity also shows a direct relationship between the two. Thus, Hill and Wedow concluded that the beds were once thicker, more porous and probably more permeable in the geologic past. The post-brecciation compaction was estimated by them to average about 29.5% of the thickness at the time of brecciation. Crushed, contorted sphalerite veinlets, which were originally straight and were formed during early Middle Ordovician time, were also used to measure compaction. The deformation was the result of compaction by ever-increasing load of younger rocks. The data indicated that 21.5% to
K.H. WOLF
550
Fig. 5-41. Diagram showing difference in thickness of a marker bed in breccia and wall rock, if brecciation took place before complete compaction of strata. Stippled area = breccia matrix; diagonal ruling = marker bed. Diagram indicates that marker bed is about 25% thinner in wall rock than in breccia block. Scale is relative. (After Hill and Wedow, 1971,fig. 4;courtesy of Econ. Geol.)
31.2%compaction of the carbonate rocks has occurred (Fig. 5-43). As mentioned earlier, one set of the measurements gave the total amount of compaction that has taken place since the deposition of the muddy sedi-
WALL ROCK
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Fig. 5-42.Scatter diagram showing relation of thickness of bed to bulk density of rock, Trio marker, Kingsport Formation, Flat Gap Mine, Copper Ridge district. Small solid circles = wall rock samples with bivariate-distribution field shown by diagonal ruling; open squares = samples from breccia blocks with bivariate-distribution field shown by stippling. (After Hill and Wedow, 1971, fig. 5; courtesy of Econ. Geol.)
ORE GENESIS INFLUENCED BY COMPACTION
BEFORE
COMPACTION
551
NOW
Fig. 5-43. Diagram illustrating amount of compaction shown by crushed sphalerite veinlet, Felknor mine, Jefferson County. No pattern = sphalerite veinlet; coarse stippled pattern = carbonate rock before compaction; dense stippled pattern = present-day light-colored dolomite with streaks and splotches of black asphaltic residue. Dashed horizontal lines are bedding planes in the dolomite, but d o not represent actual lithologic variations. (After Hill and Wedow, 1971, fig. 6;courtesy of Econ. Geol.)
0
28
100
Fig. 5-44. Diagram showing relative thicknesses of Early Ordovician clay- to silt-size sediments when (A) mudcracked shortly after deposition, (B) when brecciated in early Middle Ordovician time, and (C) at the present time. (After Hill and Wedow, 1971,fig. 7;courtesy of Econ. Geol.)
K.H. WOLF
552
ments, whereas the second set gave the measurements of the amount of compaction after breccias were formed. If the initial phase of sphalerite emplacement was essentially contemporaneous with the brecciation, then the second set of data is also a measure of post-mineralization compaction. he-breccia compaction is determined from the relationship between these two sets of data, as shown in Fig. 5-44. The average total compaction is approximately 43%, about half of which occurred before and about half after brecciation. Laboratory compaction tests were made on carbonate sediments by Fruth et al. (1966)and Robertson (1967)(see also Chapter 3, and Chapter 3 in Vol. I.) According to Fig. 5-45, the data by Hill and Wedow indicate that maximum depth of burial was probably between 2,000 and 10,000 ft, whereas depth of burial at the time of brecciation was less than 1,000 ft. The original intergranular porosity must have been approximately 35% at depths less than 11,000 f t (Fig. 5-45) and was similar to those of recentlydeposited chalk beds. Hill and Wedow concluded that the collapse breccias, which acted as host rocks for the ore mineralization, originated during early Middle Ordovician time when the aquifer zone was situated at a depth of no more than 1,000 ft. If the breccias had formed later, the rocks involved would show less post-brecciation compaction. If both the brecciation and ore precipitation occurred at such a relatively shallow depth, then
I
10
-
1
17----
100
1,000 "DEPTH" ( f t . )
10,000
Fig. 5-45. Graph showing laboratory compaction-"depth" and porosity-"depth" curves (adopted from data in Fruth et al., 1966; see also Chapter 3 on compaction of carbonates in Vol. I). Stippled zones indicate general ranges of bulk of East Tennessee compaction data presented by the authors (After Hill and Wedow, 1971, fig. 8 ; courtesy of Econ. Geol.)
ORE GENESIS INFLUENCED BY COMPACTION
553
the intergranular porosity and permeability must have been at least an order of magnitude greater than they are at present (i.e., after subsequent compaction). Thus, the geologic conditions during the Ordovician were shown to be similar to those now existing in the Tertiary limestone aquifer in the Coastal Plain region of the southeastern United States, and the latter can serve as a model available to further study. The use of isotopes in determining the source of the ore mineral components and, therefore, of the compaction fluids, if they were responsible for the transportation and precipitation of the chemical elements, was shown by Doe and Delevaux (1972).Their lead-isotope data suggested that most of the Pb in the galena ore within Paleozoic host rock of the southeast Missouri district probably was derived from a sandstone aquifer (Helgeson, 1967; Strakhov, 1969, 1970), i.e., the Lamotte Formation. Inasmuch as approximately 2 p.p.m. Pb can be easily leached from the sandstone, a large supply is available from this unit that covered 11,000 sq. miles with an assumed thickness of 200 ft, as based on calculations from the known stratigraphy. One ton of Pb would come from a block about 100 X 100 X 700 ft in size (about 5 * cubic miles), i.e., about 20,000 tons per cubic mile. The total amount of Pb available, therefore, would have been 20,000 tons from 25 sq. miles of the Lamotte Sandstone. Nine million tons of Pb then would require a leaching area of about 11,250 sq. miles (still assuming a sandstone thickness of 200 ft). All the geologic and geochemical considerations given by Doe and Delevaux (1972)supported their conclusion that the Precambrian basement with its siliceous felsite, the dolomites and limestones of the Bonneterre Formation (being the host rock of the ore), and the calcareous shale of the Davis Formation, cannot isotopically qualify to have provided more than a fraction of the Pb in the galena. Although no reference was made to compaction fluids as the metal-bearing solutions, this possibility can be excluded only with further work. Isotope data can provide information on the most likely source of the metals, independent of the mechanisms of transportation and precipitation of the ions. One should note here that if arkoses (Helgeson, 1967) supply the lead and other elements to intrastratal fluids, the latter can mineralize either sandstones (see Strakhov, 1969,1970, for example) or carbonate rocks, depending on the local geology. Boyle (1972)undertook investigations of the source of the chemical elements of some ore deposits. He concluded that much later after the accumulation of the sediments rich in chemical elements, during and after the period of severe faulting, brines pervaded the rocks and leached out Cog- anion and various elements, the most important of which were Ba, Sr, Ca, Mn, S, As, Cu, Pb, Zn, and Ag. The brines may have been of connate origin and the salt could have come from the evaporites associated with the sedimentary pile in which the ore is located. The leaching went on for a very long time,
554
K.H. WOLF
well into the Tertiary when the heat wave associated with the underlying volcanic action may have had a particularly stimulating effect on the dissolution processes. Although the process described by Boyle is not directly related to the present topic of compaction, the mere fact that he presupposes a sedimentary pile that is rich in chemical elements, as many have done before him, is of both theoretical and practical importance in the investigation of compaction. Eventually, an answer must be found to the following question: what are the precise geological, geochemical and hydrological conditions that in some cases allow very early, diagenetic removal of chemical elements by water of compaction, which after movement t o a suitable host rock can form ore deposits, in contrast to the milieu that permits the water of compaction to move out of the sedimentary pile without leaching the chemical constituents? In the latter instance, the elements would be available millions of years later for remobilization by a number of “epigenetic” processes. The latter case seems t o apply to Boyle’s study. Based on extensive laboratory experiments, G.V. Chilingarian (1973, personal communication) suggested that in case of highly hydratable clays, such as montmorillonite, large volumes of oriented water are available to do the job, whereas in the case of nonhydratable clays, e.g., kaolinite, there is insufficient amount of oriented and adsorbed water to push the elements out. Stanton (1972), in his chapter on the limestone-lead-zinc association, discussed the various genetic theories proposed for the Mississippi Valleytype ores and referred t o the concept of compaction fluids being the metalbearing solutions as “a very appealing explanation - and one that looks as if it may be close to the truth for many occurrences”. He also pointed out that Newhouse (1932) already has made this suggestion and that this mechanism has been recently supported by the research on fluid inclusions (Roedder, 1967, 1972) and by the work of Jackson and Beales (1967) from the point of view of petroleum geology. Newhouse (1932) concluded that descending meteoric waters could not have been the sodium chloride-rich and metalbearing fluids in the formation of the Mississippi Valley-type ores. The only known available sources for such concentrated solutions (outside the salt beds) are magmatic and, possibly, the first artesian flow (connate waters) from newly-tapped beds. According to Newhouse, if the ores are ever formed entirely by the connate solutions, they were formed over a very restricted geological time in the structural and physiographic history of a region, when the first flow of connate fluids took place, i.e., when compaction of the sedimentary-volcanic pile of sediments occurred. Roberts (1973) in his research of the Woodcutters deposit in Australia came to the conclusion that a combination of both syngenetic-diagenetic and epigenetic processes would explain the origin of the lead-zinc ore. He used various criteria t o indicate that the host rock underwent no more than
ORE GENESIS INFLUENCED BY COMPACTION
555
moderate burial metamorphism up to approximately 160°C, but not more than 200°C. Roberts resorted to five genetic stages, which are similar to those proposed by Jackson and Beales (1967): (1)concentration of the metals in the basin water; (2) fixation of the metals in the sediments; (3) transportation of the metal solutions; (4) precipitation of metallic sulphides; (5) later, tectonic remobilization. Considering some of the details, Roberts proposed a syngenetic-diagenetic chemical co-precipitation of the Pb and Zn, the two elements being derived from dolomite precursors (i.e., aragonite and/or Mg-calcite). This was followed by decomposition of large amounts of organic material which was accompanied by a release of metal ions into pore solutions. The ions were complexed during dolomitization, and precipitation of metal sulphide was prevented. From our present state of knowledge (Owen, 1964) (also see Chapter 3 in this book and Chapter 3 in Vol. I) it is certain that, in general, carbonate sediments become cemented or lithified much earlier than terrigenous sediments, so that the former should undergo less diagenetic compaction than the latter. Nevertheless, some minor early compaction did occur in the carbonate sediments investigated by Roberts as indicated by the deformation of diagenetic veins and sedimentary laminae. Under the condition of early cementation and minor compaction, the intrastratal fluids were retained within the unit. According to Hunt (1967), this early lithification is accompanied by the hydrolysis and solubilization of the proteinaceous matter which they contain. Abelson (1957, referring to unpublished work of Conway) reported that thermal degradation of 63% of alanine at 80"C would require lo6 years, whereas at 60°C degradation would need lo8 years. The presence of oxygen also shortens the lifetime of alanine. Based on the studies of Owen (1964) and Conway, on the carbonate lithification and stability of alanine, respectively, it can be assumed that pore fluids containing organo-metallic complexes would be preserved even under conditions of deep burial in sedimentary basins. In the case of the Woodcutters ore genesis, the fluids were expelled when tectonic folding took place after lithification. Tension fractures were produced in the anticlinal crests, whereas compression of the rock mass occurred in the deeper parts of the fold. The resulting pressure gradient allowed the pore solutions to move along fractures and solution cavities into the tension cracks (Owen, 1964), where the sulphide ore minerals were deposited due to the organo-metallic complexes becoming unstable as a result of C02 and NH3 gas release from the solution. Sulfide pressure shadows on diagenetically formed pyrite crystals indicate that the fluid movements have occurred during the post-lithification stage. One might add here that fractures and faults, which formed diugenetically in lithified limestones and dolomites, could receive and trap fluids, such as those described above, from which ore minerals can precipitate out. Consequently, there is no requirement for tectonic fissures to be
K.H.WOLF
556
formed much later in the geologic history of a limestone complex to offer open voids for ore mineralization. Copper deposits in sedimentary rocks Both Knight (1957) and Noble (1963) in their papers on the source bed concept and the formation of ores by compaction fluids, mentioned several copper deposits as examples to illustrate their discussions. The following ore districts were cited by them: Northern Rhodesian copperbelt; Katanga copperbelt, Belgian Congo; Kupferschiefer of Germany; Mount Isa, Australia; and red-bed copper deposits of the southwestern United States. Some of these deposits are discussed below. Garlick (1965&b, 1967) published several papers on the origin of the copper mineralization of the Rhodesian belt and suggested (1965a) that the ore in this area is of syngenetic origin, but was diagenetically to metamorphically remobilized. Reprecipitation took place more or less in the near vicinity, so that no long-distance transportation occurred. If it had taken place, the primary or syngenetic zoning would have been obliterated, as Garlick mentioned, whereas, in fact, the zoning was preserved. Figures 5-46 and 5-47 illustrate diagrammatically the sedimentary environment that controlled the ore localization. Rivers supplied the copper which was utilized in forming chalcocite, bornite, chalcopyrite, and pyrite, usually in that sequence, from the shore-line to the off-shore areas (Figs. 5-47 and 5-48), depending on the geochemical parameters within the euxinic bottom waters. The syngenetic sulphide precipitation must have been extremely sensitive to pH, production EDGE
OXYGENATED YRFACE WTERS NXlNlC BOTTOM WATERS H,S
AND CH,
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SUMP B&IA
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Fig. 5-46. Idealized diagram showing relation of sulphide precipitation and mineral zones t o detritial facies of Mufulira-type ore bodies. (After Garlick, 1967, fig. 30; courtesy of University of Leicester.)
ORE GENESIS INFLUENCED BY COMPACTION
557
Fig. 5-47. Shoal waters over granite hill and eddy current behind, causing iron sulphide precipitation between bays containing co p p er ir on sulphides. Location: Mufulira West, Rhodesian copperbelt, South Africa. (After Garlick, 196513, fig. 46; courtesy of MacDonald, London.)
rate of bacterially-formed H2S, amount and concentration of metals in solutions entering the sedimentary environment, and variations in accumulation of the detritus as it diluted the chemically precipitated material, thus controlling the tenor of ore (Rickard, 1973). According to Garlick (1965b), confirmation of his syngenetic theory lies in the presence of features that relate mineralization t o sedimentary structures, such as compaction folds, depositional depressions or basins, and, on a smaller scale, even hollows of ripple marks, load casts, and slump phenomena. Very strong evidence for a very early origin of the ore was supplied by washouts and fragments of mineralized indurated sediments in penecontemporaneous breccias and conglomerates. As depicted in Fig. 5-48, sedimentary transgressions and regressions took place that gave rise t o the zoning of the sedimentary facies together with the ore horizons. Although the above outline on the genesis of the Rhodesian copper deposits appears to be rather simple, Garlick in his earlier works as well as other investigators (e.g., Van Eden and Binda, 1972) found it necessary to go beyond a strictly syngenetic theory and, therefore, introduced a number of modifications by including secondary processes in their interpretations. It is
K.H. WOLF
558
OXVGENATED
IRON SULPHIDE
F."'
C
COBALT SULPHIDE
MINERAL ZONES
-OXYGENATED
TRANSGRESSION
REGRESSION
Fig. 5-48. Diagrammatic section of the formatisn of mineral zones. (After Garlick, 1965b, figs. 47, 50, 51; courtesy of MacDonald, London.)
here that compaction has been mentioned. Inasmuch as the present theme of this chapter is the relationship between ore genesis and compaction, Garlick's references to compaction should be selectively (albeit out of context) considered. The evidence of compaction ranges from a small to a regional scale. According to Garlick's syngenetic theory, metalliferous material was deposited simultaneously with the gangue components of the sediments. Shortly thereafter, jointing and fissuring occurred as a result of compaction and, later, warping and folding. These open spaces then became filled with connate water squeezed out of the sedimentary pile and ore minerals precipitated out of solutions to form the diagenetic veins observed. As to the zonation, Garlick emphasized that the metal zoning is of strictly primary, syngenetic origin, whereas the mineral zonation has been influenced by secondary diagenetic and later metamorphic reactions in the sediments. The immediate question that arises here is to what degree compaction fluids
ORE GENESIS INFLUENCED BY COMPACTION
559
were involved in these secondary processes (see discussion below). In general, it appears that research is required on the influence of subsurface fluids on the zonation of both chemical elements and minerals (cf. Chapter 3 on sandstone diagenesis). Garlick (1965b) discussed numerous structures in the sediments which he interpreted as being of secondary, compaction origin. For example, he mentioned that coarse sediments, accumulating upon hydroplastic muds and oozes, sank into the underlying material to produce load casts. Such structures may resemble those formed by compaction. If ore minerals formed during these secondary, very early diagenetic deformations, then it seems reasonable to assume that the pre-deformation ore was also of a very early origin. Garlick also mentioned that the sediments on the initial slope may have become unstable with continued accumulation so that they slid, triggered possibly by an earthquake shock, towards deeper parts of the basin to form soft-sediment recumbent folds, complex convolutions, and/or slump breccias, as depicted in Fig. 5-46. Such soft-sediment structures have been noticed by Garlick to be common amongst syngenetic sulphide deposits, possibly because the concentrations of sulphide precipitates formed soft heavy oozes. Features like these should be of diagnostic value to distinguish between syngenetic and early diagenetic ores, on one hand, and those of late diagenetic, post-compaction ores, on the other. It should be noted, however, that because load casting and slumping can occur much later after extensive diagenesis, but before the sediments were necessarily lithified, soft-sediment deformation of ore-bearing material does not unequivocally prove that the ore is of strictly syngenetic origin. Very early diagenetic compaction fluids from the basin may have precipitated the ore minerals prior to the deformation. In Figs. 5-49 and 5-50 Garlick (196513, 1967) showed compaction folds over schistose basement hills which projected up into the base of the “C” ore body. Later orogeny has accentuated the draping of the sediments over the paleo-highs. The superimposition of the graywacke “A” over that of “B” and that in turn over “C” (Fig. 5-49) indicates a control of the lithological facies by continued differential compaction folding during sedimentation, thus preserving the paleo-relief. As shown in Fig. 5-50, there are several additional good examples of this superimposition of mineralized units at Mufulira, which, according to Garlick, can be best explained by syngenetic sulphide precipitation during sedimentation and contemporaneous with compaction folding. In support of syngenetic origin, Garlick (1967) stated that at Mufulira the three superimposed ore units are mostly confined to particular lithofacies, such as chaotic slump breccia and carbonaceous graywackes located in the center of the basin, the location and persistence of these facies being controlled originally by the penecontemporaneous compaction folding. On
560
K.H.WOLF
Fig. 5-49. Superimposition of Mulfulira ‘A’, ‘B’, and ‘C’ ore bodies by compaction folding within the eastern basin. Basal and footwall beds of the ‘C’ ore body were compacted by loss of connate water, resulting in draping of thinner sediments over schist hill to the northwest and over less compactable aeolian sands to the southeast. Thus the basin of deposition controlled location of the ‘C’ graywacke and ore body. Further compaction of ‘B’ and ‘C’ ore body sediments maintained a slightly smaller lens of ore below the ‘A’ ore body immediately above the argillite. A foot or two of gritty marker at the top of the lower dolomite is of ore grade within the same lateral limits. (After Garlick, 1967, fig. 25; courtesy of University of bicester.)
the possibility of a diagenetic origin or remobilization, he mentioned that the ore minerals in the pebbly conglomerate itself and in the under- and overlying arkosic arenites should be of syngenetic origin. The infiltration of solutions over a period of millions of years into these porous and permeable beds prior to tectonic folding and metamorphism, however, is not to be excluded entirely. These solutions could be of several types of which compaction fluid is one. Although Garlick considered most of the ore t o be of syngenetic origin, he described one criterion which, according to him, is the only one that favors a secondary, diagenetic mineralization. The occurrence of ore minerals persists below the water-deposited beds of the “C” unit for a foot or two into the aeolian beds, indicating that metal-bearing solutions penetrated into the aeolian sandstone. In this case, it is an enigma as to why the solutions did not cause more widespread mineralization. An alternative mechanical migration and accumulation was suggested by Garlick, i.e., the fine particles of the sulphide precipitate may have descended from the water-laid sandstones into the aeolian sandstones. On the whole, however, Garlick maintained that surface waters were the metal-bearing fluids that gave rise to the syngenetic copper, but admitted that some contemporaneous diffusion of the solution
ORE GENESIS INFLUENCED BY COMPACTION -0
moo ~~~
Mining blocks 75
1
NW
__--
Sop0
4OpO ~
1
lOq00
8000 -=* FEET ~~~
60
561
~
~
“OpO
~~
YO00
16200 ~~
~~
30
I
0 Mining b l w b
SE
-___------
FEET
- _-_ Fig. 5-50. Longitudinal stratigraphic section, showing mineral zones and lithofacies of Mufulira ‘C’ ore body and footwall sediments. (After Garlick, 1967, fig. 26; courtesy of University of Leicester.)
through permeable sand beds have occurred en route to the basin. In general, the common absence of evidence of diffusion led Garlick to believe that the mineralized sandy sediments became closed chemical systems shortly after deposition and burial. It is rather significant to note here that he used the statement “after burial”, because there would have been sufficient time for subsurface fluids, whatever their precise origin, to penetrate into the sediments before it finally became a “closed system”. Garlick himself stated that compaction folding over the paleo-highs and other features .indicated that water was squeezed out of the sediments over a long period of time, as a result of the overburden of younger sediments. According t o him, these fluids did not dissolve the syngenetic ore removed, because otherwise the escape routes of the compaction fluids would have become mineralized. The solutions, however, appear to have formed ore veins. Although Garlick’s support for a syngenetic origin of the copper deposits is very strong indeed, the present writer would like to suggest an alternative interpretation’ based on very early diagenetic compaction, rather than on strictly syngenetic processes. In both theories, of course, the subsequent secondary modifications, e.g., due t o metamorphism, must be given full consideration. Along the coastal areas, where one not only finds rapid sedimentary facies changes but also gradational changes in the chemistry of the surface waters from fresh water through brackish to sea water, one may expect that the subsurface fluids within the sediments show similar regional chemical variations. When compaction takes place, these fluids of different chemical compositions may migrate, mix, and react with each other, which, in turn, can result in the precipitation of ore minerals. Depending on various This alternative interpretation, even if found not to be applicable to the Rhodesian ores, may well be a realistic hypothesis for similar deposits elsewhere.
