Compaction and Fluid Migration

Developments in Petroleum Science, 9 COMPACTION AND FLUID MIGRATION Practical Petroleum Geology FURTHER TITLES IN THI...

Author: Kinji Magara

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Developments in Petroleum Science, 9 COMPACTION AND FLUID MIGRATION Practical Petroleum Geology

FURTHER TITLES IN THIS SERIES

1.A. GENE COLLINS GEOCHEMISTRY OF OILFIELD WATERS 2. W.H. FERTL ABNORMAL FORMATION PRESSURES

3. A.P. SZILAS PRODUCTION AND TRANSPORT OF OIL AND GAS 4. C.E.B. CONYBEARE GEOMORPHOLOGY O F OIL AND GAS FIELDS IN SANDSTONE BODIES

5. T.F. YEN AND G.V. CHILINGARIAN (Editors) OIL SHALE

6. D.W. PEACEMAN FUNDAMENTALS OF NUMERICAL RESERVOIR SIMULATION

7. G.V. CHILINGARIAN and T.F. YEN (Editors) BITUMENS, ASPHALTS AND TAR SANDS 8 . L.P. DAKE FUNDAMENTALS OF RESERVOIR ENGINEERING

Developments in Petroleum Science, 9

COMPACTION AND FLUID MIGRATION Practical Petroleum Geology

KINJI MAGARA Associate Director Reservoir Studies Institute Texas Tech University Lubbock, Texas, U.S.A. Formerly with Imperial Oil Ltd. Calgary, Alberta, Canada

ELSEVIER SCIENTIFIC PUBLISHING COMPANY Amsterdam - Oxford - N e w York 1978

ELSEVIER SCIENTIFIC PUBLISHING COMPANY 335 Jan van Galenstraat P.O. Box 211,1000 AE Amsterdam, The Netherlands

Distributors f o r the United States and Canada: ELSEVIER NORTH-HOLLAND INC. 52,Vanderbilt Avenue New York, N.Y. 10017

Library of Congress Cataloging in Publiralion Data

Magara, K Compaction and fluid migration. (Developments in petroleum science ; 9) Includes bibliographical references and index. 1. Fluids--Migration. 2. Sediment compaction. I. Title. 11. Series. TN871.M328 553' .28 78-2004 ISBN 0-444-41654-4 Library of Congress Cataloging in Publication Data

ISBN: 0-444-41654-4 (vo~.9) ISBN: 0-444-41625-0 (series)

0 Elsevier Scientific Publishing Company, 1978. 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, P.O. Box 330, 1000 AH Amsterdam, The Netherlands. Printed in The Netherlands

To m y inspirers Professor K o z o Kawai Doctor Frank J. Moretti and m y family Tomiko, Miki and Albert J. Magara

FOREWORD

Compaction is a pressurecontrolled phenomenon in which the centres of the constituent grains of a sediment are brought closer together, usually in the vertical direction. In most cases the pressure is provided by the effective weight of the sediment overlying the rock undergoing compaction, but there are circumstances under which some compaction results from lateral pressure. Progressive burial also leads to a rise in temperature, and where such a rise causes mineralogical changes, especially the loss of constitutional or associated water from the mineral particles, the change will consequently allow compaction to take place. The composition of the pore fluids, particularly the nature and amount of solute in water, by far the commonest pore fluid, influences compaction; furthermore, it influences cementation, and the redistribution of mineral matter associated with the phenomenon of so-called pressure solution. In addition, the nature and size of the mineral particles affect compaction. Compaction varies widely in amount for a given increase in pressure. It is largely irreversible, pressure relief leading t o elastic rebound only, except in the case of the gypsum/anhydrite system in which the mineralogical change contribution to the process is reversible. Compaction causes a reduction in porosity, an increase in bulk density (because the mineral solids are denser than the pore fluids expelled as the porosity is decreased), a reduction in electrical conductivity (when the pore fluids are aqueous solutions) and permeability, and an increase in the velocity of transmission of seismic pulses. Direct measurements of some rock properties, especially under in-situ conditions, are limited, but reasonable deductions concerning their values can sometimes be made from measurements of the values of other properties, as in various types of wire-line welllogging. Organic matter and fine-grained inorganic mineral matter such as clays suffer the highest degree of compaction; the gypsum/anhydrite system allows considerable compaction; the initially coarser-grained deposits, such as sandstones and some kinds of limestone, undergo least compaction. Compaction is progressive during the course of burial for organic matter and finegrained inorganic mineral matter; the main phase of compaction is delayed in the case of gypsum, and sandstones and limestones, as well as in forming stylolites. An inevitable consequence of compaction is the movement of fluids, although this is not the only agent causing fluid flow through rocks. Such flow can bring fluids which differ in composition from those originally pres-

vii

ent into contact with the solids, with the possibility of effecting changes in the solids in terms of composition or solution/deposition. Progressive heating may change some of the solids, and mobile breakdown products may be carried away by the moving fluids. When compaction cannot proceed to the extent that is normal for the rock type and pressure and temperature conditions, a state of undercompaction arises and there is overpressuring of the fluids. Lateral variation in the thickness or nature of compactible rocks leads to differences in the amount of compaction and changes the structure of the overlying rock layers. Compaction is of great importance in geology, and also of concern in civil engineering. It causes surface settlement, a feature which can be enhanced by man’s actions. The writer’s interest in compaction was initiated over forty years ago, in connection with considering the migration and accumulation of oil and gas, and differential compaction. Quantitative studies were attempted, and later a basis for deducing differences in maximum depth of burial was outlined. Dr. Kinji Magara has paid particular attention t o quantitative aspects of the phenomenon, and especially to the amounts of water expelled from compactible rocks. He has tried t o use t o the maximum techniques which, over the years, have been developed for making measurements in wells, thereby acquiring additional information, or being able t o make useful deductions for learning more about or employing more effectively the concept of compaction.

V.C. Illing & Partners, Cheam, Surrey. England. July, 1977.

G.D. HOBSON

ACKNOWLEDGEMENT

The author thanks Imperial Oil Limited, Calgary, Canada, for permission

to publish this book. Most of the research work included in this book was accomplished while the author was employed by this company. Some of the research works in this book were supported financially by the Exxon Production Research Company as well. The author acknowledges the following people, who were the co-workers, supervisors and managers at both companies, for their valuable advice and guidance. Imperial Oil Ltd. C. Bily B. Bums E.T. Connolly J.W.L. Dick G.G. Dunbar C.R. Evans R.P. Glaister R.O. Grieve K. Gulstene T.J. Hawkings A. Heslop K. Jackson

F.H. Lane D. Milner F.J. Moretti D.J. Murphy H.W. Nelson M. Parsons D. Toderian J .H .D. Walker G. Wells R. Wilkinson J. Wishart J.B. Greig

Exxon Production Research Company D.H. Horowitz P.H. Monaghan D. Perry A. Rogers

R. Sarmiento J.P. Shannon A. Young

K. MAGARA

CONTENTS

................................................. Acknowledgement . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Chapter 1. Introduction . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Historical background . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Subjects covered in this book . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Movement of water in sediments . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . References . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Foreword

.

Chapter 2 Shale compaction and estimation of erosion and structural timing . . . . . Shale porosity-depth relationship Transit-timeporosity relationship for shale .......................... Estimating the thickness of eroded sedimentary rocks .................... Thickness of erosion and maximum burial depth ....................... Other techniques for estimating the thickness of erosion Seismic cross-section . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Interval velocity from the seismic record ........................... Sandstone porosity . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Colour of organic matter . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Change of the normal compaction slope and rate of sedimentation . . . . . . . . . . . . Analysis of structural timing using shale transit-time plots . . . . . . . . . . . . . . . . . Early Cretaceous (initial tectonism) .............................. Late Early Cretaceous (major tectonism) ........................... Late Cretaceous or later (after major tectonism) Limitations and possible problems with the techniques Use of other logs Examples of erosion estimates in western Canada References . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

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Chapter 3 Calculations of pore pressure from shale compaction data Terzaghi’s model and the subsurface model ........................... Calculation of fluid pressure . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Aquathermal pressuring . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Application of aquathermal and nonaquathermal concepts Areas of continuous deposition and burial .......................... Areas of significant uplift and erosion ....... ...................... Estimation of pore pressure by the use of charts Relationship between fluid pressure. depth and equivalent mud weight . . . . . . . . . Direct estimation of mud weight . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Calculation of fluid pressure by computer ............................ Use of other logs for pressure estimates .............................. Empirical method for estimating fluid pressure ......................... Sonicmethod

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

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

vi viii 1 1 3 6

9 Y 11 11 13 16 23 26 26 27 28 29 29 36 38 38 38 39 40 44 45

47 47 52 54 59 59 60 61 63 66 68 70 73 73

X

Resistivity method . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Example of well-log plots . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Appendix 3-1 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Appendix 3-11 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . References . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

75 79 82 82 84

Chapter 4.Causes of abnormal surface pressure . . . . . . . . . . . . . . . . . . . . . . . . Compaction . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Aquathermal effect . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Montmorillonite dehydration . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Compaction disequilibrium . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Montmorillonite dehydration . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Artesian condition . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Hydrocarbon accumulation . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Osmosis . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Fossil pressure . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Cementation . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Tectonics . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Generation of hydrocarbons . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Charging . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . References . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

87 87 98 100 101 102 110 110 111 112 112 112 115 115 115

Chapter 5 . Application of calculated pressures and porosities . . . . . . . . . . . . . . . Pressure and porosity profiles . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Drainage map . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Porosity maps . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Fluid and hydrocarbon drainage . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . References . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

119 119 125 134 136 142

Chapter 6. Concept and application of fluid-loss calculations . . . . . . . . . . . . . . . Calculation of fluid losses from shales before and after maturation . . . . . . . . . . . . The question of rebounding . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Calculation of fluid loss with time . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Examples of fluid-loss calculation . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Correction for nonclays . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Effect of organic facies . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Fluid-loss curve . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Fluid-loss mapping . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . References . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

143 143 149 150 152 156 160 160 162 163

Chapter 7 .Evaluation of pressure and capillary seals ..................... Calculation of pressure-sealing depth . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Calculation of pressure-sealing time . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Calculation of fluid losses before and after pressure sealing . . . . . . . . . . . . . . . . . Comparison of sealing pressure and excess hydrocarbon pressure . . . . . . . . . . . . . Capillary seal . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Laboratory model of a pressure seal . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Experiment A. using montmorillonite clay ......................... Experiment B. using the same montmorillonite clay . . . . . . . . . . . . . . . . . . . References . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

165 167 169 170 172 176 178 181 181 182

Chapter 8. Concept of three-dimensional fluid migration . . . . . . . . . . . . . . . . . 183 Directions of horizontal and vertical fluid migration . . . . . . . . . . . . . . . . . . . . . 183

xi Volumes of vertical and horizontal fluid movement ...................... Example of a three-dimensional fluid-flow study . . . . . . . . . . . . . . . . . . . . . . . References . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

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Chapter 9 Porosity-permeability relationship in shales . . . . . . . . . . . . . . . . . . . Fluid-pressure gradients and movement of fluids in shales . . . . . . . . . . . . . . . . . . Shale porosity-permeability relationship . . . . . . . . . . . . . . . . . . . . . . . . . . . . Comparison of logderived and laboratoryderived porosity-permeability relationships . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Volume of fluids expelled downward and upward . . . . . . . . . . . . . . . . . . . . . . . References . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

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189 198 200 201 201 204 210 213 216

Chapter 10 Changes in shale pore-water salinity during compaction . . . . . . . . . . . Calculation of pore-water salinity from well-log data ..................... Pore-water salinity of shales . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Pore-water salinity of sandstone . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Water salinity change during compaction . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Ion filtration by clays or shales . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Time of the first occurrence of iron filtration ........................ “Salinity x porosity” plot . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Application of water salinity data to exploration ....................... Proximity to bedded salt deposits . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Abnormal pressure . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Faultzone . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Permeable sandstone of significant extent . . . . . . . . . . . . . . . . . . . . . . . . . . Depositional environments . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Fresh-water contamination References . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

217 217 217 221 222 222 229 231 233 233 236 237 237 237 239 241

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243 244 244 247 248 254 255 256

Chapter 12 Estimation of oil-genesis stage ........................... Time-temperature relationship for hydrocarbon generation . . . . . . . . . . . . . . . . Chart description . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Useofcharts . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Comparison of oil-genesis chart and world oil and gas reserves . . . . . . . . . . . . . . . References . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

257 257 259 260 262 264

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Chapter 11 Importance of abnormal pressuring in shale diapiriiam Abnormal pore pressures and their significance to shale mobility . . . . . . . . . . . . . Shale compaction model without the aquathermal-pressuring effect . . . . . . . . . Shale compaction model with the aquathermal-pressuring effect . . . . . . . . . . . Movement of rocks overlying diapiric shales . . . . ....................... Significance of differential loading . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Buoyancy effect . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . References . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

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Chapter 13 Estimation of paleopore pressure and paleotemperature . . . . . . . . . . Estimation of paleopore pressure . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Geothermal gradient . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Retention of generated pressure . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Estimation of paleotemperature . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . References . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

265 265 268 268 269 273

xii Chapter 1 4 . Primary hydrocarbon migration . . . . . . . . . . . . . . . . . . . . . . . . . . Compaction fluid movement . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Aquathermal fluid movement . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Osmotic fluid movement . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Fluid movement due to clay hydration .............................. Other possible causes of primary migration ........................... Form of hydrocarbons at primary migration .......................... Migration of oil in oil phase . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . References . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

275 277 280 283 285 288 288 290 296

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Chapter 15 Oil-reserve evaluation from sandstone thickness and type and source rockquality . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Reservoirsource relationship . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Sandstone models . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Fluid-migration model . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Vertical fluid migration . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Horizontal fluid migration . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Combined fluid migration . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Use of different fluid-migration models .............................. Exploration application of the technique ............................ References . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Index

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

299 299 301 302 302 305 307 309 310 311 313

Chapter 1 INTRODUCTION

Conventional oil and gas are usually found in pore spaces and fracture openings of sandstones, limestones, and dolomite. Shales, although considered to be important in many cases as source rocks and cap rocks for hydrocarbons, usually do not contain commercial oil and gas. In other words, from the standpoint of production, shales are relatively useless rocks. If, for example, an exploratory well were drilled and a geologist tried to evaluate whether it had penetrated oil and/or gas accumulations, he would first separate the shaly sections as being of no interest, and then try to find out whether the other sections were oil-(and/or gas-) saturated or watersaturated. This rock that is so frequently ignored in conventional oil and gas studies - shale - has been the subject of my research work. I am now confident that shale is important in petroleum exploration, not only as a primary-hydrocarbon source rock, but also as a seal, a source of compaction fluids which may influence or control the direction and volume of hydrocarbon migration, a source of overpressures, a creator of structures, and as an indicator of subsurface drainage conditions, which may be related to reservoir development and more. Therefore, an intensive study of shales in a sedimentary basin can greatly benefit the exploration for oil and gas. I also believe that such a study be made in vacuo, but must be related to many other areas of study, such as structural geology, geochemistry, mineralogy, hydrology, petrology, geophysics, well logging, drilling engineering, production engineering, etc. The research has progressed toward the accomplishment of this objective, but the goal is still far away. Although general interest in the subject has increased dramatically in recent years, the number of researchers working in this particular area is still comparatively small. Since I started to investigate the problems of fine-grained clastics some ten years ago in Japan, I have benefitted from the written and oral discussions of many workers which have influenced my subsequent thinking. It seems to me, therefore, that an effective way of acknowledging my debt to these persons would be to state briefly the history and background of the evolution of my thinking with respect to the subjects included in this book. Historical background Since the early 1960’s I have become increasingly interested in shale compaction and its effect on trap development. A t that time core and cutting

2 samples were the main data source for the study of shale compaction (Athy, 1930;Hedberg, 1936;Dickinson, 1953).My previous experience with mechanical-log analysis as an operations geologist enabled me t o use such logs t o study shale compaction. At that time mechanical wire-line logs were useful only t o geologists as a correlation tool and t o log analysts as a semiquantitative indicator of hydrocarbons in the reservoir rocks. Piles of copies of these well logs remained unused, once the excitement of drilling and discovery was over. Cores were sometimes taken from reservoir sections, but scarcely ever from shale zones. Therefore, mechanical logs were the only data source that could be used for studying shale compaction. The new porosity tools, such as sonic and density logs, previously developed by Schlumberger, were extremely useful (Schlumberger, 1972). The main purpose of my earlier research with the well-log data was simply to restore the compacted sediments t o lesscompacted stages in the geological past. The restored thickness could be used t o interpret the paleostructures that might have controlled the direction of hydrocarbon migration. In 1965 several papers (Hottman and Johnson; Wallace) on the application of well-log data t o the evaluation of pore pressures were published, which drew the attention of many geologists and engineers in the oil industry. I was especially impressed by Hottman and Johnson’s paper, which helped me understand the pore-pressurecompaction relationship and encouraged me t o study further the experimental work in soil mechanics, especially that of Terzaghi and Peck (1948),the subsurface application of the same concept by Hubbert and Rubey (1959),and the work in rock mechanics, such as that of Handin and Hager (1957). Although Hottman and Johnson’s paper showed the presence of an empirical relationship between shale compaction and its pore pressure, I suspected that factors other than shale compaction might also affect pore pressure, because there was sometimes a discrepancy between the empirical compaction-pressure data and the compaction theory. The empirical relationship was the combined result of many factors, but what I wanted t o know was the effect of each factor. To discover these effects, the theory, the experimental data, and the actual subsurface data had to be combined and analysed. I continued my compaction research in Japan, using well-log and other geological data. The main producing reservoirs there are volcanic rocks (lavas and tuffs) and sandstones. Detailed study of shale porosities above and below these reservoir sections indicated that the shale porosity within a basin varies significantly, and that the traditionally accepted smooth porositydepth curve may not represent the relationship at individual locations. The porosity sometimes decreases toward the interbedded permeable sections. An easy explanation for these porosity variations would be changes in the “lithology” or “composition” of the shales. However, I found many

3 instances of significant porosity difference where there was no recognizable difference in shale composition. These situations can be explained only by an understanding of the subsurface drainage and the pore pressures in shales in the geological past. This approach could also lead us to a better understanding of the primary migration of hydrocarbons from shales into reservoirs. In 1967 I moved to Canada, where I had an opportunity t o continue the research on shales at the Geological Survey there. During the same period, clay mineralogists made progress in the understanding of abnormal-pressure occurrences and primary migration (Powers, 1967;Burst, 1969). On completion of my term with the Geological Survey of Canada, I joined Imperial Oil, where I had the good fortune t o be able t o continue my research on fine-grained clastics. Since then my work has been affected by two significant factors. The first was my association with more people of different backgrounds and, consequently, with a greater variety of subject matter. The second was my involvement with computers. Appreciation of such problems as detailed analysis of shale composition by thin-sections and X-ray, hydrocarbon maturation, heat flow, rock fractures, etc., has expanded the area of fine-grained clastics research considerably. One of the most significant contributions t o the understanding of abnormal-pressure occurrences was made during this period by Barker (1972). The aquathermal-pressuring concept he introduced has greatly increased our knowledge in this area. The calculation of fluid-loss volume has been an important part of my research since 1967, but recently the use of the computer has greatly improved the results. Calculations of pore pressure from mechanical-log data and of fluid pressures from seismic interval-velocity data, can also be made by computer. Pore-water salinity can be evaluated by analysis of the SP log, and the calculated values can be checked by chemical analyses of water samples. Such calculations for shales were new. I used a combination of sonic and resistivity (or conductivity) logs for this estimate. The change in shale pore-water salinity due t o compaction and ion filtration and its resultant osmotic fluid movement, could also have played a role in primary migration. Subjects covered by this book Why should we study fine-grained clastics? Here are my answers: (1)They constitute approximately 75% of a clastic basin fill. (2) They contain the organic matter from which hydrocarbons are generated. (3) They contain the fluids that, when expelled by compaction, carry hydrocarbons t o available traps.

4

I

.*

COMPACTION

COMPOSITION DEPOSITIONAL ENVIRONMENT

PAST (BEFORE COMPACTION1

/

J

VOLUME, DIRECTION POST-MATURATION PRE-MATURATION

-

TIME

/

SEALING TIME & DEPTH PROPERTY OF EXPELLED WATER SALINITY HYDROCARBON SOLUBILITY

*‘POROSITY COMPACTION EROSION & MAXIMUM BURIAL DEPTH RESTORATION OF FORMATION

RATE OF SEDIMENTATloN PRESSURE PRIMARY MIGRATION DRAINAGE CONDITION SANDSTONE PERMEABILITY & AREAL EXTENT DlAPlRlSM SALINITY ION-FILTRATION EFFICIENCY DEPOSITIONAL ENVIRONMENT CONTAMINATION MINERAL ALTERATION

Fig. 1-1. Diagram showing an overview of shale compaction studies.

(4) Their expelled fluids carry the dissolved ions that destroy reservoir porosity by cementation. (5) They form the top and bottom seals in clastic traps. (6) They sometimes cause such structures as shale diapirs. Details of these subjects are illustrated in Fig. 1-1. Suppose we have a fine-grained rock in the subsurface (top right-hand square, Fig. 1-1).This rock contains some amount of pore water and rock material (or minerals). The amount of pore water is usually expressed as porosity, which is a measure of the compaction state. This information may enable us to interpret erosion thickness and maximum burial depth. It may also be used to restore the thickness of a compacted formation to the precompaction stage. From the slope of the normal compaction trend, the relative rate of the sedimentation may be estimated. All the information derived from shale compaction will help us to interpret the structural timing, which is essential for a prospect evaluation. Pore pressure in fine-grained rocks can be evaluated by considering the compaction state in relation to burial depth. This information may be used to understand primary migration and drainage efficiency in shales, and to obtain some ideas on the permeability and areal extent of interbedded sandstones. Generation of extremely high pressures may cause shale diapirism. As stated previously, pore-water salinity can be calculated from well logs. Examining the ion-filtration efficiency of shales during compaction will give

5

us an idea of the original salinity or depositional environment. Fresh- or saline-water contamination and clay-mineral alteration in the geological past may also have affected the present salinity distribution. Therefore, the study of pore-water salinity may provide some ideas on these problems. All this pore-water information, as mentioned above, will be useful in understanding fluid migration and hydrocarbon accumulation. An interesting and unique research effort in this fine-grained clastics project is the study of water loss (centre rectangle, Fig. 1-1).If we know the present water volume (or porosity) of a subsurface shale, and if we can assume the porosity when the shale was deposited (this porosity may be obtained by extrapolating the subsurface normal porosity trend to the surface), we can calculate the volume of total water loss. Integration of geochemical data with this concept will produce the volumes of fluid loss, both before and after maturation. Incremental fluid losses during successive geological periods can also be calculated, and interpreted in relation to the timing of trap developments for an understanding of petroleum migration and accumulation. An analysis of threedimensional fluid migration would further improve our understanding of petroleum migration. The concept of sealing time and depth may also be important in studies of fluid migration and accumulation. The salinity of water expelled from shales in the geological past can be evaluated. (This is not the pore-water salinity in the present subsurface shales.) The salinity of the expelled water may be important in hydrocarbon migration, because the solubility of hydrocarbons in water changes with the amount of salt ions present. The following remarks apply to the fine-grained clastics study in general: (1)Our data base consists almost entirely of mechanical and seismic data. Since, in practice, we cannot count on core or even good cutting data, it is essential that we learn how to extract the geological data we need from the vast library of log and seismic data at our disposal. (2) All calculations interpretations, and plots are made by hand as well as by computer; the computer method is simple, inexpensive, gives instant results if a timeshare computer is available and, hence, is within the grasp of every operations geologist who can use such information. (3) The several types of interpretation that have been developed are checked against other types of data, i.e., water chemistry, X-ray, pressure, geochemical, and palynological data. Many problems related to shales and/or their compaction are still controversial. In this book, however, I have stressed my own ideas and the ideas of others that I favour. I have also presented many subsurface data that support these ideas. I have endeavoured, too, to explain the interrelationships of many factors associated with shale compaction, such as shale porosity, pressure, fluid loss, water salinity, sealing, etc. In other words, the different ideas presented in this book have a common base or consistency, so that

6 none of them should be in serious contradiction. This book, then, is not a reference book which contains many contradicting opinions and data. Rather, it is a guide book for the practical application of these techniques to petroleum exploration. Movement of water in sediments There are essentially two different kinds of water moving in a sedimentary basin. Their respective characteristics are as follows: Sediment-source water (1)The movement of this type of water takes place in any part of a sedimentary basin (deep or shallow). (2) The principal direction of small-scale movement is from a shale or clay t o a sandstone or other permeable bed. (3)The direction of large-scale movement is from the basin’s centre to its edges, or from the deeper parts to the shallower. (4) The amount of water is limited, because the amount of sediment in a basin is usually limited. ( 5 ) Movement of this type of water is probably important in the primary migration of hydrocarbons. (6) Most movement of this type of water took place in the geological past. Meteoric water ( 1 ) Movement of this type of water is important in the relatively shallow intervals of a sedimentary basin. (2) The direction of small-scale movement can be either from sandstone to shale or from shale to sandstone. However, most movement of this type of water may take place in sandstones only. (3)The direction of large-scale water movement is from the basin’s edges to its centre, or from shallow to deep. (4) The amount of water is unlimited. ( 5 ) Movement of this type of water is probably unimportant in primary migration, but it may affect the trapping condition of hydrocarbons in a pool. (6) Movement of this type of water is a present event and may or may not have developed in the geological past. This book mainly discusses the first type of water - sediment-source water. Fig. 1-2 shows a schematic diagram in which water is moving in an aquifer. The water pressure is measured at two points, A and B, in the aquifer. An imaginary vertical water column which corresponds to the measured pressure is made, and the height of the column above the datum level (sea level in this case) is known. This is a measure of the potential level, and is called a potentiometric or piezometric surface. Water moves in the aquifer from a higher potential point (A) to a lower one (B) as shown in Fig. 1-2. In this case it is meteoric water.

7

Fig. 1-2. Schematic diagram showing the water flow in an aquifer due to the hydrodynamic force.

If the datum level is taken at the surface, the potentiometric surface elevation can be shown as a function of excess pressure above hydrostatic pressure. Therefore, the excess pressure can also be used in determining the direction of fluid movement. Fig. 1-3depicts a schematic example in which two aquifers, A and B, have different potential levels. If there is any fluid communication route between these aquifers, the fluid will move from the higher potential point to the PRESSURE

-

I

i

Fig. 1-3. Schematic pressure-depth plot for two aquifers, A and B . The arrow shows a possible fluid-flow direction.

8 lower (or from the higher excess-pressure point, A, to the lower, B). Note that the total fluid pressure at B is greater because of its greater depth, but its potential or excess pressure is lower, so that the fluid moves from A toward B. The excess-pressure difference discussed above can be caused by the difference in elevations of the water-intake areas of the aquifers, if the water is meteoric water. If, however, the moving fluid originated in the sediments, loading of the sediment layers would be the principal cause of the excess fluid pressure. Engineers tend to use the term “pressure gradient” to express “pressure/ depth”. In the zones of abnormal pressure, the “pressure-gradient” value for a particular depth may not be the same as the “pressure/depth” value. Fig. 1-4demonstrates why. The solid curved line shows the actual “pressure versus depth” relationship in the subsurface. The pressure gradient for A is shown by the thick solid line tangential to the curve at A. The pressure/ depth relationship for point A, however, is given by the dashed straight line between A and the surface. Therefore, pressure gradient and pressure/depth are not the same at A.

PRESSURE

-

Fig. 1-4. Schematic diagram showing the difference between pressure gradient and pressureldepth at a subsurface point.

9 At B, on the other hand, the tangential pressure-gradient line is parallel to the pressure/depth line and the values are the same. Throughout hydrostaticpressure zones, of course, the value will also coincide. The pressure gradient defined in Fig. 1-4is important in the discussion of subsurface fluid migration, while the pressure/depth is essential in relating the pressure at a given depth t o drilling-mud weight.

References Athy, L.F., 1930. Density, porosity and compaction of sedimentary rocks. Bull. Am. Assoc. Pet. Geol., 14: 1-24. Barker, C., 1972. Aquathermal pressuring - role of temperature in development of abnormal-pressure zones. Bull. Am. Assoc. Pet. Geol., 56: 2068-2071. Burst, JB., 1969. Diagenesis of Gulf Coast clayey sediments and its possible relation to petroleum migration. Bull. Am. Assoc. Pet. Geol., 53: 73-93. Dickinson, G., 1953. Geological aspects of abnormal reservoir pressures in Gulf Coast Louisiana. Bull. Am. Assoc. Pet. Geol., 37: 410-432. Evans, C.R., McIvor, D.K. and Magara, K., 1975.Organic matter, compaction history and hydrocarbon occurrence - MacKenzie Delta, Canada. Proc. 9 t h World Pet. Congr., 3: 149-157 (Panel Discussion). Handin, J. and Hager, R.V., 1957.Experimental deformation of sedimentary rocks under confining pressure: test at room temperature on dry samples. Bull. Am. Assoc. Pet. Geol., 41: 1-50. Hedberg, H.D., 1936.Gravitational compaction of clays and shales. Am. J. Sci., 31: 241281. Hottman, C.E. and Johnson, R.K., 1965. Estimation of formation pressures from Iogderived shale properties. J. Pet. Technol., 17: 717-722. Hubbert, M.K. and Rubey, W.W., 1959.Role of fluid pressure in mechanics of overthrust faulting, I. Geol. Soc. Am. Bull., 70: 115-166. Magara, K., 1968a. Compaction and migration of fluids in Miocene mudstone, Nagaoka Plain, Japan. Bull. A m . Assoc. Pet. Geol., 52: 2466-2501. Magara, K., 196813. Subsurface fluid pressure profile, Nagaoka Plain, Japan. Bull. Jpn. Pet. Znst., 10: 1-7. Magara, K., 1969a. Upward and downward migrations of fluids in the subsurface. Bull. Can. Pet. Geol., 17: 20-46. Magara, K., 1969b. Porosity-permeability relationship of shale. Can. Well Logging SOC. J., 2: 47-73. Magara, K., 1971. Permeability considerations in generation of abnormal pressures. SOC. Pet. Eng. J., 11: 236-242. Magara, K., 1972. Compaction and fluid migration in Cretaceous shales of western Canada. Geol. Surv. Can. Pap., 72-18: 81 pp. Magara, K., 1974a.Compaction, ion-filtration, and osmosis in shales and their significance in primary migration. Bull. Am. Assoc. Pet. Geol., 58: 283-290. Magara, K., 1974b.Aquathermal fluid migration. Bull. Am. Assoc. Pet. Geol., 58: 25132516. Magara, K 1975a. Reevaluation of montmorillonite dehydration as cause of abnormal pressure and hydrocarbon migration. Bull. Am. Assoc. Pet. Geol., 59: 293-302. 1975b. Importance of hydrodynamic factor - discussion, Bull. Am. Assoc. Magara, K.,, Pet. Geol., 59: 890-893. Magara, K., 1975c. Importance of aquathermal pressuring effect in Gulf Coast. Bull. Am. Assoc. Pet. Geol., 59: 2037-2045. ~

10 Magara, K., 1976a. Water expulsion from elastic sediments during compaction - directions and volumes. Bull. A m . Assoc. Pet. Geol., 60: 543-553. Magara, K., 1976b. Thickness of removed sediments, paleopore pressure, and paleotemperature, southwestern part of Western Canada Basin. Bull. A m . Assoc. Pet. Geol., 60: 554-565. Magara, K., 1976c. Factors causing primary oil migration (Abstract). 1976 Annu. Meet. Geol. Assoc. Can., Prog. Abstr., 1: p. 58. Magara, K., 1977a. A theory relating isopachs t o paleo compaction-water-movement in a sedimentary basin. Bull. Can. Pet. Geol., 25: 195-207. Magara, K., 1977b. Petroleum migration and accumulation. In: D.G. Hobson (Editor), Developments in Petroleum Geology. Applied Science Publishers, Essex, pp. 83-126. Powers, M.C., 1967. Fluid-release mechanisms in compacting marine mudrocks and their importance in oil exploration. Bull. A m . Assoc. Geol. 51: 1240-1254. Rubey, W.W. and Hubbert, M.K., 1959. Role of fluid pressure in mechanics of overthrust faulting, 11. Geol. SOC.A m . Bull., 70: 167-206. Schlumberger, 1972. Log Interpretation, 1. Principles. Schlumberger New York, N.Y., 113 pp. Terzaghi, K. and Peck, R.B., 1948. Soil Mechanics in Engineering Practice. Wiley, New York, N.Y., 566 pp. Wallace, E.W., 1965. Application of electric log measured pressures t o drilling problems and a new simplified chart for wellsite pressure computation Log Anal., 60: 4-10.

