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
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-
-:
-
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80
60
60
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40
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20
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. .
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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 (
.
~
H
0
-
t
4 2 20 2
0
3
I
0
0
'
;40 -2, 0
I
,Vo
-
Ia a
I
I
I
-
FAULTED STRATA NON-FAULTED STRATA I
I
I
1
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.
188
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 (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)
>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 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 0.2 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
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* 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
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1
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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