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K.H. WOLF
factors, this mineralization could be in the form of vertical and horizontal zoning, similar to that observed in diagenetic cements in sandstones (see Chapter 3). Chemical mineral zonation need not be the result of the transgressiveregressive movements of the surface sedimentary parameters, because the subsurface diagenetic parameters will transgress and regress likewise, so that secondary diagenetic zonation is possible. As to the source of the metals, there are at least two alternatives. (1)The source of the chemical elements may have been the offshore fine sediments that released the adsorbed elements into compaction fluids. Upon moving into the porous and permeable shore-line sediments, these fluids became mixed with other types of intrastratal solutions resulting in the precipitation of the ore minerals more or less in accordance with the established sequence presented by Schurmann (1888)and thus forming the existing zonation. (2)The second source is to be sought in the land mass, as Garlick did. He suggested that river waters supplied the copper, for example, to the sedimentary environment. On the other hand, the question arises as to whether the metals could have been brought from the land mass by ground water and upon reaching the fluids in the sediments, that accumulated in the marine milieu, precipitated as ore minerals. Subsurface reducing conditions could have been provided by organic-rich compaction fluids from the euxinic basinal muds which migrated into the permeable coarser sediments. Another possibility is that precipitation of the sulphide minerals was conducive only where the sediments contained carbonaceous components and offered a reducing chemical environment in which bacterial processes produced HzS. The second theory presented above would be related t o those suggested for the origin of the red-bed copper deposits (Samama, 1973,for example), the Colorado-Plateau and Wyoming-type uranium-vanadium ores extensively discussed in the literature (e.g., Rackley, 1972; Harshman, 1972), and for other ores described by Strakhov (1967,1969,1970). Although all the evidence presented by Garlick to lend credence to his syngenetic theory is very plausible, it seems possible to suggest the compaction hypothesis as an alternative. As long as each one of the criteria used can be explained by either syngenetic or early diagenetic processes, only future research work may resolve the problem. Only recently, Van Eden and Binda (1972)came to a similar conclusion, as they stated that the processes responsible for the copperbelt ore bodies were complex and no single mechanism can be expected to explain all the observable features (see also Jolly, 1972). Migration of compaction fluids is one of the mechanisms accepted by them. Confinement of the ores to specific horizons need not be the result of syngenetic precipitation. As in a number of other instances, e.g., uranium in sandstones and copper in red beds, location of the ore could be strictly
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controlled by: (a) the presence of carbonaceous material in the host rock; (b) specific conditions like differential cementation, which lead to differential reduction of permeability and porosity; and (c) the mixing of organicrich and metal-bearing subsurface waters. In regard t o the observations made in the copper deposits of the Rhodesian belt, Newlands and Tyrwhitt (1964)(see also Amstutz and Bubenicek, 1967, p. 435), mentioned that diagenetic differential compaction of a siltstone is demonstrated where pebbles within a pebble horizon have been forced down into the underlying siltstone, thus distorting the laminae. The dispersed sulphide grains within these bent laminae are more densely packed together below the pebbles than elsewhere in the same laminae, indicating that the sulphides were present before diagenetic compaction. Both pyrite and chalcopyrite have been subsequently recrystallized contemporaneously with SiOz-overgrowth cementation on detrital quartz grains. In view of the absence of replacement textures between the sulphides, their origin was provisionally considered by Newlands and Tyrwhitt as being syngenetic with later slight diagenetic modifications. In the absence of more paragenetic data, the above reasoning that the sulphides were present before compaction does not unequivocally support a syngenetic origin, nor does it rule out that compaction fluids could have been the metal-bearing solutions, because compaction occurs over a longer period, often lasting thousands or even tens of thousands of years. It follows then, that early compaction fluids could have brought the metals required for the formation of ore minerals, whereas late compaction after mineralization may explain the distorted laminae observed by the two investigators. Lithification by cementation was the final process. The absence of replacement features is not an unequivocal criterion in establishing either the mode or time of precipitation, although it can be a useful evidence if supported by other information. Already Garlick (1964)described three localities where the mineralization in the elastic rocks (i.e., shales, argillites, siltstones and sandstones) is associated with carbonate rocks: (1)in the northern Rhodesian copperbelt where the ore is situated in the shale adjacent to barren dolomites (originally bioherms and biostromes; see discussion below), which was not mentioned in the above discussion; (2)in the Katherine-Darwin region of Australia, where the uranium mineralization occurs adjacent to, and within, silicified limestone breccias (see Fig. 5-62,in the section on uranium mineralization); and (3)at Mount Isa, Australia, where copper occurs within presumed biohermal silica-carbonate breccia, there are more extensive stratiform Pb-Zn-Fe sulphide deposits in the shales flanking the structures. In all three cases, according to Garlick, wherever the bioherms are barren the ore in the adjacent elastic beds containing extensive stratiform mineralization is of syngenetic origin, because hydrothermal fluids would have a tendency to react preferen-
564
K.H.WOLF
tially with the carbonates and the ore could preferentially replace the carbonates rather than the associated clastics. Malan (1964),who studied copper sulfide mineralization in the sediments associated with algal stromatolites at Mufulira, also agreed with Garlick (1964)(see also Paltridge, 1968) that the higher content of the ore minerals in the clastics associated with the stromatolitic carbonate bodies supports a syngenetic rather than “epigenetic hydrothermal” origin. The main reason for Malan’s conclusion was based on the assumption, also made by Garlick, that hydrothermal fluids would replace metasomatically the algal-reef complex differentially before precipitating sulphide minerals in the argillaceous and silty facies. Considering the preferential distribution of the copper sulfide, Malan observed that the stromatolite columns (Fig. 5-51)are separated by thin, dark-colored bands of silty and argillaceous sediments enriched in copper, averaging about 0.25% Cu. The clean dolomite margins of the columns adjacent t o the argillaceous beds have a very low (0.09%) Cu tenor, in contrast to the central, slightly more clay-rich cores that have intermediate Cu contents, averaging 0.12%. The inter-stromatolite material within the reef complex and the dark inter-reef silty argillite are identical in Cu content and overall composition. As shown in Fig. 5-51,an antipathetic correlation was observed between the Cu and combined CaO and MgO contents, whereas the variations in the A12O and SiOz contents is sympathetic with the variations in the Cu content. According to Malan (1964),this demonstrates that the Cu and the detrital silt and clay were deposited contemporaneously. As in the earlier discussion on the Rhodesian copperbelt, the present writer would like to raise a few relevant questions and offer an alternative interpretation based on compaction fluids, because all the observations listed above can be explained by diagenetic processes. The ore may not have been exactly contemporaneous with the associated detritus, but more or less penecontemporaneous. The difference in geologic time from a syngenetic accumulation of the detritus t o the diagenetic precipitation of the ore may be weeks only, or months, years, and tens of years - but the hiatus is very significant indeed in our contemplations on ore genesis. The following considerations and questions are pertinent: (1)Could an absence of organic matter in the limestone and its presence in the associated detrital sediments have led to the preferential diagenetic reducing environment that, in turn, gave rise to the preferential ore mineralization? It is true that algal stromatolites would be very rich in organic matter initially; whereas, on the other hand, decompositional processes affecting organic matter have been observed in reef limestones to be faster and more complete than in associated fine-grained detrital sediments, so that it would not be unusual to find a limestone unit lacking in organic matter but being surrounded by a carbonaceous shale and/or siltstone. Organic acids
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I
U
I,
Fig. 5-51. Idealized section showing the distribution of (CaO + MgO), Cu, Si02, and A1203 in biohermal stromatolites and associated argillites. (After Malan, 1964, fig. 10; courtesy of Econ. Geol. )
may seep from the shales into associated siltstones, sandstones, and reef limestones if they are permeable. If there is a direct relationship between the amounts of clay minerals and the organic material, as has been noticed in many natural environments, and if the organic components determine in turn the duration and degree of chemical reducing conditions (i.e., the Eh), then there is another alternative logical explanation for the variation in the content of the copper sulphide observed by Malan. (2) An additional factor to be taken into account is based on the adsorption of metals on clay minerals. If Cu-bearing fluids are passing through both detrital and carbonate sediments, the amount of Cu adsorbed and retained increases with increasing clay minerals content in the sediment. Later diagenetic release of the Cu may allow its reaction with bacterially freed H2S to form
566
K.H. WOLF
sulphide minerals. It remains to be shown in the future, however, whether the amount of Cu from such a source is sufficient to account for the observed tenor mentioned by Malan. (3)As discussed in the chapter on the compaction of calcareous sediments (Chapter 3, Vol. I), it is well known that carbonate accumulations are cemented and lithified easier and sooner than detrital deposits composed of silicate minerals. Thus, the possibility exists that the stromatolites were cemented very soon after formation, so that only minor amounts of fluids could have penetrated into the resulting limestone to precipitate copper. If the amount of detritus admixture within the stromatolites influenced the rate and degree of cementation and, therefore, occlusion of the pores, then this may also in part explain the correlation between the amount of argillaceous material and the copper content. (4) What are the factors that control diffusion of metallic ions? If one accepts the possibility that compaction fluids may have brought the Cu into the detrital sediments up to, but not into, the stromatolites, could the Cu diffuse into them? Could diffusion be differential (see Brown, 1974),the rate and amount being at least partly controlled by the amount of clay minerals present in the individual stromatolitic laminae? Or could the retention of the diffusing ions be controlled by the proportion of the different clay minerals, certain groups of which, e.g., montmorillonite, have a higher potential to adsorb ions? If these are plausible mechanisms, then there is no need t o assume a syngenetic deposition of the copper sulphide contemporaneous with the detritus. (5) As pointed out by Malan (1964),the stromatolites were dolomitized. Inasmuch as dolomitization of ore-containing carbonate rocks has been frequently related to the ore-forming processes, it seems wise to give this some thought here also, as it may be pertinent in determining whether a syngenetic origin of the Cu in the Rhodesian deposit is the only acceptable hypothesis. The following questions can be raised here: (a) When did dolomitization occur, i.e., during very early or late diagenetic stage? (b) Where did the Mg-bearing fluids for dolomitization come from? Could they have been compaction fluids or were they concentrated surface waters? (c) What is the paragenetic relationship between the ore mineralization and the dolomitization in the case of diagenetic origin of the ore? (Obviously, i f the ore was syngenetic, then the dolomitization occurred subsequently.) (d) Were the ore and dolomite precipitation genetically related or caused by different chemical mechanisms? (e) If the same fluids were responsible for both, is it feasible that the solutions precipitated the metallic ions in a carbon-rich milieu, whereas the Mg was utilized in dolomitizing the limy sediments? (f) Assuming for the moment that the copper sulphide was syngenetic, or originated diagenetically prior to the dolomite, is it theoretically acceptable that the
ORE GENESIS INFLUENCED BY COMPACTION
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dolomite-forming solution left the ore minerals completely unaffected during dolomitization? In a recent paper on the depositional and diagenetic environment of the Mufulira sulfide mineralization of the Zambian copperbelt, Van Eden (1974) mentioned that the ore occurs in two lithologies: (1)carbonaceous wacke and argillitic sandstone, and (2) conglomeratic or cross-bedded arenite with local intercalated argillite beds. Both types are always spatially associated with each other, as Van Eden’s geologic cross-section and vertical stratigraphic sections demonstrate. He suggested that these two varieties of ores indicate diverse genetic processes. Although some of the ore of type 1 may be of synsedimentary origin, based on several criteria, Van Eden argued (pp. 7577) that postisyngenetic mineral enrichment must have occurred. The older permeable sandstones (underlying the argillitic-carbonaceous sandstone host rock of the ore) could have been an ideal channelway for mobilized, metalcarrying pore fluids. The fluid squeezed out of the younger overlying wackes and argillites during compaction may have supplied the solution and controlled the diagenetic milieu in the host rocks. The copper sulfide in the conglomeratic sandstone host rock (type-2 mineralization) is conspicuously concentrated around argillaceous fragments and argillaceous layers. These clay-rich local concentrations offered both a permeability barrier to the copper-bearing fluids and at the same time their relatively high organic carbon content gave rise to a reducing chemical condition required for the metals to be precipitated as sulfides. If the consensus of opinion among geologists points to a syngeneticearly diagenetic origin of the ores of the copperbelt (and, possibly, similar types elsewhere in the world), then the cyclic origin of both the host and the source rocks, as well as the ore mineralization itself, must find a logical explanation in the genetic theory. Van Eden (1974,fig. 1,p. 60) and Renfro (1974,fig. 7, p. 43) have discussed the transgressive-regressive sedimentary cycles of the ore-containing formations (i.e., the “A”, “B”, and “C” ore bodies). If it is accepted that, independent of whether some of the ore was of syngenetic origin or not, secondary enrichment by compaction fluids was required, then it becomes important t o determine the relationship between the cyclicity of the host rock formation versus that of the ore mineralization itself. Future research will have to establish whether the cyclicity of both run contemporaneously in both geologic time and space, i.e., whether the diagenetic metal-bearing compaction fluids invaded the sandstone host rock soon after its accumulation and prior to the deposition of the overlying next sedimentary cycle. If this is the case, then the three above-mentioned ore bodies (i.e., “A”, “B”, and “C” in the Roan Antelope locality as depicted by fig. 7 in Renfro, 1974) formed one-by-one as successive phases in a cyclic pattern in response to transgressive-regressive sedimentation. This would
K.H. WOLF
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indicate a “very early diagenetic” origin. On the other hand, an alternative geologic possibility exists, namely that the three sedimentary cycles o f host rocks accumulated first and that the secondary ore enrichment by compaction fluids occurred more or less simultaneously much later, i.e., after the three sedimentary cycles were completed. In this latter instance, the mineralization would be “very late diagenetic to epigenetic” in nature, and the localization of the ore may have been controlled by the distribution of the carbonaceous material in the wackes and conglomeratic sandstones, as observed by Van Eden (1974). The likelihood that the Rhodesian copper deposits were formed by several processes has already been mentioned above, and this possibility will have to be pursued further in future investigations. For this reason, a brief comment is made here on the numerous transitional stages from the strictly syngenetic through early to late diagenetic, t o the burial metamorphic and, eventually, to the higher-grade metamorphic processes. As illustrated in Table 5-VII, syngenesis encompasses the stage during which detrital (= clastic) material reaches the depositional environment from the source area and, after possible transportation within the sedimentary environment, finally accumulates. Clay minerals and organic matter, for example, may bring adsorbed metal ions from the source area. Should physicochemical and biochemical precipitation in the depositional environment result in the formation of oolites, pellets and skeletons, among others, then these too may contain metallic ions (see Wolf et al., 1967, and the section on Mississippi Valley-type ores). TABLE 5-VII Syngenetic, diagenetic, and metamorphic processes in sedimentary ore genesis (the results can range from disseminations and intergranular cements to nodules, beds, veins, and coarsely recrystallized ore-mineral products) Syngenesis-
Diagenesis to burial metamorphism
-
Surface accumulation
Various primary (original) textures, fabrics and structures
::
I
+
Subsurface: very shallow to deep --FRemobilized by compaction
Adsorbed elements on clays, organic matter,‘ etc.
-+
Dissem- -+ Cements -+ Nodules inations and other openspace
I
I
Metamorphicremobilization
f
Recrystallization -+
Beds -+ Veins++ Coarsecrystals I >
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As soon as one speaks of chemically originated components, then a precise boundary between syngenesis and diagenesis becomes difficult (see Chilingar et al., 1967, among others). It is difficult to visualize a syngenetic origin of economic concentrations of copper sulphide without diagenesis being involved. As diagrammatically shown in Table 5-VII, the present writer visualizes a genetic, textural, fabric, and structural (not t o mention the more obvious compositional) continuum that is reflecting a transition from the least concentrated adsorbed metal t o a more concentrated accumulation of metal ions and minerals as disseminations, cements, nodules, and beds. All these forms of distributions may be the result of remobilization-transportation-reprecipitation of the originally adsorbed metal ions of the detrital fine material coming into the sedimentary environment from an extra-basinal source, i.e., terrigenous source. The second possible source are deep-water, euxinic, basinal sediments rich in clay minerals and organic matter, which upon compaction may release the metal ions into’the compaction fluids. The latter then migrate into a “reservoir” rock where precipitation of the ore minerals occurs, as long as the chemical requirements are fulfilled. Thus, maybe the disseminations, cements, nodules, and beds of stratabound deposits originate in this manner, because the fluid movements are predominantly controlled by the bedding and laminations. In instances of occurrence of early diagenetic fracturing, the metals may form veins. All the above can fall into the so-called diagenetic-to-burial metamorphic stages within the sediments that are undergoing compaction from near the surface to deeper portions of a sedimentary trough. Then, during the stages of metamorphism and when recrystallization may have its effects, remobilization can concentrate minerals to form another variety of stratiform or stratabound deposits, as well as veins (see Chapter 3 on sandstone diagenesis). Significant t o realize is the existence of two generations of veins, namely, an early diagenetic and late diagenetic to metamorphic (i.e., “epigenetic”) vein type. It is debatable, however, as to what degree it is possible to rnetarnorphically remobilize metal ions from well-compacted, well-consolidated (i.e., with a minimum possible porosity and permeability) rocks and move the ions ouer long distances to a site where reprecipitation occurs in a concentrated form.’ From the above discussion it should be quite obvious that the choice of the nomenclature used by each researcher is very important. If different terms are employed, then definitions are paramount so that one knows precisely what they mean, e.g., the terms “syngenesis”, “synsedimentary”, “syndiagenesis”, “diagenesis”, “burial metamorphism”, and “epigenesis”. Different investigators have used these terms in various ways, without giving the exact meaning to which they adhered, with a consequent result of misunderstanding and confusion. This may apply to many of the publications of the types of the Kupferschiefer ores, for example. The present writer doubts Even if not directly applicable to the Rhodesian ores, this discussion is of a general nature useful in the study of similar ores elsewhere.
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the existence of strictly syngenetic economic ore deposits, unless they are placer-type accumulations or formed by the erosion of older ore deposits, as has been reported for some bauxites. As long as chemical precipitation of the metals is involved, all early to late diagenetic factors and processes must be considered, including compaction and remobilization of metal ions by compaction fluids. Rickard (1973) stated that a major part of the chemical and physical evolution of a sediment occurs during diagenesis and the processes involving “synsedimentary” sulfide ore formation are mainly limited to the postdepositional stage of sedimentation. It has been difficult or impossible in the past to link genetic interpretations to quantitative considerations, but recent research, such as the one undertaken by Rickard (1973), will provide new techniques useful in evaluating several genetic hypotheses applied to one particular ore district. The quantitative techniques rest on the fact that when metals are diagenetically precipitated, a minimum amount of organic matter must be present. As to the proposal that the Rhodesian copper deposits were formed by compaction fluids, calculations eventually may determine whether a sufficient quantity of metals could have been released from the fine-grained source sediments to account for the total amount of the ore, or whether one is compelled to assume an additional source, e.g., volcanic-exhalative, as Rickard (1973) concluded for synsedimentary sulfide ores in general. When considering mineralization within fine-grained clastics versus that occurring within carbonate rocks, the following problem should be given further consideration in future investigations of both the Mississippi-Valleytype lead-zinc and the Rhodesian-type deposits. Those researchers adhering to the syngeneticrdiagenetic origin have usually proposed that the clay-rich sediments may be the source for the metal-bearing fluids (as well as for the hydrocarbons, see pp. 490-513) which moved into a suitable reservoir or host rock where precipitation of the ore minerals took place. Considering that ore mineralization has been found in both “source”, and “host rock” lithologies, the question should be posed as to a possible genetic relationship between them. An attempt should also be made to establish a conceptual model composed of end-members and several transitional examples in between, as shown in Table 5-VIII and Fig. 5-52. If the metal originated in the black-shale, euxinic environment and was precipitated there (type 1A), then it would form a Kupferschiefer-type metal deposit. Similarly, if the shale was the source of oil, but the oil remained there for various possible geologic reasons, the result is an oil shale! If the metal-bearing fluids moved from the shaly environment to a locality where there is a lithologic transition, e.g., an interbedded shale-siltstone and carbonate facies, and where the carbonate is impermeable or does not offer the geochemical conditions required for precipitation, deposition of the ore minerals may occur within the clastics
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TABLE 5-VIII Four types of ore deposits with several sub-varieties (for explanation, see text) \L
-
Typ e 1A sedimentary muddy source rock, e.g., black shale
.1
Metals remain--/ in the source rock, e.g., Kupferschiefer type deposits
__f
Type 2 Transition zone between source and host rocks
__f
Type 3 Sedimentary host rock(s)
Metals move up inLMetals move into to the shallower“reservoir” or water lithologies, host rocks; there up to the “host are several rocks”; but not into them bestones and cause they are conglomerates tight; example: (b) carbonate sulfides in rocks, e.g., detrital sediMississippi-
Type 4 * Volcanic host rock(s)
Type 1 C * Fine-grained pyroclastic source rock(s) V
-1
Metals move in- Metals remain t to other volcan- in the source ic rocks (flows, rock breccias, pyro-
-
Carbonate source rock
A
h
Two types of transition zonelore Metals remain occurrences: in the source (a) with a very 4 rock, e.g., Wood- sharp contact cutters Pb(b) with a gradaZn ore (Roberts, tional change in
*Note: Volcanic-exhalative and hydrothermal metal supply is not considered here.
around the carbonate unit (type 2, Table 5-VIII). This may have occurred in the case of some deposits in Rhodesia discussed above. If the carbonate unit is permeable and the geochemical milieu is conducive for precipitation, however, then the carbonate may become the host rock for the ore mineraliza-
K.H. WOLF
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tion (type 3b, Fig. 5-52) (e.g., Mississippi Valley-type deposits). Roberts (1973) suggested that the carbonate sediments themselves were the precursors for the metal ions that upon remobilization became concentrated to form the Pb-Zn deposits of the Woodcutter’s mine in Australia (type lB, Table 5-VIII and Fig. 5-52). It should be noted that in the above discussion no reference has been made to a volcanic source for the metals. In Table 5-VIII it is suggested that pyroclastic material, with its relatively high trace-element content, could be the source for the metal ions, as has been proposed for the Colorado Plateau types of uranium-vanadium ores, which constitute one possible variety of type 3 among several other ore types. If the metals remain in the volcanic source, the type of ore would be l C , whereas if the metals move into another variety of volcanic host rock, possibly because the latter is more permeable and more porous, the ore would be of type 4 (Table 5-VIII and aafef and fluids
Evaporites
K H WOLF(1972)
Fig. 5-52. Conceptual model showing sources for metal ions, i.e., terrigenous rocks, normal and saturated sea water, volcanic-exhalative source, and, possibly, evaporites and chloride brines derived from evaporites. Four rock varieties may serve as hosts for the ore bodies, namely, mudstones o r shaIes, sandstones, carbonates, and volcanics. The types of copper ores designated by the four numbers (e.g., 3A and 3B)are discussed in the text. One should note that this model is simplified, because it does not indicate, for example, that the detritus in sandstones (e.g., arkoses) and the limy and/or dolomitic components in the carbonates can supply chemical elements for ore minerals under favorable conditions.
ORE GENESIS INFLUENCED BY COMPACTION
573
Fig. 5-52). As the arrows indicate in Fig. 5-52, there are other possible sources and migration routes, namely, metals from the pyroclastics could be moved into sandstones or into carbonate host rocks. Although no volcanicexhalative, hydrothermal source has been considered in Table 5-VIII, it has been given a proper place in Fig. 5-52. Kobe (1960) described a Peruvian copper--silver deposit of the red-bed type in the red, coarse- to fine-grained arkosic Permian sandstones with layers and lenses of red mudstones. The ore zone is characterized by the bleaching of the red color of arkoses t o a light bluish- or greenish-gray color. The disseminated copper ore is associated with abundant disseminated carbonaceous plant remains and there is a partial replacement of wood by copper sulfide minerals. This deposit is very similar to the uranium-vanadium-copper deposits of the Colorado Plateau and Wyoming, but uranium is lacking and the vanadium occurs in only minor quantities. Kobe suggested that the deposit is of “post-syngenetic” origin and that ground water brought in the copper to the sites of carbonaceous sediments, where reducing conditions necessary for the precipitation of the sulphide minerals existed. For deposits such as the ones mentioned above, that apparently owe their origin to ground water, which has brought the chemical elements into the host rock t o be precipitated there, one may have the tendency to dismiss the influence of compaction (see also the section on uranium ores below). The following considerations demonstrate that, even if compaction fluids have not been the carriers of the metals, compaction may have had other important influences on the rock itself to make it either a likely or unlikely host for ore mineralization. Even the absence of compaction may be more conducive t o establishing a favorable geochemical milieu for the sulphide precipitation, because: (1) Very fine disseminated organic matter may be flushed out of coarse-grained sediments by compaction fluids, thus preventing bacteria to form a reducing chemical milieu. (2) Inasmuch as organic matter undergoes a gradual and gradational change during burial (called eometamorphism; see Chapter 3 on sandstone diagenesis), the absence of intense compaction and alteration would be more conducive to the preservation of the organic material. (3) In addition, under conditions of little or no compaction, the relatively unchanged organic material is more susceptible to bacterial decay to give rise to a reducing chemical environment. Also, bacterial activity decreases with increasing burial. More research is needed on establishing the relationships between the “reducing power” of various types of primary organic matter and bacterial decomposition, and the results should then be correlated with the types of organic components secondarily formed during burial. From the above considerations alone, it can be concluded that the absence of compaction, or a minimum thereof, may be
K.H.WOLF
574 A
case(a) or (b)
I
0,
case ( c )
I
f
precipitation
clay o r m u d sand or sandstone
Fig. 5-53. Diagrams illustrating various possible alternative movements of compaction fluids related to ore genesis. A. Cases a and b: Compaction fluids are squeezed out of fine-grained sediments into more porous and permeable coarser deposits. Adsorbed trace metals on the clay minerals are released into the compaction fluids. The sandstone may serve as an “aquifer” (= conduit) as well as a host rock for the ore-forming solutions and mineralization, respectively. On the other hand, under favorable conditions, the metals may precipitate within the finegrained source rock, e.g., if decaying organic matter leads to a reducing milieu and if H2S is available. T h e depleted fluids may then be pressed out and removed through the sandy aquifer(s). B1.Case c: The fine-grained, muddy sediments do not contain metals but are rich in organic matter. During fluid movements, metal-bearing fluids from adjacent and subjacent coarser-grained units may move into the fine sediments (this might be possible as long as compaction has not proceeded too far). The reducing conditions within the fine-grained sediments and the presence of H2S result in the formation of sulphides. “Membraneaction” may also be influential here. If some of the organic-rich water “leaks” into the adjacent coarser sediments containing metal-bearing interstitial fluids, then sulphide might precipitate along the contact of the fine- and coarser-grained sediments. (See publications discussed in the section on iron ores, for examples of passage of compaction fluids through a relatively impermeable material, i.e., Trendall and Blockley, 1970, and Mc-
ORE GENESIS INFLUENCED BY COMPACTION
575
significant in establishing reducing geochemical subsurface environments required for the formation of certain types of ores. At least one consequential question arises, namely, is there a certain depth of burial below which certain types of ores in sediments cannot form? Are these conditions especially fulfilled in continental, intermontane basins, where the metal-bearing fluids were surface waters that penetrated into the sediments, resulting in the copper, uranium and vanadium mineralizations in coarse fluvial deposits, for example? It has been suggested that, even in the case of these continental deposits, the associated fine-grained floodplain sediments may have been influential in ore precipitation by supplying organic matter to the coarser deposits through the transportation of compaction fluids containing suspended and colloidal organic matter derived from the clay-rich sediments. In a basin that contains a sequence of interbedded lithologies, e.g., clayand sand-rich units, one has to consider a number of possible fluid movement patterns of which some may be more common than others. It may be difficult, albeit impossible in some instances, to determine which flow pattern actually did occur; nevertheless, in basin analysis all possibilities have to be considered. Figures 5-53A to 53C outline various models. Of these, the most common fluid movement pattern may be from the finer to the coarser,
Geary and Damuth, 1973.) Case c, as given above, is probably the least common natural occurrence because the finer-grained sediments compact rather quickly and the net flowage of fluids is mostly from the fine- to the coarse-grained sediments. Nevertheless, diffusion into the clay-rich unit has t o be considered when compaction is incomplete. In this case, there is a transition t o processes involved in case ( d l ) given below. Bz. Case d l : If compaction is slight or absent, fluids may still move into the permeable aquifers due to a fluid pressure drop across the sand-clay interface, or if there is a supply of compaction fluids from elsewhere along the aquifer(s). If the solutions moving in the coarse-grained sediments bring metals from another source and if the fluid passage i s very slow, or even nil periodically, diffusion into the organic-rich mud can cause precipitation of the metals within the fine-grained sediments. Bz. Case dz: A process related t o both d l and c may occur also. If organic-rich fluids pass from the muddy deposit upwards and downwards into the sandstone and if the latter contains metal-bearing fluids, then the mixing of the fluids may result in ore precipitation within the coarse aquifer(s). As an alternative, if the mud is still relatively permeable, then part of the mineralization might extend into the fine-grained sediments. C. Case e: The compaction fluids may pass into an underlying coarse-grained aquifer and after moving along some distance, may seep out into the sea water. The compaction fluids could contain metals desorbed from the clay minerals, which are precipitated as sulphides once the fluids enter the stagnant, reducing, HzS-containing surface water. As an alternative, if the surface fluids are rich in metals, maybe supplied by the volcanic-exhalative source, then the compaction fluids could play an important role in ore genesis by bringing fine particulate organic matter to the surface to form a reducing milieu upon decomposition.
576
K.H. WOLF
more porous and permeable units. On the other hand, one should never ignore the possibility that fluids can move through muddy sediment, even if the flow rate is reduced, so that the chemical components in the fluids may be pushed by compaction forces from the coarser-grained bed into the tighter, fine-grained rock to react with the constituents in the latter. This may also occur in the case where the clay-rich unit is not composed totally of a uniform, thick deposit of clay, but where the clay minerals form small seams, lenses and paper-thin interbeds within siltstones and sandstones, or where clay occurs as a matrix in detrital sediments. Hence, a distinct separation of the source from the reservoir lithology may not always be a clear-cut one. An example of this was provided by Knoke (1968),who described a bleached zone of conglomerate and grey sandstone that is situated between the Kupferschiefer* and the thick, red, hematitic sandstone of the Rotliegendes. Inasmuch as there are no big differences between the bleached zone and the hematitic Rotliegende sandstone in mineral composition, it was suggested by Knoke that the bleaching is the result of diagenesis. The higher concentration of metal elements in the bleached zone must be interpreted as enrichment due to mobilization from below. At the top of the bleached horizon, there is a clear zone of additional enrichment just below the Kupferschiefer (Fig. 5-54).It seems that the Fe was removed (causing the bleaching) and Cu, Ni, and Pb became enriched towards the Kupferschiefer. These elements were set free from the hematitic matrix and passed into solution, which then moved towards the Kupferschiefer. Inasmuch as the sulphides of these elements have very low solubility products, the elements were precipitated upon coming in contact with the sulphur ions. In this manner, according to Knoke, the Rotliegende as well as the bleached zone supplied a part of the chemical elements to the Kupferschiefer. This example demonstrates that, although there may be an interbedded sequence of clay-rich and sandrich units, the metals may not always originate from the fine sediments but instead may come from a clay and hematitic matrix of the sandstone (see also Van Eden, 1974). In addition to the free movements of fluids in sediments, diffusion must also be considered (Brown, 1974). Several researchers have observed that secondary changes are particularly common in some instances along the boundary between the fine-grained and coarse-grained sediments. This could be caused by: (1)the mixing of compaction fluids from the clay-rich unit with the interstitial fluids in the sandy layers above, which can result in chemical reactions followed by precipitation in the lower part of the sandy unit; and (2) the transportation of chemical elements by the compaction solutions from the clay-rich unit to
* The German word “Kupferschiefer” can be loosely translated as “copper slate”, i.e., a fine-grained slaty rock containing copper.