Chapter 2 SHALE COMPACTION AND ESTIMATION OF EROSION AND STRUCTURAL TIMING

Shale compaction is the result of physical, chemical, and mineralogical phenomena in the subsurface. However, it is known in many parts of the world that the level of shale compaction is governed mainly by the burial depth (or overburden pressure), provided that the fluid pressure is near hydrostatic, or the shales are at near compaction equilibrium. If the fluid pressure is higher than normal hydrostatic, shales are compacted less than those compacted normally under the hydrostatic pressure. If the area being studied has experienced a significant uplift and erosion, the normal shale compaction trend is shifted to the direction of increased compaction at any present depth, in comparison with the trend in an area of no erosion. Therefore, we are able to estimate the amount of erosion and the maximum burial depth on the basis of shale compaction data. Shale porosit y 4 e pth relationship Fig. 2-1shows a summary of shale porosity-depth relationships in different parts of the world. Shale porosity decreases with increase of depth. The rate of porosity decrease is fast at shallow depths and slows down with greater burial. As mentioned above, shale porosity can also be influenced by subsurface fluid pressure; the higher the pressure the greater the porosity at a given burial depth or under a given overburden pressure. Therefore, many of the porosity-depth curves in Fig. 2-1,especially the ones indicating relatively high porosity values at depth, may be the result of the higher-thannormal (hydrostatic) fluid pressure. Rubey and Hubbert (1959)proposed an exponential function expressing the relationship between shale porosity and depth of the normal compaction trend, or at the compaction-equilibrium condition (fluid pressure is hydrostatic), as follows:

where @ = value of the shale porosity at depth 2, = porosity at the surface (2 = 0), e = base of the Napierian logarithms, and c = constant of dimension (length-').

$0

12

O

r

-

r +I

r

I-

n W

a

20

POROSITY ,‘I.

Fig. 2-1. Relationship between porosity and depth of burial for shales and argillaceous sediments. (From Rieke and Chilingarian, 1974, fig. 17.) 1 = Proshlyakov (1960); 2 = Meade (1966); 3 = Athy (1930); 4 = Hosoi (1963); 5 = Hedberg (1936); 6 = Dickinson (1953); 7 = Magara (1968); 8 = Weller (1959); 9 = Ham (1966); 1 0 = Foster and Whalen (1966).

The value c is the measure of dope of the normd compaction trend when it is plotted on semilog paper (logarithmic scale for porosity and arithmetic scale for depth). This equation is based on Athy’s (1930) curve derived from the Pennsylvanian and Permian shales in northern Oklahoma. Rubey and Hubbert considered that Athy’s relationship is the one closest to the compactionequilibrium conditions, because of its elapsed time since deposition. The values of &, and c determined for Athy’s curve are 0.48 (or 48%) and -4.33 . loF4 ft-l, respectively. Although relatively deeper parts of the porosity4epth curves are some-

13 times influenced by the higher-than-normal fluid pressure, the normal compaction trend is commonly developed in the shallower intervals of most young sedimentary basins. Based on Dickinson’s curve in the Gulf Coast area, Magara (1971)showed an exponential relationship (normal compaction) between shale porosity and depth above 7000 f t there. For estimating the thickness of sedimentary rocks removed by erosion, the shift of the normal shale porosity trend can be used. However, the data of shale porosity are not always on record. Shale compaction data, such as those from sonic logs, are commonly available, so that they can be used more often. Transit-time-porosity relationship for shale After numerous laboratory tests, Wyllie et al. (1956,1958) concluded that, in consolidated strata with small pores uniformly distributed, there is a linear relationship between porosity and transit time:

or :

where Atlog = transit time on the sonic log in ps/ft, Atwater = transit time of the formation water in ps/ft, and Atmatrix= transit time of the rock matrix in ps/ft. In the case of clean quartz sandstones in the subsurface, the values used for Atmatrix and Atwater are commonly 55.6 and 189-200 ps/ft, respectively. Eqs. 2-2 or 2 3 mean that in rock of uniform lithology transit time increases as porosity increases. The relationship between shale porosity and acoustic transit time was found by using data from conventional cores and the sonic log at Kambara GS-1,which drilled through Japanese Tertiary rocks (Magara, 1968). The core-analysis and transit-time data are shown in Table 2-1,and the relationship between shale porosity and transit time is plotted in Fig. 2-2. A similar relationship was established for Tertiary and Cretaceous shales in the western Canada basin, based on a study of fourteen wells in which both sonic and formationdensity logs were run (Magara, 1976).The shale porosity was calculated from the bulk density, on the basis of a shale-matrix (or grain) density of 2.72 g/cc and a water density of 1.02 g/cc. The porosity-transit-time relationship for Cretaceous shales in this area can be

14 TABLE 2-1 Data of core analysis and sonic transit time from Kambara GS-1,Nagaoka Plain, Japan Depth (m)

Core analysis

Sonic log

porosity

transit time (ccs/ft)

39.00 33.15 26.56 24.60 24.26 23.08 21.82 19.60 18.80 15.90 14.60

145 127 130 109 114 110 102 99 95 104 92

-

1029.08-1029.23 1609.40-1609.60 1808.63-1808.76 2150.65-2150.85 2296.00-2296.20 2443.46-2443.66 2607.16-2 607.33 3062.77-3062.98 3205.36-3205.53 3505.25-3505.46 3701.29-3701.49

2.00 2.11 2.13 2.27 2.22 2.24 2.26 2.28 2.32 2.35 2.42

(assumed)

expressed as :

Q, = 0.466 Atlog- 31.7

(2-4)

where

Q, = shale porosity in 96 (Fig. 2-3). The acoustic transit time of shale plotted versus depth within a zone of normal compaction will show a continuous decrease with depth as compaction progresses. The transit-time-depth plot on semilog paper does not show a true straight line, but will be expressed as a curve at depth as it

POROSITY,*

TRANSIT TIME, A

h

Fig. 2-2.Relationship between mudstone porosity, @(%) and transit time, Atlog (W/ft) of Kambara GS-1,Nagaoka Plain, Japan.

15 200

200

-t

. a

;;

100

100

w

I t In 268

40

40

0

62

50

I00

POROSITY @ (%I

Fig. 2-3. Empirical relationship between porosity and transit time of shales of Cretaceou: age in western Canada. SHALE POROSITY

/

ILOGARITHMIC SCALE)

I

/ t n

B

A

Fig. 2-4. Schematic diagrams showing transit-time-depth (A) and porosity-depth ( B ) relationships of the normal compaction trend.

16

2

rs/ft

Fig. 2-5. Example of the generalized normal compaction trend of a transit-time-depth plot of the Gulf Coast. (From Hottman and Johnson, 1965.)

approaches the matrix transit time, Atmatrix(see Fig. 2-4A). Note that the shale p o r o s i t y d e p t h relationship can be shown as a straight line on semilog paper (Fig. 24B). In the relatively shallow intervals of many sedimentary basins, however, the transit-timedepth relationship is known t o be approximated by a straight line. The pore pressure in such an interval is usually hydrostatic. An example of the generalized normal compaction trend of the transit-timedepth plot of the Gulf Coast is shown in Fig. 2-5. Estimating the thickness of eroded sedimentary rocks Fig. 2-6A shows a schematic shale transit-time-depth relationship of a relatively shallow, normally compacted interval. This normal compaction trend, extrapolated to the surface, gives a surface transit-time value of Ato, for a situation where there was no significant erosion. Fig. 2-6B shows a schematic example of the transit-time-depth relationship where the uppermost section was removed by erosion. The present surface is indicated by a wavy line. If the normal compaction trend in the subsurface is extrapolated to the present surface, the surface transit-time value Atb will be smaller than the value Ato for the case of no erosion. If the normal compaction trend is further extrapolated t o Ato, the original surface of the sedimentary section can be determined (Fig. 2-6B). The distance

SHALE TRANSIT TIMEILOGARITHMIC SCALE1

DEPTI

1

/ /

SHALE TRANSIT TIME

8

I---------

-

17

- --

I ILOGARITHMIC SCALE1

I I

I

/

I

I I

/

I

/

/ /

820

t

/ A

B

Fig. 2-6. Schematic diagrams showing the normal compaction trends of shale transittime-depth plots where there was no erosion (A) and where there was erosion ( B ) .

between the erosional surface (the present surface in this case) and the level at which the extrapolated value equals At0 is the approximate thickness of the sedimentary rocks removed by erosion. In this estimate of eroded thickness, it is assumed that there was no significant expansion or rebounding of sedimentary rocks during and after erosion, at least not enough to make such an estimate erroneous. However, even a minor amount of such sediment expansion might be expected t o reduce the pore pressure considerably. The surface transit-time value of Ato, for the case of no erosion, can be determined by two methods. The first is to determine the normal compaction trends of as many wells as possible in the sedimentary basin in question, and then extrapolate these trends to the present surface. The maximum transit-time value (the least compaction value) among these extrapolated values ( A t ; ) will be the value closest to Ato for no erosion. From studies of the normal compaction trends of over 300 wells in the western Canada basin, I obtained a maximum surface transit time of about 200 ps/ft in the area on the northeast, i.e., near the Canadian Shield, which is considered to have undergone very little erosion. It is interesting to note that the normal compaction trend in the Gulf Coast, shown in the middle of Fig. 2-7, is also extrapolated to approximately 200 ps/ft at the surface, suggesting no erosion in this case.

18 Fluid -pressure gradient

Shale resistivity

3 m 40000

-

c

Shale transit time

c

-Mud

---Formation

5ooo6000

( 1 1 FPG at 11,948 It i 0.663 based on 12.3 Ib/gal mud and -0 psi shut-in drill -pipe pressure

7ooo-

-.

8000

-

c

oooo10,m11.m

-

12,000

-

13,0000 14.000

-

A t . &s/ft

Fluid-pressure gradient. psi

Fig. 2-7. Shale resistivity Rshr shale transit time At and fluid-pressure gradient versus depth. (Redrawn from Rogers, 1966.)

The second method is based on the empirical relationship between shale porosity and transit time, as shown in Fig. 2-3. The intercept of the thick line with the vertical axis in Fig. 2-3gives a transit-time value of about 68 ps/ft, corresponding to the value for shale grains or matrix. This is the transit-time value where porosity is zero. The line is terminated at 200 ps/ft in the upper right-hand part of Fig. 2-3,because the transmit-time values of a clay-water mixture should not exceed the value for water. The transit time for formation water (+50,000 ppm NaC1) at the surface is approximately 200 ps/ft (Fig. 2-8).In the subsurface condition it is usually less. The porosity value corresponding to this termination point is about 62% (Fig. 2-3).The relationship for porosity values from 62 to 100%is shown graphically as a thick horizontal dashed line. The shale porosity-transit-time relationship shown in Fig. 2-3 may be explained as follows: The transit-time value for water, or 100% porosity, is about 200 ps/ft. Addition of a small amount (5-10s) of clay sediment to the water will produce no significant change in transit time, because the sound essentially will travel through the water, not the clay. The transit-time value will stay at almost the same level until the amount of clay becomes 38% of the total bulk volume (or 62% porosity). The transit time decreases after this stage as the amount of clay increases (or porosity decreases). This observation for shales is different from that made by Wyllie et al. (1956,

19

-.-. -----.

Tap Water NaCl

50,000 ppm NaCl 150.000 ppm NaCl

200-

; I

m

________-.

250,000 ppm NaCl

_____ __ -_ _ --__-__ 1 .

.-

I

2

;;

==150-

I

F

f

9 +

1w-

50

24

I

I

I

I

I

I

16

8

0

-8

-16

-24

SHALE TRANSIT TIME

SHALE POROSITY-

-

70 8090

IARITHMETIC SCALE1

DEPTH

A

B

Fig. 2-9. Schematic diagrams showing the shale porosity-depth relationship ( A ) and corresponding shale transit-timedepth relationship ( B ) of the normal compaction trend.

20 1958) for sandstones, in which a linear relationship is established for the entire range of sandstone porosity (or 0--100%). Fig. 2-9 shows schematic diagrams of shale porosity-depth and transittime-depth relationships in the subsurface. The porosity of clay on the sea floor is known to be 70-8076. During the early stages of burial, porosity decreases rapidly. On the basis of Dickinson's (1953)shale porosity-depth relationship in the Gulf Coast area, a porosity of 62% would be reached at about 100 ft. This critical depth varies in different sedimenatry basins. Above it, the transit time would be about 200 ps/ft, as shown in the upper part of Fig. 2-9B.In the interval below this critical depth, the transit time decreases as the shale porosity decreases (Fig. 2-9A,B). If the normal compaction trend of transit time established in this deep interval is extrapolated to the surface, we will have a Ato value slightly greater than 200 ps/ft. This difference is dependent on the depth of critical compaction (62% porosity) and the slope of the normal compaction trend. From the knowledge of shale compaction in several different basins, I believe that in most basins the value Ato will not exceed 210 ps/ft. Therefore, it may be concluded that 200 ps/ft is a good approximation for the surface transit-time value in an area where there was no significant ero-

I

I

\.

I

i

STUDY AREA

\. \

I

\,

I

-- 1 ..

..

..-

. . .

I

USA

Fig. 2-10. Map showing the area studied.

\*

..

..--100 MILES

21 SHALE TRANSIT TIME (PrlFT) 0

2000

4000

U I-

E

Em0

0

8000

'WO

Fig. 2-11. Shale transit-time-depth plot of the Pacific Amoco Ricinus 16-29-34-8-W5 well.

sion. However, using this figure (200 ps/ft) will give a minimum estimate of erosional thickness, for the reason mentioned previously. Fig. 2-10indicates an area in the southwestern part of the western Canada basin, where the above-mentioned technique was applied in practice to estimate the amount of erosion. Fig. 2-11is a shale transit-time-depth plot for the Pacific Amoco Ricinus 16-29-34-8-W5 well, showing that the shales above about 3000 f t are compacted normally. The normal trend is extrapolated to the surface at 116 ps/ft, suggesting a significant amount of erosion in the geological past. The thickness of erosion is estimated to be about 4600 ft. In other words, the maximum burial depth of each bed or present depth point can be calculated by adding 4600 f t to the present depth. Most of the shales below 3000 f t are undercompacted. This fact suggests that while continuous deposition and burial were taking place (before erosion), the deeper section was undercompacted and overpressured. Most of the overpressure may have since disappeared, because uplift and erosion can be expected to decrease pore pressure. During these events, the subsurface temperature will decrease and the pore space may expand slightly, resulting in the decline of the pore pressures. As a matter of fact, the pressures mea-

22 SHALE TRANSIT TIME (fir/FT)

Ir I

-’

I=

CARDIUM SANDSTONE

Fig. 2-12. Shale transit-time-depth plot of the Mobil et al. Ricinus 3-5-35-8-W5 well.

sured by drill-stem tests in this well are not high. The presence of the undercompacted shales in this well, however, makes it evident that there was overpressuring before the erosion. Fig. 2-12is a similar plot for the Mobil et al. Ricinus 3-5-35-8-W5 well. On the basis of this plot, 4200 f t of erosion are estimated. As there was no thick shale section at shallow depths, it was not easy to establish a normal compaction trend for this well. However, use of the data between 1500 and 3500 ft, and of the slope of the normal trend established in this general area from many other well data, made the estimate possible. As the shales below about 4000 f t are undercompacted, they are considered t o have been overpressured, at least in the geological past. The present pressure in the deep section is not known. Fig. 2-13 is a plot of the H.B. Garrington 12-8-36-5-W5 well in which about 3300 f t of erosion are calculated. The deeper section is undercompacted only slightly. Fig. 2-14 shows the plot for the Suptst. Altana HB. Caroline 10-26-366-W5well. The normal compaction trend was determined from the regional knowledge of shale compaction. The calculated thickness of erosion is about 1700 ft. The overpressure is known from a drill-stem test of the Cardium

23 SHALE TRANSIT TIME (br/FTI

Fig. 2-13. Shale transit-timedepth plot of the H.B.Garrington 12-8-36-5-W5 well.

sandstone at a depth of 7200 ft. The pressure of 3610 psi at 7200 f t is about 500 psi in excess of the hydrostatic pressure. Let us consider the possible effect of continental glaciation on the shale compaction process. If the ice sheet had been added to the sedimentary column as part of the continuous loading history, then its weight certainly would have contributed to additional compaction. If, however, the ice sheet developed after uplift and erosion, shale compaction would not have been affected, because the shales already had been “overcompacted” with refereence to their depth of burial at that time. In this area the latter is believed to be the case. The fact that the density of ice (0.9 g/cc) is significantly lower than that of average sediments (approximately 2.3 g/cc) must also be remembered. In other words, the effect of ice on compaction is believed to be insignificant and is ignored here. Thickness of erosion and maximum burial depth If the erosion surface is at the present surface, the maximum burial depth of a given bed can be calculated by summing the present depth of the bed

24 SHALE TRANSIT TIME (ps/FTI 100

2w

300

4w

0

2wo

4000

+ Y

6wo

I

t W

n

8000

1 p o

Fig. 2-14. Shale transit-time-depth plot of the Suptst. Altana H.B.Caroline 10-26-366-W5 well.

and the thickness of erosion. If the water depth of the sediment at the sediment-water interface is known or inferred from paleontological data, such water depth may be further added to obtain a more realistic maximum depth. Such an estimate of erosion and maximum burial depth would be quite important in petroleum exploration from at least two stand points: (1) petroleum maturation, and (2)structural configuration and timing. The generation and maturation of petroleum are known to be temperature and geological-time dependent (Connan, 1974), so that estimation of the maximum burial depth or the maximum temperature the bed has ever attained is an important factor in assessing hydrocarbon potential. Proper analysis of maximum burial depth would be essential in evaluating the paleostructural configuration and structural timing which are very important in a prospect evaluation. If the erosion surface is not at the present surface but in the subsurface as depicted in Fig. 2-15,the thickness of sedimentary rocks removed by erosion must be estimated differently. This figure shows that the difference between the present surface and the estimated original sedimentary surface before erosion is A, and the erosion surface is at the level whose distance or

25

----------- - - -7---' ORIGINAL SURFACE 0'0

/ / PRESENT SURFACE

NEW SEDIMENTS DEPOSITED AFTER EROSION

EROSION SURFACE

OLD SEDIMENTS

c

SHALE TRANSIT TIME

Fig. 2-15. Schematic transit-time-depth plot where the thickness of sediments deposited after erosion was less than the thickness of erosion.

A'0

t NEW SEDIMENTS DEPOSITED AFTER EROSION

/

I

/

PRESENT SURFACE

/ --

ORIGINAL SURFACE EROSION SURFACE

OLD SEDIMENTS

..,.:.:.;:..:

..:..,::c,.:2;,.::

I

-

SANDSTONE. MAXIMUM BURIAL DEPTH = c (PRESENT DEPTH)

SHALE TRANSIT TIME

Fig. 2-16. Schematic transit-time-depth plot where the thickness of sediments deposited after erosion was more than the thickness of erosion.

26 depth from the present surface is B. In other words, new sediments, whose thickness is B, were deposited after erosion. The thickness of erosion in this case is the total of A and B. However, the maximum burial depth for the sandstone bed at the present depth of C can be calculated by summing A and C. If the erosion surface is much deeper than in the previous example, and if the thickness of sediments deposited after erosion is greater than that of the sediments eroded, as shown in Fig. 2-16,the two normal compaction trends in the old and new sedimentary sequences can be plotted on the same line and extrapolated to the transit time for no erosion, Ato, at the present surface. In other words, the record of the shale compaction before the erosion was completely removed by renewed sedimentation. -Thus, if the thickness of post-erosion or post-unconformity sediments is more than that of erosion, the compaction study cannot be used for estimating erosional thickness. Other data, such as the seismic cross-section and paleontological or electric-log correlation, must then be used. Other techniques for estimating the thickness of erosion

Seismic cross-section If erosion is limited to a local area, a seismic section can be used to estimate the thickness of erosion, as depicted in Fig. 2-17.This figure shows that most of formation C is eroded at the top of the crest of the structure. The commonest method of estimating the thickness of erosion is to draw

Fig. 2-17. Schematic seismic crosssection where there was local erosion.

27

THICKNESS OF EROSION (4300 FT J

PRESENT SURFACE EROSION SURFACE

Fig. 2-18. Example of a geological cross-section through an anticline with significant truncation in northern Canada.

line A parallel to B (base of formation C) and measure the distance between A and the erosion surface. This method assumes that formation C was of uniform thickness before erosion. It is also possible, however, that formation C thinned toward the crest of the structure (see line A’), because before significant uplift and truncation, the rate of sedimentation might have slowed down at the crest. Use of a seismic section alone usually cannot indicate whether such thinning took place or not. A combined study of shale transit-time data and the seismic cross-section will provide the best solution, in that the erosion thickness derived from the shale transit time is an independent source of information based on the maximum compaction the shales have ever attained. Fig. 2-18shows an actual example of the combined application of a seismic cross-section and shale compaction. The compaction data at the crest of this structure show that the thickness of erosion is about 4300 ft, while the thickness of the equivalent section in the syncline is about 7000 ft. Significant depositional thinning toward the crestal area would thus have taken place before the erosion. If the erosion is more regional, the seismic cross-section cannot be used effectively for evaluating the thickness of eroded sediment. This is because extrapolating a paleosedimentary surface laterally for a long distance would create a tremendous error, so that the result would be ambiguous. Interval velocity from the seismic record The interval velocity of a formation can be estimated from the seismic record (Pennebaker, 1968a, b). If the reciprocal of such velocity data,or

28

Fig. 2-19. Plot of interval transit time versus depth derived from seismic velocity gathers.

transit-time data, is plotted versus depth on semilog paper, the thickness of erosion and the maximum burial depth can be estimated by using the method described previously. Fig 2-19 shows an example of such an estimate from the interval transit time derived from the seismic record. A weak point in this technique is that the effect of rock composition on acoustic velocity or transit time is overlooked. because there is no reliable method for distinguishing Iithology from seismic data alone.

Sandstone porosity

The sandstone porosity also decreases with burial, due mainly to pressure solution and reprecipitation at and around the grain contacts. Therefore, if the relationship between sandstone porosity and burial depth is established where there was no significant erosion in the geological past, the thickness of erosion may be estimated for the area where sandstones are more compacted by using a technique similar to that discussed previously. However, sandstone porosity is also affected by other factors, such as average grain size, grain-size distribution, and types of minerals that compose grains, as well as geothermal gradient, chemical composition of formation water, tectonic force, etc. In other words, the sandstone porosity is not a simple measure of the maximum burial depth, even if the pore-fluid pressure is

29 hydrostatic. Therefore, estimation of erosion from sandstone data seems to be much less reliable. The grain-to-grain contacts in shales are usually simpler than those in sandstones. The shale matrix or grains are usually weaker, so that the shales are more sensitive to the changes of overburden load, if the pore-fluid pressure is hydrostatic.

Colour of organic matter The colour of organic matter in sedimentary rock usually becomes darker with increasing burial depth due to the thermal effect (Staplin, 1969).This information is useful in evaluating the petroleum maturation stage. The maximum burial depth may not be estimated easily from the colour of organic matter only, because such colour may also be dependent on the type of organic matter and the elapsed geological time. Change of the normal compaction slope and rate of sedimentation The slope of the normal compaction trend is not always uniform, but changes within a given sedimentary basin and between different basins. This possibility may be demonstrated by the use of Fig. 2-20, which shows the shale porositydepth relationships for four different sedimentary basins: the Gulf Coast Tertiary basin (Dickinson, 1963), the Venezuelan Tertiary basin (Hedberg, 1936), the Japanese Tertiary basin (Hosoi, 1963), and the Oklahoma Paleozoic basin (Athy, 1930). By comparing the three porositydepth curves of the Gulf Coast, Venezuela and Oklahoma, Rubey and Hubbert (1959, p. 175) were able t o state that, “The fact that the curve based on Dickinson’s data for the Gulf Coast Region shows porosities higher than

DEPTH l l w 0 FTI

Fig. 2-20. Comparison of shale porosity-depth relationships in several regions: Oklahoma (Athy, 1930); Venezuela (Hedberg, 1936); Gulf Coast (Dickinson, 1953); Japan (Hosoi, 1963).

30

1.000

0.900 0.800 0.700 0.600

0.465

I-

LL \

Wv,

an 3

mfn ml-

W Z

gw 0

z

0

so00

10.000

15,000

DEPTHS IN FEET

Fig. 2-21. Reservoir pressure versus depth for Louisiana Gulf Coast wells. Solid circles = measured pressures; open circles = estimated pressures. (From Dickinson, 1953.)

those in Venezuela and Oklahoma is probably to be explained by the not uncommon occurrence of abnormally high fluid pressure there.” This finding suggests that most of the Gulf Coast shales have not reached their equilibrium condition of compaction. Abnormal fluid pressures are common in the Gulf Coast and many other relatively young sedimentary basins. Most abnormal pressures occur at relatively great depths; Dickinson’s fluid-pressure-depth relationship (Fig. 2-21) for the Gulf Coast indicates that abnormal pressures sometimes exist below about 7000 ft. Pressures above this depth are near hydrostatic. In other words, shales above 7000 f t have reached their compaction equilibrium in the Gulf Coast area. The curves for the relatively shallow parts of other basins must also represent nearequilibrium conditions, because all of these curves show lower porosity levels or more compaction at a given depth than the Gulf Coast curve indicates. It may be concluded, therefore, that all four curves probably show compaction-equilibrium conditions, at least in the relatively shallow sections. What, then, makes the difference in compaction level at a given depth in these areas? It is obvious that the idea suggested by Rubey and Hubbert (1959) that abnormal fluid pressures cause higher shale porosities in the Gulf Coast is not valid in the shallow interval above 7000 ft. Geological time is sometimes believed to be a controlling factor; the older rocks are more compacted at a given depth than the younger. This would suggest that younger rocks, such

31 as those of the Gulf Coast, are still being compacted even in the shallow intervals where the fluid pressure is already hydrostatic. This idea, however, contradicts Terzaghi’s basic concept on stress balance among total stress, effective stress, and fluid pressure. If the fluid pressure is hydrostatic, the level of compaction, which is a function of the effective stress, will not change under the constant overburden load, not even for millions of years in the future, because equilibrium has already been reached. This is what Terzaghi’s relationship means. In other words, whether the sedimentary rocks are young or old should not make any significant difference in porosity under the constant load (or depth) with hydrostatic pressure. In this case, the possible effect of cementation on shale porosity reduction is ignored. Through an examination of the relatively shallow samples of the Gulf Coast continental slope, Morelock (1967) found that lower porosity seems to be associated with a slower rate of sedimentation. Perry (1970) also concluded, on the basis of a study in the same general area, that “the state of compaction of clastic sediments in the northern Gulf slope is controlled, among other things, by sedimentation rates.” The slower rate seems to have caused more compaction or consolidation for a given depth. The intervals studied by Morelock and Perry are relatively shallow and the fluid pressures would be near hydrostatic. Other possible factors that could affect the rate of compaction with burial depth are geothermal gradient, chemical composition of formation water, mineral composition of rock grains, tectonic stress, etc. However, the significance of their influence on porosity reduction is usually very difficult to evaluate. Changes in slope of the normal compaction trend of the transit-timedepth plot are also commonly observed. Fig. 2-22shows such examples (wells A and B) in northern Canada. The trend of well A shows more compaction for a given burial depth than that of well B. Both trends are extrapolated to almost 200 ps/ft, at the surface, suggesting that there was no significant erosion at the surface of these locations. Geological time markers are indicated by numbers; rocks at a given depth are generally older in well A than in well B. The rate of sedimentation in well A was slower. The fluid pressures in the intervals depicted in Fig. 2-22are known to be near hydrostatic. Although the idea that older rocks are more compacted at a given depth point appears to apply in this case, it cannot be the principal cause of the changing slope of the normal compaction trend, for the reason mentioned above. The mathematical form of the normal compaction trend of the transittime-depth plot is as follows (the trend is assumed to be a straight line on semilog paper) : At = At&e-cz

(2-5)

32

t; Y

z I t n

Fig. 2-22. Examples of changing slopes of the normal compaction trend in northern Canada. Numbers refer to geological age; the age gets younger with increasing number.

where

At = shale transit time (ps/ft) at depth 2 (ft), Atb = extrapolated surface transit time (ps/ft), and c = constant (ft-l) indicating the slope of the normal compaction trend. Taking the natural logarithm on both sides of eq. 2-5 and rearranging it, we obtain:

1

c =--log,

2

(2)

Therefore, if the surface transit time At; and the transit time At at depth 2 are known, the slope c can be calculated. The c values for well A and well

B are shown in Fig. 2-22. A graphic solution of the slope c is also possible by comparing the actual plot with the series of slopes corresponding to different c values, as shown in Fig. 2-23.The value of c increases as the trend becomes more horizontal. The basic data used in constructing the series of lines in Fig. 2-23are listed in Table 2-11. Fig. 2-24depicts the plot of the slope value of c of the normal compaction trend (transit time) versus the average rate of burial or sedimentation

33 TRANSIT TIME ( p d f t )

50

100

150 200

2000

4000

6000

-+

8000

w w

-

I&

t

lC!oOO

12,000

lQ000

t

1q000

18#000

2 0 0

I

Fig. 2-23. Graphical presentation of different slopes of the normal compaction trend and corresponding coefficients c.

(ft/lOOO years) in northern Canada. Although the points are scattered widely, the plot shows a general tendency for c t o increase as the rate of burial or sedimentation decreases. The relationship between slower sedimentation and a larger negative c value (morehorizontal normal trend) may be explained by the use of the schematic diagrams in Fig. 2-25.If sedimentation was relatively slow, there may have been sufficient time for the (flat) shale grains t o become relatively well arranged (Fig. 2-251.1).This would cause faster porosity reduction with burial. If, however, deposition was very fast, the shale grains may not have had

34 TABLE 2-11 Transit-time values of the normal compaction trend at selected depths The value c is the mathematical expression of the slope of the normal trend Depth (ft)

At ( W f t ) *

_ _ _ ~ _ _ ~

c(ft-l):

0 2000 4000 6000 8000 10,000 12,000 14,000 16,000 18,000 20,000

0.00006

0.00008

200 177.4 157.3 139.5 123.8 109.8 97.4 86.3 76.6 67.9 60.2

200 170.4 145.2 123.8 105.5 89.9 76.6 65.3 55.6 (47.4) (40.4)

0.00010

0.00012

0.00014

200 163.7 134.1 109.8 89.9 73.6 60.2 (49.3) (40.4) (33.1) (27.1)

200 157.3 123.8 97.4 76.6 60.2 47.4 (37.3) (29.3) (23.1) (18.1)

200 151.2 114.2 86.3 65.3 49.3 37.3 (28.2) (21.3) (16.1) (12.2)

_____

* Values in parentheses are less than the matrix transit time for shale.

-0oO2-1

b

b

b

b

b

b b

- 0oOl

'bb

b

b b

b

bo 'b

bb

I

0

I

I

1

AVERAGE BURIAL RATE ft/lOOO years

Fig. 2-224.Relationship between the slope c of the normal compaction trend and the average burial rate in northern Canada.

35

A

B

Fig. 2-25. Schematic diagrams showing arrangements of shale grains when the rate of sedimentation was slow ( A ) and rapid ( B ) .

sufficient time to arrange themselves, resulting in higher porosity at a given depth (Fig. 2-25B).In either case, the grain-to-grain contact pressure would be similar under the given load within the hydrostatic-pressure zone. Reynolds (1973)reported an apparent relationship between slope of the normal compaction trend and geothermal gradient in the Gulf Coast. He suggested that the higher geothermal gradient corresponds to rapid compaction (the normal trend is more horizontal; Fig. 2-26).However, this apparent relationship must be examined carefully (see Stephenson, 1977).In an area of rapid deposition, where the shallow normal compaction trend is relatively vertical, a thicker undercompacted section may possibly have developed at depth. Heat flow through such an undercompacted area would probably be

1 2 -

3 4 -

c Y U

0

; I'

t

0

5 6 -

'8 9 10

-

11

-

12

-

13 14

-

Fig. 2-26. Apparent relationships between geothermal gradients and different slopes of the normal compaction trend in the Gulf Coast. (From Reynolds, 1973.)

36 relatively low, because of the relatively low thermal conductivity. This would result in a relatively low geothermal gradient. The rapid rate of sedimentation itself might also cause a lower temperature at a given depth, by moving relatively cool sediments to depth rapidly.

Analysis of structural timing using shale transit-time plots The transit-time-depth plots for three wells drilled in three separate structures in the Canadian east coast offshore area, are shown in Fig. 2-27. Well P is located in an area where there is no significant surface erosional feature on the seismic cross-section.The normal compaction trend is extrapolated to about 180 ps/ft, rather than 200 ps/ft at the surface. This difference is believed to be due to the difference in shale composition - the shales in this area are more calcareous than those in the western Canada and Beaufort basins. Therefore, let us assume 180 ps/ft is the non erosion surface transit

-em 180l4l . 7 w :

-6ooP

-4ooo

WELL E EROSION

9600'

-mm :

I-

1801Atl

Y

U Ly

I '

0

l0 Y

n 2000

4000

Boa)

8Ooo

'

'3ooo

f

UNC

Fig. 2-27. Normal compaction trends of three wells E , M and P in the Canadian east coast offshore area.

37 time for this area. A t about 9000 f t in this well, the seismic and paleontological data show evidence of erosion or unconformity. However, there is no significant break of the normal trends above and below this unconformity, probably because the erosional thickness was less than the thickness of sediment deposition after erosion. The record of compaction before erosion was completely removed by the loading of the new sediments, whose thickness has reached 9000 ft. There is an equivalent unconformity surface at about 2900 f t in well M. The normal trend established in the zone below the unconformity shows more rapid compaction with burial depth than in well P, and is extrapolated to 180 p / f t (no erosion transit time for this area) at about 700 f t above the surface. This would produce an estimated thickness of erosion of about 3600 ft (2900+ 700 ft). The slope of the normal compaction trend within the section above the unconformity is flatter (or more horizontal) than that of any other trend, suggesting that the rate of sedimentation there was quite slow. This slow sedimentation is also documented by the paleontological data. A t well E, the trend below the unconformity at 2600 f t has almost the same slope as that of well P, but is shifted to the left or more-compacted side. This trend is extrapolated to 180 ps/ft at about 7000 f t above the surface, indicating a significant erosion. The thickness of erosion in this well was estimated to be about 9600 f t (2600 + 7000 ft). The sonic-log data above the unconformity are not sufficient to allow any meaningful interpretation, but the transit-time values between 2300 and 2600 f t are similar to those at the same depth range in well M. Therefore, it is assumed that the compaction phenomenon and rate of burial there are similar to those in well

M. TABLE 2-111 Estimated thickness of erosion and rate of sedimentation for three wells, P, E and M, in the Canadian eastcoast offshore area

EROSION THICKNESS

SEDIMENTATION RATE

(CANNOT EST1 MATE)

FAST

9600 FT

FAST

3600 FT

INTERMEDIATE

38 The normal trend of well M, and possibly that of well E above the unconformity, are flatter than any of the deeper normal compaction trends; the rate of compaction for a given burial was greater for these shallow and young intervals. This is good evidence that geological age does not significantly affect the amount of compaction with burial, in that the older sections show less compaction for a given burial depth than the younger. There is no significant difference in shale composition above and below the unconformity. The results of the interpretations for these three locations are shown in Table 2-111. On the basis of the above-mentioned data and interpretation, the structural-timing analyses for these three locations will be made as follows:

Early Cretaceous (initial tectonism) Deposition and burial at wells P and E were relatively rapid. Rate of burial at well M was a bit slower. In other words, early structures could have been developed in the area, including well M. As a matter of fact, well M recorded some oil shows in sandstones around 9000 ft, but these were not found in any other wells.

Late Early Cretaceous (major tectonism) The most significant tectonic events occurred in the area including well E, and resulted in the removal of about 9600 f t of the shallower part of the sedimentary column. Most of the structures in that area were formed during this period. In other words, structuring there was later than at well M. The area including well M probably had formed a gentle anticline before this stage, but stronger structural events subsequently removed the top 3600 f t of the section by erosion. At well P , the thickness of sediments removed during this period was probably not great.

Late Cretaceous o r later (after major tectonism) In the area including well P, deposition continued at a relatively high rate. At wells M and E the rate and amount of burial was relatively slow and small. The loading at wells M and E during this period did not cause any significant fluid expulsion from the deeper section, because that section was already overcompacted for these new depths. On the basis of the above discussions, we may reach the following conclusions with respect t o the analysis of structural timing of this general area: (1)Structure was probably developed earliest at well M. (2) Structure at wells E and P was a relatively late event. Note that there is no significant difference in geothermal gradient in this study area.

39 WELL A

I

SHALE TRANSIT TIME

WELL B

I

RAPID SEDIMENTATION

SLOW SEDIMENTATION

EROSION

NO EROSION

LATE STRUCTURE

EARLY STRUCTURE

Fig. 2-28. Schematic diagrams showing the normal compaction trend of well A where there was rapid sedimentation followed by erosion, and well B where there was slow sedimentation without erosion.

Fig. 2-28 shows schematic transit-time-depth plots of two wells located at different structures. Both wells show the same transit-time value ( A t l )at the common depth D1, but their compaction and structural histories are different. The area including well A first experienced relatively rapid burial and deposition; then, at a later stage, the shallower part of the sedimentary column was truncated by erosion. The structure of well B, on the contrary, experienced a slower burial without late-stage erosion. In other words, the structuring was earlier there. Therefore, if other conditions for petroleum gineration, migration, and accumulation were the same for these two structures, we may be able to state that the structure of well B has the better chance of petroleum accumulation. Limitations and possible problems with the techniques Possible hazards in applying these techniques are associated with the quality of sonic-log data. Although corrections are usually made for changing hole size, the correction for the effect of shale hydration by the drilling fluid cannot be made easily. The degree of hydration would be affected by the type of shale, the type of drilling fluid, the elapsed time since penetrations,

40 SHALE TRANSIT TIME

-

Fig. 2-29. Schematic diagram showing the effect of shale hydration by drilling fluid on the transit-time plot.

and the level of natural shale compaction. Many field examples indicate that shale hydration is more significant in shallower intervals. The result could be a normal compaction trend flatter than the real trend, as shown in a schematic diagram of Fig. 2-29,which could lead to interpretations of too little erosion and too slow a rate of sedimentation. To avoid this possible misinterpretation, the sonic-log data must be checked carefully against other information, such as the well velocity survey (check-shot velocity). Use of the interval velocity or transit time from seismic data, such as shown in Fig. 2-19,has at least two advantages: (1) there is no shale-hydration effect, and (2) analysis of structural timing can be made before drilling in both anticlinal and synclinal areas. However, there are also some disadvantages: (1)the data show the averaged velocity or transit-time values for intervals of several hundred feet only, so that it is sometimes relatively difficult to draw a reliable normal compaction trend line, and (2)there is no easy way to distinguish shale values from those of other rock types.

Use of other logs The principal reason for using the sonic log in evaluating thickness of erosion and rate of sedimentation is that it is one of the commonest logs

41 that measure porosity or level of compaction in sedimentary basins. However, other porosity logs, such as formation-density and sidewall neutron logs (SNP) can also be used for this analysis. The result, however, is not as reliable as that from the sonic log, because formationdensity and SNP logs are usually more affected by hole conditions, and hole caving is quite common in shales. Induction or resistivity logs are not recommended for use in this anal-

I

I

I

......................................................... E ................................................. Location showing surface shale s a m p l e s . . ................................. Location of Strathrnore well 7-12-25-25-W4 ............................. Wells studied

Groups of wells

148

r _ _ _ _ _ _

.m

Fig. 2-30. Index map showing the location of the wells studied, and the section lines shown in Figs. 2-31, -32 and -33.

42

m

F' I

0

w

R

Y

Y

Fig. 2-31. Shale porositydepth plots of Cretaceous and Tertiary shales in the western Canada basin.

43

Fig. 2-32. Shale porosityllepth plots of Cretaceous and Tertiary shales in the western Canada basin.

44

ysis, because the conductivity or resistivity of shales reflects not only porosity, but also salinity of the formation water, formation temperature, and mineral composition. Examples of erosion estimates in western Canada The technique discussed above was applied to the western Canada basin (Fig. 2-30).Sonic logs were studied and the transit-time values were converted to porosity by using the relationship shown in Fig. 2-3.The values of shale porosity are plotted in Figs. 2-31and 2-32.The normal trends in the eastern parts of the basin are extrapolated to about 60% at the present surface, suggesting that there was no erosion. Note that 200 ps/ft corresponds to 62% shale porosity (Fig. 2-3).The interpreted thickness of erosion is shown in Fig. 2-33. UK

N

EROSION

I'

J

P

Iv

I'

I 5000 FEET

SCALE

100 MILES

0

Fig. 2-33. Geological sections of western Canada showing the estimated thickness of erosion.

45

References Athy, L.F., 1930. Density, porosity and compaction of sedimentary rocks. Bull. A m . Assoc. Pet. Geol., 14: 1-24. Connan, J., 1974. Time-temperature relations in oil genesis. Bull. A m . Assoc. Pet. Geol., 58: 2516-2521. Dickinson, G., 1953. Geological aspects of abnormal reservoir pressures in Gulf Coast Louisiana. Bull. A m . Assoc. Pet. Geol., 37: 410-432. Foster, J.B. and Whalen, H.E., 1966. Estimation of formation pressures from electrical surveys - offshore Louisiana. J. Pet. Technol., 18: 165-171. Ham, H.H., 1966.New charts help estimate formation pressures. Oil Gas J., 64: 58-63. Hedberg, H.D., 1936.Gravitational compaction of clays and shales. A m . J. Sci., 31: 241287. Hosoi, H., 1963. First migration of petroleum in Akita and Yamagata Prefectures. Jpn. Assoc. Mineral., Petrol. Econ. Geol. J., 49:43-55,101-114. Hottman, C.E. and Johnson, R.K., 1965. Estimation of formation pressures from logderived shale properties. J. Pet. Technol., 17: 717-722. Magara, K., 1968. Compaction and migration of fluids in miocene mudstone, Nagaoka Plain, Japan. Bull. A m . Assoc. Pet. Geol., 52: 2466-2501. Magara, K., 1971. Permeability considerations in generation of abnormal pressures. SOC. Pet. Eng. J., 11: 236-242. Magara, K., 1976. Thickness of removed sediments, paleopore pressure, and paleotemperature, southwestern part of Western Canada Basin. Bull. A m . Assoc. Pet. Geol., 60: 554-565. Meade, R.H., 1966. Factors influencing the early stages of compaction of clays and sands -review. J. Sediment. Geol., 36: 1085-1101. Morelock, J., 1967. Sedimentation and Mass Physical Properties of Marine Sediments, Western Gulf o f Mexico. University Microfilms, Ann Arbor, Mich., 141 pp. (Thesis, Texas A and M University). Pennebaker, E.S., 1968a. Seismic data indicate depth, magnitude of abnormal pressure. World Oil, 166: 73-78. Pennebaker, E.S., 1968b. An engineering interpretation of seismic data. SPE 2165, 43rd AIME Fall Meet., Houston, Texas, September. Perry, D.,1970. Early diagenesis of sediments and their interstitial fluids from the continental slope, northern Gulf of Mexico. Trans. Gulf Coast Assoc. Geol. SOC., 20: 219-227. Proshlyakov, B.K., 1960, Reservoir properties of rocks as a function of their depth and lithology. Geol. Neft. Gaza, 12: 24-29. Reynolds, E.B., 1973. The application of seismic techniques to drilling techniques. SPE 4643, 48th AIME Fall Meet., Las Vegas, Nev., October. Rieke 111, H.H. and Chilingarian, G.V., 1974. Compaction of Argillaceous Sediments. Elsevier, Amsterdam, 424 pp. Rogers, L.C., 1966.How Shell controls Gulf Coast pressures. Oil Gas J., 64: 264-266. Rubey, W.W. and Hubbert, M.K., 1959.Role of fluid pressure in mechanics of overthrust faulting, 11. Geol. Soc. A m . Bull., 70: 167-206. Staplin, F.L., 1969. Sedimentary organic matter, organic metamorphism and oil and gas occurrence. Bull. Can. Pet. Geol., 17: 47-66. Stephenson, L.P., 1977,Porosity dependence on temperature: limits on maximum possible effect. Bull. Am. Assoc. Pet. Geol., 61: 407-415. Terzaghi, K. and Peck, R.B., 1948. Soil Mechanics in Engineering Practice. Wiley, New York, N.Y., 566 pp. Weller, J.M., 1959.Compaction of sediments. Bull. A m . Assoc. Pet. Geol., 43: 273-310.

46 Wyllie, M.R.J., Gregory, A.R. and Gardner, L.W., 1956. Elastic wave velocities in heterogeneous and porous media. Geophysics, 21: 41-70. Wyllie, M.R.J., Gregory, A.R. and Gardner, G.H.F., 1958. An experimental investigation of factors affecting elastic wave velocities in porous media. Geophysics, 23: 459-493.

Chapter 3 CALCULATIONS OF PORE PRESSURE FROM SHALE COMPACTION DATA

Overpressures in the relatively young sedimentary basins in the world are believed to have been caused primarily by the compaction phenomena of sediments, especially shales. Hubbert and Rubey (1959) applied the soilconsolidation laboratory model by Terzaghi and Peck (1948) to the subsurface conditions. Terzaghi’s model and the subsurface model The theory of the compaction or consolidation of a water-saturated clay has been developed by soil-mechanics researchers. This concept can be explained easily by Terzaghi’s model. Fig. 3-1 shows such a schematic model, in which perforated metal plates are separated by metal springs in water in a cylindrical tube. The plates simulate the clay particles in the subsurface and the springs simulate the contact situation between the clay particles. The internal water pressure can be measured by a manometer attached to the cylinder. When the load S is applied to the uppermost plate, the height of the springs will stay unchanged unless water escapes from the cylinder. At this stage, the applied load S is supported entirely by the water pressure p, or s=p

(3-1)

(see stage A, Fig. 3-1). As some water escapes from the system, the plates move downward and the springs carry part of the load S (stage B). Stress carried by the springs is usually termed as u. As more water escapes from the cylinder, the springs become more compressed and carry a greater part of the total load. If enough water moves out of the system, compaction equilibrium will be reached (stage C). The water pressure at the equilibrium stage is hydrostatic. During this experiment, it is known that the following relationship exists: S=p+0

(3-2)

The value X shown in Fig. 3-1 is defined as the ratio of p over S. This model is analogous to a subsurface clay undergoing essentially uniaxial compaction due t o overburden pressure S (total load) at depth 2.The

48 S

I

S

U

A.1 Stage A

-

0465 < A

A =0465

1

Stage C

Stage 0

Plates

A=E

CZa Water

Fig. 3-1. Schematic presentation of clay compaction. (From Terzaghi and Peck, 1948.)

overburden pressure is given as:

where Pbw = mean

g

water-saturated bulk density of the overlying sedimentary rocks above depth 2,and = acceleration of gravity.

The stress of the springs, u, in Terzaghi’s model is analogous to the grainto-grain bearing strength of the clay particles, and p is the fluid pressure under the subsurface conditions (Hubbert and Rubey, 1959). Hubbert and Rubey stated that, “The effective stress u exerted by the porous clay (or by the springs in the model) depends solely upon the degree of compaction of the clay, with (T increasing continuously as compaction increases. A useful measure of the degree of compaction of a clay is its porosity 4, defined as the ratio of the pore volume to the total volume. Hence, we may infer that for a given clay there exists for each value of porosity 4 some maximum value of effective compressive stress which the clay can support without further compaction.” Hottman and Johnson (1965) concluded that, “The porosity 4 at a given burial depth D is dependent upon the fluid pressure p. If the fluid pressure is abnormally high (greater than hydrostatic), the porosity will be abnormally high for a given burial depth.” In the previous discussion of Terzaghi’s model, a given load S was applied instantaneously and the change of the springs associated with the water drainage with time was observed. If the load S were continuously increased while the water was being continuously allowed to escape, the model would simulate the continuous deposition of sediments more closely. In other

49

words, as long as the escape of water keeps pace with the increase in load, the springs become continuously more compressed ( u increases continuously). The internal water pressure p will remain continuously hydrostatic. This situation parallels that in the normal compaction zone, where shale porosity decreases continuously (0 increases continuously) as burial depth increases and the fluid pressure is kept at near-hydrostatic level. If the rate of escape of water does not keep pace with the continuous increase in loading, some excess water remains in the cylinder and the height of the springs is higher than in the previous normal case. The water pressure will be higher than hydrostatic. The cause of this anomaly is either that the load is increasing too fast, or the water is escaping too slowly. Under subsurface conditions, these factors respectively correspond t o a relatively rapid rate of sedimentation and a relatively low permeability or transmissibility of a rock. In other words, if either of these conditions existed in the geological past, we have a chance to find the fluid pressure in excess of hydrostatic. In some areas the development of growth faults associated with rapid deposition might also cause restricted water expulsion in lateral directions. Although Terzaghi’s model simulates the subsurface shale compaction phenomena quite well, there are several important differences between this laboratory model and the true subsurface condition, viz: (1)In the laboratory the temperature of the cylinder is usually kept constant, but subsurface temperature generally increases with depth of burial. (2) The water outlet in the laboratory model is of a constant size that allows water to escape at a given constant rate, but the permeability of sedimentary rocks decreases with burial or compaction. (3) The water pressure in the cylinder is hydrostatic as long as enough water is being drained from the system, but its value is near zero because the column is so short. Hydrostatic pressure in the subsurface, however, increases. In other words, hydrostatic pressures at depth are significant pressures. (4) The laboratory model starts out with a fixed volume of water in the cylinder, which decreases as some is allowed to flow out. Within a given block of rock in the subsurface, however, the water flow and pressure buildup could be affected by water moving from other blocks, either underlying or adjacent. In other words, subsurface flow is quite complicated. With the temperature increase associated with burial the water will tend to expand. If such expansion is restricted by a relatively closed system resulting from a combination of rapid sedimentation and low permeability, the water pressure will increase at a faster rate than it would in the laboratory model. If water is normally expelled from the sediments, the increase in temperature should not cause any significant change in pressuring. The problem of temperature effect on pressure, or the aquathermal effect, will be discussed in the next sections. Continuous decrease of sediment permeability with burial is also an important parameter in understanding subsurface fluid pressures. In other

50

words, the chances of overpressuring will increase with burial depth, even if the rate of sedimentation and the geothermal gradient stay constant. This problem will be discussed in the next chapter. The third point mentioned above is quite important in applying the results of experimental compaction in the subsurface. For example, a compaction experiment was conducted on a clay up to the total axis pressure of 10,000 psi. Water from the clay can move freely from the apparatus. If the average overburden pressure of a sedimentary basin is about 1 psi/ft, the 10,000-psi load corresponds t o the overburden pressure at a depth of about 10,000 ft. However, at 10,000 f t the hydrostatic fluid pressure is probably between 4400 and 4700 psi - much higher than the near-zero water pressure in the apparatus. By using eq. 3-2 and the hydrostatic pressure gradient of 0.44 psi/ft, the value o in each case is given as follows: Subsurface at 10,000 ft: u = 10,000 - 4400 = 5600 psi Experiment: u = 10,000 - 0 = 10,000 psi

.-

If the level of shale or clay porosity or of compaction is a function of u, these two cases would represent a significant difference in the level of compaction. In other words, this experimental result is analogous not to a shale bed at 10,000 f t , but to a bed at almost 20,000 f t with hydrostatic fluid pressure. The schematic diagrams in Fig. 3-2 show typical subsurface conditions in

s

o+p B

A SHALE POROSITY --L

DEPTH

=

PRESSURE

-

C

:LUID PRESSURE

DEPTH

DEPTC

I* S (I

p

- OVERBURDEN PRESSURE - GRAIN TO GRAIN BEARING - FLUID PRESSURE

c

STRENGTH

-

Fig. 3-2. Schematic diagram showing the relationship among overburden pressure S, fluid pressure p and effective stress u in normal compaction and undercompacted zones.

51 the abnormally and normally pressured sections. The upper part of Fig. 3-2A indicates the shale porosity relationship of the normal compaction zone. In the lower part of this diagram, the shale porosity is higher than the normal trend indicates. The shales in this lower interval are called “undercompacted shales.” Fig. 3-2Bdepicts the pressuredepth relationships for the same intervals. Overburden pressure does not increase at a uniform rate: it usually increases with depth because sediments commonly become denser the deeper they are buried. Fig. 3-3 shows the changing rate of increase of the overburden pressure in the Gulf Coast (Dickinson, 1953). Note that the line is quite close to the l-psi/ft line.. The figure of 1psi/ft is known to be a good approximation of the overburden pressure gradient in many sedimentary basins. If the overburden-pressure gradient is assumed to be constant, the overburdenpressure-depth relationship can be expressed by a straight line such as that shown in Fig. 3-2B. As Hubbert and Rubey (1959)suggested, the grain-to-grainbearing strength of a shale increases as porosity decreases (see Fig. 3-4),so that the value u in Fig. 3-2B increases with depth up to the base of the normal compaction zone. The difference between the overburden pressure S and the grain-tograin bearing strength or effective stress u is the fluid pressure p (eq. 3-2). The fluid pressure in this shallower interval increases at a uniform rate. In the undercompacted section, the value u is subnormal. Therefore, to support the overburden pressure jointly with the subnormal grain-to-grain bearing

PRESSURE OF OVERBURDEN-PSlG

Fig. 3-3. Overburden pressure versus depth in the Gulf Coast. (From Dickinson, 1953.)

.: 52

l'or SHALE POROSITY

r

>-0.1

DEPTH

k

In

P B

8000

0.01. 0 EFFECTI' OR GRAlh

2 3 STRESS ~ ( 1 0 8dyneslcm' '0-GRAIN BEARING STRENGTH

10,000

Fig. 3-4. Relationship between shale porosity f and effective stress or grain-to-grainbearing strength U. (From Hubbert and Rubey, 1959.) Fig. 3-5. Schematic shale porosity-depth plot.

strength, the fluid pressure must be abnormally high (Fig. 3-4). A schematic fluid-pressure-depth relationship in this case is shown in Fig. 3-2C. Calculation of fluid pressure Let us now estimate the fluid pressure in an undercompacted section, using a shale porosity-lepth plot as shown in Fig.3-5. The plot is made on semilog paper (porosity logarithmic and depth arithmetic), so that the normal trend is a near-straight line (see eq. 2-1). If the fluid pressure at 10,000 f t is to be evaluated, the shale porosity at this depth must be known. Then, a vertical line is drawn through this porosity value. The intercept between this vertical straight line and the normal compaction trend line is at 8000 f t in this schematic example. A t these two depth points (10,000 and 8000 ft) the shale porosity is the same, suggesting that the grain-to-grainbearing strength is also the same. If the overburden and hydrostatic pressure gradients are, respectively, assumed to be 1 and 0.44 psi/ft, the fluid pressure at 10,000 f t can be estimated by the following steps: (1)At 8000 ft: Overburden pressure S = 1X 8000 = 8000 psi Fluid (hydrostatic) pressure p = 0.44 X 8000 + 3500 psi

53 Grain-to-grain bearing strength (T = S -p + 8000 - 3500 = 4500 psi (2)A t 10,000 ft: Overburden pressure S = 1X 10,000 = 10,000 psi Grain-to-grainbearing strength at 10,000 f t is the same as that at 8000 ft, hence u .ir 4500 psi Fluid pressure p = S - u + 10,000- 4500 = 5500 psi Because the hydrostatic pressure at 10,000 f t is about 4400 psi (0.44X lO,OOO), this calculated fluid pressure is about 1100 psi in excess of hydrostatic. A more generalized equation to calculate fluid pressure was reported by Magara (1968),as:

or :

where p 2,

= fluid pressure at depth 2, = shallower depth at which the shale porosity on the normal com-

paction trend equals the shale porosity at depth 2 (see Fig. 3-5), yw = density of the formation water in psi/ft, and Ybw = mean density of the sedimentary rocks in psi/ft. The value of yw ranges from 0.433 psi/ft (fresh water) to 0.465 psi/ft (80,000 ppm NaCl solution) and the value of ybw is usually 1 psi/ft or less in relatively young sedimentary basins. In the above discussion the shale porosity plot was used, but the transittime plot can be used equally well. Eqs. 3-4or 3-5 mean that the fluid pressure at depth 2 is the sum of the hydrostatic pressure from the surface to depth 2, and the overburden pressure between 2, and 2. The same concept can be explained by the use of a continuous-burial model of undercompacted shale at present depth 2, as follows. Compaction of this shale in the earlier stages was normal, resulting in normal pore-fluid expulsion. At depth Z,, fluid expulsion was arrested completely, so that during subsequent burial to 2 there was no compaction (Fig. 3-6). In some undercompacted shales, compaction and consequent fluid expulsion may continue at a diminished rate. At this stage of discussion, however, the simple model of “normal compaction-no compaction” is used in that the effect of late-stage compaction in the undercompacted shales is removed,

54 SHALE POROSITY

I

I

DEPTH

Z,---

z

---

Fig, 3-6. Schematic shale porosity-depth plot, showing the history of the shale porosity change.

and evaluation of aquathermal and nonaquathermal effects is therefore easier. Pore pressure in this shale when it was buried t o 2, was hydrostatic. The increase in pore pressure during burial from 2, to 2 is equivalent t o the increase in overburden pressure, provided there is no temperature increase between 2, and 2. If the overburden-pressure gradient is 1 psi/ft, the increase is given by 2 - 2, psi. Aquathermal pressuring What will happen if the temperature increases between depths 2, and Z? Barker (1972) discussed this problem, using the temperature-pressuredensity diagram for water. The diagram in Fig. 3-7 shows the relationship. Specific volume (cc/g), which is the reciprocal of density (g/cc), is indicated in brackets. On the basis of the average geothermal gradient of 25"C/km (or 1.37"F/100 ft) in the Louisiana Gulf Coast, 1km (3300 f t ) of burial after complete isolation would cause a pressure increase of about 6000 psi (see points L and M in Fig. 3-7). In this case, pore water in shale is assumed t o be completely isolated at L, then buried for a 1-km (3300 ft) interval to M, which corresponds t o a temperature increase of 25°C. During this burial a given weight of water is assumed t o keep a constant volume (density or specific volume is constant). In other words, the shale pore volume stays constant during this period. The rate of pore-pressure increase in this case is about 1.8 psi/ft (6000 psi/3300 ft),which is almost twice as great as for the previous example (no temperature increase). If the geothermal gradient is

55

SPECIFIC VOLUME

1o.Ooo

L u U

! P

5 000

0

I

I

32

100

1

1

200 300 TEMPERATURE OF

1

I

4w

500

Fig. 3-7. Pressure-temperature-density (or specific volume) diagram for water. The geothermal line of 25’C/km for hydrostatically pressured fluids is superimposed on a basic diagram adapted from Barker (1972). (From Magara, 1975.)

greater than 25”C/km, of course, the rate of pressure increase due to isolation is greater. The two lines in Fig. 3-8 show the relationships between the increase of burial depth (X or Z - Z e in Fig. 3-6) after isolation, and the pressure increase in the two previously mentioned cases: (A) isolation without temperature increase (1psi/ft); and (B) isolation with temperature increase (1.8 psi/ft). The first case (A) may be termed as “nonaquathermal pressuring” and the second case (B) as “aquathermal pressuring” (Barker, 1972). Under actual subsurface conditions, temperature usually increases with burial. Therefore, the actual data would be plotted above line A , provided the overburden-pressure gradient is 1 psi/ft. The relationship indicated by line B would exist only if there was perfect isolation of pore fluids with increasing temperature. In nature there is no such perfect condition; there-

56

0

Zoo0

4000

6Mw)

BURIAL DEPTH INCREASE SINCE ISOLATION

13oo0 FT

8oo0

Z -Ze

0R X

Fig. 3-8.Relationship between increases of burial depth (2 - 2,) and of pressure ( p p e ) since the isolation of fluids. Numbers refer to the well in Hottman and Johnson‘s table 1 (1965) (see also Table 3-1).

fore, the actual subsurface data may show values between lines A and B. Hottman and Johnson (1965)demonstrated the relationship between the shale compaction anomaly detected by sonic log, and the ratio of pore pressure to depth. Table 3-1 shows the results in their table 1, which lists the depth (Z), the measured pore pressure (p), and the transit-time anomaly (Atobserved - Atnormal,see Fig. 3-9)values for eighteen wells in the Gulf Coast. The Atobserved- Atnormal value is converted to the X (or z - Ze) value as discussed above (Fig. 3-9).The normal compaction trend by Hottman and Johnson (1965,fig. 2) is used for this conversion. The hydrostatic pressure (p,) for the isolation depth (2,) is calculated with a hydrostaticpressure gradient of 0.465 psi/ft. The hydrostatic pressure (p,) is then subtracted from the measured pressure (p) for the eighteen samples given by Hottman and Johnson (1965);the value p - p e is the increase of pressure due to burial from depth 2, to 2 (or X). The porosity4epth relationship as shown in Fig. 3-6is now replaced by the transit-time-depth relationship as shown in Fig. 3-9.As the transit time is considered to be the measure of porosity for the uniform lithology (in this case shale), the increase in pore pressure due to isolation also might be examined by using the transit-time data. In Fig. 3-8the values of p - p e from the eighteen wells in the Gulf Coast are plotted against 2 - Z e or X; numbers refer to the well numbers in Table 3-1.The basic premise in constructing this diagram is that the mea-

57 TABLE 3-1 Pressure and shale acoustic log data, overpressured Miocendligocene wells (From Hottman and Johnson, 1965) Parish or County and State

Well

Terrebonne, La. Offshore Lafourche, La. Assumption, La. Offshore Vermilion, La. Offshore Terrebonne, La. East Baton Rouge, La. St. Martin, La. Offshore St. Mary, La. Calcasieu, La. Offshore St. Mary, La. Offshore St. Mary, La. Offshore Plaquemines, La. Cameron, La. Cameron, La. Jefferson, Texas Terrebonne, La. Offshore Galveston, Texas Chambers, Texas

1 2 3 4 5 6 7 8 9 10 11 12 13 14 15 16 17 18

Pressure (Psi) @)

13,387 11,000 10,820 11,900 13,118 10,980 11,500 13,350 11,800 13,010 13,825 8,874 11,115 11,435 10,890 11,050 11,750 12,080

11,647 6,820 8,872 9,996 11,281 8,015 6,210 11,481 6,608 10,928 12,719 5,324 9,781 11,292 9,910 8,951 11,398 9,422

* Formation fluid pressure gradient (pressure/depth).