ORE GENESIS INFLUENCED BY COMPACTION
;."-;Q-'
Kupferschiefer
Cu Ni
I I I
!
" -0
1 I
I
21
i
b
I
y'
n
Pt
i
t! ;
577
I
C c
i
25
I
Rot liegendes
4. I
NI
Cu
I
10 30 100 Con centrat ion (p p rn1
Fig. 5-54. Contents of copper, nickel, and lead of the borehole Drevenack (calculations made on carbonate-free basis). Concentrations in ppm. Thickness in cm. (After Knoke, 1968, fig. 6; courtesy of Contrib. Mineral. Petrol.)
the interface between the sandy deposit and overlying clay-rich bed, where deposition takes place. In considering the origin of ores in sediments of the Peko ore body, Australia, Elliston (1966) stated that it is acceptable today to think in terms of connate water, adsorbed metal ions on clay minerals, and trace components of rock-forming minerals. The Peko ore body represents an accumulation of ore mineral components in sediment (Fig. 5-55). It occurs in an elongated, steeply-dipping, pipe-like deposit which contains much sedimentary material, mineralized to various degrees. There are some accumulations of pure mineral precipitates, e.g., magnetite, quartz, jasper, pyrite, chalcopyrite, pyrrhotite, galena, and other minerals. These do not have a bedded form and were not formed at the same time as the sedimentary host rock. They were introduced later into the enclosing sedimentary rock. Elliston (1966;see also pp. 489-490) considered the various forms of transportation mecha-
K.H. WOLF
578
__
70NES OF REFLUIDISATION,
A
Fig. 5-55. Diagram showing the outflow of water from compacting and dehydrating sediments during Iithification and folding. Mineral matter is precipitated in the seepage channelways when the escaping water carries dispersed metallic substances desorbed from crystallizing colloids in the source rock. (After Elliston, 1966, fig. 1; courtesy of University of Tasmania. )
nisms of metal ions, i.e., in the form of a vapor, liquid solution (ions unite with in-situ ions to precipitate in the host rock) and as a very finely divided suspension. Elliston, by the process of elimination, concluded that the last possibility is the most likely one. He referred to the paper by Noble (1963) who suggested that the water expelled from the sediments during compaction may be an ore-forming fluid and stated, in regard to the origin of the Peko ore body, that this may apply to the Australian ore also. Dewatering of the turbidite, muddy sediments must have taken place as they contain today less than 2% water and it is assumed that an initial water content was about 77%. Three tons of water has been expelled for each ton of rock remaining. The water loss would have been asymptotic in geologic time at a progressively declining rate extending over a considerable time span. The metallization of the connate water resulted from the release of chemical elements from the sediments themselves (e.g., Saxby, 1969). Under the Peko ore body (Fig. 5-56) there is a belt of sediments totalling in volume about 2-3 cubic miles that could have yielded 750,000 tons of copper metal, if the ore-forming processes extracted and redeposited all of it by a process described by Elliston. Elliston suggested that the chemical elements from the
ORE GENESIS INFLUENCED BY COMPACTION
579
Fig. 5-56. A cross-section through the Peko ore channel illustrating its present relationship to the enclosing sediments and porphyroidal source rock. (After Elliston, 1966, fig. 2; courtesy of University of Tasmania.)
sedimentary particles, on which the elements were absorbed, were released into the subsurface fluids in the form of colloids. In the symposium entitled “Syntaphral Tectonics and Diagenesis” (Elliston, 1963), many of the arguments for and against the interpretation offered by Elliston for the Peko ore body are presented. The symposium is a recommended reading as some of the fundamentals of colloidal chemistry, thixotropy of sediments (see pp. 488-490 above), etc., related to compaction and sediment movement, are briefly summarized. An abstract of the genesis of the Peko ore mineralization as envisaged by Elliston (1966) is given below: (1)The younger beds of the turbidite and mudstone sequence were reslurried at moderate depth as a result of a sudden reduction in thixotropic yielding value of the sedimentary pile, possibly due to an earthquake shock. (2) The sediment moved as a subaqueous mud slide and flow that was accompanied by shearing, flow foliation, and the rafting of large blocks of penecontemporaneous sediment that retained cohesion. (3) At the time of movement, the sediment had the required water content so that before thixotropic yielding the intergranular fluid was a concentrated gelloid. The concentration of the colloids in the solution fulfilled the requirements for the Van der Waals attractive forces between the particles to exceed the Coulombic repulsion. Close packing of the grains was accompanied by a segregation of the material into a relatively well-sorted, clay-free granular matrix and large colloidal aggregations of the contained colloidal gels. This resulted in a texture or fabric giving a porphyry-like appearance to the rock, the origin of which has been under dispute. (4) Segregation of the metallic ions and molecules adsorbed on the colloidal particles from the porous granular matrix and their subsequent aggregation resulted in the formation of the clot-like aggregates. (5) The clots crystallized in time with further diagenesis and burial meta-
580
K.H.WOLF
morphism (e.g., compaction, dehydration, and increase in temperature). (6) The metals were desorbed and transported from the gel mesh work of the crystallizing aggregates as insoluble metal hydrosols, which were then peptized (i.e., stabilized or deflocculated by other substances in suspension or solution) by silicon and sols within the solution expelled during crystallization. (7) The fluid emerging from the lithifying sedimentary pile passed into zones of weaknesses in the overlying sediment where seepage zones or extrusion channels were established, Several episodes of reslurrying of the sediment unit in the vicinity of the Peko ore body occurred subsequently. (8)The solutions available in declining volume with time from the compacting, lithifying and crystallizing pelitic unit merged for outflow through the Peko ore channel. The fluids were successively enriched in the desorbed metal hydrosols and heated by the crystallization (an exothermic reaction) of the colloidal matter. (9) The colloidal mineral components precipitated in the ore pipe in the form of overgrowths, clots, replacements, and intergrowths, and occasionally were remobilized within the ore pipe to form massive intrusions and veins. (10) The paragenesis of the ore minerals resulted from the successive desorption of metals from the source material which in turn changed the ore-forming fluid’s composition. (For details on the above-presented stages, see Elliston’s, 1966, publication.) Harrison (1972) mentioned that anomalous quantities of copper (at least 100 p.p.m.) are present within the Precambrian Belt basin sediments in the northwestern part of the United States; this has been recorded from all but three of the formations. Similar occurrences of copper in sedimentary or low-grade metasedimentary rocks of Belt age have been found in Africa, Australia, Russia, and Canada. Certain characteristics of these stratabound deposits have led to the belief that the copper mineralization is of syngenetic or diagenetic origin, e.g., (1)the occurrence of copper along bedding planes, in minor sedimentary structures or in certain layers; (2) an extensive lateral and vertical extent interrupted only by some unmineralized beds; and (3)the world-wide distribution of copper in rocks of the same age. According to Harrison (1972), concentration of copper in a I-ft zone ranges from a background of about 20 p.p.m. or less to 10,000 to 20,000 p.p.m. (ore grade) with all gradations in between. In Fig. 5-57, Harrison presented an arbitrary classification composed of three ranges of copper tenor indicated by the thicknesses of the lines. The principal copper minerals comprise chalcopyrite, chalcocite, digenite (?), and bornite, commonly accompanied by covellite and various secondary hydrous minerals. The various anomalous copper occurrences are related to rock type or particular stratigraphic zones, chiefly in laminae, layers, or parts of the formation containing greater
581
ORE GENESIS INFLUENCED BY COMPACTION W
CHEWELAH,
WASH
E
240 MILES
PEN0 OREILLE. IDAHO
ST. REGIS- ALBERTON, MISSION SUPERIOR, MONT MTS, MONT MONT
SUN RIVER, MONT
Y
Wallace Fm
w
Prlchard Fm
I
1vry
Lower Belt
Approairnote true-scale rection
--/----
Fig. 5-57. Distribution and relative amount of anomalous copper in Belt rocks. Heavy line
= ore grade; medium line = several thousand p.p.m. is common; light line = several
hundred p.p.m. is common. "he information on copper should be considered minimal and the subdivision into three classes is arbitrary - see text. (After Harrison, 1972, fig. 13;courtesy of Bull. Geol. Soe. Am.)
amounts of sand and/or silt. There are local concentrations of copper in the thin beds of red and green argillite or limestone. Where the units containing oxidized and reduced iron minerals axe alternating, the highest copper concentration coincides with the green beds. In all of these sediments, the primary copper occurs along the bedding planes, in certain laminae or even cross-laminae, mudcrack casts, and cut-and-fill structures. Stromatolitic zones and carbonate lenses may also be favorable for mineralization. The general distribution of the copper, based on the meagre information available t o Harrison, is given in Fig. 5-57. Harrison stated that the present working
582
K.H. WOLF
Fig. 5-58. Composite of isopach maps of Revett Formation (solid lines) and lower part of the Missoula Group (dashed lines) showing relation to western Montana copper sulfide belt (stippled). (After Harrison, 1972, fig. 14; courtesy of Bull. Geol. SOC.A m . )
hypotheses for the genesis of copper are generally based on primary syngenetic-diagenetic concentration and migration and secondary reconcentration of the copper in silty or sandy units or in small fractures. Although this concept is reasonable for many copper occurrences in the Belt formations, the geometry and tectonic history of the sedimentary basin indicates a major epigenetic reconcentration of the copper in the Revett Formation in the western Montana copper sulfide belt. Figure 5-58 is a composite diagram of the Revett sedimentary prism, the isopachs outlining the Early Missoula dome, and the copper sulfide belt within the Revett sedimentary prism, which is nearly parallel to the dome. This formation is composed of at least three thick sandstone lenses in this area and is capped by a silty and argillaceous layer of the St. Regis Formation (Fig. 5-58). The latter may have acted as an impervious cap rock and trapped the ore-forming fluids. Hamson (1972) reasonably concluded that the ore localization is related to the dome and that the ore occurrence in the siltstones and quartzites is due to their higher permeability, i.e., the differential subsidence over a dome formed a stratigraphic trap. The solutions, according to Harrison, could have been connate or ground water, fluids released during burial and low-grade metamorphism, or igneous hydrothermal solutions. At the present stage of knowl-
ORE GENESIS INFLUENCED BY COMPACTION
583
edge of the geologic history of the basin, however, the possible source of fluids and the time(s) of migration subsequent to the origin of the stratigraphic trap cannot be determined and are subject to speculation. The ore must have been formed prior to the time (pre-Middle Cambrian) of formation of higher-grade metamorphic minerals (i.e., of the biotite facies), because this type of recrystallization greatly reduces the rock permeability. The occurrence of similar copper ores at the eastern edge of the basin gives some support t o the above genetic hypothesis. As shown in Fig. 5-57,copper in potential source rocks, i.e., green-beds, is present in the deeper part of the basin, but again the assumption that these rocks were the source beds is at present merely a suggestion. In a personal communication (Feb. 16, 1973, quoted with permission), Harrison stated that it is certainly his own working hypothesis that compaction fluids may have moved copper to sites of concentration in the Belt rocks. According t o him, the chemistry of the fluids seems to clearly relate t o color of the Revett (the “quartzite-type”) sediments, because anomalous copper concentrations and ore occurrence are known only in white quartzite and siltstone, whereas the bulk of these rocks away from Cuconcentrations are purple in color due to the disseminated hematite. Although not all white beds contain anomalous copper concentrations, they may be “bleached” because of fluid migration. According to Harrison, the latter hypothesis has not been tested. He also mentioned that the origin of the green coppercontaining beds are not so simply explained, and stated that at this stage of our knowledge it is hard to establish a bias as to whether the green beds are essentially syngenetic or whether they are altered red beds. A logical argument can be made for either hypothesis, depending on the data collected. The importance of compaction to the origin of copper deposits of the Belt formations is not known. Inasmuch as copper is a highly mobile element during dewatering of sediments, compaction fluids could readily move copper. Harrison pointed out, however, that this could also be done by any postcompaction ground-water circulation and, at a later stage, by further dewatering as a result of metamorphism or addition of circulating water originating from the intrusives. In regard to future studies to determine the origin of the copper, one might suggest that in cases where metamorphism has not totally obliterated the original textures, detailed petrologic work on a local and regional scale may be helpful. As mentioned elsewhere in this chapter, textural and diagenetic-paragenetic investigations as done by the petroleum geologists in establishing the time of oil invasion (see Chapter 3 on sandstone diagenesis), could be attempted in the examination of copper-bearing coarse-grained sediments. A similar investigation may be especially advisable on the uraniummineralized sandstones of unaltered Paleozoic and younger sediments.
K.H.WOLF
584
MARKER BED*
Black
thinly
laminated siltstorm
with abondont white caIcarews blobs, Ihlcknesr
Masstre-breaklng black s l l t s t o n e .
STR/PED BED.
*
Flncly lamlnalod black sillstone ulth whitm calcareous laminae.
Masslve gray s ~ l t a t o n ain graded bmds 6-18
mctms
thick; contalns coorse grained aondslone beds toward south Widely laminated siltston0 w l h reddish varrelike
c
orglllaceous partings PERCENT COWER 1
Y
3 near base. IFinel) mterlominated dark-prey siltstone
7 I
44
I
en
I
2
3
4
5
:upriferous zone
Finm- lo coarso-grained lithic sandstme. IOCallY pebbly,reddish shale and slltstone common
in middle port 01 bed.
la
cupriferous zone
1
Fig. 5-59. Stratigraphic section of the Nonesuch Shale (Parting Shale, Upper Sandstone and Upper Shale) and associated rock units. Asterisk indicates names of local usage. The "Lower Sandstone" is the upper section of the Copper Harbor Conglomerate. (After Ensign et al., 1968; also in Burnie et al., 1972, fig. 2; courtesy of Econ. G e o l . )
ORE GENESIS INFLUENCED BY COMPACTION
585
The finest recent examples of an attempt to apply elementary principles of ground water hydrology t o the formation of an ore deposit in sedimentary rocks was presented by White (1971)and Brown (1971,1974) in an investigation of the White Pine copper district in Michigan. Inasmuch as both of these authors discussed compaction fluids as being one possible source for ore-bearing solutions, at least a brief comment is advisable here. Even though the genetic ideas on that deposit are in a state of flux, particularly interesting are their attempts to make the theory quantitative. The copper, mostly as chalcocite, occurs in the lower part of the Nonesuch Shale (this is a misnomer because the formation consists mainly of sandstones and siltstones with only some shale, as shown in Fig. 5-59).Pyrite is the only sulfide that occurs in the 200-m section of the Nonesuch Shale above the cupriferous zone. In the latter, the sulfides show zonation that is recognizable even in handspecimens. At the top of the cupriferous horizon there is a narrow zone of disseminated sulfides where chalcocite changes to the bornite (transitional zone) and, then, into very fine-grained pyrite. Yellow greenockite (CdS) is commonly visible above the bornite subzone. More detailed polished-section studies reveal the more complex assemblage as presented in Fig. 5-61.The galena (PbS) and wurtzite (ZnS) occur in sparse amounts. It has been suggested that the copper is of epigenetic origin and was precipitated by fluids introduced from the underlying Copper Harbor Conglomerate (a misnomer because at least the top of this unit is a sandstone, as shown in Fig. 5-59).A thin fringe zone (Fig. 5-60),composed of bornite and chalcopyrite instead of chalcocite, marks the top of the cupriferous zone. This fringe transgresses stratigraphy and indicates that the fluids came from below. The copper ions in the solution reacted with the iron sulfide (pyrite) and the iron was replaced by the copper to form copper sulfide minerals (for the latest data, see Wiese, 1973).White (1971)mentioned that at least two mechanisms require consideration: (a) copper solutions moved upward through the Nonesuch Shale units and were completely depleted of copper at the front of the moving solutions by reaction with the pyrite; and (b) the copper diffused upward from a reservoir in the underlying deposits. White’s discussions were concerned mainly with the first possibility. According to the geologic evidence, it was not known whether the epigenetic mineralization took place when the host rock was still relatively close to the surface of deposition or much later, after the Nonesuch Shale units were buried beneath thousands of meters of overlying sediments. As White stated, however, his paleohydrologic analysis suggested that the ore deposit was formed below a relatively thick overburden, but prior to tectonic deformation. He assumed, therefore, that at the time of mineralization the host rock was relatively flat lying and that its degree of compaction was proportional only to the amount of loading by sedimentation and overburden.
K.H. WOLF
586
Copper
Sulfur
Comer Sulfur
Feet 70,
60-
5040-
30 20-
10OJ
Shaly Facles
0
1
2
3
Approximate horizontal scale
4 I
Miles
Fig. 5-60. Schematic diagram showing: (1)the close association of the fringe with sulfide-rich beds as observed across most of the White Pine district, and (2) the inverse correlation between the thickness of the cupriferous zone and the amount of mineralization in the combined No. 21 and No. 23 beds. Note the small amount of copper in the Nos. 21 and 23 beds in the left-hand profile where the fringe is stratigraphically high, and the large amount of copper in the Nos. 21 and 23 beds in the right-hand profile where the fringe is stratigraphically low. Vertical exaggeration is approximately X 300. (After A.C. Brown, 1970, fig. 1.)
In determining the likely source of the ore solutions and the nature of the driving mechanism, White (1971)offered his hydrologic analytic approach. The assumption that extensive compaction of a conglomerate could give rise to fluids is questionable, however, because of the resistance of the coarsegrained framework to deformation. One should realize, however, that the conglomerate formations usually contain units of fine-grained material, which upon compaction can release solutions. Certain assumptions had to be made for establishing a hydrologic model, i.e., copper concentration, geologic time involved, thickness of the formation, and permeabilities. Based on certain considerations, White assumed that the fluids contained 50 p.p.m. Cu at the site of precipitation and that the geologic time involved was in the order of 10 million years. The Nonesuch SKale has an assumed thickness of 150-350 m depending on locality. The underlying Copper Harbor Conglomerate is 150-320 m thick in one area but thickens rapidly to about 2000 m. One of the more difficult estimations which had to be made is that of the permeability of the rocks to water at the time of fluid movement and ore
ORE GENESIS INFLUENCED BY COMPACTION CUPRIFEROUS
PYRITIC ZONE
TRANSITION ZONE
Chalcoclte I
I ! Dgejite
Djurleite
I I
I
I I
II
aalccpyritePyrite
- -j-
- -
I I
I I I
I I
I I
I
Greenockite Galena
Wurtzite
-----I
-
-
I
~
I
-
+
-
I
Fig. 5-61. Typical zonation of sulfides at the top of the cupriferous zone, White Pine ore body. (After A.C. Brown, 1 9 7 1 , fig. 1 3 ; courtesy of Econ. Geol.)