Fig. 3-9.Schematic shale transit-timedepth plot.

FPG * W / f t1

0.87 0.62 0.82 0.84 0.86 0.73 0.54 0.86 0.56 0.84 0.92 0.60 0.88 0.90 0.91 0.81 0.97 0.78

22 9 21 27 27 13 4 30 7 23 33 5 32 38 39 21 56 18

58 sured pore pressures in the sandstones are in equilibrium with the pressures in the shales. The compaction levels in the shales may be related to the pore pressures in the sandstones in this case. The possibility that some other factor, such as “charging” of the reservoirs by aquifers at greater depth, could be responsible for the configuration shown in Fig. 3-8has been rejected, on the basis that in such circumstances the plot would be much more irregular. The premise that the shale pore pressures are in equilibrium with the sandstone pore pressures is also the general basis for pore-pressure prediction techniques using shale compaction data. All the data in Fig. 3-8 are plotted between lines A and B. The average rate of pressure increase (p - p , ) / ( Z - 2,) for the eighteen wells is about 1.4 psi/ft, a value midway between line A (no temperature increase since isolation) and line B (temperature increase of 25”C/km or 1.37”F/100 f t with perfect isolation). There may be two possible explanations for the actual subsurface data shown in Fig. 3-8: (1)There is no perfect isolation of fluids in the undercompacted shales, so that water has been expelled, resulting in reduction of pore pressures from the values shown by line B. (2) The isolation of fluids has been almost perfect, but the rock (shale) has expanded, again resulting in pore-pressure reduction. No water would have been expelled in the second case. Although both situations may have existed in nature to some extent, I think the first is the more important. Shales, even massive ones, always have some permeability, so that it is always possible for water to move out of them. The expansion of rocks at depth, on the other hand, is not easy to explain geologically, although there is no physical reason to reject the possibility. If the first assumption is applicable to the subsurface, the reduction of pressures from line B (perfect isolation) in Fig. 3-8 can be related to the amount of water expelled; the reduction for well 5, for example, is shown by M. The greater the amount of expelled water, the greater the pressure reduction. The net aquathermal pressure increase for well A is shown in Fig. 3-8 by N. If the overburden-pressure gradient in this area was less than 1 psi/ft, as suggested by Eaton (1972, fig. 5), the net pressure increase due to aquathermal effect would be more than the amount shown here. In any case, the plot in Fig. 3-8 would indicate that the undercompacted shales in the Gulf Coast have reasonably good sealing capacity, resulting in pressure development due to the aquathermal effect. Semiperfect isolation of pore water has been developed in these shales. The generalized equation for average aquathermal pressure is given as follows:

+

p(psi) y,Z,(ft) + 1.4(Z(ft)- Z,(ft))

(3-6)

59 The factor 1.4 (psi/ft) in this equation gives an approximate fluid pressure when the geothermal gradient is about 25”C/km, and the shale property is similar to that of the average Gulf Coast shales. Now let us estimate the possible aquathermal pressure at 10,000f t in the case shown in Fig. 3-5.The calculation can be made by using eq. 3-6as follows (yw = 0.44 psi/ft): p = 0.44X 8000 + 1.4 X (10,000- 8000)+ 6300 psi

The pressure value calculated here is 800 psi higher than the previous (nonaquathermal) case, although the shale porosity4epth plots are identical in both cases. Application of aquathermal and nonaquathermal concepts

A question may arise; “When should we use the aquathermal or the nonaquathermal concept?” The following sections will provide answers. Areas of continuous deposition and burial If sedimentation and burial continued in the geological past, the undercompacted section would be influenced by the aquathermal effect. Whether an area experienced continuous burial or significant erosion can be examined by studying the normal compaction trend (Chapter 2) and/or by the use of seismic and paleontological information. Fig. 3-10depicts a possible fluidpressure history of a block of shale. Compaction and fluid expulsion from the surface to depth L were normal and hydrostatic. A t depth L, semiperfect isolation of pore fluid was established and burial of the block continued to N with accompanying temperature increase. In Fig. 3-10,the rate of net aquathermal pressure increase is assumed to have been 1.4 psi/ft which corresponds t o the average for Gulf Coast shales. If the geothermal gradient is slightly lower, or if the shale’s retention capacity of aquathermally generated fluid pressure is lower, then the rate of net pressure increase would be less than 1.4 psi/ft. In any case, in an area where deposition was continuous, an aquathermal-pressuring mechanism can be expected. If, however, the shales in that area are relatively very silty or sandy, to the extent that aquathermally generated excess pressures could not be retained, we may encounter almost nonaquathermal pressuring. In this case, the rate of pressure increase is expressed by a line close to line 2 (Terzaghi line) in Fig. 3-10.In the Gulf Coast and Mackenzie Delta areas, a rate of pressure increase of about 1.4 psi/ft is known to be good approximation when continuous deposition took place.

60

I PRESSURE GRADIENT1

1.8 PSI/FT

[

ISOLATION DEPTH

-

'MAX

BURIAL DEPTH

DEPTH AFTER EROSION DEPTH

Fig. 3-10. Schematic diagram showing the generation of aquathermal pressure during continuous burial, and dissipation of pressure during uplift and erosion.

Areas o f significant uplift and erosion

As shown in Fig. 3-10, fluid pressure will drop very sharply during and after uplift and erosion (from N to P). The rate of pressure drop would be about 1.8 psi/ft, if the geothermal gradient was 25"C/km (1.37"F/100 ft), and if the subsurface temperature after erosion was adjusted relatively quickly to this gradient. As long as point P stays above line 2 there would be no further compaction of shales because the aquathermally generated pore pressure can support the overburden pressure jointly with the grain-to-grain strength of the shales. If point P drops below line 2, the shales must compact further to adjust to the new physical situation. This late-stage compaction associated with the late-stage uplift and erosion would not produce any significant fluid expulsion from shales. If the pressure point dropped below line 2 and if enough geological time was allowed for further compaction adjustment, Terzaghi's (nonaquathermal) concept would predict pore pressures reasonably well. In this case, the original net aquathermal pressure is considered to have been dissipated completely. This critical nonaquathermal stage, or the stage at which point P reaches Terzaghi's line 2 may be reached when the thickness of erosion reaches about one-half of the thickness X (burial increase since isolation of pore fluid; see Fig. 3-9).

-

61

SHALE TRANSIT T I M E OR P O R O S I T Y

DEPTH

J Fig. 3-11. Schematic shale transit-time-depth plots. Regional normal compaction trend (A) and plot after significant erosion ( B ) .

If uplift and erosion had continued further, the fluid pressure could have become much less, to the level at which it would be near hydrostatic or even lower. This final situation can be explained by the use of Fig. 3-11,where line A shows the regional or areal normal compaction trend and line B depicts a shale porosity (or transit-time)-depth plot of a well in this area. Note that the shale porosity in the undercompacted zone (deeper part of B) is compacted more than the regional or areal normal compaction trend. If the erosion reaches this level, fluid pressure in the apparent undercompacted zone would be near or even below hydrostatic. This situation is observed in the Canadian Foothills area (see Chapter 2). In summary, to make a realistic estimate of pore pressure from shale compaction data, one must know the geological events of the area - especially uplifting and erosion, which could have reduced subsurface temperature significantly. This is because the effect of temperature on pressure is quite significant in a formation with a closed fluid system. Possible expansion of the rock framework during and after uplift and unloading could also cause further pressure reduction, but it might not be as significant as the effect of fluid shrinkage associated with temperature drom Estimation of pore pressure by the use of charts The manual calculation of pore pressures mentioned in the previous section can be simplified by the use of charts. Fig. 3-12is the chart of nonaquathermal pressuring in which the average bulk density of sediments and the density of water are assumed to be 1.00 psi/ft (2.31 g/cc) and 0.435 psi/ft (1.005 g/cc), respectively. However, estimating pore pressure in an area

62

NONAQUATHERMAL h = 0 . 4 3 5 pii/lt ?b.=l

00 pri/fl

0

Fig. 3-12. Nonaquathermal-pressuredetection chart.

where water density is higher than 0.435 psi/ft (for example, 0.465 psi/ft as commonly used in the Gulf Coast) would not result in a serious error. A sample shale porosity (or transit-time) plot is shown on the top right of this figure. In this example the deeper depth 2 at which we wish to estimate fluid pressure is 8500 f t and the shallower compactionequivalent depth 2, is 6500 ft. First enter depth 2 (8500 ft) on the bottom depth scale, then proceed vertically to the hydrostatic-pressure line. From this point draw a line toward the top left parallel to the diagonal lines, until this line intercepts the vertical line at depth 2, (6500ft). Then proceed horizontally to the left to read the formation fluid pressure in psi. In this example, the estimated fluid pressure is approximately 4800 psi, which is about 1100 psi in excess of hydrostatic pressure. The corresponding mud weight can be obtained by extending the depth line (2) upward and the pressure line to the right, and finding the point of intercept. In this example, mud weight is about 10.9 lb/gal. Fig. 3-13shows the aquathermal chart in which the rate of fluid-pressure increase after isolation is assumed to be 1.4 psi/ft. The method of estimating fluid pressure is essentially the same as in the previous nonaquathermal

63

example. Although the set of two depth figures (2 and 2,) is the same as in the previous example (8500 and 6500 ft), the estimated fluid pressure is much higher (about 5600 psi). Correspondingly, the estimated mud weight is higher (about 12.8 lb/gal) than in the case of nonaquathermal pressuring. Relationship between fluid pressure, depth and equivalent mud weight As shown in the previous two examples of pressure estimation, fluid pressure/depth determines the equivalent drilling-mud weight. The relationship between fluid pressure, depth, and equivalent mud weight can be explained by means of the chart in Fig. 3-14.This chart can be used in two ways: (1)to convert from mud weight t o pressure, and (2) to convert from pressure t o mud weight. The chart (Fig. 3-14)consists of two distinct but related segments: The top part relates mud weight in (lb/US gal) to the pressure/depth ratio. As engineers tend to use the term “pressure gradient” to express “pressure/depth”, both terms are shown on the chart. The reader should

64

65 I prefer and recommend the proper term pressure/depth; pressure gradient is included in parentheses only because that term is so often used. The bottom part of Fig. 3-14 relates overburden pressure (along the top) to fluid pressure (along the bottom) by a series of diagonal depth lines. The relationships that can be derived from them are described in the next section. Suppose that the mud-weight value at a certain depth is already known. In our example, the mud-weight value is 16.5 lb/gal (at 12,500 ft). Enter this value as indicated by @ in Fig. 3-14.Move vertically upward to intercept the diagonal line at @. Move from @ horizontally to the left to read the The pressure/depth value obtained in this pressure/depth (psi/ft) at 0. example is 0.855 psi/ft. To obtain pressure values, move this pressure/depth value (0.855psi/ft) to the bottom graph and stop at the appropriate depth (12,500f t in this case) as shown by @. To read the total fluid pressure, move vertically downward from @ to @. The value in this example is 10,650 psi. If a line is drawn through @ and parallel to the equal-depth lines, the hydrostatic pressure can be read at the bottom end @ and the overburden pressure at the top end 0 .The hydrostatic pressure in this case is about 5450 psi (~0.435 X 12,500).The difference between points @ and @ is the excess pressure above the hydrostatic pressure (10,650- 5450 = 5200 psi). The overburden pressure at 12,500 ft is 12,500 psi based on 1 psi/ft gradient, and is read at 0 .The difference between the overburden pressure and the total fluid pressure of 10,650psi, read at @, is the grain-to-grain contact pressure based on Terzaghi’s concept; i.e., the proportion of the overburden pressure carried by the rock framework rather than by the pore fluids. The chart can also be used to derive a mud-weight value from a fluid-pressure value. Assume that the measured pressure (by DST, for example) or estimated pressure at 12,500 f t is 10,650 psi. Now we enter the chart at @ and pass through @, 0, @ and @ to obtain a mud weight of 16.5 lb/gal, or through @ and 0to obtain a pressure/depth value of 0.855psi/ft. Even if the hydrostatic-pressure gradient is not 0.435 psi& the chart of Fig. 3-14 can be used to calculate hydrostatic pressure as follows: Let us assume that the hydrostatic gradient in the study area is 0.465 psi/ft. On the bottom segment of the chart, draw a horizontal line through the pressure/ depth value of 0.465 psi/ft. This line now becomes the bottom margin of the chart, and the hydrostatic-pressure value is read from it. For example the hydrostatic pressure at 12,500 f t when the gradient is 0.465 psi/ft is about 5800 psi. In effect, we have simply shifted @ diagonally upward along the 12,500-ftdepth line to intersect the new gradient line of 0.465 psi/ft. The overburden pressure when the gradient is not 1 psi/ft can be calculated in the same way, by drawing a horizontal line corresponding to the correct overburden gradient, and reading the value on it rather than on the 1.0 psi/ft as shown.

66 Note that the total fluid pressure @ obtained by this chart is independent of changes in the hydrostatic and overburden pressure gradients. In other words, the “mud weight” (or pressure/depth)-“depth”-“total fluid pressure” relationship has nothing to do with the hydrostatic and overburden gradients. Direct estimation of mud weight Equivalent mud weight can be evaluated directly from a transit-time plot by the use of a series of charts, found in Appendix 3-1 *. They include a set of five nonaquathermal charts (Nos. 1-5), which correspond t o five different slopesof the normal compaction trend (c = -0.00006, 0.00008,0.00010, 0.00012, and 0.00014 ft-l), and another set of five (Nos. 1-5) for the aquathermal cases. To evaluate mud weight from a transit-time plot, the following steps must be taken: (1)Make a “shale transit-time-Aepth” plot using semitransparent semilog paper (transit time on log scale, depth on arithmetic scale), the same size as the charts. Determine the normal compaction trend by drawing a straight line through the average of the transit-time values in the relatively shallow depths. (2) Determine which set of charts to use. If significant erosion is observed by the shift of the normal compaction trend (extrapolated surface transit time is less than 200 pslft), use the nonaquathermal set. Otherwise, the aquathermal set should be used. (3) Choose the appropriate conversion chart from the five (Nos. 1-5) provided in each set; the “appropriate” one will be the one with the compaction trend that best fits the normal trend of your plot. (4) Match the depth on the chart with the depth on your plot, then shift your plot laterally, if necessary, until the “normal compaction trend” on it is superimposed on that of the chart. Never shift your plot vertically to obtain a match. ( 5 ) Read mud-weight values (lb/gal) directly from the chart. Fig. 3-15 shows an example of a transit-time plot of a well in northern Canada. The nonaquathermal chart was used because of significant erosion. The mud weight actually used for drilling and some drill-stem test pressures are also shown for comparison. The drill-stem test pressures at points 5 and 6 are higher than those of the surrounding shales. This is probably caused by charging (see Chapter 4). The pressures at points 7 and 8 are similar to those of the nearby shales.

* The ten charts comprising Appendix 3-1 are inserted separately at the end of the book.

67 TRANSIT TIME (ps/ft)

50

100

200

150

VAL1

i I

iHT

I

9 10 11 12 13 14

LBIGAL Fig. 3-15. Example of a shale transit-time-depth plot in northern Canada.

68 Calculation of fluid pressure by computer Fluid pressure can be calculated from the sonic log by means of the computer. For this calculation, a new equation derived from the old nonaquathermal eq. 3-5can be used more conveniently. The mathematical form of the normal compaction trend is as follows: Ate = At; e-cze

(3-7)

where At, = transit time at depth 2, within the normal compaction zone, and At; = extrapolated surface transit time. Rearranging this equation, we obtain:

(3-8) At the two depth points Z and Ze, the transit time is the same, or: At = Ate

(3-9)

Using eqs. 3-5,3-8,and 3-9we obtain:

(3-10) If

Ybw

and yw are 1 and 0.435 psi/ft, respectively, we will have.

0.565

p(psi) = Z(ft) - -loge C

(2)

(3-11)

This equation predicts nonaquathermal pressures. If aquathermally generated pressure exists in the subsurface, the following equation can be used :

0.965

p(psi) = 1.4 Z(ft) - -loge C

($)

(3-12)

Equivalent mud weight, MW, can be calculated from pressure/depth (psi/ ft), as follows: P(PSi) MW(lb/gal) = 19.27Z(ft)

(3-13)

69 TRANSIT TIME, Ms/sec

1po ,

?

290

?

?

2 SANDsTONES

SANDSTONES

I

f

:I.

I I.

3

.j

* '

c Lu LL Lu

NORMAL TREND

I:

.I'

::yHY D R 0 ST AT IC

PR ESS UR E

0 0

0 X

I ' c Lu Q

4

a

5

6 SANDSTONES

SANDSTONES

A

B

Fig. 3-16. Computer-generated shale transit-time-depth and pressure-depth plot of a Mackenzie Delta well.

The value of c (slope) can be obtained from a hand plot of shale transit times by establishing a normal compaction trend. Such a transit-time plot can also be made by computer. Fig. 3-16A shows an example of a computer plot, in which transit time is

70

2000-

rn

GRAVEL 4000-

0

SAND 6, SHALE 6000-

En.

0

SHALE

;800010,000-

12,000-

PRESSURE ZONES

U

NORMAL 14,000-

OVER

k l m 16,000-

ABNORMAL

Fig. 3-17. Example of a computer-generated pressure4epth plot of the Taglu well, Beaufort Basin.

plotted for shaly intervals only. Shale data can be selected by additional use of the gamma ray or the SP log, which can distinguish shaly zones. The normal compaction trend and the value c can be determined on this computer plot by calculation (eq. 2-6)or from the chart (Fig. 2-23). An example of a computer-generated fluid-pressure plot for the same well is shown in Fig. 3-16B.The advantages of using the computer for calculating pressures are (1)it is quick and accurate, and (2)it is quite easy to make a fluid-pressure profile of a well. Such a fluid-pressure profile can be used to analyse subsurface drainage, as discussed in Chapter 5. Fig. 3-17 is another example of a fluid-pressure plot made for the Taglu well in the Beaufort Basin. Fig. 3-18 is a pressure plot with an expanded depth scale. In the relatively sandy zones, rock composition and calculated porosity are shown (Appendix 3-11).Calculated pressures and equivalent mud weights are shown in shaly intervals only. Pressures in Figs. 3-17and 3-18are calculated from sonic logs. Use of other logs for pressure estimates A resistivity or a conductivity log may be used as well as a sonic log. The resistivity or conductivity of shale is affected not only by its porosity, but

71 58 PER CENT F I L T E R

DEPTH

VOLWE FRACTION FROn 6R-SONIC

HUD UEIOHT

: *

1100 :

: : :

:

*

I

*

I

:

*

:

*

*

:: * :

: : 8000 : : : :

:t

: :

:

*

*

:

8100 :

I

: *

*

** *

I

8300

I *

:*

8481

85 F B

1.1.1.1.l.l

r...~l....l....l...:l....I....I....I....1....1....1 0

4000 6000 PRESSURE FRW E L 1 4 T ( P S O 2000

8000

10000 8

10 i3 IUD U L l W (L(V6AL 1

Fig. 3-18. Example of a computer-generated pressure-depth plot. In sandy intervals, Calculated amounts of quartz (Q), clay (f) and porosity (open space) are shown.

72 CONDUCTIVITY AT lOO'F

PRESSURE, X 1000 psi

mmho

10

2

1000

.100 ,

, '

7

SANDSTONES

4

e

SANDSTONES

3 I-

w

!

Lu Y

0 0

i

0

4- NORMAL

X

I- HYDROSTATIC PRESSURE I

I '

t w

I

TREND

4 i

0

I

1

5

6

SANDSTONES

A

SANDSTONES

B

Fig. 3-19. Computer generated shale conductivity (at 100°F)-depth and pressuredepth plots.

also by pore-water salinity and temperature. Correction for a salinity change is not easy, because usually there are no salinity data available in a shale zone, and the salinity of nearby sandstones may not represent that in shales. However, a correction for changing temperature can be made relatively easily, especially when the computer is used for calculation. Fig. 3-19A shows an example of a shale conductivity plot by computer. This is the same well as that shown in Fig. 3-16.The conductivity values at 100°F are plotted

73 versus depth on semilog scales. A near-straight-linerelationship is established between 3000 and 4700 f t , but the conductivity values above 3000 f t are lower than this trend indicates. Probably they result from the freshening of formation water by near-surface water. Pore pressure was calculated and plotted by computer (Fig. 3-19B). In other words, possible salinity changes in the shale pore water are ignored, and shale conductivity is assumed to indicate shale porosity. Fig. 3-19Bshows that the calculated fluid pressure below 4700 f t is higher than hydrostatic. The pressures in this zone show slightly higher values than those estimated by sonic log (Fig. 3-16B),probably because of the higher conductivity due to the higher formation-water salinity there. Other wire-line logs such as formation density and neutron can also be used for pressure estimation. However, they are not run as frequently as sonic and resistivity or conductivity logs, and are usually more affected by borehole condition, so that their use for evaluating fluid pressures is more limited. Empirical method for estimating fluid pressure The above-mentioned methods are essentially theoretical methods based on Terzaghi’s concept and the laboratorydocumented aquathermal concept. In applying these techniques in an actual basin, we must check their validity by comparing the resultant estimates with the actually measured fluid pressures. In other words, they are not entirely theoretical methods, but methods that combine theory with real data. There is another method which is based entirely on empirical data on shale compaction and measured fluid pressures. In 1965,Hottman and Johnson proposed techniques using sonic and resistivity logs.

Sonic method Hottman and Johnson first established the normal trend such as shown in Fig. 2-5 (Chapter 2), for the Texas and Louisiana areas. If an overpressured section is encountered, the transit-time values deviate from the normal compaction trend toward higher values as shown in Fig. 3-20.The density of shale in the overpressured section is usually subnormal. The amount of deviation at a given depth was related to the measured pressure in the nearby reservoir rocks. Fig. 3-21 shows schematically how the transit-time deviation is measured. Fig. 3-22depicts the relationship between the transit-time deviation (At,,,, - At,) and the pressure/depth for Miocene and Oligocene formations of the area Hottman and Johnson studied. The data used to make Fig. 3-22are shown in Table 3-1.The standard deviation from the curved line representing the data of Fig. 3-22 is 0.020 psi/ft, which corresponds to approximately 0.4 Ib/gal mud weight.

74 Well "H" Jefferson Co, Texas 2000 O[

I

4000 -

6ooor c

$

n

-10.000

-

12.m

-

I I I

dt(Sh)

2

' .

ps/ft

Pb(sh) *

dCc

d t ( s h ) , p s /ft

Fig. 3-20. Example of a shale transit-time-depth plot of a Gulf Coast well. Right-hand side shows a plot of the bulk density. (From Hottman and Johnson, 1965.) Fig. 3-21.Schematic diagram showing how the difference between observed (At,) and normal (At,) transit times are read. (From Hottman and Johnson, 1965.)

%b(sh)

- dtnlsh)*

PSIft

Fig. 3-22.Relationship between transit-time difference (At,b - At,) and pressure/depth (or FPG) in the Gulf Coast. (From Hottman and Johnson, 1965.)

75

-

4000- Well " R " Cameron PH. LA

6000 -

k

8000 -

c

1

Trend of hydrostatic

I I

.*I* \'TOP

Of

over--

+

Est. FPG from shale traveltime x Mud gradient oPressures from tests

10.000 -

14,m

t1

16,000

A

tB

I

I

I

Fig. 3-23. Example of a shale transit-time-depth plot of a Gulf Coast well. (From Hottman and Johnson, 1965.)

The method for estimating fluid pressure in a new well drilled in the same general area is as follows: (1)A shale transit-time-depth plot is made. (2)The normal compaction trend is established by using data from relatively shallow intervals. (3)The deviation of the shale value from the extrapolated normal compaction trend value is measured at the depth where the pressure is to be estimated. (4)k o m Fig. 3-22,the pressure/depth corresponding to At,, - Atn is found. (5) The pressure/depth value is then multiplied by depth to obtain pressure. Fig. 3-23shows an example of such a pressure estimate from a sonic log. Resistivity method

Possible factors influencing the resistivity of water-saturated rocks are (1) porosity, (2)temperature, (3)salinity of the formation water, and (4) mineral composition. In relatively clean shale sections of a given area, the shale composition may be assumed to be relatively uniform. The effect of changing temperature can be corrected as shown in Fig. 3-19A.However, the effects of porosity and salinity need some explanation. The relationship between the resistivity of a water-saturated rock, R, the formation-water resistivity, R, , and the formation-resistivity factor, F, is as

76 follows (Schlumberger, 1972):

R

= FRw

(3-14)

According to Archie (1942),F is given as:

(3-15) where a and m are coefficients; rn is called “cementation factor”. Introducing eq. 3-15into eq. 3-14,we obtain:

(3-16) If a and m are assumed to be 1 and 2 (the most standard values), respectively, we obtain:

(3-17) This equation means that the rock resistivity decreases as the water resistivity decreases (salinity increases) and/or the porosity increases. In undercom4ooo-

5

6000

-

8000

-

10,000

S-

c

a

2

12,000 14,000

\

\

i

I

1 ) Average Oligocene-

16.000

‘-,

miocene, SW, L A 2 ) Miocene, Jefferson PH, LA ’\ \ ‘x 3 1 Miocene, Iberia PH, LA

R(,,,,

‘

Ohm-meters

Fig. 3-24. Several normal compaction trends of shale resistivity in the Gulf Coast. (From Hottman and Johnson, 1965.)

77 O4

r

I

101 10

1

I

l

15

I

l

20

30

I

I

40

I

I

50

Normal -pressured R(5h)/observed R(5h)

Fig. 3-25. Relationship between the ratio of normally pressured shale resistivity over observed shale resistivity and pressure/depth (or FPG) in the Gulf Coast. (From Hottman and Johnson, 1965.)

pacted shale sections, the pore-water salinity is known to be subnormal and the porosity is abnormally high. However, the effect of porosity could override that of salinity, because porosity has a square effect in eq. 3-17.As a matter of fact, it is known that in most young sedimentary basins resistivity of an undercompacted shale is less than that of a normally compacted one Est. FPG from shale resistivity x Mud gradient

8000 -

i

0

pressures f r o m tests

10,000 -

c P 0

12,000-Test

'

d.4

' d6'd.S'k~ '

R(Sh)J

A

Rm

210 0.4

0.6

0.8

10

1.2

FPG, psi/ft

B

Fig. 3-26. Example of a shale resistivity-depth plot of a Gulf Coast well. (Fom Hottman and Johnson, 1965 .)

TRAJ4SIT TIME 0

JCTlVlTY

la

)

I

p

12

EPTH

MUD

13

SG

1.5

14

500

i

f

...-

-

lo00

\

i-

i'

I IF I

I500

. . .

L

.. .

2000

I i

b

a

-.

IJ

rJ

2500

1

1

100

3

tb

zi,

36

Lo

DRILLING RATE

Fig. 3-27. Mudstone conductivity, transit time, drilling rate and specific gravity of drilling fluid of Shiunji SK-121. (From Magara, 1968.)

79 at equivalent depth, although the pore-water salinity in undercompacted shale is usually less (water resistivity is higher). Fig. 3-24 shows several normal compaction trends of the resistivity plots determined by Hottman and Johnson. The relationship between the resistivity ratio (Rn/Rob) and the pressure/depth is shown in Fig. 3-25. The reason for using the resistivity ratio instead of the resistivity difference is not usually explained clearly, but most empirical resistivity methods use the ratio. Fig. 3-26 is a resistivity plot for a well in the Gulf Coast. Subnormal resistivities are recorded below about 10,500 ft. The right-hand diagram shows the mud weight used for drilling this well, some drill-stem-test pressures, and the pressures estimated from the resistivity plot. These values are shown in pressure/depth (psi/ft). The method of estimating pressure and/or mud weight from the resistivity plot is essentially the same as for the sonic plot. Examples of well-log plots Fig. 3-27 shows a plot of the conductivity, sonic log, drilling rate and drilling-mud weight of Shiunji SK-21 in Nagaoka Plain, Japan. In the interval between about 2250 and 2500 m, where the maximum conductivity anomaly is recorded, the drilling rate increased. This means that these undercompacted shales are relatively soft so that drilling was faster. Another example of resistivity and transit-time plots is shown in Fig. 2-7 (Chapter 2). Empirically estimated pressure/depth values are indicated on the right-hand side of this figure. Fig. 3-28 shows an example of a bulkdensity plot from a formation-

5-

,

6-

r c

0 0

-

2

7-

I c :

.

2

8-

9-

1 0 -

I

1

I

1

1.0 2.0 3.0 2.0 2.2 2.4 2.5 Linear Logarithmic Shale bulk density,g/cm3

Fig. 3-28. Logderived shale density plots on arithmetic and semilog paper for the same well. (From Fertl, 1976 p. 205.)

80 Density (g/cc)

230 240 250 260

Conductivity (millirnhos)

1

2 3 4

11 000'

11 000'

11500'

11 500'

12000'

12000' 12 500'

13 000' 13500' 14 000' 14 500'

15000' 15500' 230 240 250260

1

2 3 4

Fig. 3-29. Shale bulkdensity variations in normal and over-pressured zones plus welllog conductivity curve and mud-weight requirements. (From Boatman, 1967.)

P O R E - W A T E R S A L I N I T Y (1000 p p m l F R O M S P LOG

SHALE POROSITY

Fig. 3-30. Shale porosity-depth and sandstone water salinity-depth plots of a Gulf Coast well. (Adapted from Overton and Timko, 1969).

81

Fig. 3-31. Average temperature and pore-pressure profiles. (From Lewis and Rme, 1970.)

density log, and Fig. 3-29is a plot of cuttings density. This example is from the Gulf Coast. As mentioned earlier, formation-water salinity in undercompacted sections is known to be subnormal, as shown in Fig. 3-30. The SP log is used for estimating water salinity in sandstones. The geothermal gradient in an area of rapid deposition is usually lower than where sedimentation has proceeded slowly. Because undercompacted zones, which commonly occur in an area of rapid deposition, are good insulators, the rate of heat flow becomes less. Therefore, the geothermal gradient in a rapid-sedimentation area is usually lower. The rapid deposition also tends t o move cool sediment rapidly, so that the temperature at a given depth and hence the geothermal gradient, may become lower. These conclusions are based on results from two fairly adjacent areas of rapid and slow deposition. However, if we compare the normally compacted and undercompacted (overpressured) intervals in a given area, the situation is different. In this case, heat flow through the two intervals, is considered to be uniform. Because the thermal conductivity of undercompacted sediments is usually less, their geothermal gradient must be higher to keep a constant heat flow vertically. This situation can be explained by the use of the following basic heatflow equation: dT H=CT-

dz

(3-18)

82 TABLE 3-11 List of specific physical features in abnormal-pressure zones SHALE

POROSITY

I S ABNORMALLY HIGH

TRANSIT TIME

HIGH

(VELOCITY)

(LOW)

CONDUCTIVITY

HIGH

(RESISTIVITY)

(LOW) LOW

BULK DENSITY DRILLING RATE WATER SALINITY TEMPERATURE GRADIENT

FAST LOW HIGH __

where

H =heatflow, CT = thermal conductivity, and dT - = geothermal gradient. dZ

Fig. 3-31 shows an actual example of subsurface temperature and pressure (in terms of mud weight) profiles in the Gulf Coast area. Several logging and other parameters to be encountered in abnormally pressured or undercompacted zones are summarized in Table 3-11. Appendix 3-1: see loose leaves at the end of the book Appendix 3-11 The rock composition and porosity are calculated from gamma-ray and sonic logs. The gamma-ray log measures the natural radioactivity of sediments. In clastic sequences it usually reflects the clay content, because clays contain high concentrations of radioactive elements. Pure sands or sandstones, and carbonate rocks, usually have very low levels of radioactivity. As gamma rays pass through sediment or rock, they are absorbed and their energy levels decline. The amount of absorption depends on the formation bulkdensity. Two formations having the same amount of radioactive material per unit volume but different bulk densities will show, on the gamma-ray log, different radioactivity levels; the less dense formations will appear more radioactive (Schlumberger, 1972). If a formation is composed of two materials, clay and nonclay, each with its own radioactivity level, the following relationship exists:

83 where

GR = gamma-ray reading (API unit), Pb = density of the formation (glcc), VClay = volume fraction of clay, Vnon~ay= volume fraction of nonclay, and A and B = coefficients. GRpb is called the “normalized gamma ray” and is directly related to the volume fraction of clay or nonclay. The coefficients A and B are the respective values of GRpb in the zones of 100% clay and 100%nonclay. The normalized gamma-ray technique described above has been applied experimentally to calculations of clay and nonclay contents in many Beaufort wells. Cuttings and core samples of the same wells have been analyzed by X-ray diffraction. Analysis by X-ray shows the volume fractions (or per cent) of different minerals such as montmorillonite, illite, chlorite, quartz, feldspar, calcite, etc., in rock samples. All clay minerals and all nonclay minerals were grouped separately, and the volume fraction (or per cent) of each group calculated. The result was then compared with the log-calculated values mentioned previously. This trial was not altogether successful. While the results for deep or compacted intervals showed good agreement, for relatively shallow or less compacted intervals the logcalculated clay content was almost always less than that derived from X-ray analysis. This discrepancy probably comes from ignoring the effect of compaction on gammaray values. Previously, the two formations were assumed to have the same amount of radioactive material per unit volume but different bulk densities. However, this assumption is probably not valid when the increase in bulk density is due mainly to compaction. In other words, the amount of radioactivity per unit volume of rock should increase with compaction, or the increase of bulk density. The effect of compaction on the gamma-ray log must, therefore, be evaluated and corrected. It is interesting to note that we commonly observe a trend of shale base line on gamma ray log which tends to increase in API with depth or compaction. This is the result of shale compaction. Suppose that a rock has a porosity of @ and a given mixture of clays and nonclays. The normalized gamma ray, as shown in the left-hand side of A3-1 is the total radioactivity of the rock material and the pore water, in this case: GRPb = (AVclay + Bvnonclay)(l - $1 + w@

(A3-2)

where W = radioactivity of the pore water. The radioactivity of pore water is usually very low compared with that of clays. If the pore-water radioactivity is assumed to be zero, eq. A3-2 can be reduced to:

or : (A3-3) This equation means that the value GRpb/(1- @)can be directly related to the volume fraction of clays or nonclays at any compaction level. Because the porosity is given in the

84 following general form as:

(A3-4) where

pm = matrix density of the rock (glcc), and p f = density of the pore water (glcc). Eq. A3-3 can be rewritten as:

Pb - 1

(A3-5) '

"--J

when p f is assumed to be 1 glcc. Eq. A3-5 can be further simplified as follows:

GRpb = A'VCIay+ B'Vnonclay

(A3-6)

Pb-1

where A' and B' are new coefficients respectively equal to A/(pm - 1 ) and B/(pm - 1) (Pm is constant). Eq. A3-6 is simpler to use than eq. A3-3, because its left-hand side contains only the two terms, GR and pb, instead of GR, pb and 9 as in eq. A3-3. The value GR&,/(pb 1 ) or GRpbl(1 - 9 ) in these equations may be called the compactioncorrected normalized gamma ray, which is directly related to the proportions of clay and nonclay contents in sediment or rock. To calculate the volume fraction of clay or nonclay from well-log data, the coefficients A' and B' must be determined. For this purpose, about 160 samples of cores and cuttings from wells drilled in the Beaufort Basin have been analysed by X-ray diffraction t o obtain actual volume per cent of clay and nonclay minerals. These total volumes from X-ray analysis were introduced into Vclay and Vnonclay in eq. A3-6, and the value GRpb/(pb 1) was also calculated !or the depth at which the sample had been collected, t o determine the coefficients A and B'. In the Beaufort area these respective values were found to be about 200 and 5 0 API units. For application of this technique, a combination of gamma-ray and sonic logs is more practical than the gamma-ray-density log combination, because the sonic log is usually more available. The following empirical relationship between sonic transit time, A t , and bulk density, &, of shales in the Beaufort may be used to obtain pb from At:

~b = 2.99 - 0.00616 At

(A3-7)

When goodquality density and sonic logs are both available, it is of course recommended to use the former, rather than calculate the bulk density from the transit time. The sandstone porosity is calculated from sonic log and corrected for the amount of clays which was estimated previously from the gamma ray and sonic combination.

References Archie, G.E., 1942. The electrical resistivity log as an aid in determining some reservoir characteristics. Trans. A m . Znst. Min. Metall. Eng., 146: 5 5 - 6 2 .

85 Barker, C., 1972. Aquathermal pressuring -role of temperature in development of abnormal-pressure zones. Bull. A m . Assoc. Pet. Geol., 56: 2068-2071. Boatman, W.A., 1967.Shale density key to safer, faster drilling. World Oil, 165:69-74. Dickinson, G., 1953. Geological aspects of abnormal reservoir pressures in Gulf Coast Louisiana. Bull. A m . Assoc. Pet. Geol., 37: 410-432. Eaton, B.A., 1972. The effect of overburden stress on geopressure prediction from well logs. J. Pet. Technol., 24: 929-934. Fertl, W.H., 1976.Abnormal Formation Pressures. Elsevier, Amsterdam, 382 pp. Hottman, C.E. and Johnson, R.K., 1965. Estimation of formation pressures from logderived shale properties. J. Pet. Technol.. 17: 717-722. Hubbert, M.K. and Rubey, W.W., 1959. Role of fluid pressure in mechanics of overthrust faulting, I. Geol. SOC.A m . Bull., 70: 115-166. Lewis, C.R. and Rose, S.C., 1970.A theory relating high temperatures and overpressures. J. Pet. Technol., 22: 11-16. Magara, K., 1968. Compaction and migration of fluids in Miocene mudstone, Nagaoka Plain, Japan. Bull. Am. Assoc. Pet. Geol., 52: 2466-2501. Magara, K., 1975. Importance of aquathermal pressuring effect in Gulf Coast. Bull. A m . Assoc. Pet. Geol., 59: 2037-2045. Overton, H.L. and Timko, D.J., 1969. The salinity principle - a tectonic stress indicator in marine sands. Log Anal., 10: 34-43. Schlumberger, 1972. Log Interpretation, 1. Principles. Schlumberger, New York, N.Y., 113 pp. Schmidt. G.W.. 1973. Interstitial water composition and geochemistry of deep Gulf Coast shales'and sandstones. Bull. A m . Assoc. Pet. Geol., 57: 321-337. Terzaghi, K. and Peck, R.B., 1948. Soil Mechanics In Engineering Practice. Wiley New York, N.Y., 566 pp.

This Page Intentionally Left Blank

Chapter 4 CAUSES OF ABNORMAL SUBSURFACE PRESSURE

A number of causes have been proposed for the generation of abnormal pressures: (1)Compaction (2) Aquathermal effect (3) Montmorillonite dehydration (4) Artesian condition (5)Hydrocarbon accumulation (6)Osmosis (7)Fossil pressure (8)Cementation (9)Tectonics (10) Generation of hydrocarbons (especially gas) (11)Charging Compaction Terzaghi’s model and its application to subsurface conditions as discussed in the previous chapter are very simple and effective explanations for the generation of abnormal pressures. However, real subsurface conditions are more complicated than those simulated by the models. In the case of these simplified models, fluid expulsion from a given block of rock is considered. In the subsurface, however, fluids expelled from the other blocks below and beside a particular block will influence the fluid-flow conditions within that block. In other words, the cumulative effect of fluid migration will be threedimensional. There are several more complicated mathematical models proposed by different researchers (Bredehoeft and Hanshaw, 1968;Smith, 1971;Kojima et al., 1977). Although estimating such a threedimensional flow condition is very complicated, a twodimensional model in the vertical direction is not difficult to make. I feel this type of twodimensional model simulates the subsurface conditions in many sedimentary basins reasonably well. Suppose a clay or shale sequence in which the clay or shale has reached a compaction equilibrium and within which the fluid pressure is hydrostatic (stage A of Fig. 4-1).Additional sediments, whose thickness is lo, are added above this sequence under water. If the entire shale reaches a new equilibrium condition of compaction within the time interval t, a porosity distribution such as shown by stage B in Fig. 4-1would be established. An exponen-

88

1,NEW LAYER

OLD SEDIMENTS

Fig. 4-1. Schematic diagram showing stepwise compaction of shales and fluid expulsion due to loading of layer 20.

tial function between shale porosity and depth is established at stages A and B (compactionequilibrium conditions). Suppose that the outlet for fluid expulsion exists only at the surface, and the fluid is expelled vertically upward. Shale compaction from stage A to stage B in this case occurs from the shallower part to the deeper part of the sequence in a stepwise manner (Fig. 4-1). The increase of fluid pressure ( p l 0 ) due to the instantaneous loading of the new sediment lo is given as:

where SI, = overburden-pressureincrease, Pbo = water-saturated bulk density of the new sediment layer l o , and g = acceleration due to gravity. When the fluid-pressure increase is resolved into two components (Hubbert and Rubey, 1959,eqs. 133 and 134):

89

or : (4-3)

where = normal

or hydrostatic-pressure increase, and

= excess-pressure increase.

The value (Pzo),in this case is shown as:

where pw = density of formation water. Introducing eqs. 4-1 and 4-4 into eq. 4-3, we obtain:

or :

where dpaldZ = excess-pressure gradient. This excess-pressure gradient generated at near-surface is considered to be carried downward in this stepwise model as compaction progresses. The vertical fluid movement in this case may be shown by Darcy’s equation:

where q = volume of fluid crossing unit area normal to the flow direction in unit

time,

k

= permeability of shale, and I.( = viscosity of the fluid.

In this case we assume one-phase (liquid) and one-component (water) fluid movement in the shales. The waterdensity term is not included in this equation because the effect of changing water density on the total fluid flow is relatively very small. The volume of fluid, q l , passing depth Z1, in unit

90 time in a shale column is expressed from eqs. 4-6and 4-7as:

where

Z1=depth, q1 = volume of fluid passing through the shale at depth Z1, and kl = permeability of the shale at depth Z1.

The volume of fluid, Q1,passing in time interval t is given as:

Assuming that shale compaction occurs simply by the expulsion of fluid from shales, the porosity difference in Fig. 4-1would indicate the amount of fluid that must be expelled between stages A and B for the new compaction equilibrium to be reached. Supposing the direction of fluid expulsion in this case is upward, the amount of fluids that should pass through the shale at depth Z1can be calculated. An exponential relationship between shale porosity and depth at the equilibrium condition of compaction, proposed by Rubey and Hubbert (1959)and based on Athy's (1930)porosity-depth curve, is as follows:

(4-10)

@ = @o e-CZ

(see eq. 2-1).The line of stage A in Fig. 4-1can be shown mathematically by eq. 4-10.Suppose that sediment deposition and subsidence of lo occurs and the shales reach a new compaction equilibrium (stage B in Fig. 4-1)within the time interval t. The porosity-depth relationship in this case is shown as: @+@I

= @ o -c(Z+lo) e

(4-11)

where = porosity difference between stages A and B at depth 2. From eqs. 4-10and 4-11,we obtain: = @o e-cz(e-c'o

- 1)

(4-12)

The total amount of porosity decrease in a shale column with the unit base area in this case is given by the integral of eq. 4-12as follows: J@2dZ= @o(e'-c20 - l)Je'zdZ

(4-13)

91 This amount of porosity decrease is here considered to be the volume of fluids that should be expelled from the shale column for the new equilibrium to be reached. As this volume of fluids is considered to go upward in this case, the volume of fluids that should pass through depth Z1for a new equilibrium condition to be reached is expressed as:

(4-1 4) In this case 2 in eq. 4-14is now considered to be the bottom depth of the shale sequence, and an impermeable bed exists below 2. The effect of the fluid (water) volume change due to temperature change is not taken into account. When this volume is equal to or less than Q1 in eq. 4-9,enough fluids would have been expelled from the shales and a hydrostatic-pressure environment would be established. When the volume is greater than Q1, on the other hand, some fluids would remain and an abnormal-pressure condition would occur. When the volume is balanced with Q1 in eq. 4-9,the following relationship would exist:

):(-

Q1=

z

(pbo-pw)gt = r)o(e-czO- 1)

1e-czdZ

(4-15)

Z1

kl in this equation is considered to be the minimum permeability for the new compactionequilibrium condition reached, and is given as:

(4-16) If the average subsidence or sedimentation rate in the time interval t is given as A l , the following relationship exists: 1, = tAl

(4-17)

When we take a certain time interval (for example, 1 s) for t in eqs. 4-16 and 4-17,we can calculate kl at several different sedimentation rates. In order to discuss the fluid-migration conditions during continuous sedimentation in the Gulf Coast area, Dickinson's shale porositydepth curve is used (Fig. 4-2).Fig. 4-3 shows the same curve on semilog paper. Athy's (1930)porositydepth relation is plotted in these figures as well. Dickinson also reported data concerning reservoir pressure versus depth for the Gulf Coast area (Fig. 4-4).According to this figure, abnormal pressures occur below about 7000 ft. Above this depth, in other words, the fluid

92

Fig. 4-2. Comparison of shale porosity-depth relationships of Oklahoma (Athy, 1930), and the Gulf Coast (Dickinson, 1953).

Shale porosity

Fig. 4-3. Shale porosity-depth relationships of Oklahoma (Athy, 1930) and the Gulf Coast (Dickinson, 1953) on semilog paper.

93

Fig. 4-4. Reservoir pressure versus depth of the Gulf Coast. Solid circles = measured pressures;open circles = estimated pressures. (From Dickinson, 1953.)

pressure is hydrostatic and the shales are considered to have reached the equilibrium condition of compaction. An exponential relationship between shale porosity and depth at the compactionequilibrium condition is determined by drawing an average straight line above 7000 f t (actually the interval from 2000 to 7000 ft) on Dickinson’s curve (Fig. 4-3).The function at the equilibrium condition of compaction (normal trend) is as follows:

9 = 90e--cZ - 0 39 e-0.0000952(ft)

(4-18)

where Z > 2000 ft, or:

4 = 0.39 e-0.0000031Z(cm) where

Z > 610 m.

(4-19)

Dickinson’s curve above 2000 f t does not fit the function shown as eqs. 4-18or 4-19.Because the possibility of abnormal pressures above 2000 f t seems to be very small, we will not consider them at these shallowest depths.

94

2 = 1 5 . h ft

EX

Ir g n

15,000

10.000

I

I

5,0000

I

2,000 I

I

I

Depth of normal trend I

I

5000 ft, 10,000 ft, 1971.)

...,

-

in the Gulf Coast; A1 is the sedimentation rate. (From Magka,

By the use of eqs. 4-16and 4-18or 4-19,the minimum permeability kl for the new compaction equilibrium reached can be calculated. In the Gulf Coast region, some 6 mi (10km) of sediments have been deposited in about 60 million years, or at a mean rate of about 1.7 cm/year (Rubey and Hubbert, 1959,p. 181),corresponding to A1 = 5.4 cm/s.

-

The viscosity of the formation water would change mainly with temperature. Assuming that the average geothermal gradient in this area is 1.4"F per 100 f t of depth, and the average surface temperature is 74°F (see Nichols, 1947),the viscosity of the water is determined (Pirson, 1963).The values of the minimum permeability kl for the compaction equilibrium in this case are shown as curved solid lines in Fig. 4-5(pb0 = 1.4 g/cc, p w = 1.08g/cc). Fig. 4-5 shows that, when the rate of sedimentation is constant (5.4* 10-l' cm/s), the minimum permeability kl necessary for the compactionequilibrium condition to be maintained increases with increase of the total thickness of sediment, 2, If the actual permeability in the subsurface is greater than or equal to kl,enough fluids can be expelled upward and the hydrostatic-pressure condition will be established. If the actual permeability

95

I Y.

Z = 5,OOOft

z:

10.000

/J

$1 'J

/

I-

.

-

+/---/---I-

i

i

000'

.

Fig. 4-6. Relationship between minimum permeability kl for compaction equilibrium and depth (or shale porosity), at several sedimentation rates; A1 is sedimentation rate. (From Magara, 1971.)

is less than kl, some fluids would have to remain in the shales and abnormal pressure would occur. The possibility of abnormal pressure when the sedimentation rate is constant would, therefore, increase with increase of the total thickness of sediment 2, as shown in Fig. 4-5. In other words, the possibility of abnormal pressure at any depth would increase at the later stages of continuous and constant sedimentation (2is large). As the rate of sedimentation increases, more fluids must be expelled in unit time, or the minimum permeability kl for the equilibrium must increase. Fig. 4-6 shows the values of k l at several values of the sedimentation rate. It is concluded that the possibility of abnormal pressures would increase with increases in the rate of sedimentation and the total thickness of sediment. If the actual permeability is greater than or equal t o kl, abnormal pressures never occur at any depth. As long as this kind of normal compaction and normal fluid-expulsion situation is maintained, the minimum permeability kl in Figs. 4-5 and 4-6 is necessary for further maintenance of the compaction

96

equilibrium. Once abnormal pressure occurs at a certain depth (some fluids remain in undercompacted shales), however, the volume of fluids being expelled upward would decrease. According to Figs. 4-3 and 4-4, abnormal pressures exist below 7000 f t in the Gulf Coast area, and the shale porosity at those depths is higher than the normal trend would indicate (undercompacted shale zone). The volume of fluids moving upward in this area would, therefore, have been smaller than that assumed in the previous discussion

i I TUAL VINLRLA

c I

CAI11 QUlLl !SON'!

V b ' i FOR IUM :URVL

Fig. 4-7. Comparison of estimated actual permeability and minimum permeability ki, when undercompacted shales exist at depth in the Gulf Coast. k = kaolinite; b = bentonite; m = montmorillonite; other solid circles = clay or shale.

97

(Figs. 4-5, 4-6). The value of k; when undercompacted shales exist at depth should be calculated from the actual porosity-depth relationship rather than the extrapolated normal trend. The value of 12; is calculated on the basis of Dickinson's curve and shown as a curved solid line in Fig. 4-7. As mentioned previously, the actual permeability in the subsurface would be greater than or equal to k; in the hydrostatic-pressure zone (shallower than 7000 ft), and less than k; in the abnormal-pressure zone (deeper than 7000 ft). In addition, it would be reasonable to assume that, at 7000 ft or the boundary surface between the hydrostatic- and abnormal-pressure zones, the actual permeability equals k; . This permeability value is indicated as p in Fig. 4-7. Data on the permeability-porosity relationship of shales are still relatively scarce. Bredehoeft and Hanshaw (1968) have compiled a certain amount. The Geological Survey of Japan measured the permeability and porosity of the mudstones of several stratigraphic test wells drilled in Tertiary formations in Japan. Those data are shown as solid circles in Fig. 4-7. Although obtained from different formations in different areas, most of the permeability-porosity data in Fig. 4-7 are plotted in the relatively narrow shadowed zone. On the basis of the previous assumption that k; at 7000 f t equals the actual permeability, a broken line parallel to this zone is drawn through point p in Fig. 4-7. The values indicated by this broken line are, therefore, considered to be the estimated actual permeability in the Gulf Coast area. Because the estimated actual permeability is greater than k; above 7000 ft and less than k; below, the previous assumptions are considered to be verified. In other words, the actual permeability is greater than k; above 7000 ft and less than k; below. This low permeability would have caused abnormal pressures in the deep subsurface there. Fig. 4-8, constructed from Figs. 4-6 and 4-7, shows the relationship between the rate of sedimentation and the possible top of the overpressured section in the Gulf Coast. The total thickness is assumed to be 33,000 ft. It is interesting t o note that in Fig. 4-8 the depth range of the top of the overpressured interval is from about 3000 f t to about 20,000 ft. This range corresponds to the range of sedimentation rate from 5.4 10-l' cm/s ($ of the average for Gulf Coast sediments) to 5.4 . lo-' cm/s (10 times the Gulf Coast average). The fact that this calculated depth range covers the actual depth range of the top of overpressuring in most parts of the Gulf Coast suggests that the assumptions used for this calculation are not too unreasonable, and that the rate of sedimentation there was probably within the range from 5.4 * cm/s t o 5.4 lo-' cm/s. In the previous discussions, the sequence was assumed to be composed of shales only. However, most sedimentary basins contain other rock types that are more permeable. If any permeable rocks were interbedded in the

-

-

98

O

7

33,000 FT (IOKM)

7000 FT

I

5.4~10-~

5.4X10’10 SEDIMENTATION RATE (Cm/sec)

5.4~10 -”

Fig. 4-8. Diagram showing estimated top of abnormalpressure zone with changing rate of sedimentation.

shale section, no significant overpressuring would occur. As a conclusion, overpressuring is considered to be the result of compaction phenomena caused by the combined effects of three factors: (1)Lack of a permeable bed (e.g., sandstone). (2) Rapid rate of sedimentation. (3) Thick accumulation of sediments. The actual rate of pressuring, however, may also be affected by other factors. Aquathermal effect During continuous burial, temperature commonly increases, which tends

to increase the volume of a given weight of water. If the water is in an open system, it can expand freely and some water will migrate. If the system is closed, as might be expected in a massive shale section, the water cannot expand to the same degree, and the pressure in this section will therefore increase (aquathermal effect). The rate of pressure increase under the aquathermal condition is more than that under the nonaquathermal. In the case of nonaquathermal pressuring, the fluid pressure will increase at the rate of overburden-pressure increase (“1.0 psi/ft) if pore fluid is sealed in the rock. The rate of aquathermal pressuring depends on the geothermal gradient and the shale’s cap-

99 EQUIVALENT TRANSIT TIME (pS / FT)

5 e

1

160

150

140

130

I

I

I

I

5 z

110

100

I

1

I

012

v) W

9

120

6.

1.6-

*' 0 18

W

10.

K

3

3

1.4-

a13 a15

09

a 17

Fig. 4-9. Relationship between gradient of pressure increase (p- p , )/(Z -2, ) and isolation depth 2, for 18 sets of well data derived from Hottman and Johnson (1965). Transittime value corresponding to isolation depth is shown at top. (FromMagara, 1975b.)

function of the composition and level of compaction of undercompacted shales. As mentioned in the previous chapter, the rate of fluid-pressure increase with a geothermal gradient of 25"C/km (or 1.37"F/100 ft) within a complete closed system is about 1.8 psi/ft. However, shales are not complete seals, so that the average net gradient of pressure increase in the Gulf Coast area, after semiperfect isolation, is about 1.4 psi/ft. This figure is the midvalue between 1.8 psi/ft (perfect aquathermal pressuring) and 1.0 psi/ft (nonaquathermal pressuring). In other words, about 50% of the aquathermally generated pressures have been dissipated from the average Gulf Coast shales. The degree of pressure dissipation may depend on the permeability of the undercompacted shales. If so, the compaction levels of such shales may be important, because permeability would decrease with compaction for a given shale composition. Fig. 4-9 is a plot of the gradient of the pressure increase since isolation versus isolation depth for the eighteen sets of well data from the Hottman and Johnson paper (1965; see also Chapter 3). Generally, the gradient

100 increases with increase in the isolation depth; an exception is seen at well 9, (Fig. 4-9), where the gradient is not very high for the relatively deep isolation depth. This general trend would mean that, the deeper the isolation depth, the more pore pressure is retained. All the shales discussed here are undercompacted at present. Among undercompacted shales, any that are relatively more compacted tend t o retain more abnormal pressures generated mainly by aquathermal effect. The scatter of the points in Fig. 4-9 suggests that factors other than the shale compaction level are also effective; the variable composition of shale may be significant in this respect. In conclusion, the significance of aquathermal effect on pressuring is that, if the three factors discussed above (lack of permeable bed, rapid sedimentation rate and thick accumulation of sediments) created the conditions to cause overpressures, the rate of pore-pressure increase would be higher. Therefore, the presence and relative strength of aquathermal effects would result in a significant difference in the final fluid pressures. This pressuring requires neither any special physical change of water, such as from solid to liquid phase, or any special minerals, such as montmorillonite. In other words, it can happen during burial if water exists in pore spaces that are reasonably well sealed. Montmorillonite dehydration In 1967 Powers, using Gulf Coast data, showed that alteration of montmorillonite to illite begins at a depth of about 6000 f t and continues at an increasing rate to a depth, usually about 9000-12,000 ft, where there is no montmorillonite left. This alteration offers a mechanism for desorbing the last few layers of bound water in clay and transferring it to interparticle locations as free water. As the last few layers of bound water have a greater density than free water, this released water increases its volume as it is desorbed from between unit layers. If the water expands and cannot escape, it will increase the pore-water pressure to abnormally high levels. Powers further stated that, “Abnormally high fluid pressures may easily be caused by a volume increase associated with the desorption of the last few monomolecular layers of water from montmorillonite during its diagenesis to illite.” This mechanism is called “montmorillonite dehydration”. It is known in many parts of the world that abnormal fluid pressures are usually associated with undercompacted shales. This association can be explained very well by fluid expulsion and shale compaction phenomena. However, montmorillonite dehydration may also provide an explanation. In this chapter, therefore, I wish to compare these mechanisms by using schematic shale porosity-depth plots. The approach is to try to explain the known association of abnormal pressure and undercompaction by using each hypothesis individually, as though no other cause of abnormal pressure existed.