genesis, and during the different stages of compaction. White, therefore, used the data for a relatively unconsolidated and unmetamorphosed Tertiary sedimentary sequence from California that resembles in composition the Upper Keweenawan sedimentary units of the White Pine area. He assigned hydraulic conductivity values for the Precambrian sediments by comparing them with the slightly compacted, unlithified Tertiary equivalent lithologies. White then calculated the amount of water and its rate of flow needed to form the copper mineralization, and discussed the horizontal flow in the Copper Harbor Conglomerate on a quantitative basis. Two models 'were considered, involving: (a) lateral migration of saline compaction fluids derived from the underlying sediments to the site of mineralization, and (b) inflow of surface water from its northern flank into the basin. White's c'alculations related to model a indicated that the volume of water available provided the most stringent limitations. The amount of water produced by compaction of the units of the underlying older units in the center of the basin, however, was just adequate. If future data should suggest that of the two proposed models model (a) is the more likely one, then it will be necessary to assum that considerable funnelling of solutions toward the White Pine deposit has occurred. If this model is factual and if calculation procedures can be refined, then the maximum volume of ore that can be expected to be present in the
588
K.H.WOLF
various parts of the White Pine region can be determined, because a basin cannot supply more than a particular maximum amount of compaction water. Although model (b) is hydrologically feasible, White (1971)considered it to be less acceptable for several geologic reasons. The surface water entering the Copper Harbor Conglomerate in the north would be an adequate source of fluids if the point of entry was significantly higher in altitude than the water table that existed at that time near the White Pine region. The geology, however, suggests that this was not the case, so that White’s compaction model seemed to be the more reasonable one. He concluded, however, that his percolation models will have to be quantitatively compared with others, e.g., models of infiltration and diffusion as proposed by Brown (1971), whose paper appeared at the same time. The concept of density- or gravity-stratification had been described by White (1968)as well as others, and was briefly discussed by A.C. Brown (1970, 1974). The former proposed that the upper limit of the White Pine ore minerdization was defined by the position of a horizontal density-controlled interface between metal-bearing brines below, and near-surface ocean waters above. The cross-cutting relationship of the top of the cupriferous zone is attributed to inclinations of the strata at the time of mineralization (Fig. 5-60).Several reasons for this interpretation were provided by A.C. Brown (1970,1971),who believed that the flow of the fluids was upwards from below. He encountered certain problems and proposed that they might be solved by a combination of the gravity-stratification model with the infiltration and/or diffusion model, which he discussed in 1971. The upward advance and ultimate position of the copper-bearing solution front was determined by the supply of copper from below. According to Brown, the diffusion model could be combined with the other genetic theories (e.g., natural chromatography, Brown, 1974), in particular if the compaction of the clay-rich units impaired the free flow of ore solutions, because the density interface would cease to influence the ore mineralization process. The brine below the interface could continue to be active in metal complexing and in exchange reactions among the mineral and organic components of the clayey sediments and the interstitial solution. Brown (1971,1974)suggested that the metals more soluble than copper were apparently swept upward ahead of the copper front as indicated by the presence of Cd and Pb immediately above the cupriferous zone (Fig. 5-61). He considered the simple process of infiltration (i.e., the transportation of copper in solution by bodily flow of ore solution from the conglomerate unit into the overlying clayey ore-containing units) and diffusion (i.e., the migration of dissolved copper through a stationary pore solution due to a chemical potential gradient) models to test the above hypothesis. As to the source of the copper, Brown listed: (1) latent copper-bearing, volcanic-exhalative solutions generated
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during the volcanic activity that ceased only shortly before sedimentation of the Nonesuch deposits; (2)conventional magmatic hydrothermal solutions; and (3) extraction of the metals from older extrusives and/or sediments by reactions with intraformational waters (e.g., Helgeson, 1967).The transport of the copper from these possible sources to the site of precipitation occurred through the permeable Copper Harbor sediments either by compaction and/or by recycling of copper-containing fluids by heat-induced convection. As to the time of mineralization and the pressure-temperature environment, Brown noted that the major episode of mineralization occurred after the origin of the early diagenetic pyrite and chlorite, as the latter minerals are replaced by copper-bearing sulfides. The copper precipitation has taken place during the early history of the Nonesuch sediments, when the porosity and permeability were greatest to enable the solutions to reach the earlierformed iron sulfide and to replace the iron by copper, for example. The maximum thickness of the overburden was approximately 8000 ft, which is equivalent to a maximum lithostatic pressure of about 650 bars and a minimum of about 240 bars. These pressures would have had a negligible effect on the mineralization processes, but would have been sufficient for expulsion of the connate fluids during compaction. In his discussions of his two models (i.e., infiltration and diffusion), Brown (1971)referred several times to the importance of compaction and stated that the copper-bearing fluids infiltrated the Nonesuch sediments during the early history of the fine-grained deposits when the permeabilities and pore water contents were high. As the effective permeabilities were reduced by progressive Compaction, the copper infiltration may have been succeeded by diffusion and natural chromatography (Brown, 1974). During compaction of the clayey sediments, the downward infiltration carried soluble organic matter from the lower Nonesuch Shale into the older Copper Harbor coarse sediments. Under intense compaction the pore spaces could have been severely restricted to such an extent that the diffusibility could have decreased sufficiently to arrest the process of mineralization. Brown estimated that the episodes of mineralization by the infiltration and diffusion processes averaged 879,000 and 2,010,000 years, respectively. Thus, they were of sufficiently short duration t o assume that mineralization of the lower Nonesuch deposits could have been completed before the permeabilities of the clayey sediments reached prohibitively low values under compaction. Should the earlier-mentioned gravity-stratification model be applicable, the degree and rate of compaction as part of the evolution of the sedimentary basin must be considered in this case also, because the salt-sieving phenomenon becomes operative in the clay-rich sediments at a certain low porosity. Although the precise mechanisms involved in the formation of the White Pine copper deposits have not been unequivocally established, several aspects
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in regard to compaction seem clear. In any genetic theory of copper precipitation, compaction must be considered; it could have had a very fundamental direct influence, only a very indirect influence, or was unimportant altogether. If one can determine through detailed investigation and by using the process of elimination that either one of these three possibilities applies, then one has affirmative and supportive information for a specific genetic theory. In an absence of this data on Compaction, any proposed theory would be weakened. The mineralization was either pre-, syn-, and/or postcompaction in origin. If it can be demonstrated in the future that the ore is post-compaction in origin, then one would not have to consider the significance of compaction fluids, which have already been pressed out at the time of copper precipitation. Nevertheless, even in this case the compaction solutions’ influences on ore formation have to be considered, because they were involved in the various diagenetic processes that determined the pre-ore porosity and permeability of the numerous beds in the sedimentary pile. These mass properties, in turn, would have had a definite control on the invasion rate and pattern of the subsequent ore-forming solutions. If the ore was pre-compaction in origin, which is the least likely possibility as deduced from all the geologic data, except for the early diagenetic pyrite mentioned above, then questions arise about the secondary influences of compaction fluids on the ore minerals, e.g., crystallization, recrystallization and other alterations, and remobilization. The third alternative is the syn-compaction origin. Situated between the pre- and post-compaction stages (Table 5-I),the syn-compaction stage extends over thousands and even millions of years in the evolution of sedimentary basins and has been accepted as the most applicable stage during which the copper of the White Pine ore was formed. Consequently, the influences of compaction waters have found their place in the theories proposed by most investigators: (1)compaction fluids may have supplied metal elements; (2) they may have brought in dissolved organic constituents, catalysts, etc.; (3)these fluids may have assisted in the removal of unused chemical components that otherwise could have prevented chemical precipitation of the ore minerals; and (4) related t o the above, the compaction waters may have mixed with other solutions, resulting in pH, Eh, temperature and concentration changes, which could have been periodically conducive and non-conducive to ore mineralization. A classification of ores as based on their host rock lithology has been proposed by Bogdanov (1968) (see also Bogdanov and Kutyrev, 1973) for stratified copper ores and is given in Table 5-IX. Although much of the White Pine ore occurs in fine-grained sediments or their metamorphic equivalents, a considerable amount of the metalliferous components is present in siltstones and fine-grained sandstones. It is not recommended, therefore, to group the White Pine deposits with those ores occurring either in shales-
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TABLE 5-IX Classification scheme of stratified copper deposits (after Bogdanov, 1968, p. 407) A. Sedimentary deposits: 1. Kupferschiefer type (sedimentary-diagenetic) 11. Coppersandstone type (catagenetic-diagenetic) ( 1 ) Dzhezkazgan subtype (paralic, deltaic); ( 2 ) Ural subtype (continental, alluvial) B. Metamorphosed sedimentary deposits I. regionally metamorphosed 11. contact-metamorphosed
argillites or sandstone host rocks. Even if it would appear to be convenient to do so, it may be the first step in making a serious error in genetic interpretations, because many concepts are developed from an extrapolation and comparison of better known deposits to those for which a genetic theory has to be found. As many hypotheses proposed for ores in siltstones and sandstones differ from those in shales-argillites, this distinction should be maintained at all times. Could the White Pine ore deposits be an example of a transitional type? Will one eventually have to find a theory that borrows from those applied to ores occurring in both clay-rich and coarse-grained sediments? For example, it may be inappropriate on first consideration to suggest that one should take into account many of the genetic ideas used in explaining the origin of the Colorado Plateau and Wyoming uranium-vanadium ores in sandstones in the search for a theory on the origin of the White Pine deposits. Comparative approaches, however, are highly recommended techniques in such cases. By establishing the differences and similarities, one will find that although the tectonic setting (e.g., control of tectonism on sedimentation and facies distribution) and depositional environment (e.g., non-marine fluvial versus marine, deltaic) may be different, there are numerous similarities as t o some of the aspects of paleohydrology (e.g., the relationships between permeability and mineralization) and geochemical factors controlling localization and chemical precipitation of the ore minerals. URANIUM DEPOSITS IN SEDIMENTS
In contrast to certain occurrences of the Mississippi Valley-type leadzinc-barite-fluorite deposits, the concept of compaction and ore mineral precipitation from compaction fluids has not found general acceptance as a plausible process of formation of the uranium deposits in sediments. On the other hand, a number of publications have treated the effect of porosity and
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permeability on uranium deposition within sandstones. Although the main theme of this volume is on compaction, the interdependency among porosity, permeability, fluid movements, and compaction is so close, that at least two publications are discussed here that do not directly refer to compaction. One is on the uranium in the Colorado Plateau sandstones and the other is on ore deposits of the Russian Platform. Many of the discussions presented above on the origin of copper in sediments may also be applicable to the genesis of uranium and vanadium in sandstones. In some cases it has been stated that very high permeabilities and porosities are most conducive for uranium-ore formation - a deduction possibly based on “logical reasoning”. When a detailed, quantitative study was performed, however, it was discovered that not the largest values but the intermediate or even the lowest values of porosity and permeability were most favorable for the ore precipitation. In addition t o the relative and absolute porosity and permeability, other factors, such as grain size, sorting, grainmatrix/cement ratios, particle shape and orientation, and distribution of organic matter also were important. All these parameters would control the compactional history and compaction would, in turn, control some of them. Compaction must have been very significant in preparing certain horizons of sediments to act as conduits and reservoirs for uranium-bearing fluids and to become host rocks to the ore minerals. Compaction in combination with many other factors determined porosity and permeability. Inasmuch as both relative and absolute values of porosity and permeability are important, the following should all have been of significance: (1) the paragenetic compactional history (e.g., time relationship of cementation to compaction, as discussed in Chapter 3 on sandstone diagenesis); (2) the time available prior to the ore mineral genesis; (3) the rate of increase in overburden pressure as a function of sediment supply; and (4) the amount of compaction after ore genesis. McKelvey et al. (1955) listed at least four types of sediments in which uranium occurs, namely: (1) sandstones; (2) coal and carbonaceous shale; (3) marine black shale; and (4) phosphorites. Among the source rocks of uranium, they mentioned tuffaceous beds and granitic and arkosic rocks; however, coal, carbonaceous shale and marine black mudstones could also act as sources if the chemical elements could be released into the compaction fluids during diagenesis. On the other hand, certain geochemical conditions that prevent the release of these elements into interstitial fluids, must have prevailed in numerous instances, because so many of the potential “source” rocks are found t o be rich in uranium and other elements (see Krauskopf, 1955, for example). A similar situation has been mentioned in this chapter in the section on the Mississippi Valley ore deposits and other sulfide occurrences. In regard to the uranium in sandstones, insufficient consideration has
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been given t o the concept of compaction. At present, the majority of investigators suggest that some variety of “ground’’ or “subsurface” fluids was involved in the precipitation of the ore minerals, and the precise nature of these solutions vary from one deposit to another according to the particular author. For example, McKelvey et al. (1955) considered pyroclastic material as a source for various chemical elements and stated that field studies and microscopic examination of the ores suggest a relationship between the ore minerals and montmorillonite clay formed by divitrification of volcanic glass. During divitrification, ground water may have leached alkalies, uranium, vanadium and other elements from the ash. It is not impossible to envisage that this ground water was actually mobilized through compaction (see Chapter 6 on compaction of pyroclastics). Consequently, numerous questions arise, such as: (1) What is the relationship between the diagenetic change of volcanic glass t o montmorillonite and the rate of release of uranium, on one hand, and the rate of compaction, the rate of fluid movement and the amount of compaction fluid, on the other? (2) Is the total volume of pyroclastic material, with a theoretical maximum tenor of chemical elements, sufficient to account for the known amount of ore? McKelvey et al. stated that the large uranium accumulations in sandstones and coal deposits served to emphasize the importance of aquifers in channeling ore solutions through sedimentary rocks and in localizing their precipitates. They also pointed out that quartz overgrowths are abundant on the sand grains, particularly along the edges of the ore bodies and in the wall rocks, whereas within the ore many of the overgrowths are etched or ragged in outline (Waters and Granger, 1953). In reviewing the main hypotheses of uranium ore genesis within sediments, McKelvey et al. mentioned five: (a) a placer origin, followed by recrystallization and redistribution; (b) derivation from surface or ground water at or shortly after deposition of the sediments, followed by recrystallization and redistribution; (c) derivation from the associated volcanic tuffs, as mentioned above, or other sediments elsewhere in the stratigraphic column by ground water action; (d) derivation from petroleum or petroleum source rocks; and (e) derivation from hypogene solutions from igneous rocks (= igneous-hydrothermal), which may have been injected into circulating ground waters. Although each of the above-listed theories appear to be based on independent processes, each can accommodate the concept that compaction fluids may have been involved. Even in cases where the ore minerals were not precipitated initially by compaction solutions, subsequent recrystallization and redistribution may have been caused by compaction. In regard t o the origin of uraniferous, marine, black shales, McKelvey et al. (1955) (see also Vine and Tourtelot, 1970) made some statements that included direct or indirect references to the significance of compaction. In
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vertical stratigraphic sequences, the uranium-rich black shales are associated with phosphatic sediments, especially nodular phosphorites, as well as with chert, carbonates, glauconite, and sandstones. When thick, high-grade phosphorite units are present, however, the black shales are not appreciably uraniferous, and most of the uranium is associated with the phosphorites themselves, rather than being present in the shales. The uranium and other metallic elements were syngenetically-diagenetically precipitated with the enclosing sediments. On the other hand, as a result of the capacity of colloidal and carbonaceous materials to adsorb chemical elements, it is possible that black muds adsorb uranium from circulating waters as long as the sediments maintain permeability after accumulation. Under relatively very slow rates of sediment accumulation, the muds are exposed to the sea water for a longer than normal period of time and, therefore, could adsorb uranium and other metals. A large volume of fluid must be expelled from the black muds on compaction (3, 7, and 15 volumes of fluids if their water content at the time of deposition was 80, 90 or 95%by volume, respectively.) * All metals, including uranium, in these fluids might be removed on passing through layers having environment favorable for their precipitation. Similar to the black shales, much of the uranium in the phosphatic sediments may have been derived from the sea water at, or shortly after, the time the sediments were formed. As a result of the phosphatic sediments’ greater permeability, the uranium may have undergone redissolving and reprecipitation, which was less likely in the case of the clay-rich sediments after some degree of compaction. An important question could be raised here: Could the primary adsorbed uranium in the clay-rich units be released into the compaction waters, which can move into the more porous phosphatic units where the secondary reprecipitation of the uranium can then occur? McKelvey et al. (1955) did not entirely rule out such a mechanism, because they stated that variations in the uranium content in the sediments may be the result of variations in (a) the uranium content of the sea water at the time of deposition; (b) pH, Eh, bicarbonate ion concentration of sea water; and (c) length of time the sedimentary particles were exposed to the sea water prior t o burial. The largest variation in uranium content, however, may be the result of the greater permeability of the phosphatic beds, permitting the passage of percolating solutions, e.g., compaction fluids, after burial from which the uranium is removed by adsorption. The possibility that the secondary remobilization of uranium from clayrich sediments may have been a prerequisite to the origin of the KatherineDarwin ore district was suggested by Condon and Walpole (1955). They proposed a syngenetic origin for this Australian example of uranium mineralization and demonstrated a dominant and widespread lithologic, stratigraphic, and environmental control. The geologic relationships, as depicted in Fig.
* R.F. Beers and Heroy, personal communication, 1951, by McKelvey et al. (1955).
ORE GENESIS INFLUENCED BY COMPACTION Rg. 5-62. Fkconstruction of the sedimentary environment showing position of the uranium occurrences of the KatherineDarwin region, Australia. (After Condon and Walpole, 1955, plate 3; courtesy of Bureau of Mineral Resources, Canberra.)
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5-62, raise the question as to what extent compaction fluids from the euxinic facies, for example, could have supplied the metals to the area of precipitation. One may note here that the biohermal limestone reservoir rocks in this area are surrounded by this facies. It is of some significance that subsequent exploration efforts, guided by their syngenetic theory, helped in locating new ore deposits. Gruner (1956), as did others before him (see previous discussions, p. 513), discussed the possibility that volcanic deposits could have been a likely source of the chemical elements of the Colorado-Wyoming uranium deposits. He believed that pyroclastics in their early diagenetic stages have a high permeability that permits an easy passage of fluids accompanied by rapid leaching and alteration, which result in the removal of uranium even when present in small amounts of only 4 t o 15 p.p.m. Gruner also believed that the removal of the chemical elements probably would not be possible after the formation of illitic and/or bentonitic (montmorillonitic) rocks. This differs somewhat from the opinion expressed by McKelvey et al. (1955). If pyroclastics are to be considered as likely source rocks for uranium and other metallic elements, then future studies will have to determine the precise mechanisms involved in releasing the metals into fluids, because a number of genetic explanations hinge on a plausible process of release. The following questions arise at this point: (1) Can a release of uranium, vanadium, copper and other metals from pyroclastics occur by leaching without a prior transformation of pyroclastics into clay minerals, i.e., is a divitrification of the volcanic glass shards to micro-sized feldspar and other minerals sufficient? The difference may be rather significant because the clay-mineral neoformation may produce an impermeable rock, whereas if the shards are replaced by pseudomorphs then the overall texture and permeability of the rock are not altered appreciably. Thus, in the latter process of alteration, the possibility of fluid movements through the rock could be maintained, while the volcanic glass changes and releases the uranium into the interstitial fluids. On the other hand, during the process of transformation of volcanic glass to clay minerals, the pore space may be progressively reduced until the rock becomes impermeable. Consequently, in this case, the removal of metallic elements can occur mainly during the initial stage of diagenesis while the permeability of the rock is still sufficient for fluids to pass through. (2) To what extent can the montmorillonite, formed as an alteration product from vitric shards, retain adsorbed metallic ions? Of all the clay mineral types, montmorillonite adsorbs the greatest amount of water and metal ions (see Weiss and Amstutz, 1966, and Wolf, 1959, for example), so that upon later compaction, with or without concomitant clay mineral transformation, a release of the metal ions into the compaction fluids can be expected.
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(3) From the above considerations, the obvious question arises as to whether the processes proposed in questions 1 and 2 could succeed each other in time, so that there are at least two major stages of release of metals into subsurface fluids? Further release stages of uranium from the montmorillonite may be controlled by the several stages of clay mineral transformation (see Chapter 3 on sandstone diagenesis). If the release of metal ions from the source rock occurred in stages, one can assume that the formation of the ore body itself proceeded in pulses also, namely, in the form of solution-migration-precipitation, which forms the basis of Gruner’s (1956) “multiple migration- accretion” theory. Inasmuch as the compaction history of a sedimentary basin extends over a considerable span of time and because the metal ions are possibly released in stages from the source sediments, Gruner’s theory should find accommodation within the framework of any concept of ore genesis that is based on the important role played by the compaction fluids. Getseva (1958), for example, mentioned three geologic stages required t o form the uranium in altered sediments, and her information indicates that compaction must have been significant during the early stages of basin evolution. Danchev and Ol’kha (1959) presented information on the effect of porosity on the accumulation of uranium in carbonate sediments. They adopted the concept that the primary concentration of uranium compounds occurred during the syngenetic stage and the amount of uranium accumulation increased during the subsequent diagenesis. The accumulation of uranium in sediments was assisted by a paleogeographic environment which promoted leaching of uranium compounds upon destruction of source rocks on the continent. This resulted in the fixation of uranium in littoral zones, where reducing conditions prevailed because of the concentration of organic matter. In their study of the effect of porosity on mineralization, Danchev and Ol’kha (1959) determined the principal types of carbonate rocks that are ore-bearing, namely, oolitic, skeletal (i.e., fragments and well-preserved shells of gastropods, pelecypods, foraminifera, and fragments of echinoderms, bryozoa, algae, and others), and crystalline dolomite rocks. Of these, the most common ore-containing rocks are the oolitic and skeletal types. A comparison of the effective porosity for the three types of host rocks is given in Fig. 5-63. The average effective porosity for these groups increases from oolitic t o the skeletal and, then, to the crystalline varieties. Inasmuch as intermediate rocks do exist (skeletal-oolitic, ooliticskeletal, etc.), the porosity ranges widely. The most typical representatives of each group, however, are marked by considerably narrower ranges. The results of the investigation showed that the principal ore-bearing carbonate units (i.e., the oolitic and skeletal types) have a marked lower effective porosity in contrast to the crystalline dolomites and calcareous dolomites. A similar relationship has
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been shown to exist in some of the mineralized sandstones of the Colorado Plateau, i.e., the greatest amount of ore occurred in those units that have intermediate values of porosity and permeability, as discussed below. Figure 5-64 shows the quantitative relationship between the effective porosity and the uranium content. The specimens can be divided into three groups, i.e., low, intermediate and rich in uranium content. Zone I contains points from all three types of ore-bearing carbonate rocks, the effective porosity of which ranges from 1 to 15% with crystalline rocks having the highest porosity. These points are chiefly concentrated in the lower right portion of the curve. The ore-bearing skeleton limestones are intermediate in position between the crystalline and oolitic carbonates. Zone 11, located in the middle section of the graph, is represented predominantly by oolitic host rocks with intermediate ore tenor. Only two points in Zone I1 represented one skeletal and one crystalline carbonate rock. Zone 111, which comprises rich mineralized units, is characterized by oolitic mineralized limestones, with an effective porosity that varies considerably between 1and 9%. Figure 5-64 shows a change in the effective porosity of ore-bearing car15-
- 1
-
-
2 X
3
-
-oolitic
0""
1
51
'
1
1
1
10 1
15
structures
20 %
Effective porosity (%)
Fig. 5-63. Effective porosity ( $ J e ) of the principal petrographic varieties of carbonate rocks, i.e., oolitic, organogenic and crystalline. 1 = range and extreme values of $Je for each petrographic variety of carbonate rock; the level of the line corresponds to the average value of $Je for each variety; 2 = range and extreme values of @e for 75% of all specimens anatyzed of each petrographic variety; 3 = average effective porosity for 75% of specimens from each petrographic variety. Arrow shows the change in effective porosity from one variety to another. (After Danchev and Ol'kha, 1959, fig. 2.)
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o 1 oolitic 0 2 organic A 3 crystalline
1
0
1
1
I
I
3
1
1
5
I
I
7
1
1
9
"
1
11
1
13
1
1
15
Effective porosity (%)
Fig. 5-64. Relationship between uranium content and effective porosity of ore-bearing carbonate rocks. 1 = oolitic limestone; 2 = organic limestohes; 3 = crystalline dolomite and calcareous dolomite. I = lean ores; I1 = ores with an average uranium content; I11 = rich ores. (After Danchev and Ol'kha, 1959, fig. 3.)
bonate rocks with an increase in their uranium content. For Group I and the lower portion of Group 11, the lower half of the graph demonstrates an inverse relationship between uranium content and average porosity. Then there is a break in the graph, and for the upper part of Groups I1 and l'II the porosity increases somewhat with increasing uranium content. Danchev and Ol'kha (1959)believed that this peculiar distribution is due to the certain genetic features of the uranium in the carbonate host rocks. The more or less even distribution of uranium minerals in the oolitic limestones is evidently determined by the micro-textural features of the carbonate host rock. The fine uranium components within the oolitic granules constitute the ore, and the distribution of the uranium is closely related to that of fine organic matter rather than being controlled by the original porosity of the three rock types. The oolitic limestones had the highest organic content (5.2%),which probably played a major role in the precipitation (or adsorption) of uranium. It was also shown that the uranium disseminations predominate where the rock is least affected by the processes of recrystallization and leaching.
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An important question that should be asked here is: What is the influence of compaction on the origin and distribution of uranium in the abovedescribed rocks? Although Danchev and Ol’kha (1959) have not mentioned this aspect, one may be allowed to speculate and extrapolate to other geologically related situations. Assuming a carbonate rock with more or less uniformly distributed organic matter, one might expect a corresponding uniform distribution of uranium, if the latter is syngenetic. On the other hand, if the carbonate host rock sequence underwent pre-ore diagenetic alterations, such as recrystallization, leaching, and dolomitization (possibly enhanced by mechanical and chemical compaction processes), then the originally uniformly-distributed organic matter is affected. As a result of differential diagenetic changes from unit to unit, the post-diagenetic rock sequence will have varying amounts of organic matter. Subsequent introduction of uranium by subsurface fluids, possibly compaction solutions, will result in a uranium distribution that should be more or less parallel to the content of the remaining post-diagenetic organic matter. There are several other ways in which compaction fluids could have an influence: (1)Compaction waters passing through the carbonate sediments could remove part or all of the organic matter adsorbed on the grains, thus reducing the total amount of carbonaceous material. Only the organic matter within the grains and, therefore, not easily accessible to the interstitial fluids, unless recrystallization and other diagenetic processes will release the organic components, may be retained for subsequent reaction with the uraniumbearing solutions. Consequently, the paragenetic relationships of numerous possible pre-ore diagenetic alterations are important in determining the final amount of primary organic matter available to the ore-forming fluids. (2) If compaction and reflux solutions caused an increase in porosity and permeability of certain carbonate units as a result of dolomitization and/or dissolution, then the permeability may have been increased above the optimum value for ore mineralization. As mentioned earlier, it is not the maximum, but the intermediate values of permeabilities that are most conducive for ore formation in both carbonate and clastic (or detrital) sediments. (3) Compaction fluids may bring into the host rock sequence the organic matter, which is subsequently available for reaction with the uranium-bearing solutions (Perel’man, 1967). (4)Compaction fluids may have brought the uranium from the source rock (e.g., black, euxinic, basinal muds). The above chosen examples demonstrate that compaction processes and compaction fluids can influence positively or negatively uranium mineralization in sediments. Although this chapter selectively discusses the possible relationships between compaction and ore genesis in sedimentary-volcanic piles, the importance of igneous processes are not to be ruled out. Page (1960), for instance,
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mentioned a widespread relationship between the occurrence of mafic and intermediate igneous rocks and uranium-bearing districts. This suggests that emanations from subjacent igneous masses entered into ground water, which then moved along fractures and faults upward in non-porous rocks and laterally in permeable sedimentary horizons to the sites of ore mineral precipitation. If such igneous emanations occur in a comparatively young sedimentary basin, then both igneous-hydrothermal and compaction solutions may have contributed to the ore genesis. The geothermal gradient, whether controlled by igneous intrusions or not, may mobilize subsurface fluids, including compaction solutions, and determine the temperature of these fluids. In some instances, basinal waters are comparatively hot without having been exposed to igneous emanations and constitute one variety of “hydrothermal” solutions. A detailed regional investigation by Jobin (1962)demonstrated the control of the sandstone “reservoir” rocks on the distribution of uranium in the Colorado Plateau sediments. Although Jobin did not discuss the influence of compaction, others have considered the possibility that compaction fluids may have been the ore-forming solutions. For example, Hostetler and Garrels (1962),in presenting their interpretation of the mechanisms involved in the transportation and precipitation of uranium and vanadium at low temperatures, found it quite plausible that in Late Cretaceous or Early Tertiary time, or even earlier, the heavy overburden of younger deposits squeezed out uranium- and vanadium-bearing solutions from deposits rich in volcanic ash, e.g., Chinle Formation and Brushy Basin Member of the Morrison Formation of Late Triassic and Late Jurassic age, respectively (see also Waters and Granger, 1953). Although the precise nature of the proposed ore-forming solutions differs somewhat from investigator to investigator and it is not known to what degree compaction fluids were involved, the data of Jobin (1962)are presented here to serve as an example of a regional study of porosity, permeability and transmissivity* trends in controlling ore occurrence. When combined with techniques, such as those outlined in Chapter 3 future investigations should result in the desired data that may indicate to what degree mechanical and chemical compaction, as well as other diagenetic processes like differential lithification, have had an influence on the properties of the host rock. The latter in turn would have controlled the differen-
* Transmissivity is expressed quantitatively by the coefficient of transmissivity which is defined as the product of mean permeability and total thickness of the transmitting medium. It is expressed in units of darcy-ft. Transmissivity was used, instead of permeability, to permit comparisons of areal variations of several different units within a specific area. Transmissivity characterizes the capacity of an entire rock unit to transmit fluids.
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tial precipitation of the ore minerals. For instance, compaction waters may be either responsible for cementing sediments or not; depending on the fluid chemistry, or may cause differential cementation over a wide region, as shown by examples presented in Chapter 3. Especially in the latter case, there would be regional variations in porosity and permeability that, in turn, would influence regional ore mineralization, as shown by Jobin. He mentioned that transmissivity decreases towards the source area (i.e., paleolandmass) in the Moss Back Member but increases in the Shinarump deposits. These regional trends in transmissivity, which are a function of sorting and grain-size variations, controlled the amount of cementing material precipitated. TABLE 5-X Comparison of the transmissive character of different sandstones of the Colorado Plateau (after Jobin, 1962, table 6, p. 102; courtesy of U.S.Geol. Surv.) Sandstones of eolian or marine origin
1. Moderate to high mean permeability 2. Small range in local permeability, vertically or horizontally 3. Slight, uniform gradients in regional permeability 4. Generally thick, chiefly one stratum
Sandstones of fluvial origin
1. Low to moderate mean permeability 2. Large range in local permeability, vertically and horizontally 3. Moderate, uniform gradients in regional permeability 4. Each stratum generally thin; units with several strata may attain moderate to great total thickness 5. Generally slight, uniform gradients in 5. Large erratic gradients in local thickness; local and regional thickness generally moderate to large, uniform gradients in regional thickness 6. External geometry simple, commonly 6. External geometry complex; consists of blanket-shaped or wedge-shaped one or more vertically and horizontally contiguous lenticular strata; individual layers are characterized by numerous small areas of local non-deposition 7. Internal geometry is relatively simple, 7. Internal geometry is relatively complex; generally consisting of large-scale, unsysusually consists of one or more lenticular tematic, tangential cross-beds with few, imbricated strata composed of festoon through-going erosional planes and and planar-type cross-beds frequently truntorrential-type cross-beds cating one another and separated by mudstone seams; mudstone pellets and chert granules and pebbles frequently coat lower surface of festoon 8. High local and regional transmissive 8. Low to moderate regional transmissive capacity with slight, uniform local and capacity with large non-uniform local graregional gradients dients, but with slight uniform regional gradient
ORE GENESIS INFLUENCED BY COMPACTION
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108'
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18"
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-
80lsopach. in feet
+
Grid point on arbitrary grid system vdth origin at the point where Colorado, Utah, Arizona. and New Mexico meet (center of circle with 40-mile radius within which the thickness measurements were averaged to obtain a mean value for the grid point)
34'
Fig. 5-65.Isopach map of lower sandstones of Chinle Formation. (Compiled from Gregory, 1917, 1938, Harshbarger et al., 1957,and Stewart et al., 1959,by Jobin, 1962,fig. 9; courtesy of U.S.Geol. Surv., Washington, D.C.)
604
K.H. WOLF
According to Jobin (1962), two types of aquifers account for most of the regional transmissive capacity of the Colorado Plateau sediments: (a) sandstones of eolian and marine origin, and (b) sandstones and conglomerates of fluvial origin. Their characteristics are given in Table 5-X. Most of the known uranium deposits occur in the two major host rocks of fluvial origin, namely, the lower part of the Triassic Chinle Formation and the lower part of the Jurassic Morrison Formation. The data selectively presented here from Jobin (1962) are mainly concerned with the Chinle Formation to serve as an example of his approach. Figures 5-65 t o 5-75 give: (1)isopach maps of the lower sandstones; (2) isopach map of the mudstones and siltstones; (3) isopermeability and isotransmissivity maps with and without superimposed localities of closely-spaced uranium deposits; (4) one map of horizontal transmissivity characteristics as related to the average character at the most productive uranium mining areas; (5) two maps illustrating standard deviation of sandstone thickness and permeability; and (6) two maps classifying the Chinle and Morrison formations, respectively, comparing the horizontal and vertical transmissive character with the average transmissive character of the most-productive mining areas. Thick mudstones and a few thin limestones are relatively abundant, are interbedded with the coarser-grained aquifers, and have little or no transmissive capacity (practically impermeable). They confine fluid movements within both overlying and underlying aquifers. According to Jobin, conduits for vertical fluid movements through any considerable thickness of the stratigraphic sequence are restricted to strongly folded and fractured zones. Maps of the two major host rocks, i.e., Chinle and Morrison formations, illustrate the relationship of the horizontal transmissive character of each host rock to the mean horizontal transmissive character of its most uraniferous parts (e.g., Fig. 5-71). A relatively small portion of the lower sandstone unit of the Chinle Formation is hydrologically similar to its most uraniferous parts, whereas large portions of the Morrison sandstones show such similarities. The preparation of a classification map of the two sandstone formations, based on the mean horizontal and vertical transmissive characters of their most uraniferous areas, resulted in outlining much smaller favorable target areas, where, as suggested by Jobin, exploration should concentrate in the search for new ore deposits (Figs. 5-74 and 5-75). Comparing the distribution of the uranium deposits of the Colorado Plateau with the veriation in transmissive characteristics yielded the following generalizations concerning the major host rocks: (1)they are thin; (2) they had moderate t o low permeability (Fig. 5-67) and low transmissivity (Fig. 5-68), probably throughout their geologic history; (3) they have steep local gradients in thickness (compare Figs. 5-65 and 5-71, for example), as well as steep local gradients in permeability and transmissivity; and (4) they are
ORE GENESIS INFLUENCED BY COMPACTION
605
Fig. 5-66. Isopach map of the mudstone and siltstone unit of the Chinle Formation. (Data compiled by W.L. Newman and E.M. Shoemaker, 1954; from Jobin, 1962, fig. 12; courtesy of U.S. Geol. Sum., Washington, D.C.)