101 Compaction disequilibrium

Fig. 4-10A is a schematic diagram of a shale porosity-depth relationship (broken curved line) in a typical abnormal-pressure area. The porosity scale is logarithmic and the depth scale arithmetic. At relatively shallow depths, shale porosity decreases at a constant rate. In the relatively deep subsurface, the shale porosity is abnormally large (undercompacted zone) and, as mentioned previously, the pore pressure is abnormally high. This departure from a normal compaction trend is what is expressed as compaction disequilibrium. Let us consider the porosity-reduction history; Lee,the compaction history of the two shales A and B in Fig. 4-10A. Shale A is normally compacted and shale B is undercompacted. The best guess on the compaction history of shale A is that it had, when deposited, the original porosity expressed by Go in this figure, and experienced the history shown by the normal compaction trend, indicated by a solid straight line with an arrow. The history for the undercompacted shale B may be postulated as follows: This shale had an original porosity of Go when deposited, and experienced normal compaction at relatively early stages (shallow burial depth). At a certain compaction level thereafter, shale permeability was reduced t o a critical level, below which normal water expulsion could not occur. Alternatively, one may postulate that a restricted drainage system caused by faulting or a lack of permeable beds did not permit the water to escape after a certain stage. Whatever the real cause of restricted water expulsion, the porosity-reduction rate (depending on sedimentation rates of overlying beds) after this stage must have been much less than before. The possible com-

B

A SHALE POROSITY

4

SHALE BULK D€NSITY -c

Qa

COMPACTED

DEPTH

.1

PA I

Fig. 4-10. Schematic diagram showing changes of porosity and density during burial of normally compacted shale (A) and undercompacted shale ( B ) , when compaction disequilibrium is the cause of undercompaction.

102 paction history of shale B, therefore, can be indicated by the solid curved line with an arrow in Fig. 4-10A. (A similar assumption of compaction history also is made by Chapman, 1972.) An important corollary of the previous discussion is that the undercompacted shale B has never attained the present degree of compaction of shale A. Since the time of original deposition, shale B has released a much smaller volume of water than shale A. Because shale B has abnormally high porosity, or its framework strength is abnormally low, the pore water it contains must support a greater proportion of the total overburden. Consequently, the fluid pressure in B should be abnormal. Temperature increases after the isolation of pore water also might increase the pore pressure in this case (Barker, 1972; Magara, 1974b). Fig. 4-10B depicts the shale bulkdensity-depth relationship corresponding to the porositydepth relationship in Fig. 4-10A. As bulk density increases with decrease in porosity, schematically we can expect Fig. 4-10B t o be a mirror image of Fig. 4-10A. The corresponding compaction histories for shales A and B are shown similarly by the solid lines. In this case shale B has never reached the present density level of A. (Shale compaction is an irreversible process; rebounding is rejected as an explanation of this reversal in shale porosity and bulk density with depth.)

Montmorillonite dehydration Now let us assume that montmorillonite dehydration is the only cause of abnormal pressure and undercompaction, and again trace a postulated compaction history. In the schematic porositydepth diagram shown in Fig. 4 - l l A , shale A has experienced the same compaction history as in the disequilibrium case (Fig. 4-10A), because it has not reached the montmorillonitedehydration level. The undercompacted shale B experienced a normal compaction history until it reached the level at which the bound water of montmorillonite could be released to become free (dehydration level). Then dehydration caused the shale t o rebound t o the present stage. The compaction history in this case is shown schematically by the solid line with an arrow in Fig. 4-llA. If montmorillonite dehydration is the only cause of abnormal pressure and undercompaction, all shales below the dehydration level once must have been compacted to the minimum-porosity levels (fluid expulsion from the shale up to this stage must have been normal and therefore relatively large), and then rebounded to the present porosity level as a result of the expansion associated with montmorillonite dehydration. The respective bulkdensity changes for shales A and B in this case are shown by solid lines in Fig. 4-11B. Shale B must once have reached the maximumdensity point just above the dehydration level and then rebounded to the present density of dehydration. During this expansion process, there

103

A

B

SHALE POROSITY ----c

SHALE BULK DENSITY

-

z -\ DEPTH

I

ZONE

LEVEL

Fig. 4-11.Schematic diagram showing changes of porosity and density during burial of normally compacted shale (A) and undercompacted shale ( B ) , when montmorillonite dehydration is the cause of undercompaction.

must have been a significant volume increase of the dehydrated formations. In any case, if montmorillonite dehydration were the main cause of the abnormal pressure and the abnormally high porosity and low density, the abnormally pressured shale must have experienced a compaction history different from that in the case of compaction disequilibrium. The proposed montmorillonitedehydration process involves the following two physical changes in the water: (1) water phase change (bound water-free water) and (2) water expansion. (1)Porosity of shales will increase in response t o water phase change; i.e., when the water is bound in montmorillonite clay it is a part of the clay matrix, but when it becomes free its porosity increases. Bulk density, however, will not be altered by this phase change, provided all released water stays in the pore space. If the released water is flushed out of the shales, as suggested by Burst (1969), bulk density actually should increase after dehydration. It is apparent, therefore, that water phase change cannot explain the reduced bulk density of shales in zones of abnormal pressure - a characteristic observation in many areas in the world. (2) To expand after clay dehydration, bound water must have a higher density than “free” water. Burst (1969) showed a detailed calculation of interlayer waterdensity values in his table 1,in which the density of the second water layer is calculated to be as high as 1.15 g/cc *. In his second dehy-

* Powers (1967)suggested the average density of

the first four water layers to be about

1.4 g/cc, ,which is derived from Martin (1962). Cebell and Chilingarian (1972)cited a value less than that for ordinary water (or less than 1 g/cc) from Anderson and LOW (1958). Because of such a large discrepancy in the interlayer waterdensity values obtained by different investigators, I will use first the value by Burst in this chapter.

104

Dapth, 1000 I

Fig. 4-12. Shale density-depth and salinity-depth relationships in a Gulf Coast well.

(From Overton and Timko, 1969.)

dration stage, this water layer can be released and its amount is, according t o Burst, approximately 10-15% of the bulk volume of the sediment at that time. If the highest values (1.15 g/cc and 15%) are used, we can expect about 2% expansion of bulk-shale volume from water expansion associated with dehydration. Fig. 4-12 is an actual plot of (2.55 minus shale density g/cc) versus depth in the Gulf Coast area, taken from Overton and Timko (1969).The shale density in this case is derived from the formationdensity log. Because this plot shows the (2.55 minus shale density g/cc) value instead of shale density, the original shale density was recalculated and the scale for it is shown on the right-hand side of the figure. The shale density below about 10,000 f t is abnormally low. Pore-fluid pressure below about 10,000 f t is abnormally high (Fig. 4-12).If interlayer water dehydration were the only cause of the

105 reduced shale density below 10,000 ft, the shales in the deep section must have rebounded from the maximum density at about 9000 f t to their present lower values. The shale density at 9000 f t is about 2.38 g/cc, and the minimum density in the undercompacted zone in this case is about 2.25 g/cc. These values suggest at least a 6% bulk-volume increase at dehydration. An expansion of 2%, based on Burst's waterdensity figure, is not sufficient here. In fact, if, as proposed by Burst, clay dehydration also increases shale permeability and flushes some water out of the shales, the actual expansion of shales after dehydration must be even less than the 2% of the previous calculation. It is, therefore, quite difficult t o believe that montmorillonite dehydration could have been the single cause of observed abnormal pressures and subnormal density in this case. Fig. 4-13shows an example of the cuttingsdensity-depth plot in the Gulf Coast (Boatman, 1967). For shale at the maximum-compaction level (12,500 ft, 2.46 g/cc) to rebound to a reduced density of 2.33 g/cc such as is found

Bulk density (g/cc)

Foet/hour

drilling time

l0,soo'

10,500'

llpoo'

11,OOO'

ll.500'

11.500'

12.000'

12,OOO'

12,500'

12,500'

13,000'

1 3 , ~ '

2.20 230 2.40

2.w

50

30

l0

Fig. 4-13. Shale density-depth and drilling time-depth relationships in the Gulf Coast. (From Boatman, 1967.)

106 Bulk density ( g/cc 1 2.10

2.20

2.30

2.40

2.50

10,000'

10,500'

ral area

1l.OO0

-

'

11,500'

12,000'

12,500'

2.10

2.20

2.30

2.40

2.50

Fig. 4-14. Shale density-depth relationship in a Gulf Coast well. (From Boatman, 1967.)

at about 13,000 ft, a 6% bulk-volume expansion would be required. Again, if some water is flushed out of the shales by dehydration, an expansion of more than 6% would be necessary to explain the actual subsurface conditions. The example in Fig. 4-14,from Boatman (1967), indicates that the shale density at the bottom of the compacted zone is about 2.45 g/cc (11,700 ft) and the minimum density in the abnormal-pressure zone about 2.12 g/cc (12,300ft). If clay dehydration were the cause of the low density in the abnormal-pressure zone, the shales must have rebounded at least 16%in volume - a fact quite difficult to believe. Fig. 4-15shows another example, by Rogers (1966), in which the maximum density of about 2.45 g/cc at about 11,500 f t would have to be reduced to 2.21 g/cc at about 12,500 ft. A t least an 11%shale-volume expansion would be required in this case.

107

m8000 Normal trend

9ooo-

10,000L +8

5‘

11,ooo12.000 -

Transition zone

13,000 14.W -

L

’6’W2.20

2.30

2.40

L

2.50

260

Density, g/cc

Fig. 4-15. Shale d e n s i t y d e p t h relationship in a Gulf Coast well. (From Rogers, 1966.)

In all these examples, therefore, it is impossible t o explain the observed subnormal density values by using Burst’s numbers for montmorillonite dehydration. As an alternative, let us now consider the use of Powers’ (1967)boundwater density figure of 1.4 g/cc, which is derived from Martin (1962).Powers stated that “. . .the volume increase in mudrocks of the Texas Gulf Coast would range from a probable low of 2.5 percent for sandy mudrocks t o a high of probably 20 percent for relatively pure clay rocks.” He uses an interlayer water density of 1.4 g/cc and the equation, Vi = PC, where Vi is the increase in bulk volume of the sediments, P is the difference in density between the last four water layers and normal water (0.4g/cc in this case), and C the percent of montmorillonite that on dehydration collapses to 10 A (Powers, 1967, p. 1249). The value C should, however, be the volume per cent of the montmorillonite interlayer water t o be released at dehydration, rather than the percent of montmorillonite. Powers assumed that the volume of the last four water layers is equal t o the volume of dry clay in the mont-

108 morillonite (i.e., 50% of montrnorillonite clay is the interlayer water). In other words, the sediments contain 100% montmorillonite, and there is no free pore water in sediments before dehydration in this case. This is not true, because even relatively clean shales contain some amount of nonclay material; they also may contain clay minerals other than montrnorillonite, even before dehydration. Sediments also should contain some free pore water (probably at least 15-20%), which is not related to clay dehydration. Therefore, his 20% expansion figure does not seem possible in real subsurface sediments; it should be reduced at least to about one-third *, even if the large figure of 1.4 g/cc is used for interlayer water density **. If a waterdensity value of less than 1.4 g/cc is used, the expansion figure of course will be reduced further. In addition to the previous reasoning, the rebounding of a large sediment mass presents a difficult problem in geoIogic understanding. For exampIe, if the 6% expansion figure is used, the thickness of a 30,000-ft section should increase by 1800 f t on dehydration, and that of a 20,000-ft section by 1200 ft. If the 20% expansion figure suggested by Powers is used, the thickness increase should be tripled. It may be concluded that dehydration of montmorillonite interlayer waters is not sufficient to explain the reduced bulk density in the abnormally pressured shale zones when Burst’s interlayer waterdensity figure (1.15g/cc) is used. With values greater than 1.15 g/cc, some subsurface conditions may be explained, but the waterdensity values must be verified carefully; those proposed by the various investigators differ quite widely. It, therefore, is concluded that montmorillonite dehydration probably is not adequate as a single cause of abnormal-pressure generation; it could, however, be a secondary cause. In other words, if abnormal pressures already have been generated by some other cause such as compaction disequilibrium, dehydration could increase them further.

* If the free pore water occupies 20% of the sediments, and if minerals other than montmorillonite (other clay minerals + nonclays) occupy 30%, the volume percent of montmorillonite is 50 (or 0.5). According to Martin’s (1962)figures of adsorbed (interlayer) water density used by Powers, for montmorillonite clays having a water content of more than 0.284 g HzO/g dry clay, the water density drops sharply to values less than 1.32 g/cc. Volume per cent of adsorbed water in the clay having 0.284 g HzO/g dry clay is about 30,assuming that the density of dry clay is about twice that of the adsorbed water. The 50% figure used by Powers seems t o be too high, if one wants t o use the adsorbed-water density figure 1.4 g/cc. Hence, the volume percent of the adsorbed water in bulk sediments in this case would be about 15% or 0.15 (0.5X 0.3). If the entire adsorbed water is. released by dehydration, we obtain Vie=0.4 X 0.15 = 0.06, of 6% expansion of the sediments. This would be the value for relatively clean shales, and therefore probably the possible maximum when the 1.4 g/cc figure is used. ** In his calculation, Powers assumed a waterdensity change from 1.4 g/cc in montmorillonite t o 1.0 g/cc in illite, but Martin (1962)showed a water density of 1.36 g/cc in illite (table 2), which is calculated from DeWitt and Arens (1950).Hence, the waterdensity difference of 0.4 g/cc caused by dehydration also would be too large.

109 A t this point the question arises as to why there is a decrease in the amount of montmorillonite in many abnormally pressured zones in the deep subsurface. According to Burst (1969),clay dehydration depends mainly on subsurface temperature, with the average dehydration temperature being 221"F. Inasmuch as subsurface temperature tends to increase with depth, it eventually must exceed dehydration temperature. The depth at which this critical subsurface temperature is reached will vary with differences in geothermal gradient, both within a sedimentary basin and between one basin and another. That is, there must be local and regional variations in the depth of dehydration. An important factor controlling the geothermal gradient is the amount of water in the rocks. Because the heat conductivity of water is much less than that of the rock matrix, when rocks contain large quantities of water (high porosity), the geothermal gradient becomes greater (Lewis and Rose, 1970; Reynolds, 1970).As already mentioned, shale porosity in abnormal-pressure zones is abnormally high; consequently the geothermal gradient also is abnormally high. Where there are abnormal pressures, then, a high subsurface temperature is reached at relatively shallow depths. We thus are led to the hypothesis that a combination of abnormal pressure, abnormal shale porosity, and subnormal bulk density initially may be created by compaction disequilibrium. These conditions will produce a high temperature gradient which, in turn, will increase the chances for montmorillonite in the abnormal-pressure zone to release its interlayer water and be transformed to illite. Certain chemical conditions for potassium fixation also are required for this conversion. Clay dehydration and consequent water expansion can further increase the pore-water pressure. If there is no initial compaction disequilibrium and, consequently, no abnormal pressure or porosity in the subsurface, the temperature gradient will be lower, and the threshold level for montmorillonite dehydration will not be reached until a much greater depth. Dehydration could produce some degree of abnormal pressure, but it would be less than that caused by compaction disequilibrium and montmorillonite dehydration combined. Furthermore, in this case the top of the dehydrated zone could coincide with the top of the abnormal-pressure zone. If the initial abnormal pressure resulting from compaction disequilibrium occurs at an extremely shallow depth such as 3000 f t (the sedimentation rate may have been very fast and/or initial subsurface drainage may have been very poor), montmorillonite-illite conversion may occur at very shallow depths, but not necessarily at the top of the initial abnormal-pressure zone (3000 ft). The hypothesis just described is based on the previous reasoning that montmorillonite dehydration is not adequate as a sole cause of abnormal pressure, whereas compaction disequilibrium is.

110

Fig. 4-16. Schematic diagram showing artesian conditions.

Artesian condition If the subsurface system is relatively open, fluid pressure equals the weight of the water column above a given point. Therefore, if the surface in an area is relatively flat, the pressure caused by this effect is reIatively uniform and there would be no significant pressure anomalies in this area. Fig. 4-16shows that aquifer A is exposed at a high elevation where meteoric water is supplied. If a well is drilled at a low elevation point, pressure at that point will be abnormally high. Most abnormal pressures encountered by the oil industry are not of this type. Abnormal pressures associated with a closed fluid system are usually more serious than those caused by an open system such as shown in Fig.

4-16. Hydrocarbon accumulation This type of abnormal pressure cannot cause a regional problem, but only a local problem where there is hydrocarbon accumulation. As depicted in the schematic diagram of Fig. 4-17,the excess pressure at point A due to hydrocarbons increases as the density difference between hydrocarbons and water increases and also as the height of the hydrocarbon column increases. The mathematical expression is as follows:

Fig. 4-17. Cross-sectional view of hydrocarbon accumulation.

111 where

Pa

= excess pressure due

to hydrocarbons,

pw and P h = densities of water and hydrocarbons, and h = height of hydrocarbon column.

Osmosis Clays and shales are known to act as semipermeable membranes that will permit osmotic and electro-osmotic pressures t o develop wherever there is a marked contrast in the concentrations of the dissolved salts on either side of the clay or shale. The relationship between salinity difference and osmotic pressure is shown in Fig. 4-18. Osmotic pressure does not rank with the other causes such as compaction and aquathermal pressure, which could generate anomalies up to several thousand psi; the osmotic pressure difference resulting from a salinity difference of 50,000 mg/l, NaC1, is only about 600 psi.

Salinity dif terence, in milligrams per liter

Fig. 4-18. Relationship of pressure differential to water salinity difference across a clay barrier. (From Jones, 1967.)

112 Fossil pressure Suppose a deeply seated, closed aquifer were uplifted due t o late-stage erosion which removed the shallower part of the sedimentary column. If the original fluid pressure in this aquifer were to remain constant during and after this uplift, the pressure would be abnormal for the shallower depth. However, for the fluid pressure to remain constant during uplift would be quite difficult, because the temperature would be decreasing at the same time, so that the fluid pressure would be bound to drop significantly. As the rate of pressure reduction during erosion is probably more than that of overburden-pressure reduction, subnormal pressures could result, but abnormal pressures would be very unlikely. Cementation If reservoir rocks underwent cementation, the cementing minerals would have plugged the pore spaces, thus causing pressures to increase. Cementation would probably have been more active if formation water containing mineral ions in solution was supplied continuously. In other words, there would have been more cementation if the system was open. Near-normal (hydrostatic) pressure is usually developed in the zones of an open system. Even though some abnormal pressures might have been generated by the effect of cementation, I believe the degree of overpressuring could not be very high. Tectonics Lateral tectonic compression is sometimes used as an explanation for abnormal pressures in areas close t o tectonically complicated areas. A good example is the Foothills of the Canadian Rockies, where some overpressured reservoirs such as the Cardium have been found. Most of these overpressures can be explained by compaction and aquathermal phenomena as shown in Chapter 2. Whether the tectonic effect contributed to these overpressures can be determined by examining hydrofracturing data. Fig. 4-19shows the locations of the fields in Alberta that had fracturing data. Fig. 4-20 is a plot of fracture-pressure gradient (psi/ft) versus depth. Most of the fracture-pressure gradient values are less than 1psi/ft, suggesting that the fractures are mostly vertical to open fracture spaces against horizontal compression in the rocks. In other words, fracture pressure measures horizontal tectonic pressure. If the fracture orientation was horizontal, the fracture-pressure gradient would be approximately 1 psi/ft or more, because the overburden must have been lifted to cause the horizontal fractures, and because the overburden pressure in this area probably increases at approximately 1psi/ft.

113

Fig. 4-19. Index map showing locations of fields where fracturing pressures were obtained in Alberta.

FRACTURE PRESSURE GRADIENT 0.1

0.1

0.b

0.7

0.1

0.9

1.0

-

PSI/FT 1.1

1.1

1.3

1000.

2000.

3000

~

IOOO-

I-

#

JOOO-

6000’

7000.

8000

9000

~

-

Fig. 4-20. Plot of fracture pressure gradient versus depth for Alberta wells. Numbers refer to wells studied.

114 COEFFICIENT K I

looo,

2000

3000

I

0.4

0,5

0,6

0,7

0,8

0,9

IqO

1,l

l,2

l,l

~

4000

-

5000

-

tY 6000-

7000

-

8000

-

9000

-

10000

-

Fig. 4-21. Plot of-coefficientKi versus depth for Alberta wells.

Relationships between fracture pressure gradient, F, fluid pressure, p , vertical effective stress, uv, and horizontal, effective stress, o h , and depth, D, are shown as follows:

F =p / D + u ~ / D

(4-21)

where Ki = a constant relating the horizontal effective stress to the vertical defined by Matthews and Kelly (1967). The value Ki usually ranges from 0.3 to 1.0. Fig. 4-21shows the plot of the coefficient Ki versus depth for the Alberta data. It is interesting to note that the fracture-pressure gradient generally decreases toward the Foothills area (Fig. 4-20).The value Ki, which is the ratio of horizontal effective stress to vertical, also decreases westward (Fig.

4-21). In other words, the fracturing data do not support the idea that the present tectonic pressure in the Foothills area is more than that in other areas. If there are significant tectonics, the fracture-pressure gradient and Ki should increase westward. There is no doubt about the fact that in this area strong tectonic forces caused foldings and faultings in the geological past. However, whether such

115 tectonic forces still exist or not must be a different story. Berry (1973)recently discussed the possibility that tectonic forces caused abnormal formation pressure in the California Coastal Ranges. Unfortunately, Berry provided no data on either presentday horizontal stress or shale compaction. In other words, there seem to be no real physical data presented to document Berry’s hypothesis. By making plots of shale transit time versus depth for a few wells in this general area, I found typical undercompacted shales which probably originated from subnormal fluid expulsion during sedimentation. Most of these abnormal pressures can, therefore, be more satisfactorily explained by compaction phenomena than by tectonic forces that in any case have not been documented. In summary, I personally feel that the significance of tectonic forces in generating abnormal fluid pressure in many areas, has been overstressed. In most cases these hypotheses have been made from geological interpretations only, rather than actual physical data. Reexaminations of these areas are strongly recommended. Generation of hydrocarbons Generation of natural gas by thermal processes at deep burial may cause significant overpressures; This problem was discussed recently by Hedberg (1974) in relation to the diapiric movement of overpressured shales. It is, however, not easy to quantify such effect at present. Charging Shallower reservoirs are sometimes charged through faults or fractures by high fluid pressures that originated in deeper sections. Fig. 3-15shows an example of such charging, in which shale transit time is plotted with equivalent pressures from drill-stem tests. The drill-stem-test pressures at points 5 and 6 are higher than those indicated by nearby shales. Charging is considered the best explanation for these pressure anomalies. References Anderson, D.M. and Low, P.F., 1958. Density o f water adsorbed by lithium-, sodium-, and potassium-bentonite. Soil Sci. SOC.A m . Proc., 22: 97-103 Athy, L.F., 1930. Density, porosity and compaction of sedimentary rocks. Bull. A m . Assoc. Pet. Geol., 1 4 : 1-24. Barker, C., 1972. Aquathermal pressuring - role of temperature in development of abnormal-pressure zones. Bull. A m . Assoc. Pet. Geol., 56: 2068-2071. Berry, F.A.F., 1973. High fluid potentials in California Coast ranges and their tectonic significance. Bull. Am. Assoc. Pet. Geol., 51: 1219-1249. Boatman, W.A., 1967. Shale density key to safer, faster drilling. World Oil, 165: 69-74.

116 Bredehoeft, J .D. and Hanshaw, B.B., 1968.On the maintenance of anomalous fluid pressures, I. Thick sedimentary sequence. Geol. SOC.A m . Bull., 79: 1097-1106. Burst, J.F., 1969. Diagenesis of Gulf Coast clayey sediments and its possible relation to petroleum migration. Bull. A m . Assoc. Pet. Geol., 53: 73-93. Cebell, W.A. and Chilingarian, G.V., 1972. Some data on compressibility and density anomalies in halloysite, hectorite, and illite clays. Bull. A m . Assoc. Pet. Geol., 56: 796-802. Chapman, R.E., 1972. Clays with abnormal intersitial fluid pressures. Bull. A m . Assoc. Pet. Geol., 56: 790-795. DeWitt, C.T. and Arens, P.L., 1950. Moisture content and density of some clay minerals and some remarks on the hydration pattern of clay. Trans. 4th Znt. Congr. Soil Sci., 2: 59-62. Dickinson, G., 1953. Geological aspect of abnormal reservoir pressures in Gulf Coast Louisiana. Bull. A m . Assoc. Pet. Geol., 37: 410-432. Hedberg, H.D., 1974. Relation of methane generation to undercompacted shales, shale diapirs, and mud volcanoes. Bull. Am. Assoc. Pet. Geol., 58: 661-673. Hottman, C.E. and Johnson, R.K., 1965. Estimation of formation pressures from logderived shale properties. J. Pet. Technol., 17:717-722. Hubbert, M.K. and Rubey, W.W., 1959. Role of fluid pressure in mechanics of overthrust faulting, I. Geol. SOC.A m . Bull., 70: 115-166. Jones, P.H., 1967. Hydrology of Neogene deposits in the northern Gulf of Mexico Basin. Proc. 1st Symp. Abnormal Subsurface Pressure, Louisiana State Univ., Baton Rouge, La., pp. 91-207. Kojima, K., Ikeda, K. and Kawai, K., 1977. Mud compaction and the mathematical model for “compaction current” in sedimentary basin. Jpn. Assoc. Pet. Technol. J., 42: 100-106. Lewis, C.R. and Rose, S.C., 1970.A theory relating high temperatures and overpressures. J. Pet. Technol., 22: 11-16. Magara, K., 1971. Permeability considerations in generation of abnormal pressures. Soc. Pet. Eng. J., 11: 236-242. Magara, K., 1974a. Compaction, ion-filtration and osmosis in shale and their significance in primary migration. Bull. A m . Assoc. Pet. Geol., 58: 283-290. Magara, K., 1974b. Aquathermal fluid migration. Bull. A m . Assoc. Pet. Geol., 58: 25132516. Magara, K., 1975a. Reevaluation of montmorillonite dehydration as cause of abnormal pressure and hydrocarbon migration. Bull. Am. Assoc. Pet. Geol., 59: 292-302. Magara, K., 1975b. Importance of aquathermal pressuring effect in Gulf Coast. Bull. Am. Assoc. Pet. Geol., 59: 2037-2045. Martin, R.T., 1962. Adsorbed water on a clay: a review. Clays Clay Miner., 9 (Proc. 9th Natl. Conf. Clays and Clay Minerals, 1960),Pergamon, New York, N.Y., pp. 28-270. Matthews, W.R. and Kelly, J., 1967. How to predict formation pressure and fracture gradient. Oil Gas J., 65 (8):92-106. Nichols, E.A., 1947. Geothermal gradients in Mid-Continent and Gulf Coast Oil fields. Trans. A m . Inst. Min. Metall. Eng., 170:44-50. Overton, H.L. and Timko, D.J., 1969. The salinity factor: A tectonic stress indicator in marine sands. Oil Gas J., 67: 115-124. Pirson, S.J., 1963. Handbook of Well Log Analysis. Prentice-Hall, Englewood Cliffs., N.J. Powers, M.C. 1967. Fluid-release mechanisms in compacting marine mudrocks and their importance in oil exploration. Bull. Am. Assoc. Pet. Geol., 51: 1240-1254. Reynolds, E.B., 1970. Predicting overpressured zones with seismic data. World Oil, 171 (5): 78-82.

117 Rogers, L., 1966. Shaledensity log helps detect overpressures. Oil Gas J., 64 (37): 126127,130. Rubey, W.W.and Hubbert, M.K.,1959. Role of fluid pressure in mechanics of overthrust faulting, 11. Geol. SOC.Am. Bull., 70: 167-206. Smith, J.E.,1971. The dynamics of shale compaction and evolution of pore fluid pressure. Znt. Assoc. Math. Geol. J., 3: 239-263.


Chapter 5 APPLICATION OF CALCULATED PRESSURES AND POROSITIES

The pressures and equivalent mud weights calculated from well-log data by the methods discussed in Chapter 3 are useful information for safe drilling. Pressure and mud-weight mapping based on old well data will provide important information for drilling a new well in the same general area. Predictions of top of overpressured interval and mud weights, or pressures to be used for drilling this interval, are valuable information before drilling any new well. Pressure and porosity profiles As discussed in Chapter 2, abnormal pressures generated in the western Canada Foothills were reduced during and after the erosion that removed the uppermost sections of sedimentary rocks. The abnormal pressures there today are much lower than those encountered in the Gulf Coast, but they still can be significant. Fig. 5-1 is an index map of the study area in western Canada. Figs. 5-2, 55-3 and 5-4 show comparisons of shale porosity and reservoir pressures measured by drill-stem tests. Calculated pressures can also be used for interpreting the subsurfacedrainage which is, in most cases, influenced by the permeability, thickness and areal extent of the interbedded sandstones and carbonates. Fig. 5-5 shows the pressure profiles of two wells drilled in the Mackenzie Delta area. Pressures are calculated in the Cretaceous shales only. In Fig. 5-5A,the shales are underlain by porous dolomite. The calculated fluid pressure drops toward the underlying dolomite. The inferred directions of fluid migration are shown by arrows. In Fig. 5-5B,ori the other hand, the Cretaceous shales are underlain by low-permeability Paleozoic shales. No downward fluid flow is indicated in this case. Fig. 5-6 is a plot of Miocene mudstone porosity instead of fluid pressure, calculated from the sonic log of a well in the Mitsuke oil field, in Japan. Calculated porosity decreases toward interbedded sandstone and volcanic rock, indicating that fluid pressure decreased in the same direction. It is apparent that these sandstones and volcanic rocks provided good drainage conditions. As a matter of fact, hydrocarbons were found in both types of reservoir rock. Fig. 5-7 shows another example of shale porosity4epth and pressure-

120

I

I \

I

I \,

\, \,

\

........................................................ E Groups of w e l l s ................................................. Location showing surface shale s a m p l e s . . ................................ Location of Strathmore well 7-12-25-25-W4 ............................. Wells studied

.I48

r _ _ - - _ _

.#

.m

Fig. 5-1. Index map showing the wells studied and section lines in western Canada. (From Magara, 1972.) A

B+C

Fig. 5-2. Comparison of shale porosity distribution and subsurface fluid pressure in western Canada. (From Magara, 1972.)

121

Fig. 5-3. Comparison of shale pososity distribution and subsurface fluid pressure in western Canada. (From Magara, 1972.)

depth relationships for a well in Japan. The shale porosity trend is interpreted to be near normal level to 2200 m. Below this depth the shales are undercompacted, but they tend to return to normal toward the underlying volcanic tuff bed below 2750 m. The inferred direction of compaction fluid flow is downward below 2450 m and upward above. Fractured parts of this tuff bed contain commercial gas. Fig. 5-8shows the porosity cross-section of the area including the Mitsuke oil field, Nagaoka Plain, Japan. At Mitsuke SK23,there is a sharp shale porosity 0

m

P

Q f R

rn’

Fig. 5-4. Comparison of shale porosity distribution and subsurface fluid pressure in western Canada. (From Magara, 1972.)

122 FLUID PRESSURE, X 1000 psi 1

2

3

4

2

-

3

u, Y 0 0

x

I' u n., 0

I

4

d

5

Fig. 5-5. Shale porositydepth plots for two wells in the Mackenzie Delta, Canada.

Mudstone porosity - 0.3 0.40.5

" i 0 5

looor 1000

5

2

n

1500m

1500mb----

L,

0.1

0.3 04 0.5

Fig. 5-6. Mudstone porosity-depth plot of Mitsuke SK-23,Japan. (From Magara, 1969.)

123 SHALE POROSITY

0

0I

02

0.3

0.4

0.5

FLUID PRESSURE

0.6

0.7

A Fig. 5-7. Mudstone porositydepth and calculated pressure-depth plots of Shiunji SK21, Nagaoka Plain, Japan.

decline toward the Mitsuke tuff. Here the volcanic rocks are fractured and contain commercial hydrocarbons. In the area including Tsubame and Nakanokuchi the tuff is relatively tight, where porosity of the overlying shales stays relatively constant. The fluid-pressure crossTsection in Fig. 5-9 shows the fluid-flow directions in the shales clearly; there is no downward fluid movement in the latter area. Figs. 5-10 and 5-11 indicate the porosity-depth and pressuredepth crossaections for another area not very far from the previous one (Nagaoka Plain, Japan). Sharp porosity and excess-pressure reductions toward the two

124 MITSUKE

m SK-2".,

KUROSAKA SK-3.

SH IRAVAMA SK-1

TSUBME R-1

1.

NWWOKXHI SK-2

~"IIoL G€PlW 1M O

2000

Fig. 5-8. Mudstone porosity profile in the Mitsuke tuff region, Nagaoka Plain, Japan. (From Magara, 1968a.)

Fig. 5-9. Calculated fluid-pressure profile in the Mitsuke tuff region, Nagaoka Plain, Japan. (From Magara, 1968b.)

125

HlNWt VO111

Fig. 5-10. Mudstone porosity profile in the Fujikawa-Kumoide region, Nagaoka Plain, Japan. (From Magara, 1968a.)

volcanic reservoirs are indicated. Both reservoirs are porous and permeable, and contain commercial amounts of natural gas. Fig. 5-12shows more examples of fluid-pressure profiles in northern Canada. Drainage map About 300 wens, all with sonic logs, have been studied in northeastern British Columbia and northwestern Alberta (Fig. 5-13).Shale porosity of the Cretaceous formations has been determined from the sonic logs by using the empirical relationship established for this area. The vertical shale porosity distributions along lines A-A', B-B', C-C', D-D', E-E', F-F' and G-G' (Fig. 5-13)are shown respectively in Figs. 5-14to 5-20.

1500

2000

2500 KrI

Fig. 5-11. Calculated fluid-pressure profile in the Fujikawa-Kumoide region, Nagaoka Plain, Japan. (From Magara, 1968b.)

126 PRESSURE-psi

PRESSURE-psi

PRESSURE-psi

9" DST '011 STAIN

SANDSTONE LIMESTONE

\

PRESSURE-psi 4000

PRESSURE-psi 1000

I00

1

I

k*

i,.

300

3000

I

5000

6000

PRESSURE-psi 4000 5000 I

I

I00

PRESSURE-psi 9000 10.000

I

i.

.

1000

I

METAMORPHIC ROCKS

0

I

10,000'

11.000

3000 7000

K 'i 12.000

6ooo

\

Fig. 5-12. Examples of calculated pressure profiles in northern Canada.

The curved dashed lines in all seven figures indicate the boundary surfaces between the upward and downward fluid migration, determined by applying the same concept as used in Fig. 5-7. The thickness of the downward migration zone commonly increases westward. Abrupt porosity decreases in the shales close to the reservoir rocks occur mainly in the western part of the area, suggesting that the underlying rocks have relatively high permeabilities and the fluid would hence have migrated t o theqrelatively easily from the overlying shales. In the eastern parts, such porosity decreases

127 N \

1

BC.

Wells Studled.

........................................*

0

MUes. 100

200

d

Fig. 5-13. Index map showing the wells studied and section lines in northeastern British Columbia and northwestern Alberta, Canada. (From Magara, 1972.)

are not clear, suggesting that the underlying rocks have relatively low permeabilities. Fig. 5-21 shows an isopach map of the downward migration zone in the shales, as well as the locations of oil and gas pools in the underlying rocks (Lower Cretaceous, Jurassic and Triassic formations). As seen in Fig. 5-21, most of these oil and gas pools are in the area where the zone of downward migration is thicker than 500 ft. A recent geochemical study (Powell, 1977) has shown that the source of most of these oils is considered to be the overlying Cretaceous shales. The next four diagrams (Figs. 5-22 to 5-25) indicate the results of detailed fluid migration in the Beatton River area in northeastern British Columbia. Fig. 5-22 is an index map of the area. Fig. 5-23 is a diagrammatic northwestsoutheast cross-section. The Notikewin sandstone is interbedded with Cretaceous shales in the eastern part of the study area but disappears to the

.

y*l

mow-

low

-

StA L t V t l A

-low-mno'-

Fig. 5-14.Shale porosity profile along A-A'. (From Magara, 1972.)

SDK C A L S l A N

1000'-

PERMIAN

Fig. 5-15.Shale porosity profile along B-B'. (From Magara, 1972.)

I, I,

Fig. 5-16. Shale porosity profile along C-C'. (From Magara, 1972.)

3--m-=mo

.l"" "a111

o

m

Fig. 5-17. Shale porosity profile along D-D'. (From Magara, 1972.)

.

I,

.".I I .

....

-

.......

OIL

.Imo

-

I I

OIL

Fig. 5-18. Shale porosity profile along E-E'. (From Magara, 1972.)

3ow

-

loo0'-

StA I€V€l

-F

-1000'

-

.zoo0

I

Fig. 5-19. Shale porosity profile along F-F' (From Magara, 1972.)

-

OIL

m-

-1410'-

?40*

-

I 011

Fig. 5-20. Shale porosity profile along G-G'. (From Magma, 1972.)

132

BC

ALTA

Contours showing shale thlclmess (feet) d downward mlgratlon 011 (L gas pools Ln Lower Cretaceous, Jurassic & Triassic formations

-100-

’

Miles

50

0

50

100

Fig. 5-21. Map showing the thickness of the downward fluid-migration zone in Cretaceous shales. (From Magara, 1972.)

west. Sonic logs of approximately forty wells studied showed a trend of decreasing shale porosity and excess fluid pressure toward the sandstone in the eastern part. The drainage envelope, defined as the interval from the highest excess-pressure level in the overlying shales t o the highest excesspressure level in the underlying, is relatively thick in the eastern part but thins toward the west as the Notikewin thins and disappears. The potentiometric elevation of the sandstone and its equivalent was also estimated. Fig. 5-24 is an isopach of the drainage envelope just discussed, which shows it thinning toward the northwest. Fig. 5-25 is the potentiometric map, which shows the general fluid flow t o be southeast from shaly areas to sandstone areas. Because the Notikewin sandstone is exposed in the east without any significant closure, no commercial hydrocarbons are found in it. It is interesting to note that, according t o Fig. 5-25, compaction fluids have moved for at least thirty miles within the shales - a relatively long distance for fluids t o migrate in these rocks. The next example is derived from the study of Devonian shales in the Northwest Territories in Canada (Willow Lake area). The index map is shown in Fig. 5-26. Fig. 5-27 presents the pressuredepth plots along sections A-A’, B-B‘ and C-C’ (see Fig. 5-26). At well 43 along section C-C’,a distinct downward fluid migration is observed. The underlying limestone

133 1000'

I

039 S

40.

37. 032

0 3 3 34.

3.

38.

035

36.

031 KT 041

*26

.IS

028

020

014 013

IS .1 042

4.

-3

-2 0

r6

011.

I

l2l*W

30

I 30

57-00 00

Fig. 5-22. Index map showing wells studied in the Beatton River area, northeastern British Columbia, Canada.

1

DIRECTION OF FLUID MIGRATION

N.W.,

4

FLUID PRESSURE d

Fig. 5-23. Schematic d-iagram showing the drainage envelope and potentiometric surface estimated from logderived fluid-pressure plots.

134

!-

IT --

Fig. 5-24. Map showing the thickness of the drainage envelope for Notikewin sandstone in the Beatton River area, northeastern British Columbia, Canada.

reservoir at this point is relatively porous and fractured. On the other hand, at wells 41, 47 and 46 in the synclinal area, there is no recognizable downward component of fluid flow, which suggests that the underlying limestone is relatively tight. It is interesting to note that relatively old shales such as those shown in Fig. 5-27retain the memory of undercompaction generated a long time ago. Figs. 5-28and 5-29are isopachs of downward fluid-migration zones in the Nagaoka Plain, Japan. The porosity and pressure profiles for the same area are shown in Figs. 5-8to 5-11. Porosity maps The shale porosity of a given geological horizon can also be mapped, as shown in Figs. 5 3 0 and 5-31.The porosity values tend to be reflected by the maximum burial depths. Usually such maps show the significance of burial

I

Fig. 5-25. Potentiometric surface map of Notikewin sandstone and its equivalent (shale) in the Beatton River area, northeastern British Columbia, Canada.

74

i

65

I'

c 116

-

-

I

-~~

-

_-

I10

Fig. 5-26. Index map showing the wells studied and section lines in,the Willow Lake area. Northwest Territories, Canada.

136

I

-C

Fig. 5-27. Calculated fluid pressure profiles of Devonian shales in the Willow Lake area, Northwes, Territories, Canada.

on shale porosity reduction; the direction of lateral fluid flow may also be inferred from them. Fluid and hydrocarbon drainage The shale porosity profile may be tied to geochemical data to indicate primary migration of hydrocarbons. Fig. 5-32 compares a shale porosity4epth

137

Reprnmhlin d l o Conlrd

g praunl mdrlom IMckr*01.1 rhich have caused Ih. , domward cmpaclion Conlw shorr*lp , s i d a r Ihickmr.r v k h acrlie Ih. Fqikawa vokanic RHrWir

,

//

,I’

well

Conlour m

Fig, 5-29. Map showing the thickness of the downward fluid-migration zone in the Mitsuke tuff region, Nagaoka Plain, Japan. (From Magara, 1968a.)

MINAMI-YOITA

I

0

T

1

2

3

4

p

m

Fig. 5-30. Map showing the porosity distribution of the mudstones overlying the Nagaoka agglomerate region in the Nagaoka Plain, Japan. (From Magara, 1968a.)

I

/

-

141

ch/%

Q01

005

Fig. 5-32. PorosityAepth and chic,, (ratio of carbon in hydrocarbons over total organic carbon)-depth plots of the MITI-Yoshida well, Niigata, Japan. (From Fujita, 1977.)

plot with the ratio of carbon in hydrocarbons over total organic carbon in rocks (ch/co). These data are from a well in Niigata, Japan. Hydrocarbon concentration is expected to increase with depth due to maturation as shown by the line “Normal increasing trend” in the right-hand side of this figure. However, the actually analysed data shows the decline toward the 3500-4000 m interval where the drainage condition was interpreted to be excellent. Fig. 5-33 is another example of reducing c h / C o in the drainage zone in the Niger Delta, suggesting primary hydrocarbon migration.

Fig. 5-31. Map showing the porosity distribution of the mudstones overlying the Mitsuke Tuff in the Nagaoka Plain, Japan. (From Magara, 1968a.)

142

Fig. 5-33. Porosity-depth and ch/c,, (ratio of carbon in hydrocarbons over total organic carbon)-depth plots of a Niger Delta well, Nigeria. (From Fujita, 1977.)

In summary, plots of shale pressure or porosity can be used to indicate the nature of the subsurface drainage as well as to whether the interbedded rocks are permeable and/or have a large areal extent. Maps of the drainage envelope, potentiometric surface and porosity can be used to interpret regional fluid-flow directions. References Fujita, Y., 1977. The role of shale porosity anomaly in hydrocarbon exploration. Jpn. Assoc. Pet. Technol. J., 42: 107-116. Magara, K., 1968a. Compaction and migration of fluids in Miocene mudstone, Nagaoka Plain, Japan. Bull. Am. Assoc. Pet. Geol., 52: 2466-2501. Magara, K., 1968b. Subsurface fluid pressure profile, Nagaoka Plain, Japan. Bull. Jpn. Pet. Inst., 10: 1-7. Magara, K., 1969. Upward and downward migrations of fluids in the subsurface. Bull. Can. Pet. Geol., 17: 20-46. Magara, K., 1972. Compaction and fluid migration in Cretaceous shales of western Canada. Geol. Surv. Can. Pap., 72-18: 81 p. Powell, T.G.,1977. Origin of petroleum in the western Canadian sedimentary basin, Alberta -A geochemical study (Abstract). Can. SOC.Pet. Geol., Res. Bull., 4: 1-2.

Chapter 6

CONCEPT AND APPLICATION OF FLUID-LOSS CALCULATIONS

This chapter describes a method of calculating fluid loss from shales before and after maturation * and includes several actual examples of such calculations by computer. Although several problems associated with these methods still have to be solved through further research and field tests, the concept used in the programs is likely to interest many geologists and geophysicists. The main purpose of this chapter, therefore, is to describe the concept. Calculation of fluid losses from shales before and after maturation Fig. 6-1shows a schematic diagram of a shale porosity-depth plot. It indicates that, at shallow depths, porosity decreases at a constant rate (normal compaction). From empirical studies of shale porosity in many sedimentary basins, the porosity-depth relationship of this normally compacted zone is known to approximate an exponential function. For convenience, the equation given by Rubey and Hubbert (1959)already discussed in Chapter 2 (eq. 2-1) is repeated here:

If the porosity is plotted on a logarithmic scale and the depth on an arithmetic scale, the porosity-depth relationship in the normal compaction zone can be expressed by a straight line; the intercept of this line with the surface (or zero depth) gives the do value, and the slope of the line is directly related to the c value. Deeper in the section depicted in Fig. 6-1,the porosity is abnormally large. This zone is the so-called “undercompacted zone”, in which the fluid pressure is abnormally high (see Chapter 3). The maturation zone - that is, the zone within which potential source rocks are mature enough geochemically to yield oil - is also indicated on this diagram.

* A source rock must be mature before it can generate oil. Young and/or immature sediments, since they contain only methane (C,) and no gasoline-range hydrocarbons (C4C7), may be sources of dry gas but not of oil. As temperatures increase, maturation begins with generation of wet gases and gasoline-range hydrocarbons, until the rock reaches a mature state when it can be a source of oil. The state of maturation can be evaluated from geochemical data.

144 SHALE POROSITY-

-r UNDER-

C 0 MPACTED

p"'

'

Fig. 6-1. Schematic porosity-depth relationship of shales.

The porosity4epth relationship in Fig. 6-1represents the compaction state of the section at present. It can also, however, be used to represent the compaction history of shales in the subsurface. The shale at A, for example, may be assumed to have had an original porosity (at time of deposition) equivalent to the surface porosity, Go, extrapolated from the subsurface normal compaction trend as shown in Fig. 6-1,and may also be assumed to have experienced a compaction history close to the normal compaction trend during continuous burial. (Porosities at several different stages of burial are shown schematically in Fig. 6-1by open rectangles.) Finally, it reached the present porosity level shown by the solid rectangle at A. During this compaction process, large volumes of fluids must have been expelled from the shale. In order t o calculate the volume of fluids expelled - that is, fluid loss - the shale at A must first be restored t o its original uncompacted state. This restoration of shale volume can be made, as follows: Assuming that compaction of shales occurs simply by expulsion of fluids, and that there is no mineralogical change in the shale matrix during compaction, the following relationship will result:

or :

vo=v(-) 1-4 1-40

145 where V = volume of shale after compaction, $I = porosity of shale after compaction, Vo = volume of shale before compaction, and

$Io

= porosity

of shale before compaction.

Eq. 6-2means that, if the volume V and the porosity $I of a shale at a given depth (after compaction) are known, and the original porosity $I0 (before compaction) is assumed, the original volume of the shale, VO (before compaction), can be calculated. Then the fluid loss, W,due to Compaction, equals the difference between Vo and V, or:

When the values V , $I, and $Io are, for example, 1cu ft, 10%(or 0.1) and 60% (0.6),respectively, the original shale volume Vo can be calculated (see eq. 6-2)as: V,=lX--

1 - 0.1 0.9 - 1 X - = 2.25 cu f t 1 - 0.6 0.4

The original volume of 2.25 cu f t has been reduced to 1 cu f t in this case. Hence, the volume of fluid loss is given as (see eq. 6-3):

W

=

Vo - V = 2.25 - 1 = 1.25 cu f t

This is the total fluid loss that has occurred since the time of deposition. In order to simplify the above manual calculations, a convenient chart has been developed as shown in Fig. 6-2.The bottom scale is the present porosity ($I) of the shale. The left-hand scale is the original shale volume ( V O and ) the right-hand scale is the water volume (W).Several diagonal lines are drawn to show original shale porosities ($Io), ranging from 90 to 10%. The previous calculation can be simplified by plotting a point that corresponds to $Io = 60% and $I = 10%(see point A in Fig. 6-2),and by reading the value on the right-hand scale at 1.25 cu ft. This chart shows the volume of water expelled from 1cu f t of shale (present volume). Suppose we have geochemical data by which we can recognize the top of the mature zone, as indicated by M in Fig. 6-1.If it is assumed that this “maturation threshold’’ has been at the same depth during burial and compaction a , ~it is at present, the fluid-loss volumes before and after maturation can be calculated. If the shale porosity at the top of the mature zone M is given as $Im, the post-maturation fluid loss Wm is the difference between the

146

Fig. 6-2. Diagram relating original and present shale porosities to original shale volume and expelled water volume.

shale volume Vm at the top of the mature zone, or M in Fig. 6-1, and the shale volume V at A at present, or:

because Vm = V k s ) (see also eq. 6-2) -4m Similarly, the fluid-loss volume before maturation, W,, is given as:

(see eq. 6-5).

(6-5)

147 The total of W , and W , is equal to W as:

to:)

w = w, + w, = v -

(6-7)

(see eq. 6-3). Now let us continue the sample calculation of fluid loss, using 6 , = 15% or 0.15.In this case, the fluid loss after maturation is calculated from eq. 6-4 as:

W,=lX

0.15 - 0.1 - 0.05 - 0.06 cu f t 1-0.15 0.85

The fluid loss before maturation is given from eq. 6-6as:

W,=lX-

1-0.1 0.6-0.15 - -0.9 x--0.45 - 1.19 cu f t X 1-0.6 1-0.15 0.4 0.85

These calculations show that 1 cu f t of shale at A lost 1.19 cu f t of fluids before it reached geochemical maturity, and only 0.06 cu f t afterwards. The total volume is the same as previously calculated from eq. 6-3. Let us calculate the fluid losses of the shale at B in Fig. 6-1.This shale is undercompacted, that is, it has an abnormally high porosity for its present depth of burial. As in the case of the shale at A, it is quite reasonable to assume that this shale had a porosity of @ J ~when deposited, and that its early compaction history followed the “normal compaction trend” line. Some time before it reached its present burial depth, however, compaction of this shale, and consequent fluid expulsion, seem to have been arrested or at least severely suppressed *. In such cases, we do not know with certainty what happens; let us consider two alternatives: (1)The shale at B experienced a normal compaction history until porosity was reduced to a value represented by C (Fig. 6-1);at that point compaction was arrested completely, so that with further burial the shale experienced no further porosity decrease and no further fluid loss. This behaviour pattern is indicated by the vertical dashed line from C to B. If this in fact did happen, these shales have experienced no fluid loss since they entered the geochemically mature zone. (2) The shale at B experienced a normal compaction history until porosity was reduced to a point represented by D (Fig. 6-1);fluid loss during this

* Later in this chapter I will discuss the question of whether a shale can reach a given state of Compaction and then, with further burial, become “uncompacted” by rebounding with a consequent increase in porosity.

148 phase can be termed primary fluid loss *. From that point on, with further burial, porosity decrease and consequent fluid loss (which we can call secondary fluid loss **) continued, but at a severely diminished level, as represented by the inclined dashed line from D to B. In this case we can assume that there has been some fluid loss from the shales now at B since they reached the geochemically mature zone. We can be sure that secondary fluid loss takes place at a slower rate than primary fluid loss, but the rates are very difficult to quantify. Controlling factors will include, (1)how long the rate of compaction has been slower, (2) the pressure gradient in the undercompacted zone, (3)shale permeability and water viscosity. The latter two factors will change with time as well. As mentioned previously, we do not know which of these alternatives more accurately describes the compaction process in nature. We suspect that secondary fluid loss, where it has occurred, has been very small compared with primary fluid loss. On this basis, and in view of the difficulties in quantifying secondary fluid loss, our calculation programs are based on the first alternative presented - that normal compaction proceeds to a certain depth and is totally arrested. We feel the errors inherent in this assumption are acceptable at our present level of understanding. One obvious important observation t o be made from Fig. 6-1is that the shale at A, which is normally compacted, has lost more fluids (both before and after maturation) than the one at B that is undercompacted. Therefore, from a fluid-loss standpoint, other conditions being equal, normally compacted sequences may be more attractive for exploration than undercompacted ones. Now let us consider the importance of the position of the geochemically mature zone with respect to the pattern of fluid loss during compaction. In the next example, (Fig. 6-3), the shale at B is undercompacted, but the top of the geochemically mature zone is shallower than in the previous example. In this case, some post-maturation fluid loss can be calculated even for that shale. Fig. 6-4depicts a different situation. The shale at B is undercompacted, but its porosity is less than that of the shales at the base of the normally compacted zone (point X in Fig. 6-4). In this case, the program assumes the $o-X-B trend as the compaction history of the shale at B. Note that this trend is composed of two straight lines ($o-X and X-B) which, on semilog paper, show a sharp break-over point at X, but if the trend is plotted on arithmetic paper the change is gradual (Fig. 6-5). Also, in this example, although the top of the geochemically mature zone is below the zone of normal compaction, there has been post-maturation fluid loss from the shales at

B.

* The corresponding porosity loss can be termed primary porosity loss. ** The corresponding porosity loss can be termed secondary porosity loss.

149 SHALE POROSITY-

X

I

ZONE

UNDERCOMPACTED

Fig. 6-3. Schematic porosity-depth relationship of shales when the top of the ge0chen.ically mature zone is relatively shallow.

The question of "rebounding" In the example shown in Fig. 6-3, the porosity of the undercompacted shale at B is greater than that at the base of the normal compaction zone (or at X ) . If we were to assume the C-X-B trend instead of the C-B trend to ;HALE POROSITY

-

SHALE POROSITY-

DEPTH

DEPTH I

7

I

UNDERCOMPACTED MATURATION

UNDERCOMPACTED MATURATION

B

I Fig. 6-4.Schematic porosity-depth relationship of shales when the porosity of the undercompacted shale at B is less than that at the base of the normal compaction zone (point XI.

Fig, 6-5. Schematic porosity-depth relationship of shales plotted on arithmetic paper.

150 describe the compaction history of the shale at B, it must first have undergone a porosity decrease to the point represented by X, then a porosity increase by reverse compaction (rebounding) from point X to B. But rebounding or reverse compaction of formations in the deep subsurface is not an easy process to explain geologically. Therefore, the current program assumes the &-C-B trend to be the compaction history in this case. If one insists on this rebounding concept, one may calculate the volume of the fluid losses between points C and X, and points X and B. If the B-C line is vertical (porosity at B is the same as that at C), the fluid-loss volume between C and X is the same as that between X and B, but the latter has a negative sign (reverse Compaction). This means that while the shale at B could have lost some amount of fluids between points C and X, exactly the same amount of fluids has been returned to it during the reverse compaction between points X and B, so that the net fluid loss from C to B is zero. Therefore, the calculation based on the rebounding concept yields the same result (see line C-B in Fig. 6-3) as obtained from the current program. If, however, the return to higher porosities with depth (or apparent rebounding) is caused by some other factor, such as montmorillonite dehydration as proposed by Powers (1967) and Burst (1969), the situation becomes more complicated. The current program cannot evaluate this effect. An important point in this respect is that the volume of water released from the montmorillonite on dehydration may be calculated if the amount of dehydrated montmorillonite is known, but all this water could have been trapped in the shale pores. The volume of water squeezed out of the shale (fluid loss) must differ from the volume released from the montmorillonite. Therefore, evaluation of the montmorillonite dehydration alone cannot significantly improve the fluid-loss calculation. The volume of squeezed water in this case may be related to pressure gradient, permeability, viscosity, and time, as in the case of the secondary fluid loss mentioned above. The fluid-loss calculations described in this chapter, then, are based on several assumptions: (1) The original porosity of a shale can be derived by extrapolating “porosity versus depth” plots of the normally compacted zone. (2) Fluid loss from the shales is predominantly and directly related to compaction history. (3) “Rebounding” of shales is not a realistic explanation for the excessive porosity in undercompacted shales. (4) Undercompacted shales can be considered to have undergone a twostage compaction history - an early stage of normal compaction, and a later stage of little or no compaction, despite deeper burial. Calculation of fluid loss with time

So far we have described the concept of fluid-loss calculation, and the technique for calculating not only total fluid loss but also the proportions of

151

r-

3

2

STAGE 1

UNIT

i

UNIT

I

UNIT

Fig. 6-6. Schematic diagram showing the method of calculating fluid-loss volumes during successive depositions of increments of sedimentary cover.

loss before and after maturation. Now we can proceed to describe a further refinement of the technique, which is t o calculate the fluid loss of a shale sequence during successive time intervals, as it is buried progressively deeper. In this way, fluid-release history can be related in a quantitative way not only to geochemical maturity but to other timedependent factors such as trap development. Fig. 6-6 illustrates schematicalIy the burial depths and compaction states of the three units A , B and C at three different points in time. The column at stage 3 shows the three at present. If the uppermost formation, C, were removed, we would have a column such as that at stage 2, representing a point in geological time before C was deposited; similarly, stage 1 represents a point in time before B was deposited. Because the overburden load over formations B and A is less at stage 2 than at stage 3, the bulk volumes of these formations at stage 2 (B2 and A , ) must be greater than those at stage 3 (B3 and A 3 ) . The difference between volumes B2 and B3, or A2 and A3, as shown in Fig. 6-6, is equal t o the volume of fluid loss between stages 2 and 3. Stage 2 may be considered the end of the sedimentation of formation B , or the beginning of the sedimentation of formation C. Stage 3 may be similarly considered as the end of the sedimentation of C. Therefore, the volumes Bz - B3 and A 2 - A 3 can be considered the respective volumes of fluid loss from formations B and A , during sedimentation of the youngest formation, C. Similarly, the volume A -A is the fluid-loss volume from formation A during sedimentation of formation B (see columns at stages 1 and 2, Fig. 6-6). The sum of fluid losses during different increments of time should, of course, be equal t o the “total fluid loss” described in the previous section.

152 We can calculate the fluid losses before and after maturation within successive time increments based on this concept. Examples of fluid-loss calculation Fig. 6-7 is a cross-section through five wells drilled in northern Canada. It shows gross time-correlation lines along with the calculated thickness of section eroded at the presentday surface (the method of calculating the thickness of eroded section is described in Chapter 2). The maturation zone for each well, interpreted from geochemical data, is indicated by the shaded area. In well 1, for example, the top of the maturation zone, based on sapropel (amorphous organic matter), is at 6100 ft. If we take into account the erosion of 3500 f t at this location, we can conclude that 9600 f t of burial was necessary for maturation of sapropelic shale. Figs. 6-8 and 6-9 are examples of a fluid-loss-composition plot by computer. In the relatively sandy intervals, the "sand + silt versus clay" composition derived from a gamma-ray log is shown with porosity. In the shaly intervals only fluid loss is plotted. For calculation of fluid loss, shale porosity data are necessary. They are obtained from the sonic log by using the follow-

WELL 1

2

3

5

4

1600

-.

^^

-I

1800-

LOSS ZONE

--

I-

I

E1-L-

TIMING OF FLUID LOSS

'MAXIMUM

FLUID LOSS TIME

Fig. 6-7. Geological cross-section of five northern Canada wells. L-LK = late-Late Cretaceous time; ET = Early Tertiary time; LT = Late Tertiary time.

TOP OF MATURATION ZONE I S 6100 FT (FOR SAPROPELIC ORGANIC MATTER)

EK

TOTAL FLUID LOSS

FLUID LOSS

E-LK

FLUID

Loss

( C U FT / C U F T OF S H A L E )

10000

-% lo

.'

.O1

L O S S AFTER MATURATION

F I t FLUID LOSS BEFORE MATURATION

l2O0O

i

'F-

t .

0

.

.

.

i .

. 100

ROCK COMPOSITION 1'.

Fig. 6-8. Fluid-loss plots of well l ( 1 ) . EK = Early Cretaceous time; E-LK = early Late Cretaceous time.

TOP OF MATURATION ZONE IS 6100 FT IFOR SAPROPELIC ORGANIC MATTER)

FLUID LOSS

L-LK

ET

FLUID LOSS

FLUID LOSS

LT

( C U FT / C U FT O F SHALE)

.Ol

1 ,

.1

.Ol

.I

1

.1

.Ol

1

,

a

100

:-.

3

10000

10000~

lOO0Ol

!

MATURATION 11000

MATURATION

b

1200 ! 0~

12000~

#

12000;

z i !I 0

ROCK

!

,

"

100

0

,

I

,

COMPOSITION X

Fig. 6-9. Fluid-loss plots of well l(2). L-LK = late Late Cretaceous time; ET = Early Tertiary time; LT = Late Tertiary time.