K.H. WOLF
606
EX P L A NA T I 0N
-6-
lxlpleth of permeability, in millidarcy
40'
38"
36"
0
Location of computed permeability profile
+
Grid point on arbitrary grid system with origin at the point where Colorado, Utah, Arizona. and New Mexico meet (center of circle with 4Wmile radius within which the permeability measurements were averaged to obtain a mean value for the grid point)
34"
Fig. 5-67. Isopermeability map of the lower sandstones of the Chinle Formation. (After Jobin, 1962, fig. 10, and table 31; courtesy of U.S. Geol. Sum., Washington, D.C.)
pp. 607-609
I
I
i
#
+
Gri point on arbitrary grid system with origin the point where Colorado. Utah. Arizona. and New Mexico m n t (avorage grid point usad as control)
I
PO'
0"
38'
8'
$6'
16"
14'
34s
Area of closely r'lircr waud uranium daporits
@ %I&
0,
I
t
.
.
50 (
100 MILES I
2
Fig. 6-68.Isotransmieeivity map of the lower sandstones of the Chinle Formation. (After Jobin, 196.2,fig. 11;courtesy of U.S. Geol. Surv., Washington, D.C.)
Fig. 6-69.hopermeability map of the lower sandstones of the Chinle Formation showing dmtribution of areas of closely-spaced uranium deposita. (After Jobin, 1962, fig. 44; courtesy of U.S. Geol. Surv., Washington, D.C.)
Fig. 6-70. Isotranemiesivity map of lower randstones of the Chinle Formation, showing distribution of area of closely spaced uranium deposits. (After Jobin, 1962, fig. 46; courtesy of U.S. Geol. Sum.,Washington, D.C.)
PP. 610-612
114"
I
I
j
i.____ ____
i
wyPMING_-
1120
110'
114'
l0V
I
110"
112'
108'
I
-- _ __ ___ _
UTAH1 COLORADO
LIMIT
EXPLANATION
-
OF STUDY
/
I ?
lsopleth of transmissive character. expressed in natural logarithms; 0 equals identity with standard of references
40'
10.
38'
38'
36'
36'
34"
34'
I
i
A-
Area ot closely spaced uranium deposits 0
-L
50
100 MILES
I
Fig. 6-71.Lower sandstones of the Chinle Formation classified as to similarity of horizontal transmissive character to the average transmissivity at the most productive mining areas. (After Jobin. 1962,fig. 46;courtesy of U.S. Geol. Sum., Washington, D.C.)
Fig. 5-72.Standard deviation of thickness of lower sandstones of the Chinle Formation showing distribution of areas of closely-spaced uranium deposits. (After Jobin, 1962,fig. 47;courtesy of U.S. Geol. Sum., Washington, D.C.)
Fig. 5-73.Standard deviation of permeability of lower sandstones of the Chinle Formation showing distribution of areas of closely-spaced uranium deposits. (After Jobin, 1962, fig. 48; courtesy of U.S. Geol. Surv., Washington, D.C.)
ORE GENESIS INFLUENCED BY COMPACTION
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100 MILES
'
I
100
18"
16'
lsopleth of horizontal transmissive character, expressed in natural logarithms; 0 equals identity with standard of reference 3 -
14'
Area of closely spaced uranium deposits
Area of greatest vertical transmissive capacity; closely stippled area with solid outlines indicates coincidence of horizontal and vertical transmissive characteristics considered most favorable for ore deposits
1
,
1
Fig. 5-74. Lower sandstones of the Chinle Formation classified as to similarity of horizontal and vertical transmissive character to the average transmissive character of its most productive mining areas. (After Jobin, 1962, fig. 59; courtesy of U.S. Geol. Surv., Washington, D.C.)
K.H.WOLF
614
1120
114O
1100
i
i
108"
I
LO'
18"
I 6"
DEWSITIONAL
EXPLANATION
-5-
Isopleth of horizontal transmissive character, expressed in natural logarithms; 0 equals identity with standard of reference
~
14O
closely stippled area with solid outlines indicates coincidence of horizontal and vertical transmissive characteristics considered most favorable for ore deposits
Fig. 5-7 5. Sandstones of the Morrison Formation classified as to similarity of horizontal and vertical transmissive character to the average transmissive character of its most productive mining areas. (After Jobin, 1962, fig. 60; courtesy of U.S. Geol. Surv., Washington, D.C.)
ORE GENESIS INFLUENCED BY COMPACTION
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almost invariably overlain by thick impermeable mudstones. With an increase in the mean volume of uranium deposits and a decrease in the number of deposits, there is a concomitant increase in the mean transmissivity of the host rock. Considering one host rock unit, the range in both the size and the total number of uranium ore localities vary directly with the variations in local horizontal transmissivity and inversely with distance from zones most likely to have vertical transmissive capacity. Noble (1963),in his paper on the formation of ores by water of compaction, mentioned uranium-vanadium mineralization in sandstones of the Colorado Plateau type. According to him, hundreds of these deposits in the Jurassic Morrison Formation are apparently genetically related to compaction fluids. Ores were, most likely, formed from the large volumes of formation fluids, derived during compaction of fine-grained sediments, that passed through the associated sandstones and conglomerates. In Fig. 5-76,Noble (1963)showed the hypothetical source rocks for the major uranium deposits; the latter are totally surrounded by these source rocks. The patterns of
I
S O U R C E
j
1
i
I
\
BEDS
i
n //Belt
I I
I
Fig. 5-76. Relationship of major uranium mining areas to hypothetical source beds of the Colorado Plateau. (After Noble, 1963, fig. 2; courtesy of Econ. Geol.)
616
K.H. WOLF
heavy mineral alteration in the sandstones and the distribution of color bleaching of adjacent mudstones (Fig. 5-77), which is more intense in some restricted sandy zones of greatest permeability, and the relationship between these alterations and the occurrences of ore mineralization, was thought by Noble to be indicative of the passage of the compaction solutions. It should be pointed out, however, that fluids of other modes of origin could have caused similar secondary modifications of the sediments. Nevertheless, this is not sufficient to negate the influence of compaction waters in the origin of the ore. Noble observed that the relatively few large uranium deposits originated mainly near the conduits of greatest permeability and exhibiting greatest alteration. This deviates from Jobin’s (1962) observations given above. According to Noble, the large deposits were precipitated only in places where the compaction waters moved in large quantities. Similar solutions, that permeated less permeable and thinner sandstones, formed the thousands of minor ore concentrations distributed over wide areas. It is difficult to visualize that fluids from an external source, permeating the sediments after compaction and lithification, could have been responsible for such large numbers of widely-distributed small uranium deposits. On the other hand, just
Fig. 5-77. Relation of the Morrison Formation ore bodies to thick sandstone and bleached mudstone. Generalized from an actual occurrence in the Uravan mineral belt, Colorado. (After Noble, 1963, fig. 3 ; courtesy o f Econ. Geol.)
ORE GENESIS INFLUENCED BY COMPACTION
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such an external source has been proposed by many investigators of which the more recent are Rackley (1972)and Harshman (1972),for example. Adler (1964),in his paper on the conceptual uranium ore roll*, suggested that the superposition of a geologically recent or present-day ground water drainage on a paleodrainage system incised in the sediments may have been responsible for the origin of the uranium concentrations. He did not consider “water of compaction” as a likely source for the chemical elements. On the other hand, Shawe and Granger (1965)pointed out that there are certain differences between the Wyoming and Colorado Plateau types of uranium deposits, a distinction which Adler (1964)did not make; these dissimilarities were then further treated by Fischer (1970).It is important to note here that the Wyoming-type ores are relatively younger (i.e., they formed long after the formation of host rock) and originated near the surface after major structural deformation, uplift, and denudation. They are, therefore, more likely related to the surface conditions in that the ore has been precipitated adjacent to a body of oxidizing water of immediate meteoric derivation. Redistribution of earlier syngenetic-diagenetic uranium, possibly related to
* Rolls, solution fronts or geochemical cells, as they have been variously called, are epigenetic alteration features with a definite configuration, and are present in permeable sandstone units. The rolls are the product of moving ground water and are tongue-shaped in vertical cross-section, with a C-shaped (or compressed crescent) termination, separating altered and unaltered sandstone. From the convex side of the smooth and sharp edge of the C-shaped contact, which points in the direction of fluid movement, the tongue-shaped oxidized alteration zone “follows behind” and the narrow front points down the hydrostatic gradient. The rolls extend vertically through a sandstone unit and the front is conspicuously discordant to the bedding, indicating their secondary origin. Such a roll or cell is in effect a dynamic geochemical system with advancing oxidizing zones or fronts with marked changes occurring in Eh, pH, mineralogy, chemistry, and type and amount of microorganisms. A cell originates essentially at a point in a sandstone and expands to form a continuous, three-dimensional, finite body. The shape of such a roll and associated ore bodies is determined by gross permeability and the availability of pyrite and carbonaceous material. The ore and altered rocks probably were formed by the action of ore-bearing solutions that contained oxygen and uranium. These solutions percolated through the permeable host rocks and reacted with the pyrite and organic carbon compounds already present in the sandstone. The dynamic process continued down-dip in the sandstone and the migration of uranium occurred in stages in the order of oxidation .+ dissolution +. precipitation +. reduction. The richest uranium content occurs near the edge and decreases gradually to trace amounts in front of the concave side of the cell. The ore occurs in relatively unaltered sandstone in contact with the edges of the altered tongues and invariably is enriched in iron sulfide minerals, as well as minerals that contain U, Se, V, Mo, Cu, Ag, and Cr in minor amounts. The sandstone unit, with its trace to minor amounts of metals, therefore, is the source of the components, that upon enrichment through the action of “ground water”, form the ore concentrations. The uranium is found at the sides and ends of the tongue-shaped zones of altered rock. (For detailed descriptions, see Fischer, 1970; Warren, 1971; Rackley, 1972.)
618
K.H.WOLF
compaction water movements, may or may not have occurred. In contrast, the Colorado Plateau uranium precipitates are relatively old (i.e., formed shortly after the host rocks accumulated and before or during maximum burial, but prior to major structural deformation) and are not genetically related to uplift and late denudation (Shawe and Granger, 1965). This age relationship suggests that water of compaction could not have been responsible for the uranium ore accumulation of the Wyoming type, but must be considered as a possibility in the Colorado Plateau ore varieties. Fischer (1970) listed the possible ultimate sources that have been suggested for both of these two types of uranium ores, which are similar or nearly identical: (1) Solutions from rocks, e.g., granite, in uplands. These rocks were also the principal source of the clastic material in the feldspar-containing sandstones constituting the host rocks of the uranium ores. (2) Solutions from the arkosic material within the sandstone host sediments. (3)Solutions from the volcanic ash, most of which is associated with the host sediments. For both the Colorado Plateau and Wyoming types of ores, therefore, external and internal sources have been advocated. Although Fischer, as late as 1970, did not indicate a preference, he did point out that in the four uranium-bearing areas investigated by him, of the three suggested source rocks presence of volcanic debris is the only lithologic characteristic common among them. Nevertheless, Fischer concluded that more data on the possible source or sources and, therefore, on the mode of fluid mobilization and derivation, might be helpful in guiding the search for new mining districts. The influence of compaction in the formation of uranium ore, therefore, has not yet been ruled out, as can be deduced from the above considerations. In&much as the C-shaped (occasionally also S- or E-shaped) uranium rolls are indicative of the direction of fluid movement (exceptions have been described by Shawe and Granger, 1965),a detailed regional investigation should be conducted on the relationships between the directions of the ancient subsurface fluid motions, on one hand, and (a) the structural attitudes of the host rock, (b) the general configuration of the sedimentary basin, and (c) the location of possible source beds, on the other. From the data thus obtained, one should be able to determine whether water of compaction or descending fluids from the surface, after tectonic movements, for example, gave rise to the ores. Rackley et al. (1968) mentioned that burial protected the uranium-containing sediments from erosion and provided a cover capable of maintaining the hydrodynamic system which developed in the host rock following tectonic adjustment subsequent to burial. When considering the influence of tectonism and the inception of the geochemical cell, Rackley et al. proposed that tectonism subsequent to burial was required in the evolution and development of a uranium district. The event may be folding as a result of igneous intrusion, or orogenic or epeirogenic folding which gave rise to ground water
ORE GENESIS INFLUENCED BY COMPACTION
619
movements through the aquifers. A complex physical, chemical, and biochemical reaction caused alterations of the rocks; the zone of reaction is a solution front or roll which is called a geochemical cell (see footnote on p. 617).The uranium and other elements are concentrated by such a cell moving along in the host rock bed to concentrate the ore bodies. Thus, there appears to be no need to resort to compaction and compaction fluids in the search for a plausible genetic theory. It is significant to note, however, as has been done in many other studies, that only the geologic features which are observed today have been interpreted, i.e., exclusively the lust geological processes and parameters have been considered. For example, it is assumed that tectonism was responsible for the inception of the geochemical cell and the origin of the uranium ore roll. Although this is the most plausible and traditional way of approaching geologic studies, one should not ignore Gruner’s (1956)multiple migration-accretion theory which offers a whole sequence of events of which only the last stage left a visible imprint accessible to direct visual investigation. If, however, an attempt is made to understand not only the last stage but also the first and intermediate episodes of basin evolution and sedimentation, then one must take the possible role of sediment compaction under scrutiny. During diagenesis, water of compaction may have supplied the uranium which was precipitated as primarily uneconomic disseminations in the sandstones, where the uranium remained for a relatively long period of time until tectonism folded, and possibly fractured, the sedimentary rock sequences in the basin. A new hydrodynamic system was then established which was fed by surface waters, as suggested by Rackley et al. (1968),for example. As to the supply of the uranium under these latter conditions, there are several possibilities (see Fig. 5-79):(1)surface waters from the source rocks outside the sedimentary basin may have extracted uranium from the regolith and soil during surface weathering; (2) the clastic, often arkosic, debris may have supplied the uranium which was released into the meteoric subsurface water; (3) the fluids passing through volcanic rocks may have leached out the metals for the ores; and (4)the earlier-formed, syngenetically and/or diagenetically disseminated uranium, which was precipitated from compaction solutions, was subsequently reworked by the ground water. The sources 1 through 3 have already been mentioned earlier, whereas source 4 appears to have been less often given its due attention. In the case of any one of these four alternative explanations, or a combination thereof, uranium may have undergone dissolution, transportation and retransportation in cycles that resulted in the economic concentrations of ore deposits. The episodic advance of the mineralization may have occurred according to Gruner’s “multiple migration-accretion” theory as rolls. One might mention here the study by Warren (1971)in which he merely mentioned “ground water”, without being more specific, as being the
K.H. WOLF
620
Sandstones and siltstones
8[ T I r:
,
,
Coarse-grained elastics and tuffs partly altered (hydrothermally)
F I
Basement rocks
(L
0
'
-+
1
I
5
10
15
20
25
2
4
6
8
10
I
1
a Intensity of uranium mineralization (in relative units)
--+ b Effective porosity
in per cent
Fig. 5-78. Relation between intensity of mineralization and porosity of rocks in a uranium deposit confined to Mesozoic complexes, U.S.S.R. (After Pelmenev in Vol'fson, 1968; from Ruzicka, 1971, fig. 23; courtesy of Geol. Surv. Canada.)
agent responsible for the uranium mineralization. He has gone beyond the results of others, however, by presenting criteria by which it is possible to discriminate between biogenically and chemically formed ore-stage pyrite associated with the roll-type uranium. Ruzicka (1971),mentioning the work of Pelmenev (1968),showed a correlation between the intensity of mineralization and physical properties of the rocks, i.e., a relationship between intensity of mineralization and porosity (Fig. 5-78). The most favorable host rock occurs near and above the granitoid basement. In this particular case, according to Ruzicka, it seems that the uranium is of hydrothermal origin and the fluids apparently moved into the porous sandstone. It must be pointed out, however, that unless unequivocal evidence for an igneous-hydrothermal origin is available, the possibility of hot, non-igneous subsurface fluids, such as compaction water from the associated fine-grained sediments and volcanics moving in an environment with a high geothermal gradient, must also be considered in having played a part in the origin of these ore deposits. The localization of the ore in tuffaceous sandstone lying above an igneous basement, may be explained by the ore-bearing fluids travelling within the permeable clastic unit above the relatively impermeable basement. Basinal fine-grained mud may have supplied the compaction water, whereas both the mud as well as the pyroclastic debris could have furnished the uranium upon leaching. The
ORE GENESIS INFLUENCED BY COMPACTION
621
siltstone-rich upper unit shown in Fig. 5-78, in this instance, acted as a “tight” cap rock. Similar situations have been observed by the writer from Queensland, Australia, during drilling operations, where evidence of fluid movements within a paleo-karst in Cambrian carbonates above the Precambrian basement were noted and sulphide precipitation appeared to be confined along this unconformity. Numerous references can be found in the literature to the movements of hydrocarbons that were controlled similarly, i.e., where the basement rock was fractured and had an upper unit of paleosoil (e.g., “granite wash”), the hydrocarbons penetrated downwards. In each of the above cases, the basement controlled the localization of the economic deposits, because it formed an impermeable barrier to horizontal and downward-moving solutions. Gabelman (1970), in considering the metallotectonic and, therefore, the regional control of uranium distribution, mentioned that because compressive orogeny is mostly crustal, the connate and ground waters and water of crystallization released during neomorphism of minerals, are mobilized by differential tectonic pressures as well as temperatures. These mineralizing solutions have sufficient corrosive power to extract mobile elements from the rocks and, under favorable geochemical conditions, may serve as oreforming fluids. Thus, although compaction per se is commonly considered to be a diagenetic process, its extension in both space and time, to include the influence of compressive tectonism and certain styles of metamorphism (see Table 5-1, and Chapter 3) even millions of years after accumulation of the sediments and volcanics, is warranted in the study of ores formed by “water of compaction”. Finer subdivisions have to be made in the future, based possibly on an approach similar to that offered in Table 5-1. In the International Atomic Energy Agency symposium entitled “Uranium Exploration Geology” (1970), Polo (1970, p. 243) suggested that the opposite of compaction and compressive tectonism, namely, “tectonic relaxation”, could result in “decompaction” and decementation of rocks. The porosity and permeability thus created, in addition to that formed as a result of fracturing, could make certain rock horizons available to fluid movements and, therefore, have an influence on the localization of ore deposition. A multiple source for the uranium in the sandstones of Wyoming was suggested by Harshman (1970, 1972), which included the granite in the cores of the mountain ranges surrounding the basins and the pyroclastic or tuffaceous beds that occupy or once occupied the eastern two-thirds of Wyoming. The water from the surface moved down the flanks from the cores of the mountains into the basin where precipitation of the metal took place. The delivery of the uranium from the tuffaceous material may have been accomplished by water of compaction. As a way of comparison, Harshman mentioned that in the past few years roll-type uranium deposits were found
622
K.H. WOLF
in Eocene t o Miocene littoral and fluvial sandstones of the Gulf Coastal Plain area of Texas that contain considerable amounts of tuffaceous debris (Eargle e t al., 1971). These latter uranium deposits may have had an origin similar to those in Wyoming; the pyroclastic debris supplied the uranium to compaction fluids, which in turn precipitated the metal in concentrated form elsewhere in the sandstone aquifers. Some possible fundamental differences between the two ore districts, however, are considered here. In comparing the Wyoming with the Texas uranium deposits, one must take into account the regional environmental and the tectonic settings of each in establishing a likely source for the uranium. If the volcanic components constituted the source in both cases, then the ground water and/or compaction fluids transported the uranium from the source to the sandstone to be concentrated there. If an additional source is to be sought for, then the leaching of the uranium from the granite and the subsequent precipitation of this element from ground water may be a plausible explanation for the origin of the Wyoming-type deposits; this may not necessarily be the case in the Texas deposits. In the latter instance, however, another factor should be considered that is not applicable to the continental Wyoming uranium mineralization, namely, that the offshore compaction fluids may have moved into the littoral and fluvial sandstones t o precipitate the uranium. If these offshore compaction waters did not supply the uranium, then one at least should question the effect they may have had on the host rock and on the subsurface fluids’ chemistry because of mixing of various intrastratal waters, especially in nearshore marine sediments. In addition, the chemistry of compaction fluids changes upon migration (G.V. Chilingar, pers. comm., 1974). If uranium rolls are formed by descending meteoric waters in the Wyoming districts, then one might assume that ascending compaction fluids could also produce roll-type uranium structures. As pointed out earlier above, the shape of such rolls are directional indicators of the movement of both descending and ascending subsurface solutions. Thus it is hoped, that in future detailed investigations the overall movement patterns of subsurface fluids can be established by using directional indicators, which, in turn, would provide some insight as to the locality and the type of source rock(s). The information could also be useful in determining whether compaction fluids were the ore-forming solutions. It is suggested that future research be directed towards finding a number of indicators reflecting fluid movement patterns. That within one particular region uranium deposits can occur in both older and younger host rocks, has been pointed out by Mittenpergher’s (1970)study. In the European deposits studied by him, uranium occurs in the deeper levels in the thicker parts of the post-Hercynian sedimentary and volcanic rocks, and has an epigenetic character. On the other hand, in the much thinner Permo-Triassic sequence of the Pennide zone, uranium concen-
623
ORE GENESIS INFLUENCED BY COMPACTION
URANIUM-CONCENTRATION EPISODE I
L
URANIUM-CONCENTRATICX EPISODE
II
-
POST- DEFORMATIONAL STAGE
STAGE
I
1 \\I-
Fig. 5-79. Conceptual model of the origin of uranium in sandstones (K.H. Wolf, 197273). For details, see text.
trations occur mostly in the upper parts of the continental sediments, at the transition to the epicontinental marine lithologic formations. Based on the above discussions, the present writer has prepared a simplified conceptual model (Fig. 5-79) of the uranium genesis of the sandstonetype ores. The model has the following features: (1) There are two possible uranium-concentration episodes, i.e., one occurring early during the filling of the sedimentary basin and prior to tectonism (i.e., intra-basinal stage) and the second one occurring after the deformation of the sedimentary and volcanic sequences (i.e., post-deformational stage). (2) Each sedimentary depocenter, receiving detritus, has an extra-basinal (= outside), terrestrial source (box I) that upon weathering supplies lithic grains and pebbles, various sand and silt grains (boxes I11 and IV), and, depending on the degree of chemical weathering, various clay minerals to
624
K.H. WOLF
the basin of accumulation (boxes VI and VII). Volcanism may take place on land (box 11) and pyroclastic debris would then be moved directly through the air, or indirectly by rivers, into the sedimentary basin (box V). All the types of detritus described above may contain a certain amount of elements, such as uranium and vanadium, for example. (3) In addition to the non-volcanic (= epiclastic) and volcanic (= pyroclastic) debris delivered from the land, these same sources may provide chemical elements to surface and subsurface fluids (line 1A) whenever deep weathering and leaching takes place. These elements may or may not be precipitated upon reaching the sedimentary milieu. (4) The deposits, which accumulated in the depocenter (ie., boxes I11 to VII), could have been composed of grains that contain traces of uranium (e.g., in arkoses, lithic sediments, and pyroclastics) and/or clay minerals with adsorbed uranium (e.g., in muds and claystones). ( 5 ) Compaction waters (plus, possibly, some meteoric fluids and, during late diagenesis, also water of crystallization released as a result of clay mineral transformations) moving from the fine-grained muds interbedded with the sandstones and conglomerates in a fluvial sedimentary complex, as well as from deeper basinal sediments, may carry uranium released from volcanic ash and clay minerals. Additionally, compaction fluids moving through permeable sandstones, conglomerates and coarse pyroclastics may leach out uranium. Within the coarser-grained aquifers, these compaction solutions may precipitate uranium wherever organic matter is present and provides a local reducing environment. The uranium mineralization is either in the form of fine disseminations or as ore-grade concentrations (box VIII). The process may have been cyclic in nature, namely, dissolution, transportation and reprecipitation may have occurred many times before the flow of subsurface water was terminated for a number of possible geologic reasons. Theoretically, this would complete the ore-concentration episode 1 in the idealized conceptual model presented here. ( 6 ) The same coarse-grained sediments that upon leaching can supply the uranium to the subsurface fluids, can also act penecontemporaneously as “reservoir” rocks for the ore mineralization. For this reason, a clear separation of “host rock” (boxes I11 and VIII), on one hand, and “source rock”, such as arkoses (box IV), for example, on the other, is not possible. In other words, the separate boxes should not mislead the reader to assume that the host rock of the uranium ore is different from the coarse-grained sediments listed in the model. (7) The scheme in Fig. 5-79 is by necessity very generalized and incorporates at least two varieties of uranium deposits, namely, Colorado PlateauWyoming and the Texas uranium types depicted on the left-hand side in the diagram. The former were formed in a continental, intermontane basin with-
ORE GENESIS INFLUENCED BY COMPACTION
625
in the deposits of a fluvial-lacustrine complex, whereas the latter originated within sediments that accumulated under nearshore and marine conditions (Eargle e t al., 1971). Thus, the types of source and host rocks are controlled by the total geologic setting. (8) Although syngenetic ore concentrations in general are known in sediments, the uranium deposits discussed here are of a “secondary” origin. Even early diagenetic precipitates in sand are secondary in comparison t o the detrital sand grains which accumulated prior t o the introduction of the chemical precipitates. Late diagenetic products in sediments, as known well now, show all the characteristics of both early diagenetic and epigenetic processes (see, for example, Vine and Tourtelot, 1970).For this reason, it is often difficult, if not impossible, t o distinguish between diagenetic material and that formed much later (usually called “epigenetic”). Nearly all uranium deposits in sandstones have epigenetic features, but the precise time of formation as related to the host sediments (i.e., whether or not early diagenetic and, therefore, nearly contemporaneous with the sand grains) is not discernible in many investigations, if textures, fabrics, local structures, and stratigraphic properties are considered. (9) As a consequence of folding, jointing, faulting, and subsequent subaerial erosion, a new geologic situation is established, which is different in most respects from those that existed under the original sedimentary milieu. A new surface and subsurface hydrologic system is initiated, with meteoric water playing the major role. There are several alternative geological settings that are significant in uranium genesis: (a) The tectonism resulted in folding and fracturing of the sediments that already contained some diageneticallyconcentrated uranium. The meteoric subsurface fluids passing through the folded beds (lines 11,13, 14, and 15 in Fig. 5-79) dissolved, transported and reprecipitated the uranium, forming “roll” structures. Repeated, cyclic movement of the uranium eventually resulted in an economic ore deposit. (b) As a result of tectonism, intermontane basins are formed that were fringed by crystalline rock outcrops, such as granite. Chemical weathering may supply uranium to the surface meteoric water which, upon reaching the sedimentary rocks, penetrated the latter (box IX, Fig. 5-79). The subsurface meteoric water may have picked up additional uranium from the arkosic sandstones and pyroclastics while passing through them. Hence, there is a transition to the conditions mentioned in case a. Upon reaching a chemically reducing milieu, these meteoric waters precipitated the uranium, possibly in the form of “rolls”. (c) If the mechanical erosion of the crystalline rocks (box X) fringing the intermontane basin formed a new sedimentary complex, then these fine- to coarse-grained deposits would have been exposed to the diagenetic processes which were already mentioned under items (1)through (8). All the discussions presented earlier, therefore, are applicable here again.
K.H. WOLF
626
SEDIMENTARY IRON ORES
Dimroth (1968) and Dimroth and Chauvel(l972,1973), in their study of Precambrian oolitic Superior-type ironstones, discussed the sedimentary-tometamorphic textures, diagenesis, and environments of formation of these deposits. Dimroth employed the nomenclature and classification scheme of limestone petrographers (Folk, 1957 and 1959) and adopted it to the iron sediments: (1)orthochemicals: a-femicrite, b-matrix, chert, and c-cement chert; (2) allochemicals: a-pellets, bintraclasts, c-oolites and pisolites, and d s h a r d s (but no fossils or skeletons for the Precambrian iron ores). As shown in Fig. 5-80, the constituents of the iron formations were precipitated from the sea water and the composition of the iron and silica components was related to the Eh and pH of the depositional medium. Subsequent intrabasinal reworking produced the second- and third-cycle intraclastic sedi-
t
silicanel
iron oxide
siderite
hydrate
mud
I f
I
/
1
\/
hematits dust
I
precursor si
1
micro-pranrlir sidjtite
specularite
t
megnetlto c-recrystallized siderite
t
micropolygonal quartz
1
specularite
recryslallized quartz
specularite
I
quartz
I
racrvstall ized siderite
reaction with At-silicates
reticulated magnstita
recrystallized quartz
1 recrystallized
magnetite
nltite
maghemite
recrystallized
magneti to
and martite
dissolulion
Fig. 5-80. Epigenetic processes in the Precambrian Sokoman Iron Formation, Labrador Trough, Quebec, Canada. (After Dimroth and Chauvel, 1972, fig. 5 ; courtesy of Geol. Rdsch.)
ORE GENESIS INFLUENCED BY COMPACTION deposition of allochsm grains (inlraclisls and ooliths)
627
t
desiccation cracka
I /
silicael
0
primocfyrtallization of wmiz
by chalcsdoy and columnar quartz
I cemented and crystalline cherts
I
I
Fig. 5-81. Lithification of the Sokoman Iron Formation. (After Dimroth and Chauvel, 1972, fig. 6; courtesy of Geol. Rdsch.)
ments. As diagrammatically shown in Fig. 5-81, lithification of the iron formation was accompanied, or even controlled, by compaction. Fluids driven by the compaction mechanism supplied the chemical components required for lithification and may have removed others. Figure 5-82 illustrates the paragenetic sequence of the Precambrian Iron Formation of the Labrador Trough in the Precambrian Shield of Quebec, Canada. Dimroth and Chauvel (1972) were among the few investigators who have examined the effects of compaction: (1)Desiccation, shrinkage, cementation, and compaction are part of the early diagenetic stage. The stylolites can form both during the early and late diagenetic stages. (2) Crystallization of silica gel in the following sequence took place near the end of the early diagenetic stage and extended into the late diagenetic episode: a-primocrystalline quartz and chalcedony (normal and lengthslow), b-micropolygonal quartz, and c-recrystallized quartz. The latter extends into and may be modified during the metamorphic stage. (3) Similar sequences, as given above, have been established for the crystallization of hematite dust, migration of iron and crystallization of specula
K.H.WOLF
628
Silicate-Carbonate Iron Formations
Hamatite Iron Formations
dropr
1
adsorbed iron oxide hydrate
aggregation
t
pllets
1
femicrite
Matrix" chert
I _ INIRABASIN EROSION
I d l i c a t e and carbonate)
(chert)
I
First
Matrix" chert
cycle rock
aggregation
of silicagel
I
c /
Second cycle rocks
c
INl,QA#ASIN EROSION
Compler aggregation of
siliT1 complex
Fig. 5-82. Deposition of the Sokoman Iron Formation. (After Dimroth and Chauvel, 1972, fig. 7;courtesy of Geol. Rdsch.)
ORE GENESIS INFLUENCED BY COMPACTION
629
hematite and of magnetite, and (4) for the crystallization and recrystallization of siderite, minnesotaite, stilpnomelane, riebeckite and talc. The migration of iron and the recrystallization of specularite, magnetite and siderite took place essentially before the complete compaction of the sedimentary units and before and during the primocrystallization of quartz. The recrystallized quartz, minnesotaite, and possibly, stilpnomelane and talc are essentially data diagenetic minerals, whereas riebeckite originated during metamorphism. A certain amount of recrystallization occurred during regional metamorphism. From the above information one can deduce that the discussions on the compaction of limestones (Chapter 3, Vol. I) and sandstones (Chapter 3, Vol. 11) may well be applicable to oolitic, pelletic, and related iron ores of the Clinton, Superior, Minette, and similar types (as well as to texturally comparable phosphatic ores), possibly with minor modifications of the concepts to adopt them to sediments of an overall different composition. All coarser-grained sediments, including the ores considered here, may be composed of various grain types as well as different proportions of matrices and cements, which, in turn,control the porosity, permeability and the compactability of these deposits. Consequently, with a change in compactability there is a concomitant alteration in the amount and rate of flow of compaction fluids passing through the sedimentary units. The whole diagenetic history, therefore, may change from horizon to horizon, so that the trace and minor element composition of the iron oxide, for example, as well as the degree of oxidation and reduction of the iron, will be related to the compactability. Dimroth’s contributions, among others, can serve as a stimulation and guide in the future mineralogical, textural, and geochemical investigation of the influence of compaction on sedimentary ores. Ellison (1955), in his publication on economic applications of paleoecology, demonstrated the use of various techniques used in the fields of “soft rock geology” in the search and genetic interpretation of ores within sedimentary and volcanic piles. Although neither he nor Hunter (1970) discussed the influence of diagenetic compaction on the iron ores of the Silurian Clinton Group, and although nothing could be found in the published literature, the present writer would like to make a few related comments and, especially, to raise a few pertinent questions. Figure 5-83 illustrates the regional pattern of the iron ore environment which grades eastward into a clastic and westward into a carbonate environment. Four cross-sections and one map (Figs. 5-84 and 5-85) show the more detailed lithologic and stratigraphic relationships between the iron ore formations and the associated units. Such a varied lithologic assemblage would lend itself particularly well to a detailed regional investigation of diagenetic facies, which, of course, would include the study of the influence of compaction fluids. Here too, one could apply
630
K.H. WOLF
Fig. 5-83. Distribution of Clinton (Silurian) iron ores as related to the Silurian environments. (After Caster and Caster, 1950, McCallie, 1908, Lowenstam, 1950; from Ellison, 1955, fig. 2.) Cross-sections by Hunter (1970) are superimposed on the diagram;see AA' to E-N'; courtesy of John Wiley and Sons, New York, N.Y., and Econ. Geol.)
the techniques and concepts described in Chapter 3 (on sandstones) and in Vol. I (on limestones). The following questions related to compaction can be asked here: (1) What was the effect of the compaction fluids on the origin of the various orthochemical grains and their possible subsequent diagenetic alter-
ORE GENESIS INFLUENCED BY COMPACTION
631
Fig. 5-84. Lithofacies map showing the occurrence of iron minerals in the upper part of the Lower Clinton and the lower part of the Middle Clinton. Most of the ironstones occur in or at the contacts of the Wolcott Limestone of New York and its approximate equivalents elsewhere. For details, see Hunter, 1970, and Fig. 5-85below. (After Hunter, 1970, fig. 7; courtesy of John Wiley and Sons, New York, N.Y.)
ations, and on the formation of interstitial diagenetic-authigenic cements and matrix, on both the local and regional scale? Considering that: (a) the regional variation in lithology is from argillaceous through sandy, glauconitic and chamositic limestones, hematitic limestones, oolitic chamosite and
UPPER
LOWER
ClRBOhaTE
IREINALES
SHALE
YEM
I LS I
500 FEET
EXPLANATION -400
- 300
-
0
SCALES
n
-200
KM
60
40
20
@
100 20
4 0 MILES
-
SANDSTONE
SMILE CAROONATE ROCK IRONSTONE
Fig. 5-85. Stratigraphic cross-sections of the Clinton Group and equivalents in the Central Appalachian Basin. Locations of the lines of cross-sections are shown on Fig. 5-83. (After Hunter, 1970, fig. 4; courtesy of John Wiley and Sons, New York, N.Y.)
ORE GENESIS INFLUENCED BY COMPACTION
633
oolitic hematite to more pure limestones as shown in Fig. 5-84; (b) the fact that each one of these facies must have had intrastratal fluids of different chemical composition (cf. results of studies on Recent unconsolidated sediments), possibly with different assemblages of chemical-producing bacteria; and (c) the complex stratigraphic interbedding of different lithologies, as depicted in Fig. 5-85, the following other questions arise. (2) What were the relative fluid movements during diagenetic compaction? (3)What were the chemical reactions, if any, during mixing of compaction and interstitial fluids? (4) What changes occurred in the trace and minor element contents as a result of release into and adsorption from the passing compaction fluids? (5) What is the relationship between the iron ore facies and the hematite cement? Were the orthochemical grains composed of iron oxide, etc., of syngenetic origin in equilibrium with the Fe-rich interstitial fluids, and were the grains formed where they are found now in the stratigraphic position, or were they transported there? Did the Fe-rich intergranular fluids upon compaction move into the sandstones to precipitate there as iron oxide cement? The answer to these types of questions could be extrapolated to any of the other facies of the Clinton Group. A hint of the existence of diagenetic facies is given in Fig. 5-84, where a change from the silica- to the hematitecemented sandstones is presented. Also, the hematite grains would have formed under oxidizing conditions, whereas the chamositic and glauconitic types originated under a slightly reducing milieux. The subsurface fluids in these different environments would also differ, so that, first, the early diagenetic processes would be different and, second, the compaction fluids from each of these sediments would also show various chemical dissimilarities. Upon reaching other units, the compaction fluids probably would precipitate dissolved material depending on their chemical composition. Studies, such as those suggested above, related to the mineralogic, textural-fabric, and paragenetic facies, as done by Warner (1965) on cements, are essential here. These must be supplemented and supported by information on: (a) the geochemical milieux required for the particular minerals to form as indicated by the phase diagrams (e.g., Garrels and Christ, 1965); (b) the compositional ranges of surface fluids in various Recent sedimentary environments, and compared with (c) the compositional ranges of the early to late diagenetic interstitial waters within Recent sediments; (d) the results of chemical reactions which could occur during mixing of fluids of different composition; and (e) the hydrologic factors, e.g., models of fluid movement patterns must be established. Attempts should be made to supply the sedimentologists and stratigraphers with “conceptual models” on subsurface fluid movements; first with simple ones, but which can be modified with any increase in the available data.
634
K.H.WOLF
Trendall and Blockley (1970) studied the banded iron formations (BIF) of the Precambrian Hamersley Group in northwest Australia. Their work was concentrated on one stratigraphic member of one of the eight formations. This member was subdivided into 33 numbered macrobands, which are up to 82 f t in total thickness and are recognizable in the field. Accurate regional thickness measurements were made of thirty-one sections. An isopach map showed, in general, an ovoid pattern. Within the BIF macrobands there are mesobands, which average less than an inch in thickness and are composed of chert, chert-matrix* (mixture of fine-grained quartz, carbonate, and iron oxides), magnetite, stilpnomelane, carbonates and miscellaneous types of other minerals. The first three components make up more than 90% of the macrobands. Most of the chert mesobands are internally laminated or microbanded, each being about 1mm thick. One of the mesobands has certain characteristics that define the Calamina cyclothem. The macro- and mesobands have basin-wide continuity and have been influenced by compaction as described below. The microbands within the chert mesobands are composed of regularly alternating Fe-rich and Fe-poor laminae within finely crystalline (average size of crystals is 20 microns) quartz mosaic. The ironbearing minerals are ankerite, hematite, magnetite, and stilpnomelane. Microbands vary in thickness up to a maximum of approximately 4 mm and show wide variety in textural and compositional detail. The microband interval increases and decreases systematically t o define regular strips which can be correlated over a distance of 185 miles. The mesobanding mostly can be considered as geometrically ideal planar layering only. Lateral discontinuity in chert mesobands occurs as several types of spatially confined chert bodies, i.e., random pods, cross-pods, small spheroidal nodules and macules (=, abrupt local thickening of a limited stratigraphic thickness of the BIF). Slumping and other minor collapse structures are present as well as rippling. Trendall and Blockley (1970) have alternative genetic interpretations for the microbands, and the interested reader may consult the original publication. Only the information related to compaction is selectively discussed here. The striking regularity, even in thickness, and lateral continuity over 20,000 square miles suggest that the microbands are seasonal, possibly an-
* The dark grey, greenish, or brown (when fresh) chert-matrix is the characteristic enclosing material of the chert pods, and is also the common dark separating material between successive pale buff, grey or pink (when fresh and unweathered) chert mesobands. The chert-matrix is composed of 30--50% quartz in contrast to the chert that usually contains more than 60% quartz. The chert-matrix is finely crystalline and lacks the regular microlamination of cherts. Instead it has a vaguely-defined and ghost-like irregular striping due to variations in the magnetite content. As a result of decrease in the quartz and an increase in the iron oxide content, there is a complete gradation between the magnetite mesobands and the chert-matrix mesobands.
ORE GENESIS INFLUENCED BY COMPACTION
635
nual, varves. This concept, however, does not explain the presence of the mesobands of chert-matrix and magnetite in between microbanded layers. As the number of microbands above and below the chert-matrix is not constant, the chert-matrix must have formed at erratic intervals, assuming that each microband is a seasonal or annual indicator. To explain the discontinuous presence of the chert-matrix mesobands, Trendall and Blockley simply assumed that the chert was once continuous laterally and that the neighboring chert-matrix was formed by the compaction of originally thicker layers of micro-laminated chert. This was the first step in a proposed modification of the basic hypothesis that the micro-laminae of chert are seasonal varves. As a second step, it was suggested that all chert-matrix owes its origin to the compactional modification of chert. If this is factual, then the third step follows, namely, that all cherts were derived from compaction of originally thicker silica precipitates. The fourth step in finding a plausible theory, includes the assumption that all three mesoband types (i.e., chert, chertmatrix, and magnetite), which together constitute 90% of the volume of the BIF macrobands investigated in detail, owe their origin to compaction and diagenesis of an evenly and regularly laminated comparatively coarselyvarved parent precipitate. This hypothesis accommodates all the observed textural details of the microbands. It requires precipitates of credible chemical composition (i.e., that could occur as a natural precipitate), especially in regard to its iron content, and no drastic changes in the geochemical parameters within the sedimentary environment with time need to be invoked. But one main question remains unanswered, i.e., why did the different varve sequences respond to compaction so differently? Trendall and Blockley offered the following explanation. The accumulation of iron as a thin skin over the basin floor each year appears to have occurred without interruption, as suggested by the lateral continuity and the absence of primary discordant contacts. This process occurred for a very long period of time without extrabasinal elastic grains being supplied into the sedimentary basin. It would seem that the reason for the differential reaction of the groups of microbands to compaction lies in the slight, primary chemical or physical differences, controlled by changes in the depositional environment. As a result of these changes, occurring periodically rather than haphazardly, there is some regularity, which in turn permitted the definition of cycles or cyclothems, exemplified by the Calamina cyclothem. All the features mentioned above were accommodated by a number of salient points in the hypothesis by Trendall and Blockley (1970): (1)A second-order climatic control, giving rise t o the cycles or cyclothems was superimposed on the annual, seasonal, climatic cycle. The latter controlled the varving of the parent material, which upon compaction became microbanded chert, chert-matrix, or magnetite.
636
K.H. WOLF
(2) Superimposed on the second-order cycle was a longer climatic cycle, or possibly a double cycle, of the order of 1000-3000 years. It is of less significance here in the discussions on compaction. (3) A threshold value for some climatic factor (designated as x ) existed, above which compaction was disproportionally inhibited, i.e., when during a second-order cycle this climatic factor was raised above the threshold value, the compaction process was such that microbanded chert was formed, A
D
B
.
Chertmagnetite group ,Mixed grow
.Chertmagnetite group Mixed 'grow
<
Chert>magnetite
Chert mesobonds Striped lacies Chert mesobands Striped lacies Chert mesobands
I Fig. 5-86. Diagrams illustrating a possible cyclic climatic control of mesobanding. For detailed explanation, see text. (After Trendall and Blockley, 1970, fig. 71; courtesy of Geol. Surv. Western Australia.) Abscissa (x) in graph A represents some unspecified factor of mean annual climate; numerical value and sense are irrelevant. Ordinate represents geologic time in years. Line 1 = second-order rhythm, here arbitrarily shown as 40 years; dashed line 3 = threshold values for x; columns B and C = depositional stratigraphic results (see text); black = material liable to extreme compaction; white = material showing potential chert.
ORE GENESIS INFLUENCED BY COMPACTION
637
whereas if the material was precipitated when this climatic factor was below the threshold value, there was subsequent compactional collapse (see discussion below for details). (4)Compaction of the varved material, which precipitated under conditions below the threshold, was intensified in proportion t o the abundance of potential chert, i.e., the varved deposit tended t o compact evenly in the absence of potential chert. This hypothesis is applied in Fig. 5-86.All varved material deposited at times when the climatic factor x was above the threshold value, was highly resistant to compaction and may have become chert, whereas material deposited when x was below the threshold value was severely compacted t o either chert-matrix or magnetite. The stratigraphic column B shows the uncompacted depositional result of these suggested climatic and sedimentary depositional conditions. Stratigraphic column C represent column B after compaction, with 80% of the black material and 20% of the white material removed. The proposed equivalents to one of the cyclothems (i.e., the Calamina cyclothem) is given on the right-hand side of column C (compare with fig. 13 in the original publication). In the lower part of column A, a simple climatic fourth-order cycle of a period of about 1300 years is given, whereas the same cycle in the upper part has an intermediate central irregularity superimposed. This is a possible mechanism for the absence or presence of chert mesobands in the cyclothems. If one raises the threshold value to a level indicated by line 4 in column A, column D in Fig. 5-86 would be the result. The same results would be obtained if the whole curve moved down relative to a stationary threshold. The possible course of compaction of sediments in column D is shown by column E. The degrees of compaction are assumed t o be similar to those from columns B to C. The basic hypothesis by Trendall and Blockley suggested that Fe-bearing and SiO 2 - r i ~ hvarves were, after accumulation, compacted differentially in groups to form mesobands, i.e., degree ofcompaction varied from one group of varves to the next. The general trend of the chemical changes that occurred during compaction is shown in the curve in Fig. 5-87,i.e., from the upper left down to the lower right. As shown here, the Fe-rich chert-matrix lost, among other compounds, a great deal of Si02, whereas the finely microbanded chert lost smaller amounts of SiO 2. More thickly microbanded chert lost the least amount of SiOz during compaction. The SiOz was then redissolved, and was transported upwards by the expelled compaction fluids t o the waters of the sedimentary basin. In order t o overcome (1)the passage of fluids through a rather dense material, (2) the redissolving of S i 0 2 in interstitial water shortly after it has been chemically precipitated, and (3) the return of S i 0 2 to the same Si02-saturated water from which it was derived, another modification of the basic hypothesis pas offered by Tren-
K.H. WOLF
638
+
.
\ % .
.-. -------_ me
\
X
1
I+-
SILICA
PRECI?ITATED
SILICA
STABLE
,
5 E0 IM EN1 S I L I C A - FREE WATER
LXPfLLED
h
Fig. 5-88. Diagram illustrating the suggested movement of silica during the diagenesis of iron formations. (After Trendall and Blockley, 1970, fig. 73; courtesy of Geol. Sum. Western Australia.) ,
ORE GENESIS INFLUENCED BY COMPACTION
6 39
dall and Blockley (see Fig. 5-88). According to their quantitative calculations, the original wet precipitate contained 8%Si02, 4% Fe203, 2% FeO, 1%C02, and 85% H20. The latter includes combined and interstitial water, assuming that half of the water was combined in the particulate gel, while half was interstitial. According to the modified hypothesis, each varve group was subjected to increasing pressure on burial and underwent compaction. Some groups, however, reacted abruptly, at an unknown threshold level mentioned above, which is marked by the transition from chert to chertmatrix. During this transition, a proportion of the S i 0 2 was redissolved and streamed upwards through the less compacted and highly permeable overlying sediment (Fig. 5-88). This sediment bathed by the ascending SiO2-rich solution may have absorbed more SiOz, thus raising the Si02 content per microband above that originally precipitated. The movements of this compaction water may be responsible for the partial disturbance of the internal laminations of some microbands. The rate of compaction (i.e., rapid versus gradual) probably had a marked effect on the results of compaction and the chemistry of expelled solutions (see Rieke and Chilingarian, 1974, p. 253). The second difficulty mentioned above, namely, the dissolving of Si02, is less severe, inasmuch as the above discussions on fluid movements show that it is not necessary to assume that the SiOz was carried in solution right back to the region of its original precipitation. Instead, Trendall and Blockley (1970) suggested a zone of change in Si02 stability situated between levels A and B in Fig. 5-88. The reasons for this change are not known, but three suggestions were made: (1)C 0 2 is another constituent which may be removed in the upward streaming fluids during compactional diagenesis. According to the physicochemical nature of the precipitated material, C 0 2 may be lost at a higher level than SiO2 and the stability of the latter may be affected by consequent Eh and pH changes. One must keep in mind, however, that as fluids move upward, COz may evolve as a result of decreasing pressure and carbonates can precipitate. Escape of COP may result in higher pH, which is conducive to Si02 solution. (2) It is possible that a biochemical process was involved in controlling precipitation, so that an assumption could be made that the zone AB in Fig. 5-88 represents the depth of burial below which the chemical balance of SiO is not subject to biological (photosynthetic?) control. (3) The precise chemical make-up of the precipitate is unknown, i.e., whether the silica was in an uncombined or simple hydrated form. There is even the possibility that the initial precipitate was, wholly or partly, sodium silicate, because the formation of chert from sodium silicate (= magadite) has been reported by Hay (1968), Eugster (1969), and O’Neil and Hay (1973). The source of the Na required to form riebeckite, which originated by meta-
K.H. WOLF
640
somatic replacement of chert at a fairly late stage of compaction, as described by Trendall and Blockley, may be, therefore, explained also (see below). The originally precipitated sodium silicate may have broken down under a certain pressure of burial, liberating both S i 0 2 and Na simultaneously, because colloidal sodium silicate is known to be pressure-sensitive. The widespread occurrence of riebeckite in the iron formation was explained by Trendall and Blockley (1970) as being due to two possible diagenetic mechanisms: (a) either massive riebeckite has been formed by replacing BIF, or (b) both formed a common parent material. In order to be able to choose the most likely process, it is fundamental to determine the paragenesis (= time relationships) between the riebeckite and compaction (Fig. 5-89)of which there are three alternatives. (1)In row B in Fig. 5-89, the original thickness of material (column 1)is locally compacted (column 2) and the uncompacted part is replaced by riebeckite (column 3). (2) In row C, the same original thickness is uniformly compacted and subsequently replaced by riebeckite with expansion.
C
t
a
1
C
e/.
1
3
.:. ........,...: :. t . .... , , ,:i,.'.!.:;.'!': b
.
A:..: . . . . ...... ... .
.....
-------- _ _ _ _ _ _ _ d
d
Fig. 5-89. Diagrammatic stratigraphic relationship of massive riebeckite and iron formation, and three theoretically possible sequences in its development; see text for details. (After Trendall and Blockley, 1970, fig. 75; courtesy of Geol. Sum. Western Australia.) Columns 1, 2, and 3 = time sequence of three alternative genetic possibilities (B, C, D); rows B, C, and D = three alternative genetic possibilities; b = original thickness; a = thickness after compaction; stippled area = riebeckite of secondary origin.
ORE GENESIS INFLUENCED BY COMPACTION
64 1
(3) In row D, the original layer is partially converted to riebeckite, followed by compaction. Alternative 2 is discarded because subsequent swelling of a riebeckite body should show evidence of distension at the sides, top and bottom, with micro-folding and micro-jointing of the adjacent BIF, and such features have not been observed. Thus, 1 and 3 remain as possible alternatives. If the sequence in row D is correct, then one would expect some marginal fracturing and bending of the earlier-formed riebeckite, but none is present. That alternative 1 is the most probable one is supported by observations on duplicate swells which may or may not contain riebeckite, i.e., the replacement by riebeckite preferentially affects those parts of some chertmagnetite groups which have compacted less than the average. The last example of the relationship between compaction and iron-bearing mineral genesis is provided by McGeary and Damuth (1973) and Norton et al. (1974) (see also Mothersill and Shegelski, 1973), who have described an iron-rich crust containing a number of metal trace elements (Mn, Cu, Ni, Co, Zn, Cr) from a deep-sea environment. The crust separates terrigenous hemipelagic lutites and turbidites from overlying pelagic foraminiferal oozes and lutites on the deep-sea abyssal plains investigated. There is a marked change in color and oxidation-reduction state of the sediments: (a) the terrigenous lower part is generally gray and reduced; (b) the crust is yellow or rustcolored and oxidized; and (c) the pelagic lutite above is tan and oxidized. McGeary and Damuth (1973) considered four genetic mechanisms, i.e., origin similar to that of manganese nodules, volcanic emanations, terrigenous deposition and solution of iron in intrastratal fluids and precipitation after expulsion to the surface. These investigators concluded that the solution of iron from within the unconsolidated sediment with reprecipitation at the sediment-water interface appears to be the most probable explanation. The turbidites contain the iron-bearing minerals such as biotite and chlorite. (In this case it seems to be detrital chlorite, not to be confused with that formed during burial metamorphism, so common in geosynclinal greywackes of turbidite origin.) Inasmuch as the reduced iron is much more soluble than oxidized iron, any plant material decaying in the hemipelagic lutites and turbidites could reduce and, therefore, dissolve the iron from the terrigenous minerals (an example of “halmyrolysis” or subaqueous weathering). The reduced iron would be carried upwards by the compaction waters during sediment compaction. Upon mixing with the sea water lying above the sediment, the iron would be oxidized and reprecipitated, forming a cement that would bind the sediment into a hard iron-rich layer (Fig. 5-90). The fluid movement pattern in this case corresponds to that shown in Fig. 5-53,Bl. That dissolution during burial of earlier-formed, syngenetic chemical precipitates is possible, has been mentioned by Tucker (1971) in his study of Recent and Devonian ferromanganese nodules.
K.H. WOLF
642
NORMAL PELAG I C SEDI MEN TATION
p I
a
*
P.
..
. . 1
4
Fig. 5-90. Mechanism of crust formation. 1. Plant-bearing turbidite is deposited. 2. Decaying plants reduce iron, which dissolves; compaction fluids carry reduced iron upward in solution. 3. Reduced iron comes in contact with sea water, oxidizes, and precipitates, cementing pelagic ooze. 4. Compaction is over; normal pelagic sedimentation buries crust. (After McGeary and Damuth, 1973, fig. 6; courtesy of Bull. Geol. SOC.A m . )
CONCLUSIONS AND SUMMARY
Those readers requiring a quick summary of the relationships between compaction and ore genesis, will find a comprehensive overall view in Figs. 5-52, 5-79, 5-91, 5-92, and 5-93 and in the eight Tables 5-1 to 5-VIII, and 5-XI. Also, the section entitled “Delineating and defining some compaction problems” offers information in summary form. Only a few supplementary statements are made here supported by two comprehensive genetic conceptual models. (1)There is little doubt that at the present time the theory of ore genesis by compaction fluids is a plausible one. Among the many publications discussed here, those by Beales and Jackson (1966, 1968), Jackson and Beales (1967), Jackson and Folinsbee (1969), and Beales and Onasick (1970) on the Pine Point ore field in northwestern Canada typify a trend that is developing at an increasing rate in the study of ore bodies, that is, to view them as a part of the total stratigraphic history and as a normal result in the evolution of a sedimentary basin. Intraformational solutions, for example, can collect and transport metals in ionic or complexed ionic form. The brines,
ORE GENESIS INFLUENCED BY COMPACTION
643
which can be derived from evaporitic minerals or rock sequences, are mobilized and moved by compaction. The following question must be answered by future research: What are the relative proportions of the metals derived from the evaporite deposits and those supplied by the fine-grained basinal sediments, which are more commonly suggested as being the source rocks? Brines can be derived from the compacting basinal sediments after "maturation" has occurred, as well as from evaporites. (2)In general, the sources of the metals constituting ore minerals in sediments and associated volcanics (Figs. 5-28to 5-30,5-52,5-91,5-92, and 5-93) are either extra- or intrabasinal: (a) Terrigenous rocks that upon weathering change to a soil rich in clay mineral and organic matter absorbed* and/or adsorbed trace and minor chemical elements. The sand-sized grains of the soil may also contain trace elements. Upon erosion, transportation and accumulation of the soil components, fine-grained and coarse-grained sediments are formed that contain varying amounts of elements such as copper, lead, zinc, and uranium. Dissolution of the elements and their salts in the terrigenous area may result in solutions bringing them to the sedimentary environment where precipitation of ore minerals may take place (Fig. 5-92).
* In the opinion of one of the authors (G.V.C.) based on research work, very minor amounts of metals are present in "absorbed" form in clays and only about 10% are "adsorbed''. The remainder of metals mobilized by compaction waters are present. in finegrained sediments as precipitates, such as carbonates. EXTRA-
INTRAB/\SINAL
PRIMARY PROPERTlES CF SEDIMENTARI IN40 VOLUNlCl PlLE
wnicu AND
RI'6ICOMEMICU PROPERTIES
7
CHEMICAL COMPONENTS
1
Figure 5-91. Simplified flow-chart showing the sequential factors in the extra- and intra-basinal environments that influence the physical properties of a pile of sediments, e.g., compactability, and therefore, the remobilization of chemical elements. (See Fig. 5-93for a more comprehensive model.)
K.H.WOLF
644
INTRABASINAL
I‘i L
SOURCE - A
EXTRkondior INTRABASINAL SCURCE
-c-
... -5-
. ,
~. -8-
=
Fig. 5-92. All-inclusive conceptual model illustrating the source and host rocks of ores commonly considered, and the numerous genetic relationships between them (see “Conclusions and summary” for details).
(b) Normal and saturated sea water and lake water have been proposed as possible sources of chemical elements required for the formation of ore deposits. The ultimate source of these elements within the surface waters may be either terrigenous rocks drained by rivers or volcanic-exhalative processes (Fig. 5-92). These surface waters, of course, can become interstitial
PP. 645-650 ~~
:POSITIONAL TERRIGENOUS SOURCE AREA
PRIMARY PROFERTIES OF SEDIMENTARY ( 2 PYROCLASTIC) PILE
ENVIRONMENT
ARINE ENVl RONMENT iysicAL a BSTRACT
PRIMARY ( r O R I G I N A L ) STRATIGRAPHIC PROPERTIES (l0COl torcglonol) e g . r a t i o of impermeable to permeable beds. c han nels, lensing. Interbedding. ~ o n t l n u l t y . b o s m w o r dond sourceword t r e n d s cyclic n o l u r e
3
I PRIMARY ( : O R I G I N A L ) SEDIMENTARY STRUCTURES t h ~ c k n e s s . r o t l o sof Iitnologies.cross-bcdding.
grading.sole-mark:lng,etc
I
t
I
MARINE WATER PHYSICOCHEMICAL B IO L O G IC A L
CLASTICS C l a y - to pebblesized
======i
====D
PRIMARY ORE DEPOSITS IN SEDIMENTARY-VOLCANIC PILE AFTER COMPACTION (subsequent metamorphism not considered here). IF%
I
REA BETWEEN 80URCE A N D E PO S ITI 0 N A L
I -
SYNGENETIC-DIAGENETIC
PROCESSES
a
OF
I
FACTORS
PHYSICAL cOmpa~t~on.des~CCat~On.shr~ n koye. pencontemporoncars intern01 deformation COrrosion.mechon~CaI internal s e d i m e n t o t i O n . g O s - b U b b l ~ ~ ~ .ond ~ ~ ~re-transs~o~ portotion and redeporttion !n n-cycICS.odSOPptlOn and obsorp110n.elc BIOPHYSICAL gaS-bUbbltng. borlng.burrowlng.mlxlng and homogenlzlng of sediments.corrosion.a~~~ct~on ond surfc aggrrgotion.porticle- s i z e reduction.etc B I O C H E M I C A L accretion and aggregotion.corrosion.gasbubbllng.breohlng down and synthesizing of Orgonlc and Inorganic comoounds.ctc PHYsCOCHEMICAL s o I U t i o n . o c ~ i p i t o t ~ o ~ , ~ o ~ ~ a ~ i o n ,
proc
///a I l '-
PROPERTIES
REMOBlLlZATlON OF TRACE AND MINOR METALS IN SEDIMENTS CONTEMPO RANEOUS WITH COMPACTION.
545
17
leac~ing.bleaching.oxldo(~on-reductlon.~nwerslon, ~ ~ ~ y s 1 a 1 1 ~ 2 0 t 1 0 n . n C O m o r ~ h ~ s m . c e m e nond tot~on decementa1iOn.~~10c~ment,chem ical Internal
I
scdlmcntat~on.occretlon.aggrcgation.ctc
7
I
Fig. 5-93. Comprehensive model presenting the numerous factors and processes (and their complex interrelationships) that directly and indirectly control the compactability of a pile of sediments and associated volcanics. Included are the factors of the extra-
6
I
subsurface processes basinal source arca and the intrabasinal depositional environment that, in turn, determine all the primary ploperties of the sediments and, through them, the compactability. (After K.H. Wolf, 1972-73.)
ORE GENESIS INFLUENCED BY COMPACTION
65 1
TABLE 5-XI Various types of ore deposits that had their metal components derived from a terrigenous source rock (the metals were eroded and transported by chemical and mechanical processes )
ores, and land-deposited volcanics)
-+ B.
4 1. higher concentrations
muddy sediments, clays with adsorbed metallic ions, i.e., 1st-cycle concentration (usually trace to minor elements only)
2 C. volcanic detritus with trace metals
of metals in mudstones, shales; 2ndcycle concentration -+ 2. ditto in sandstones d -+ 3. ditto in carbonates 4 4. ditto in pyroclastics ’
% I. volcanics (pyroclastics)
9 2. sandstones % 3. carbonates
b
-+
~-
~
~
D. sandstones: arkoses, lithic sandstones, and muddy sandstones with trace metals
5 1. pyroclastics 5 2. sandstones $ 3. carbonates
~~
a = transportation in solution and as a colloid; b = erosion and transportation; c = diagenesis with dissolution, transportation, and reprecipitation; d = metals are released into compaction fluids, transported, and reprecipitated; e = metals are released into “ground water”, which may or may not include compaction fluids, transported and reprecipitated. Only occasionally are economic ore deposits formed in this case (A), because they occur mainly as disseminations. These trace and minor metal concentrations, however, can be later released from the sediments into subsurface waters (B, C, D), including compaction fluids, which then move into the permeable host rocks. If the milieu is conducive there, the metal is precipitated. Three lithologic types of host rocks have been listed in each case. The table is only an example and similar ones can be prepared t o supplement Fig. 5-92 for the sources 11,111,IV, and V.
fluids within the sediments and, when mobilized during burial in the basin, constitute the compaction waters. (c) A number of intra-basinal deposits may contain trace and minor elements (boxes A t o E, Fig. 5-92)”, that upon leaching and corrosion can be released into subsurface fluids, e.g., into compaction solutions, and upon
* It is significant to realize that the sediments in boxes A to D (Fig. 5-92) can act as both source and reservoir (= host) rocks, depending on numerous geological and geochemical parameters, as discussed in the text of the chapter.
652
K.H. WOLF
transportation and precipitation can form ore deposits. In particular, muddy sediments have been commonly considered as a source of elements (e.g., see Vine and Tourtelot, 1970). Pyroclastics and certain varieties of sandstones, in particular arkosic types, also have been considered as possible source rocks. It seems less in vogue to suggest carbonates as source rocks, but recently it has been proposed that limestones and dolomites could also supply trace elements for ore minerals (e.g., see Roberts, 1973). In addition to chloride brines, it seems that evaporites should be given more serious consideration as a likely source of metallic elements. Vine and Tourtelot (1970), for example, suggested that black muds can act both as a source of and as a local “sink” for trace elements. Under conditions of very slow sedimentary accumulation, the clay minerals are in direct contact with surface waters for a long time and elements could be adsorbed progressively. Suggestions have been made that petroleum may supply the constituents necessary for the formation of ore, but the degree to which this is occurring under natural conditions remains to be tested in the future. Although the main topic in this chapter is the ore genesis as a result of non-igneous processes, one should not lose sight of the fact that in particular the theory of volcanic-exhalative source of metals has gained wider acceptance during the past few years. The “exhaled” metal-rich fluids may reach the surface and become intermingled with sea and lake water. The volcanic solutions may also become mixed with normal ground water and compaction fluids. One always has to count on a hybrid origin of the oreforming solutions. Inspite of the real, even overwhelming, importance of volcanic-exhalative source of the metals for ore formation, the main concern in this chapter remains the establishment of the likelihood of a direct sedimentary basinal or a direct terrigenous source of the metals, as proposed, for instance, by Jackson and Beales (1967)and Strakhov (1969)and Samama (1973),respectively. One could always argue, that on going back far enough in the geologic history of the earth, all metallic components were derived from a crystalline, terrigenous source (aside from the volcanic-exhalative) by subaerial weathering and transportation to the bodies of water (lake or ocean). But it is the more immediate stages of the chemical evolutions or cycles of Cu, Pb, Zn, U, Mo, Fe, etc., that concern the practical geologists. It is important to know whether the components of the ore minerals in the sediments came directly in solution from the offshore basinal sediments or from the terrigenous source. It is possible that the fine-grained basinal sediments originally were derived from a land-mass also, and brought the metals adsorbed on the clay minerals to the basin. But the latter extrapolation, although important, is not quite as significant in determining the final, or latest, geological situations under which the ore components were transported and precipitated. It is the youngest geological and geochemical
ORE GENESIS INFLUENCED BY COMPACTION
653
factors that determine the localization of the ore deposits. If metals were derived from the offshore basinal sediments, then compaction would be important. In the case of the direct terrigenous source, compaction fluids may also remobilize the metals. (3) The various processes of mechanical and chemical compaction (Table 5-1) begin to take effect very early in the history of sediments, i.e., as soon as particles accumulated and formed an overburden. Compaction, in its broad-> est sense, continues for thousands or millions of years and may even pass into the stages of tectonism and metamorphism (Fig. 5-7). Different compaction processes may proceed at varying rates in different geologic settings, which accounts for many of the variations between ore bodies. A number of compactional mechanisms may be operative contemporaneously or successively, sometimes accompanied by other diagenetic processes. For example, the mere physical squeezing out of fluids during tectonism may be associated with transformations of clay minerals, so that in that case two types of fluids are mixed, e.g., pore fluids per se and the chemically-bound water released from the minerals. In regional basin analyses or in detailed petrographic work, i.e., whatever the scale of the study, one should attempt to find clues for determining the compactional history and subsurface paleo-fluid migrations. (4) Fluid movements in the subsurface (for simple examples, see Fig. 5-53) may change throughout the history of a sedimentary basin, as a result of a variety of geologic factors. It has been demonstrated, for example, that differential compaction may reverse the direction of fluid movements in the subsurface. Important also would be the time relationships between the compactional history of the sediments and the occurrence of fracturing, faulting, and folding of the stratigraphic sequence. Differential changes in porosity and permeability of the different fine- and coarse-grained sediments, as a response to chemical and mechanical compaction, may determine the interrelationships between the permeable “aquifers” and tight “cap rocks”. Tectonism has a drastic-effect on the overall subsurface plumbing system of large regions of sedimentary rocks. It may: (1)cause the reversal of the directions of fluid movements; (2) change the rates of migrations of fluids that may decrease or increase; (3) cause fracturing which may create new vertical and horizontal conduits; (4) result in partial to complete dewatering of whole sedimentary units; and (5) influence the formation of karst systems. In addition to the aqueous solutions lost during the dewatering, as a response to tectonism, hydrocarbons may also be flushed out. A whole spectrum of hydrologic conceptual models is required to assist geologists in unravelling the paleo-plumbing system of sedimentary basins responsible for the origin of metalliferous concentrations (and hydrocarbon accumulations).
654
K.H. WOLF
The investigations of younger basins, e.g., Tertiary basins in Florida, California, and Japan, as has been done already to some degree, will be most helpful in the development of these hydrologic conceptual models. Attempts should be made to find directional indicators for subsurface fluid migrations. Whenever the regional geology in the vicinity of the ore body is known (and assuming that the ore originated from compaction fluids and that the source was intrabasinal), geochemical mass balance calculations should be undertaken to ascertain that the types and volumes of lithologies (e.g., muds versus sands) can account for the amount of fluids and chemical elements required for the formation of the ore deposits. A few case histories available have suggested that basinal sediments could have been the source for all the metals of an ore. An attempt should be made to answer the following questions: (a) What are the various possible evolutionary trends in subsurface fluids with geologic time, such as diagenetic fluids + petroliferous fluids -,highsalinity metalliferous brines + meteoric fluids? (b) What are the relationships between clay-mineral transformations and types of fluids released? (c) What are the relationships between the mineral neomorphisms in general (including clay-mineral transformations) and the possible sequence of metallic ions released into subsurface fluids? (d) What are the interrelationships among the clay-mineral transformations, types of fluids released, sequence of metal ions released, and the paragenesis of ore mineral precipitation? (e) What are the influences of shock (e.g., earthquakes) on the rheological properties of thick sedimentary and volcanic piles and on the sudden release of large volumes of metal-bearing solutions, as proposed by Elliston (1963)? ( 5 ) A number of investigators have proposed and some have actually observed a relationship between the origin of hydrocarbons and metalliferous concentrations in sedimentary basins (Table 5-VI, and Figs. 5-2 to 5-7, and 5-9 to 5-15). So far, this relationship appears to be “observational” and mainly interpretive, although some acceptable genetic hypotheses have been offered. As in many other problems related to ore genesis, however, the precise details remain to be established in the future. Although there is always a chance, never to be excluded totally, that the oil and ore are coincidentally associated and, therefore, genetically unrelated, it seems that there is either a direct, or maybe somewhat indirect, relationship between the occurrence of hydrocarbons and ores in a basin. A brief exercise in logic, however, will demonstrate that the problem of genetic relationship is not quite as simple as it may appear on first sight, because the key-words are “direct” and “indirect”. (a) If the localization of ore mineralization is controlled by the same
ORE GENESIS INFLUENCED BY COMPACTION
655
factors as the accumulation of hydrocarbons, namely, that both metals and hydrocarbons are “trapped” in stratigraphic and structural traps (Figs. 5-2 to 5-7), then one can hardly consider this as proof of a direct genetic relationship between the two economic deposits. Inasmuch as both were transported t o the traps by fluids, the conduits or routes controlled their migration pattern and the rate of movement. There could be basins with oil and no ore, and basins with ore but without oil, because either one can form without the other as a result of an absence of a genetic interdependency and differences in composition of compaction fluids. (b) If the source rock, for example black basinal muds, supplied both hydrocarbons and metals to the compaction waters, then these fluids could have supplied both the oil and ore constituents to their respective host rocks. This does not prove a genetic relationship beyond both having had the same source, i.e., the muds contained both the hydrocarbons and metals prior to their release into the compaction fluids. There are muds, however, that contain only one or the other. It seems then that there may be a prerequisite in the search for a direct genetic relationship between hydrocarbon and ore genesis, namely, that there is a common source for both, e.g., basinal black muds. The constituents of oil and ore are not necessarily released from the muds simultaneously, but possibly according to an, as yet unknown, paragenetic geochemical timing that depends on pressure, temperature, chemical laws affecting desorption, mineral transformation, fluid maturation, and other factors during basin evolution (see discussion below and Fig. 5-16). (c) There are other relationships between hydrocarbon and ore formation. For example, invasion of oil field brines by fluids containing some organic matter can result in a reducing milieu along a particular stratigraphic horizon. This establishes one of the several possible geochemical “barriers” that give rise to the chemical precipitation of sulfide minerals. Bacterial decomposition of the organic matter may result in the formation of the required H2S. (d) To what degree oil itself can transport metal elements that subsequently are made available for ore formation, remains to be investigated in greater detail. Whatever the final answers may be to the questions on the genetic relationships between hydrocarbons and ores in sediments, there is little doubt that their terminal stage of migration and accumulation are controlled by similar factors. Thus, the various exploration techniques developed by the oil industry are to be considered in the search for ores in sediments. For this reason, the first most obvious step to be taken in establishing the degree to which there is a connection between the genesis and occurrence of hydrocarbons and ore, would be t o summarize all known factors that control the origin, migration, and occurrence of oil and gas (this was done by Weeks, 1961, as presented here in Table 5-VI), and then determine theoretically and
656
K.H.WOLF
pragmatically to what extent the same factors influence the origin of ores in sediments. In other words, the establishment of the similarities and differences of the factors should be one of the first more fundamental studies to be undertaken. (6) Independent of, or rather additional to, the above-mentioned possible relationships between the origin of hydrocarbons and ore mineralization (especially those requiring a reducing milieu), there is the often-quoted need for the presence of carbonaceous material in the host rocks. Compaction fluids can move organic particulate matter, organic acids, and hydrocarbons into otherwise clean quartzites. If some of the organic compounds are retained in the sandstone, bacterial processes can result in a reducing subsurface environment which is one of the requirements for sulphide mineral formation. On the other hand, compaction waters may physically “flush out” the fine carbonaceous matter present within the coarse-grained permeable sediments, so that no reducing chemical milieu can be established. Whatever the exact origin of the organic or carbonaceous matter may be, i.e., whether present as detrital vegetable hash or as adsorbed microscopic particles on clay minerals and silt plus sand grains, or whether brought in as minute particulate matter or as oil droplets, it appears to be one prerequisite for the origin of certain types of ores in sediments (see, for example, uranium in sandstones). Several researchers have attempted to show a relationship between the evolution of organic compounds since the Precambrian or Early Paleozoic time and the first appearance of certain types of ores (e.g., Watson, 1973). Mixing of fluids with varying chemical and physical properties can result in the precipitation of ore minerals (e.g., Runnels, 1969), and this particular field of investigation in geochemistry is in its infancy. Thus, one finds only occasional reference in the published literature to this specific cause of ore mineral precipitation. As so often required in geological studies, both laboratory and field work have to be combined to establish the various “geochemical barriers”” (Lebedev, 1967b), location of which depends on compaction fluid movements. The field observations without the necessary theoretical data for a reasonable explanation are not very useful, because as long as the latter is lacking, the natural observations cannot be properly extrapolated to other geological areas and situations. Consequently, the search for ore deposits must often be done on the trial-and-error basis rather than using reasoned predictions. For this reason, a comprehensive summary of the numerous geochemical and physical barriers is needed by the practicing geologist and a thorough comparative study should be undertaken.
* Geochemical barrier can be defined as a zone where some change in chemical and/or physical property or properties causes the precipitation of minerals.
ORE GENESIS INFLUENCED BY COMPACTION
657
As to the alteration of subsurface fluids, the logical starting point of a comprehensive investigation would be the selection of a sedimentary complex of young deposits that accumulated in different environments. The different facies (e.g., deltaic vs. off-shore; reef vs. clastic sediments; shallowwater vs. deep-water) may contain different connate fluids inherited from the original environments of sedimentation. Subsequent diagenesis may either increase or decrease these compositional variations. If compaction led to movement and mixing of the fluids, chemical reactions may have been possible that otherwise could not have occurred. At any given point during the diagenetic history of sediments, the chemical composition of the subsurface fluids can vary from one extreme to another. This sort of investigation should then be extended from the younger to the older sedimentary units to determine the mechanisms of “aging” or “maturation” of the subsurface waters under specific geologic conditions. (7) There are numerous other indirect influences of compaction on ore genesis (Tables 5-1, 5-111 and 5-V; and Figs. 5-8, 5-16, 5-19, 5-20, 5-21, 5-91, 5-92, and 5-93), in addition to those mentioned above. (a) There is a direct correlation between porosity and permeability of sedimentary units, on one hand, and ore mineralization, on the other. Consequently, logic dictates that inasmuch as there is often a correlation between the degree of compaction and porosity (or permeability), there must also be a connection albeit very indirect, between compaction and ore mineralization. (b) Differential compaction can result in differential reduction of porosity and permeability* which, in turn, will determine preferential ore mineralization along specific beds. It is not necessary for the sedimentary unit to have a very high permebility to be mineralized, instead mineralization occurs in beds having some intermediate values of permeability (see section on uranium ores). (c) Differential compaction can form structural or stratigraphic traps into which the metal-bearing compaction solutions can move and precipitate ore (see section on copper ores). (d) Variations in lithology of the host rocks and the chemical milieu within them may influence ore precipitation. (e) The evolution of compaction fluids within the sedimentary basin (see pp. 574-575) may determine the sequence of release of metals from the muddy sediments into fluids that, in turn, would influence the paragenesis of ore minerals.
* A comparative investigation should be undertaken of those diagenetic processes that do not lead to lithification and reduction in porosity and permeability vis-i-vis those mechanisms that do.
658
K.H.WOLF
(f) Numerous diagenetic processes and variations in compaction fluid chemistry would control the isotopic composition of the ore minerals. (g) The earlier diagenetic history of the sediments may determine their susceptibility t o ore mineralization. For example, it has been suggested that the iron in early diagenetic pyrite was partly replaced by copper, supplied by compaction fluids, to form cupriferous ore concentrations. It is not known what would have been the fate of the copper if the sedimentary host rock had no earlier-formed pyrite. (h) Ore minerals, whether syngenetic-diagenetic or late-diagenetic and epigenetic in origin, or whether related t o compaction fluids or not, can be influenced by mechanical and/or chemical compaction processes and compaction fluids (see section on iron ores, for instance). Oolitic, pelletic and other granular varieties of ores may undergo mechanical compaction, as revealed by textural and fabric features, and their cements may have chemically precipitated from compaction fluids (e.g., silica in intergranular spaces of pellet iron ores). The quantitative data on the relationships among the various properties of sediments and ore locaiization is very meagre, and the information is usually confined to direct correlations between two parameters, such as porosity and ore tenor and burial depth and permeability. As soon as one or two other factors are added, their complex relationships to one another become apparent. Little data has been obtained in the past on the relationships between more than two or three parameters. Nearly all the information on the influences of indirect geologic controls is still qualitative and, therefore, one has to be satisfied with non-numerical conceptual relationships, such as those depicted in Figs. 5-91, 5-92, and 5-93. (8) As to the compaction fluids, there are several aspects t o be considered: (a) The origin of ores from basinal fluids has to be viewed in the light of the total evolution of the sedimentary basin (Jackson and Beales, 1967, and Beales and Jackson, 1968, for example), because ore-forming solutions are available from the earliest to the latest stages of geologic development (Tables 5-1 to 5-IV, 5-VI and 5-VII; and Figs. 5-52, 5-91 to 5.93). Surface and syndiagenetic fluids would be followed by: (1) early t o late compaction fluids; (2) water released during clay-mineral transformations; (3) fluids suddenly made available by possible earthquake shocks that result in rheological changes in the pile of sediments; and (4) solutions mobilized by tectonic squeezing (Table 5-1). It is well known from the observations made by petroleum geologists that in many cases the subsurface waters become more saline with depth, with specific deviations from normal sea water, so that one can speak of a “maturation” of the intra-formational solutions. Maturation is a time- and depth-related phenomenon, being part of basin evolution.
ORE GENESIS INFLUENCED BY COMPACTION
659
(b) The influence of compaction, burial metamorphism, and tectonism on evaporites requires more consideration as they can supply the chloride brines necessary for the dissolution and transportation of metals in ionic and complexed form (Figs. 5-52 and 5-92). Though a general relationship between the presence of evaporites and ore genesis has been proposed by a number of investigators, the exact connections between the two remain to be examined in greater detail. (c) Concomitant with the maturation of basinal subsurface fluids, there is also a time-, depth-, temperature-, and pressure-dependent evolution of the carbonaceous material in the sediments and the origin of oil and gas within the basin (see Chapter 3). It seems appropriate to determine whether there is a maximum depth of origin of certain types of sedimentary ore deposits that depends on specific varieties of organic matter, because: (1)organic matter is often required for the precipitation of sulphide ores; (2) there is a possible relationship between the origin of hydrocarbons and ore in sedimentary basins; (3) possibly, there is a minimum depth at which oil and gas can originate (Fig. 5-16) and migrate (see also the section on the changes of the types of hydrocarbons with depth in Chapter 3); and (4) there appears to be a maximum depth at which both carbonaceous matter and hydrocarbons can exist, all these being dependent on the physical and physicochemical factors that change progressively in a basin that undergoes tectonic subsidence. It is believed, therefore, that a possible correlation among depth of burial, the amounts and types of organic matter, and the origin of particular varieties of ores, should be investigated. One of the major problems that can be envisaged lies in the present difficulty of distinguishing the products of ore-forming processes, which are dependent on the sedimentary milieu, from those that are relatively independent. Laboratory experiments, for example, could be performed on different types of organic matter, corresponding to various depths of burial and degrees of alteration, to determine their respective effects on the susceptibility to mineralization through biological and non-biological mechanisms. Syngenetic-diagenetic mechanical and chemical compaction processes should be studied theoretically and experimentally in the laboratory and by observations on naturally-formed rocks. The results of these investigations should then be extrapolated into the low- to high-grade metamorphic zones t o investigate the changes in the release of burial-metamorphic and highergrade metamorphic solutions, as compared to diagenetic fluids as well as the alteration of organic matter (Table 5-VII). (9) When it is considered that the ores in sediments owe their origin to a continuous, or more likely a continual, process that can be operative at any time during the evolution of a sedimentary basin, and in fact their origin is directly linked to the origin of the sedimentary and associated volcanics and
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their subsequent history, then in many instances there may be no basic difference between the so-called syngenetic-diagenetic and epigenetic ores (Figs. 5-28 and 5-29). This has been pointed out by many researchers studying copper and uranium ores in sandstones and Mississippi Valley-type leadzinc ores in carbonates. Ores that are not stratabound and cut across all primary sedimentary features, therefore, are not necessarily of a very late origin and unrelated to the original primary sedimentary environment, such as would be the case in many igneous-hydrothermal deposits. The ore in this instance may have been precipitated from early to late compaction fluids into fractures of a diagenetic origin (e.g., Fig. 5-28). On the other hand, stratabound ores following faithfully a particular bed and being possibly controlled by the permeability of the rock, could have been precipitated by solutions of a multitude of alternative derivations, including late-stage igneous-hydrothermal ones. The conclusion that there is no definitive difference between syngenetic-diagenetic and “epigenetic” (in the more classical sense of the ore petrologists) ores, has to be maintained. Precise definitions have to be given whenever such terminology is employed. Compaction solutions can give rise to both stratabound ore bodies as well as to those that cross-cut primary sedimentary structures. From a strict sense of geologic time, all ores formed from compaction fluids would be either early or late diagenetic and epigenetic, because these solutions are being mobilized mainly after syngenesis. One of the main difficulties in their study rests on the fact that these ores can exhibit features of both syngenetic-diagenetic and igneous-hydrothermal mineralizations. (10) Metalliferous concentrations in sediments and associated volcanics must be considered multi-cyclic in origin (e.g., Fig. 5-79, related to uranium mineral concentrations). This concept is mentioned only occasionally and is frequently ignored. One episode of geologic processes may be enough in some instances to form a sufficiently concentrated metallic deposit of economic interest. Under many geological circumstances, however, the primary, dispersed trace elements must undergo episodic geochemical reworking (i.e., dissolutionietransportation-reprecipitation)by one or several related processes to lead to a progressive increase in tenor of the metalliferous constituents. The uranium-roll in sandstone host rocks is one of the most obvious examples, but there are probably many other examples of ores formed by multiple migration-precipitation processes. Compaction fluids can be influential in this episodic concentration of originally dispersed metals. (11)A number of controversial placer-type ore deposits exist which may have been formed from intraformational fluids rather than having been accumulated as detrital grains. It is quite possible, for example, for specks of organic material, which are distributed throughout the intergranular spaces in a sandstone or conglomerate, to control mineralization that results in a
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spotty precipitation of ore. Upon remobilization and recrystallization, the specks of ore may then take on textural features that resemble recrystallized original clastic (i.e., placer-type) grains. (12) As an extension of points 9, 10, and 11 discussed above, it can be concluded that it is not always necessary to establish separate hypotheses to explain the origin of ores that on first sight exhibit sufficient differences to classify them into separate categories. As has already been pointed out above, ores formed by compaction fluids can exhibit either syngeneticdiagenetic or “epigenetic” features, or show both. The host rock can vary from mudstones or shales to sandstones and carbonate rocks (Figs. 5-52 and 5-92). One could subdivide the ores originating from compaction fluids into categories based on: (a) time relationships of ores with the host rock; (b) the host rock types; (c) textures, fabrics and structures of ores; and (d) ore composition. Whichever approach one prefers, it is of paramount practical and theoretical importance not to ignore a more fundamental genetic factor, common to a whole group of ore bodies, in classification schemes and in the preparation of genetic conceptual models. In other words, one should make certain that the classifications and models are not established on the basis of less important differences between the ores. (13)Both mechanical and chemical compaction processes must find their deserved place in the all-encompassing field of diagenetic investigations, which are applied in unravelling the history of the host and source rocks of the ore, in order to determine the reasons for the localization of the metalliferous mineralization in particular stratigraphic units. As difficult as it may be in general (as a result of expenditure required for manpower, time and equipment, and not necessarily because of unsurpassable geologic difficulties), if it can be established that compaction fluids were or were not influential in forming the ore, then either conclusion will open alternative routes in the search for the most plausible genetic hypothesis. A process of elimination is useful in certain types of ore studies. So far, compaction studies have found their deserved place in the investigation of: (1)the Mississippi Valleytype lead-zinc ores; (2) the copper and uranium deposits in the coarsegrained detrital sedimentary rocks; and (3)certain types of iron ores. One should be reminded that although many of the above considerations are based purely on geological extrapolations rather than on direct observations, they are difficult to nullify because of an absence of contradictory criteria and because they stand up well to the scrutiny of pure logic. The “degree of extrapolations” vary according to how far one is willing to push the various interpretations. One can easily argue that ores in sediments were formed from “subsurface fluids” (that can include a number of types of solutions, including telescoped igneous-hydrothermal), but it is quite difficult to prove that these subsurface fluids were actually compaction waters.
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More recent investigations, supported by new stratigraphic, hydrologic, mineralogic, and geochemical data, have become sufficiently refined to show the feasibility of the hypothesis that many ores in sedimentary piles were formed from compaction waters. Both researchers and explorationists should consider this theory in all interpretive work on ore genesis, and determine whether or not compaction fluids were directly responsible for ore genesis. (14) A conceptual model that puts mechanical and chemical compaction, especially diagenetic compaction (Table 5-I), into its correct sedimentological context is presented in Fig. 5-91 (for other models, see Wolf, 1973a,b). Figure 5-91, which is a birdseye view in the form of a flow-chart of the more detailed model in Fig. 5-93, demonstrates that the environmental factors in the source area may very indirectly, yet importantly, control the compactability of the sediments in a basin. It is rather self-evident that investigations of all factors and parameters in the conceptual models have to be undertaken with the realization that the difficulties vary from one geologic problem t o another depending on scale (e.g., thin section versus regional stratigraphic studies) and on actual evidence collected as well as on the degree of extrapolations made (e.g., interpretations related to observable textures versus hypothesizing about origin). The above difficulties, however, will not deter progress, because independent of the reliability of the interpretations at any given point during investigations, the interpretations are scientifically legitimate and permissible. (15) The more immediate factors and processes responsible for the formation of ores in sedimentaryvolcanic piles are summarized in the generalized conceptual model of Fig. 5-92, which has been applied in a simplified form in the section on copper mineralization (Fig. 5-52). This scheme comprises “boxes” in the upper part (I to IV), depicting the four major sources for the metals, and boxes listing the four groups of lithologies that commonly act as hosts for the ore bodies (A t o D). One of the sources is extrabasinal, two are intrabasinal, and one is volcanic, which can be both extrabasinal and intrabasinal as visualized in relationship to the area of sedimentation and ore mineralization. It is important to realize that any of the so-called “host” rocks can at the same time also serve as a source of primary dispersed trace and minor metals, which can be reworked and concentrated. The chloride brines derived from evaporites are depicted at the bottom of the model as another source (box E). The numerous genetic interrelationships, ranging from simple t o complex, are illustrated by the arrows. (16) If the concensus of opinion supports and upholds the hypothesis of compaction processes as being the best explanation for the origin of a certain ore deposit belonging to a specific “type”, e.g., Mississippi Valley-type leadzinc ores, caution is still in order in making extrapolations. For example, the mere fact that the Pine Point ores in Canada are believed t o have been
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formed by compaction fluids should not be taken as a suggestion that all lead-zinc ores in carbonate rocks are of an identical origin. Good scientific judgement must still be applied in studies of the other deposits, now grouped together to constitute the Mississippi Valley type, because with a further use of the “multiple hypotheses” it may be possible in the future to establish a number of sub-types among these lead-zinc ore deposits. One may find that although the compaction hypothesis is acceptable in some cases, another genetic concept is the more plausible alternative in other instances. ACKNOWLEDGEMENTS
This chapter is part of a broader, more inclusive effort by a selected group of investigators to bring together the concepts on ore genesis within sedimentary and related volcanic environments (i.e., on stratiform and stratibound ore deposits) to be published separately in the future. Funds received from the National Research Council of Canada and the Geological Survey of Canada covered part of the expenses involved, whereas Laurentian University provided the facilities. Dr. Frank W. Beales, University of Toronto, read the first draft of the manuscript and his suggestions and constructive comments are appreciated. Particular appreciation goes to Professor George V. Chilingar for his assistance in the preparation of this contribution. My wife, Margaret, had assisted in some of the typing and related work. Reviews and summaries, such as the present one, could not have been written without the indirect and direct contributions of numerous individuals who have made their experience available in widely distributed journals in different languages. Many authors have supplied over the past 15 years reprints to the writer, which have enabled him to reduce the time involved in writing this synthesis and have been valuable in preparing the various conceptual models in ore genesis and in formulating some of his own ideas. REFERENCES Abelson, P.H., 1957. Some aspects of palaeobiochemistry. Ann. N . Y . Acad. Sci., 69: 276-285. Abu Amr, A.R., 1971. Spatial and temporal anchimetamorphism and petroleum occurrence in the Algerian Sahara. Neues Jahrb. Geol. Palaontol. A b h . , 138(3): 259-268. Adler, H.H., 1964. The conceptual uranium ore roll and its significance in uranium exploration. Econ. Geol., 59: 46-53. Ahrens, L.H., 1966. Ionization potentials and metal-amino-acid complex formation in the sedimentary cycle. Geochim. Cosrnochirn. Acta, 30: 1111-1119. Amstutz, G.C., 1963. Bemerkungen zur Genese von kongruenten Blei-Zink Lagerstatten in Sedimenten. Geol. Ges. D.D. R., Ber., Sonderheft 1: 31-42.
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Chapter 6
COMPACTION OF ASH-FLOW TUFFS MICHAEL F. SHERIDAN and DONAL M. RAGAN
INTRODUCTION
Although there are several types of tuffs, the discussion presented in this chapter is restricted to welded ash-flow tuffs. These units are emplaced by hot subaerial turbulent flowage and are characterized by welded textures that result from compaction during the slow cooling from the high emplacement temperatures. The welding process is controlled by the rate of viscous deformation of glass fragments which is in turn temperature dependent. Thus, the essential textures of these units are formed within a few tens of days after emplacement. Excellent summaries of ash-flow features were presented by Smith (1960a,b) and Ross and Smith (1961). The compaction of ash-flow tuffs results from a number of different mechanisms operating at several stages. This chapter deals principally with the compaction that occurs within the first few days after the subaerial eruption and emplacement of the hot ash. This compaction is due to a progressive decrease in the pore space of the deposited ash. Compaction is measured by changes in bulk density and by deformation of objects of known original shape. The first technique provides an accurate measure of the average compactional strain. The second allows the orientation and shape of the finite strain ellipsoid to be determined. Because the compactional strain is uniaxial, the resulting strain ellipsoid is oblate and, therefore, the essential features of this strain can be fully depicted in two dimensions. The characteristic foliation of welded tuffs is perpendicular to the direction of maximum shortening. Measured strain is inhomogeneous on at least four levels. Single cooling units, of the order of a few meters to a few tens of meters thick, show a regular variation in compactional deformation that is referred to as welding zonation (Smith, 1960b). Strain variation is also related to small-scale irregularities in the inhomogeneous compacting materials. Strain is concentrated at the tops and bottoms of rigid bodies such as lithic fragments and crystals; pressure shadows develop at the sides. The area within the pressure shadow zone shows a complex deformation history that depends on the viscous or brittle properties of the material. Compactional strain is also affected by the large initial porosity and viscosity contrast between the ash matrix and pumice or scoria inclusions. Superimposed on these effects are variations due to
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local irregularities in basement topography. These irregularities can produce deviations from the general horizontal foliation to a nearly vertical foliation adjacent to buried cliffs. During the deflation and degassing stages, pumice behave as essentially rigid bodies that rotate and change relative positions in response to the dynamic flow conditions in the fluidized state. Their relative position is fixed and they begin to rotate toward the foliation plane during the stage of mechanical compaction. A highly significant level of alignment and a great degree of flattening is attained only during the stage of welding compaction. At this stage glassy particles respond by permanent deformation that is related to the original temperature of the sheet. The pumice has a low relative viscosity and a high initial porosity and as a result the pumice deforms much more rapidly than the matrix. Rarely, deformation of the still hot tuff may occur within the central densely-welded zone over areas of steep basement topography after all or nearly all of the pore space is eliminated. This would result in equal volume deformation by essentially pure shear within the established foliation plane due to sagging and stretching of the welded zone. Even less frequently, slumping in the welded zone is accompanied by more general types of deformation. Recognition of two basic types of units is of critical importance in understanding the compaction process in ash-flow tuffs. The ash-flow is the unit of emplacement, and the cooling unit represents one complete welding regime. These terms are defined by Smith (1960a, p. 800) as follows: “Ash flow : The basic unit of ash-flow deposits is the ash flow that is analogous to, or perhaps the same as, the deposit resulting from the passage of one nu6e ardente. Cooling unit: A single or multiple ash-flow deposit that can be shown to have undergone continuous cooling is a cooling unit. . . A simple cooling unit is an ash flow or sequence of ash flows that has had an essentially uninterrupted cooling history. Such units remain either nonwelded and noncrystalline during cooling or form a pattern of textural zones.” In any discussion of compaction, a meaningful method of expressing the amount of compaction of the material is needed. Three interrelated terms are useful in describing compaction at various stages: density (p), solidvolume fraction ( G ) , and pore fraction ($). Density is the mass concentration in g emw3. The product of the solid volume fraction and powder density (pp ) gives the bulk density (Pb ):
Pore fraction is equal t o 1.00 minus the solid volume fraction. Bulk pore
COMPACTION O F ASH-FLOW TUFFS
679
fraction of tuff samples can be calculated by using the following equation:
6 = lSoo -
(6-2) Compactability (C), the total compaction strain resulting from the removal of all pore space, is the inverse of the solid fraction: = (pp - p b )/pp
(6-3)
C = 1/G
The relationship of bulk density to compactability is shown in Fig. 6-1, where the powder density of silicic tuffs (pp = 2.45 g cm-3) is assumed. This graph can be used to determine the compactional strain assuming an original density, or to calculate the bed thickness at some previous state using the known bed thickness at an assumed density. As the solid fraction of non-yielding inclusions ( Gi) increases, the amount of compaction due to closure of pore space during welding decreases. This effect is shown in Fig. 6-2 for tuffs with Gi = 0.25 and Gi = 0.50. The bulk density still yields the value for matrix strain in these tuffs, and their fluidized compaction remains the same. Figure 6-3 allows the true volume frac30 20
m
+J
U
m
E" 0
0
5 3 2
1
.I
.2
.3
5
1.0
2.0 3.0
Density, g ~ r n - ~ Fig. 6-1:The relationship between bulk density and compactability for silicic tuffs. The compactability scale can be considered to be the strain ratio required to eliminate the pore space. The columns to the right indicate the range of properties of pumice and ash at the various stages (I = degassing; 2 = deflation; 3 = mechanical compaction; 4 = partial welding; 5 = complete welding).
M.F.SHERIDAN AND D.M. RAGAN
680
1.0
1.5
2.0
2.5
3.0
Bulk density, gcrn-3
Fig. 6-2. Compactability related to volume fraction of non-yielding inclusions; curve A : Gi = 0; curve B : Gi = 0.25; curve C: Gi = 0.50.
tion of inclusions to be calculated for the totally compacted equivalent (pb = 2.45 g cm-') from measurements made on thin sections or outcrops of partially welded to nonwelded tuff. This value should be cited in reports of mineral or lithic percentages so that nonwelded and densely-welded portions of the same cooling unit can be directly compared. 0
10
20
30
Volume percent rigid particles Fig. 6-3. Percent increase in crystal-lithic volume during welding.
COMPACTION OF ASH-FLOW TUFFS
68 1
For solid tuff bodies it is convenient to use bulk density as a measure of compaction; for fluidized masses, the solid fraction is more useful. This dual system has the advantage that a stationary defluidized mass has an initial bulk density of 1.00 g ~ m - so ~ ,that the bulk density is approximately equal to the compactional strain ratio for most tuffs (Ragan and Sheridan, 1972). EMPLACEMENT AND DEFORMATION
Several theories have been presented on the transport and emplacement of ash flows (Smith, 1960a). Although the mechanism of transport flowage and emplacement is critical to an interpretation of textures in terms of postemplacement deformation, few hard facts are available concerning the details of actual observed flows and their undisturbed textures. It is generally agreed that the ash flow moves as a fluidized bed. Whether the moving flow is a dense-phase fluidized bed (G 0.4) or a dilute-phase fluidized bed (G < 0.1) of Leva (1959), is currently being debated (see McTaggart, 1960; O’Keefe and Adams, 1965; Fisher, 1966; Lindsay, 1972; Pai et al., 1972). Data on flow thickness, viscosity, and solid concentration are usually developed in terms of theoretical considerations rather than field data. Of critical importance to the present discussion is the effect of primary flow on the orientation and shape of particles in the fixed bed. While the tuff is fluidized, particles are not in continuous contact, and adjacent neighboring particles are constantly changing orientation and position. The orientation varies as particles freely rotate in the completely yielding fluid medium. The shape remains nearly the same, except for rounding due to breakage. After degassing, the stationary sheet is a fixed bed (G 0.4) completely supported by the close-packed framework of particles. From this point, adjacent neighboring particles are fixed, although these bodies may change their shape and orientation. Compactional strain textures must, therefore, develop after the sheet has come to rest. Textures associated with the moving fluidized stage should be considered primary; those associated with the fixed-bed compaction are secondary. The emplacement and deformation of a welded tuff cooling unit can be considered in a series of partially overlapping events related to the progressive closing of pore space. The general order is: (1)deflation of one or more turbulent fluidized clouds; (2) degassing of a stationary fluidized bed; (3) mechanical compaction (rotation and compression of particles without a change in shape); (4)welding compaction; and (5) equal volume deformation with or without shear. Table 6-1 shows the relative removal of pore volume through these stages for a typical cooling unit. The critical stage in developing compactional textures is stage 4, i.e., weld-
-
-
M.F.SHERIDAN AND D.M. RAGAN
682 TABLE 6-1 Changes in pore space in compacting ash-flow tuffs Stage
Bulk density (Pb)
Pore fraction Solid-volume (@I concentration
ReIative thickness
(GI Before deflation Before degassing Loose packing Partly welded Densely welded N o pore space
0.25 0.86 1.00 1.35 2.20 2.45
0.90
0.65 0.592 0.45 0.10 0
0.10 0.35 0.408 0.55 0.90 1.00
10.00 2.85 2.45 1.82 1.11 1.00
ing compaction. This chapter deals principally with processes that occur during such welding. The role of other stages, however, must be assessed in order to understand and properly interpret tuff textures. There is little information in the literature on the relative density or solid concentration of a flowing cloud as it moves outward from the vents toward the position of emplacement. Estimates of solid concentration range from those of incipient fluidization (G S= 0.4) t o those of a rarified cloud (G < 0.1). Certain facts related to observed nu6e ardentes support dilutephase fluidized flow for most tuffs: (1)the moving cloud is very sensitive to underlying topography (Anderson and Flett, 1903; Perret, 1937); (2) ashflow sheets have initial depositional slopes away from the vents in the range of 3-6" (Smith, 1960a); and (3) the average velocity of flow is near 60-80 km hr-' (Lacroix, 1904; Perret, 1937; Moore, 1967; Moore and Melson, 1969). Due to the nature of the eruption, accurate values of cloud height that can lead to solid concentration estimates are difficult to obtain from historic observations. In the famous 1902 eruption of Mt. Pelhe that destroyed St. Pierre, an eruptive cloud of a few meters in thickness left only a 35 cm thick ash deposit (Lacroix, 1904). Assuming a cloud thickness of 8 m and the emplaced ash concentration of 0.4, the approximate deflation on emplacement was 100 : 1. The Aira caldera, Kyushu, Japan (Matumoto, 1943, 1963) is an example of a large-scale prehistoric ash-flow center that also sheds some light on the nature of the fluidized cloud. The Ito pyroclastic flow, one of the large units produced from this center, is approximately 22,000 years old (Kigoshi et al., 1972). This sheet is essentially a single cooling unit with a maximum thickness of about 170 m and an estimated volume of more than 150 km3. In a detailed study of the distribution of the Ito pyroclastic flow, Yokoyama
683
COMPACTION OF ASH-FLOW TUFFS
Vertical Exaggeration 5 x
5
0
10
15 kilometers
Om
Fig. 6-4. Topographic cross-section showing the distribution of the Ito pyroclastic flow in southern Kyushu, Japan. (After Yokoyama, 1972.) The cloud height must have been great enough to clear the mountains (600 m) and deposit the thick ash (shown in black) beyond. The initial height of the cloud is estimated to have been about 1000 m. The direction of flowage is perpendicular to the page.
(1972)has shown that ash flows, originating from a caldera rim at sea level, have surmounted ridges in excess of 400 m in several places as much as 50 km from the vent. Figure 6-4 is a topographic section of the Ito sheet showing that the initial transport cloud at this location must have been greatly in excess of 600 m thick. If the resulting 170 m thick deposit has an average solid volume frac-
1 Fig. 6-5. Fumarolic pipes in the Ito pyroclastic flow, Kyushu, Japan. (After Yokoyama, 1970.) Relative compaction of the stationary bed before welding is shown. A is the zone of fumarolic pipes; B is an intermediate level of massive concentrations of lithic fragments; C is the basal lensoid concentration of lithic fragments; D is the amount of compaction during degassing. Thickness in meters is an arbitrary intermediate value for the emplacement of this tuff sheet.
684
M.F.SHERIDAN AND D.M. RAGAN
tion, G, of 0.6 (pb = 1.47), the completely compacted thickness would be 102 m. The solid volume concentration of the flowing cloud would have been about 0.11, and the emplaced fixed bed (G = 0.4)thickness would have been 250 m. Calculations based on the highest ranges surmounted and the lowest passes not overcome, bracket the possible initial concentrations between 0.09 and 0.18. When the emplacement cloud is deflated to the extent that it no longer moves laterally, it may still behave like a stationary fluidized bed. If the escape of gases is concentrated in zones, slugging or channeling is possible. There is an intense gas escape from subsurface fumarolic pipes where fine particles are blown upward and the coarse lithic fragments and pumice are concentrated (Fig. 6-5). Fumarolic steam pipes have been reported from the Ito pyroclastic flow in Kyushu (Yokoyama, 1970) and the Shikotsu pumice flow in Hokkaido (Yagi and Hunahashi, 1970), and such structures are common features of several other pyroclastic flows (Walker, 1971). The preservation of these structures requires that the sheet remained stationary during their formation. If the sheet underwent lateral displacement during or subsequent to channeling, the pipes would be sheared out in the direction of flow. During the degassing stage, the ash flow is fairly compact and probably behaves as a fixed bed in some parts. Although the pore fraction of fluidized uniform spheres is 0.52 (Leva, 1959), the very poorly sorted irregular-shaped particles of the ash flow should have a minimum fluidizing value near 0.65. The solid concentration is close to the critical fluidizing value for ash flows (G 0.4). Measurements of loosely-packed ash yield a bulk pore fraction at the end of the degassing stage near 0.59-0.60 that corresponds to a bulk density of 1.00 g ~ m - ~ . Mechanical compaction takes place essentially without a change in the shape of the particles (bubbles, shards, or pumice), except for possible breakage. Adjacent particles maintain their relative position, but elongate fragments are slightly rotated toward the horizontal. Although mechanical compaction can significantly reduce porosity and produce a compact rock, its effect on the resulting texture (grain orientation and shape) is relatively minor. Pumice fragments in vertical sections from most nonwelded tuffs affected by this process show a random orientation on testing with the chi square statistic at the 99% confidence level. The ratio of long to short axes ( x / z ) in these tuffs, after correction for the orientation factor, is the same as for pumice lapilli collected from nearby outcrops of unconsolidated ash. The density of tuff during this stage varies from 1.00 to 1.35 g ~ r n - ~although , some tuffs with density in the range of 1.20-1.35 g cm-3 show signs of incipient welding in thin section. Rigid-body mechanical compaction unaccompanied by welding deforma-
COMPACTION OF ASH-FLOW TUFFS
685
tion must occur at an early stage in all ash-flow sheets. A significant foliation of pumice due solely to mechanical compaction without welding, however, is rarely observed. Welding of less viscous components may well proceed simultaneously with mechanical compaction of more rigid elements, so that a broad field of overlap is possible. A significant level of orientation does not occur unless there is a measurable level of viscous deformation of particles due to welding. Welding compaction is the result of slow viscous deformation of glassy fragments (Riehle, 1973). The controlling factor is the residence time at various temperature levels above the threshold for welding. At temperatures below about 550°C there is negligible deformation, but above this level the rate of deformation is a function of temperature (Smith et al., 1958). Relatively low emplacement temperatu