155 ing empirical relationship between shale porosity, @, and sonic log transit time, At, (ps/ft) for this area:

9 = 0.00374At - 0.206 The left-hand side of Fig. 6-8 shows total fluid loss, both before and after maturation, for the zone 9700--12,200 ft, in cubic feet of expelled fluids per cubic feet of present-day shale. Note that the fluid-volume scale is logarithmic. The right-hand margin of this plot shows the per cent of total fluid released after maturation. The total of the fluid loss before and after maturation is about 1 cu f t per cubic foot of shale. This means that an original 2 cu ft of shale was reduced by compaction to about 1 cu ft, having lost 1cu f t of fluids. Post-maturation loss has generally been less than 10%of the total. The volume of fluid loss after maturation generally increases toward the sandy zones, which suggests that favorable compaction drainage existed near them. The two other computer plots in Fig. 6-8, and all three plots in Fig. 6-9, show fluid losses before and after maturation of the same interval (970012,000 f t of present depth) during sedimentation of each of five successively younger slices of the overlying section : (1) Lower Cretaceous (corresponding geological time is Early Cretaceousshown as EK in the figures), (2) lower (part of) Upper Cretaceous (early Late Cretaceous time -E-LK), (3) upper (part of) Upper Cretaceous (late Late Cretaceous time - L - L K ) , (4) Lower Tertiary (Early Tertiary time - E T ) , ( 5 ) Upper Tertiary (Late Tertiary time - L T ) . These plots show that (1)most of the total fluid loss occurred during Early Cretaceous and early Late Cretaceous times, but (2) this interval then had not yet been buried to maturation depth. It reached maturation depth at some point during late Late Cretaceous time, and has released post-maturation fluids only since then. As mentioned previously, the sum of these five successive fluid losses is the same as the total fluid loss shown on the lefthand side of Fig. 6-8. This example well was drilled on a domal structure that probably originated from shale diapirism. The time of formation of the structure is interpreted to have been Late Tertiary, according to geological and geophysical analysis of the area. Therefore, most post-maturation fluid losses seem to have occurred before the structure was formed, and the fluids have moved elsewhere. No commercial hydrocarbons were found in this well. Note that the plots in Figs. 6-8 and 6-9 show the fluid losses of only part of well I; i.e., the interval from 9700 to 12,200 ft. The major post-maturation fluid-loss zones of the entire well are shown in Fig. 6-7 by thick vertical bars, and it is apparent that there are three that lost significant volumes of fluids. The right-hand side of each bar shows the time of fluid loss for each

156 of these three zones. The uppermost, for example, below 7200 f t in well 1, lost its fluids during Late Tertiary time (LT). The middle zone lost its postmaturation fluids during Early Tertiary-Late Tertiary time (ET-LT), with maximum loss during Late Tertiary time (as indicated by the horizontal bar in LT).The lowermost zone, most of which is shown in Figs. 6-8and 6-9,lost its post-maturation fluids during late Late Cretaceous-Early Tertiary time (L-LK-ET), but the main loss occurred during late Late Cretaceous time. As we can see from this figure, the time of major post-maturation fluid loss becomes earlier as we go deeper, because the deeper section reached maturation depth earlier. If the time of formation of a structure is late, as interpreted in well 1, a deeper section may not be too attractive for exploration because most of the post-maturation fluids may have been lost before the structure was formed. A shallow section, with geochemical maturity, conversely, could be attractive because most post-maturation fluid loss may have occurred late, after formation of the structure. The second column from the left in Fig. 6-7shows a similar plot of well 2. This well was drilled on a flank of the same structure as was well 1. The upper two zones are interpreted to have lost post-maturation fluids during Late Tertiary time. The timing of the structural development is the same as for well 1 - Late Tertiary. The upper zones, therefore, are interesting (at least from the point of view of fluid-loss timing) but no commercial hydrocarbons were found in this well. Most of the geochemically mature zones in three other wells (3,4 and 5 ) are in the Lower Cretaceous section (Fig. 6-7),and the time of post-maturation fluid loss for these wells is interpreted to have been Tertiary. According to geophysical data, trap development had begun as early as the end of Early Cretaceous time, and most structures were completed during Late Cretaceous time. It is therefore possible that most fluids have migrated effectively toward available structures. Hydrocarbon reservoirs have been found in the Lower Cretaceous section of this general area. At present, play and prospect evaluation based on fluid-loss history can be only very tentative. In addition to the many problems as yet unsolved, we do not know how much fluid loss per unit of presentday shale volume is necessary for efficient hydrocarbon migration. To acquire some understanding of this problem, we must gather compaction and fluid-loss data from many basins where hydrocarbons do occur, and where fluid-loss history can be related especially to trap timing and maturation. Correction for nonclays Shales are commonly composed of a large proportion of clays and some nonclays (probably of silt size). According to X-ray analysis of cores and cut-

157

CORRECTED FOR NON-CLAY CONTENT TOTAL FLUID LOSS (CU F f /CU FT OF SHALE)

.Ol

.l

1

TOP OF MATURATION ZONE I S 6100 F T (FOR SAPROPELIC ORGANIC MATTER)

EK FLUID LOSS

E-LK FLUID LOSS

1

FLUID LOSS BEFORE MATURATION

Fig. 6-10. Fluid-loss plots of well I ( I ) , corrected for nonclay contents.

tings samples of this area, relatively clean shales contain at least 20% nonclay minerals, most of which is quartz. Amounts of nonclays, of course, vary widely within shale zones. In the previous fluid-loss plots, the most sandy intervals were separated from the shaly, based on a quantitative evaluation of the nonclays by gamma ray. These plots show the fluid-loss calculation only for the relatively shaly zones. The wide variability in composition of the shaly zones may affect the quality of the expelled fluids as hydrocarbon sources. Assuming that most organic matter is associated with clays, fluids expelled from very silty or sandy shales containing a large percentage of nonclay materials may not be as good as fluids from relatively pure shales. With this concept, a correction

158

CORRECTED FOR NON-CLAY CONTENT TOTAL FLUID LOSS

TOP OF MATURATION ZONE I S 6100 F T (FOR SAPROPELIC ORGANIC MATTER)

EK FLUID LOSS

(CU FT /CU FT OF SHALE)

.01

.l

1

-

E L K FLUID LOSS .01

10000

10000

11000

11000

12000

12000

.1

1

FLUID LOSS BEFORE MATURATION

MATURATION

3

Fig. 6-11. Fluid-loss plots of well 1(2),corrected for nonclay contents.

for the nonclays can be made, as follows: Wcorr

=

WVclay

(6-8)

where

W,,,, = fluid-loss volume corrected for nonclays, and Vclay = volume fraction bf clays in rock. If Vclay is 100%or 1, W,,,, and W are the same. Figs. 6-10 and 6-11show corrected fluid-loss plots for the same interval as in Figs. 6-8 and 6-9. Vclay was calculated from the gamma-ray log. The fluidrelease volumes in the corrected plots are generally less than those in the un-

TOP OF MATURATION ZONE IS 9600 FT (FOR CUTICULAR ORGANIC MATTER)

1 - L K FLUID LOSS

TOTAL FLUID LOSS (CU FT /CU FT OF SHALE) .01

.I

E l FLUID LOSS .Ol

1

10000

.1

1

L T FLUID LOSS .Ol

I 3

.1

1

100003

;

1

2

->

11000

MATURATION

12000

:>i

8

120001

+

f

0 " "

ROCK

160

0

100

0 " "

100

o

n

o

a

1 '0 0

*

COMPOSITION %

Fig. 6-12. Fluid-loss plots of well I before and after maturation, based on maturation of cuticular organic matter,

160 corrected plots (Figs. 6-8 and 6-9), because shales always contain some amount of nonclay minerals. When the correction is applied, therefore, most of the calculated post-maturation fluid losses during Early and Late Tertiary times fall under the minimum scale (0.01 cu ft) and cannot be shown on the plot. Effect of organic facies The depth to the maturation threshold varies with organic facies; sapropelic (amorphous) organic matter usually becomes mature at a lower temperature and therefore at a shallower depth than nonsapropelic (cuticular) organic matter of the same geological age. A t well I, the top of the maturation zone for cuticular organic matter is interpreted to be at 9600 ft, compared with 6100 f t for sapropelic. The left-hand column of Fig. 6-12 shows total fluid loss (before and after maturation) for the maturation threshold of 9600 ft. (If the eroded thickness of 3500 f t is taken into account, this means that 13,100 f t of burial would have been necessary to mature the cuticular shales at this location.) No correction has been made for nonclay materials. Again, it can be demonstrated that the zone of interest (9700-12,200 ft) had not reached maturation depth during Early Cretaceous and early Late Cretaceous times, so the corresponding pre-maturation fluid-loss volumes are the same as those in Fig. 6-8;they are, therefore, not included in Fig. 6-12. The three right-hand plots in Fig. 6-12 show respective fluid losses for late Late Cretaceous, Early Tertiary and Late Tertiary times. As the maturation threshold is displaced downward - in this case because of difference in organic facies - the volume of post-maturation fluid loss becomes less, and the time of important fluid loss later. Fluid-loss curve The fluid-loss history discussed above can also be presented in the form of a curve, as shown in Fig. 6-13. Fig. 6-13A shows the fluid loss (cubic feet/ cubic feet of shale) versus burial depth * (feet) of the average shale now buried to between 11,000 and 12,000 f t in well I. The portions of fluid loss occurring before and after maturation are indicated. In this example the top of the geochemically mature zone, based on sapropel, is 6100 ft, and no correction has been made for shale composition. The geological ages corresponding to the burial depths are shown in the right-hand side of Fig. 6-13A.

* Note

that the burial depth shown in this case is the depth before surface erosion of 3500 ft, as indicated previously. Therefore, the shale at 11,000 ft (present depth) had once been buried to a depth of 14,500 ft.

161

A

B

FLUID LOSS

(cu n m n

FLUID LOSS (T.BBL /lo00 Fl X lo00 SQ MILES OF SHALE)

OFSHALE)

1 32

loo00

POST-MATURATIO

E"T

El

+LT

E l + 11

I t Fig. 6-13. Fluid-loss curves before and after maturation for the interval of 11,00012,000 ft at well 1.

An example of the fluid-loss curve of a large block of shales can be seen in Fig. 6-13B,in which the fluid-loss volume from a block 1000 ft thick (11,000-12,000 ft) and 1000 square miles in area is shown in trillions of barrels. Fig. 6-13Bshows that this volume of shale has lost about 5 trillion barrels of fluids since deposition. In this example, the dimensions of the block of shale were selected arbitrarily to serve as a sample calculation. In a real situation, if one can define the limits of drainage envelopes within compaction sequences, at least quantitative estimates can be made of the compaction fluids they have released with time. Then, by applying a hydrocarbon/water ratio for compaction fluids, one can calculate the volumes of hydrocarbons that could have been carried to available traps from the potential source beds within the drainage envelope. It is quite apparent, then, that several factors are interacting:

162 (1)Compaction history, whether normal or modified by the development of undercompacted zones, provides the basic fluid-release history of the zone of interest. (2) The character of the organic matter in the zone of interest (organic facies) will determine the depth (and time) at which it becomes geochemically mature and capable of releasing hydrocarbons to the compaction fluids being expressed. (3) Traps may develop at any time during the compaction history of the zone of interest. The relative timing of compaction-fluid release, onset of maturation, and formation of traps is unquestionably a key consideration in the attractiveness of an exploration prospect. Fluid-loss mapping Fluid-loss volume in cubic feet/square feet was calculated for the downward migration zone in Cretaceous shales, shown in Fig. 5-21 (Chapter 5 ) . Fig. 6-14shows a map of such fluid-loss volume, as well as the locations of oil and gas pools in the underlying older reservoir rocks. Most oil and gas pools are concentrated in the area where the greater volume of fluid is considered to have been expelled.

-

N W T B C

ALTA

Contours showing volume of flulds (cu ft /ft2) expelled downward from Cretaceous ahales Oil & gas pools In Lower Cretaceous, Jurassic & Triassic formations

Miles

50

0

50

100

I I

* * I. I Fig. 6-14. Volume of fluids expelled downward from Cretaceous shale in western Canada.

163 References Burst, J.F., 1969. Diagenesis of Gulf Coast clayey sediments and its possible relation to petroleum migration. Bull. Am. Assoc. Pet. Geol., 53: 73-93. Powers, M.C.,1967. Fluid-release mechanisms in compacting marine mudrocks and their importance in oil exploration. Bull. Am. Assoc. Pet. Geol., 51: 1240-1254. Rubey, W.W. and Hubbert, M.K., 1959. Role of fluid pressure in mechanics of overthrust faulting, 11. Geol. SOC.Am. Bull., 70: 167-206.


Chapter 7

EVALUATION OF PRESSURE AND CAPILLARY SEALS

As shown in Chapter 5 , there are many examples of oil and gas accumulations overlain by slightly undercompacted shales. In most cases the shale compaction tends to return to normal toward the reservoir sections. For those intervals composed of slightly undercompacted (overpressured) shales and normally compacted shales and sandstones, Evans et al. (1975)proposed the term “mixed compaction facies”. These slightly overpressured shales are considered to have restricted the vertical escape of the fluids in the sandstones and are called “pressure seals”. About 90% of the accumulated hydrocarbons (mostly gas) in the Beaufort Basin are found in the mixed compaction facies. Fig. 7-1 shows an example of a pressure--depth plot in the Beaufort Basin. Another example, in the Gulf Coast, is shown in Fig. 7-2. The shales between 7000 and 9000 f t , which are overpressured, overlie the normalIy pressured sandstones between 9000 and 10,000 ft. These pressure seals commonly occur in an intermediate depth range in many young sedimentary basins. The existence of capillary seals has been known to engineers and geologists in the oil industry for many years, but the pressure-seal phenomenon is relatively new. A comparison of the respective properties of pressure and capillary seals is of interest. Pressure seal (1)Seal for any form of hydrocarbons, whether in solution in water or in a separate phase. (2) Developed during the intermediate stages of shale compaction. (3) May be more important for gas than for oil, since gas is more soluble in water. Capillary seal (1)Seal only for hydrocarbons in the hydrocarbon phase. (2) May be more significant during the later stages of shale compaction, when it becomes more effective. (3) May be more important for oil than for gas, since gas is more soluble in water. For a pressure seal to be effective, the excess pressure due to the buoyancy of the hydrocarbon column must be less than the excess pressure of the overlying shale above hydrostatic pressure. As pressure seals are usually associated with capillary seals, the combined effect may control the trapping condition.

166 P O R E PRESSURE-

I CALCULATED PRESSURE

SAND

'-

ROSTATIC PRESSURE

Fig. 7-1. Example of a calculated pressure-depth plot in the Beaufort Basin, Canada.

Although pressure sealing plays only a part in overall sealing effectiveness, the depth and timing of a pressure seal can be interpreted from shale compaction data. In other words, we are able to tie this information t o other factors, such as structural timing and maturation timing, which are important in petroleum assessment. However, the timing of capillary-seal development is not readily evaluated.

PERCENT SAND

PAN AMERICAN NO. A-5 FARMERS LAND AND CANAL MANCHESTER FIELD. LOUISIANA

MEASURED PRESSURES I N MANCHISTER FIELD

13

-

NORMAL HYDROSTATIC GRADIENT

y./

7, \

PERCENT SHALE

FORMATION FLUID PRESSURE IPS1 X lwOl

Fig. 7-2. Example of a calculated pressurdepth plot in the Gulf Coast. (From Schmidt, 1973.)

167 Calculation of pressure-sealing depth In the current fluid-loss program, the compaction history for the shale at B (Fig. 6-1)was assumed t o have been &,-C-B (normal compaction and normal fluid release up t o point C, and no compaction below point C, Fig. 7-3).This means that the shale now at B stopped releasing fluids at the time it became buried t o point C. The shale at D in Fig. 7-3,which is more compacted than the shale at B , can be similarly assumed to have terminated compaction at point E, which is deeper than point C. This suggests that when the shale at B had reached point C, the depth at which compaction and fluid release stopped, the shale at D was still releasing its fluids because it had not yet been buried to its compaction termination depth, point E. From the time when the underlying shale at B had reached depth C, therefore, the fluids expelled from the shale at D had t o move upward or horizontally, because fluid movement in the shale at B was restricted. The same concept can be applied to the shale at F, which was still losing fluids when the overlying shale at B reached the termination point C. From then on the fluids from the shale at F must have moved downward, or horizontally, because the shale at B formed a pressure seal above. The shale at B, therefore, which has maximum porosity in the D-B-F interval and reached the termination point of fluid release first, determines the direction of compaction-fluid movement - upward or horizontally above and downward or horizontally below - and is considered to be a seal for the underlying rocks. SHALE POROSITY-

I/

DEPTH

1T -

MATURATION ZONE

-r UNDER-

COMPACTED

v

UNDERCOMPACTED

"

Fig. 7-3. Schematic diagram showing the compaction history of shales when there is no secondary fluid loss. Fig. 7-4. Schematic diagram showing the compaction history of shales when there is some secondary fluid loss.

168 As discussed earlier, it is possible that, rather than ceasing abruptly, compaction merely slows down as some critical depth is reached; this would allow for some secondary fluid loss as depicted schematically in Fig. 7-4.

A

B

NO SECONDARY POROSITY LOSS

1/10 SECONDARY POROSITY LOSS

SEALING DEPTH (1000FT.)

P

?

?

j SEA1 : Tertiary

t lp

Q

G TIME Late

SEALING DEPTH (I000FT.)

j !

Early

-.

I

c r e t a e

..

.

.. ..

.. 9

.

?

t

SEA1 Tertiary

..... ....

1000

,

...

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

.

.... .. .. ... . ..... .._.. . .

.*.. .... .

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

. . *.. . ..

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

. ..

... ... -. ..

..

11000~

* G TIME

.. .

... ... *

'P '?

...- ....

.. .. .. .. *

?

.?

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

.

., .. .. ..

10000'

?

1p

......__ . ....

llooo;

*.

LAYS

OROSITY

-...

.. ..

.. . . .. f

...

...

12000;

12000

.

.

6""

ROCK 1' COMPOSITION II.

SEALING-DEPTH

(

...

I AND SEALING-TIME

...)

PLOTS

Fig. 7-5. Sealingdepth and sealing-time plots of well 1 in northern Canada.

169 However, note that if the lines C’-B, E’-D, and GI-F are assumed to be parallel, we reach the same conclusion with regard to sealing effects as we did in Fig. 7-3. Shale B, having maximum porosity in the interval D-B-F, reached the termination point of fluid release first, and from that time acted as a seal for underlying compacting shales. Determination of the sealing depth is quite easy when one assumes that no secondary fluid loss occurred. An example of this kind of plot for well 1 is shown in Fig. 7-5A, in which the sealing depth is indicated. The sealing depth for the shale at 10,000 ft, for example, is calculated as 6800 ft. Fig. 7-5B shows a similar plot, assuming some secondary fluid loss; in this example 10% or & of the shale porosity is assumed t o have been lost by secondary fluid expulsion. With this assumption, the sealing depth is shallower than in the previous case (no secondary fluid loss), but the difference is not great. As already discussed, quantifying the effect of secondary fluid loss from shales is quite difficult, and I feel that the porosity loss due to this effect will not exceed 10%of the total shale porosity. We do not yet know whether this degree of accuracy justifies the additional effort required to peform the calculation. Calculation of pressure-sealingtime Once the sealing depth has been calculated, the point in geological time when sealing first began may also be determined. This calculation can be explained with the aid of a schematic diagram, Fig. 7-6. Suppose that shale 1 at 8000 f t of present depth (see left-hand column) is calculated to have reached sealing depth, i.e., to have become sealed, at 4000 ft. Let us restore this shale to a burial depth of 4000 f t , as shown in the right-hand column of Fig. 7-6, by using the concept described in Chapter 6 (eq. 6-2). Note that, since at the shallower depth the shale would have had greater volume, the restored column shows it as thicker.

PRESENT SHALE COLUMN

RESTORED SHALE COLUMN

4000 FT

I

Fig. 7-6. Schematic diagram showing the method of calculating sealing time.

170 Take shale 2 immediately above the first shale (see left-hand column) and restore it also t o a shallower depth (see right-hand column). Continue these restorations until the cumulative thickness of the restored section is equal to the sealing depth (4000 ft). In this case, the last shale t o be restored is shown by n in Fig. 7-6. From the restoration described above, we conclude that when shale 1 was sealed, shale n was either at the surface or still being deposited. If we know the geological age of shale n, we know the time when shale 1 became sealed. Fig. 7-5A includes a computer plot of sealing time, in which no secondary fluid loss is assumed, and Fig 7-5B shows one assuming 10%porosity reduction due t o secondary fluid loss. The sealing time for the shale at 10,000 f t in well 1 is Middle-Late Cretaceous. Calculation of fluid losses before and after pressure sealing If we can determine the point in time at which a clastic sequence containing potential reservoir beds was sealed, we can also calculate the amount of compaction fluid released by that sequence before and after sealing. Before explaining the method and showing examples of such calculations, we should discuss the usefulness of the information. During the earliest stages of compaction, movement of released fluids is almost entirely vertical; i.e., the newly deposited sediments are, in effect, continuing t o settle within the water column. With progressively deeper burial and consequent greater compaction, upward vertical movement of fluids through the compacting sequence will become increasingly difficult: lateral movement, and downward vertical movement, will become increasingly important. The precise pattern to be established will depend on many factors, such as the lithologic character of the sequence in terms of the number, quality and continuity of drainage intervals, and the rate of burial of the compacting sequence. In the case of a shaly sequence undergoing a normal compaction history, we can say intuitively that, at some point, it can act as an effective reservoir seal. Its effectiveness will depend on its thickness, lithology (clay versus quartz), and state of Compaction, as well as the character of the reservoir fluids (oil versus gas). Unfortunately, we do not know how t o define that sealing effectiveness quantitatively so that, even though a given exploration prospect may measure up in all other respects, there may remain some suspicion that reservoir seal is a problem - especially if the sequence is geologically young. On the other hand, if it contains zones of undercompaction overlying prospective reservoir zones, we can say, with certainty, that a seal has existed from some point in time. To be able to establish that point in time, and assess the effect of sealing on the direction and amount of potentially hydrocarbon-bearing compaction fluids making their way into the reservoir facies, should sharpen our understanding of the key factors in hydrocarbon occurrence.

171 NO SECONDARY POROSITY LOSS

1/10 SECONDARY POROSITY LOSS

SEALING DEPTH (1000 FT.)

SEALING DEPTH (1000 FT.)

P.?

? ? t

? . ? t. 0 ?

: SEALING T I M E Tertlary PLIO' Late PLEIST Middle Early Cretaceous

I

-

-

PLIO:

iALING T I M E Tertlary Late Early Cretaceous

I

-

1

PLEIS'

---

75OOF

d

, .'

8000;

,POROSITY

NON

..... ... .:.. h

LA YS

8 *...

.... *.., .:, ...:.

ssoo-

9000

h .... ..,. ._..

-. ... ..::..

:.:.I

ij

... .....

:y 9SOO;=

I ,I

-

.

FLUID LOSS A I O V E SOOOFT

Fig. 7-9. Fluid-loss plots before and after sealing, well A in northern Canada, corrected for nonclay content (l), LK = Late Cretaceous time; ET = Early Tertiary time.

w

4 0

E f - 3 FLUID LOSS

M I - 1 FLUID LOSS

MT - 2 FLUID LOSS

PLIO-PLEIST FLUID LOSS

(CU.FT./CU.FT. OF SHALE) .l

01

8000 7 5 --0

0

1 ,

~ 8000 '

8500

9000

2

7

3 8500

8500

9000

9000

i2

8500

9000

f b

9500

9500

9500

5

3 9500, I

>

LEGEND FLUID LOSS BELOW SOOOFI

FLUID LOSS A l b V F SOOOFT

Fig. 7-10. Fluid-loss plots before and after sealing, well A in northern Canada, corrected for nonclay content (2), MT = Middle Tertiary time.

PRESSURE

175

-

DEPTH

I .'... ...

\

Fig. 7-11. Schematic diagram showing maximum sealing pressure Psh and excess hydrocarbon pressure Ph.

I '

'B/

'*

0 I

0

1

/

200 3W 400 500 Mx) 700 EXCESS PRESSURE DUE TO HYDROCARBONS

100 I

/

0

I

1

1

500

1

8

1

1

1

"

I

"

1000 EOUIVALENT GAS COLUMN

'

"

EM) PSI

'

1500

FT.

Fig. 7-12. Plot of maximum sealing pressure versus excess hydrocarbon pressure (and equivalent height of gas column) for northern Canada wells.

176 diagram in Fig. 7-11shows the excess hydrocarbon pressure, Ph, at the top of the reservoir and the maximum excess pressure (here called “maximum sealing pressure”), calculated in the overlying shales, P s h . Fig. 7-12is a plot of maximum sealing pressure on the vertical scale and excess hydrocarbon pressure (and equivalent height of gas column) on the horizontal scale for northern Canada wells. At the line marked critical sealing pressure, both pressures are equal; if the sealing pressure is less than the excess hydrocarbon pressure, the reservoir will lose hydrocarbons. All the data are plotted above this critical-pressureline, suggesting that, in this area, the pressure seal alone can retain hydrocarbons (mostly gas) in the reservoir. Group B (Fig. 7-12)which is clearly separated from group A, has much higher sealing pressures for given excess hydrocarbon pressure. This higher pressure means that the shales in group B have a better sealing capacity, and are probably cleaner, than in group A. Capillary seal As mentioned earlier, hydrocarbons in their phase in the reservoir can be retained by capillary seal as well as pressure seal. The maximum (critical) height of oil column 2, retained by capillary seal is given by Berg (1975), as follows:

where = interfacial tension between oil and water, rt = pore-throat radius of overlying cap rock, rp = pore radius of reservoir rock, g = gravity acceleration, pw = density of water, and po = density of oil.

y

Berg considered a well-sorted, fine-grained, sandstone with a porosity of

26%. Such a natural aggregate may approximate a rhombohedral packing of uniform spheres in which pore sizes are 0.154 D, 0.225D and 0.414D D being the sphere diameter (Graton and Fraser, 1935).Fig. 7-13shows the maximum height of oil column as estimated by these assumptions: if the column exceeds this critical height, the oil will move; otherwise, it will

not. IDmis model, the reservoir rock, whose grain ghe, D,is 0.2mm,is over-

lain by the same, or finer, rock. The critical height of oil column (vertical scale) is shown for a given density difference (Ap) between water and oil.

177

Reservoir

X

Qx5

Dt

-

=

0 2 mm

in = 2O6 % I

7 = 35 d/cm

Mean grain size

-

rnm

B a r r i e r facies

a = 0,

n=26%

Fig. 7-13. Height of oil column, Z,, that can be trapped by barrier rock of mean grain size, D t , in a reservoir rock of grain size, D , = 0.2 mm where both rocks are composed of uniform spherical grains in rhombohedra1 packing and porosity, R, is 26%.Interfacial tension, y , is assumed t o be 35 dyn./cm. (From Berg, 1975.)

If, for example, this reservoir is overlain by a rock with 0.01 mm grain size (silt), an oil column of about 150 f t can be held by the capillary sealing capacity, when the fluid-density difference (Ap) is 0.2. If a finer rock overlies the reservoir, a longer oil column can be held. In the case of gas accumulation, the maximum column that can be held by the capillary seal is generally less, because the fluiddensity difference for gas is more than that for oil. If water is moving as a result of hydrodynamic force in the reservoir, the critical height of oil column, Z,, ,can be expressed differently (Berg, 1975):

178

Fig. 7-14. Diagram of oil stringer held in aquifer by downdip flow of water. (From Berg, 1975.)

where dh/dx = inclination of potential surface, and = horizontal width of the oil accumulation, (see Fig. 7-14). X, The optional sign in eq. 7-2 refers t o flow directions: the positive sign corresponds to dowdip flow and the negative to updip flow. In other words, more of a hydrocarbon column can be retained if there is a downdip fluid flow or downdip potential gradient - other conditions being equal. Laboratory model of a pressure seal Katz and Ibrahim (1971) demonstrated an interesting model of shale compaction, shown in Fig. 7-15. This model has more metal plates separating the springs than Terzaghi's, and has two water outlets at top and bottom. In the outer compartments, water can be expelled easily as the stress, S, is applied. The highest water pressure will remain in the central compartment, which contains the maximum amount of water (or porosity). This model simulates a shale bed 'intercalated by permeable sandstone beds, undergoing compaction and fluid expulsion. Magara (1972) reported the result of a similar experiment with natural clay (montmorillonite). Fig. 7-16 shows the apparatus he devised. It is composed of a metal base ( a ) and a brass tube 2 inches in diameter ( b ) divided

179

Perforated disc (low permeability) Shales

Springs (analog clay aggregates)

r

j Sands

Fig. 7-15. Schematic representation of shale compaction, porosity and permeability relationships, and creation of abnormally high pressure. k = permeability; 6 = porosity; t = time; p = pore pressure; T~ = specific weight of water; h = height t o which fluid will rise in the tubes; W = water; G = gas. (From Katz and Ibraham, 1971).

into ten 2-inch segments. Brass plates ( c ) are welded t o the top and bottom of each segment: the brass bolts ( d ) , which hold the segments together, permit removal of individual segments, and rubber rings between the segments prevent water leakage. Metal plates (e) are also placed between the segments and are used for slicing the clays. Above the ten segments is a cylinder ( f ) containing a brass piston (g) with a rubber tip. The piston may be systematically loaded at (h) to a maximum weight of 400 lbs. In experimental compaction, the cylinder and the ten segments are filled with water-saturated clays. During progressive loading, water is expelled through the top (i) and bottom (j)outlets. Two sandstone cores are placed - one at the top and one at the bottom of the clays. The inner volume and empty weight of each segment (b) are first measured. After a segment has been filled with water-saturated clay, it is weighed again; the difference is the weight of the water-saturated clay. The bulk density Pbw of the clay can then be determined by dividing this weight by the internal volume of the segment. Porosity is calculated by using the following equation: '

180

Fig. 7-16. Compaction apparatus used for compaction experiment. (From Magara, 1972.)

or : Pg

- Pbw

Pg-

(7-3)

Pw

Where P b w , pw and pe are bulk, water and grain densities of the clay and 4 is porosity. The pg of the clays is assumed t o be 2.65 g/cc. The initial porosity of the clays in each segment is calculated by eq. 7-3.Then the ten segments, a metal base, a cylinder and a piston are put together. The piston is loaded t o a weight of 400 lb. When compaction has taken place and water is expelled from the clays, the ten segments are removed and the weight of each, including the clays, is measured. Porosity distribution at this stage, is determined. By repeating such measurements, porosity patterns can be determined for different stages of compaction. New clays are added at the top (in the cylinder), and compaction experiments proceed.

181 A. MONTMORILLONITE +1.5 N NaCl WATER B U L K D E N S I T Y ( p m cc1 OPEN

OPEN

10 STAGE

'

STAGE4321

100

B. MONTMORILLONITE +1.5 N NaCI WATER

P o = 2.65 gm/cc P O R O S I T Y 1%) 50

100

BULK D E N S I T Y f g m / c c ]

20 -~ STAGE? 2-1

_..

p g = 2 . 6 5 gm/cc 1%) 50 0 STAGE 3-2

POROSITY

0

\ \

~

Fig. 7-17. Bulk density and porosity at different stages of compaction of montmorillonite saturated by 1.5 N NaCl water. A = top and bottom outlets are open; B = only top outlet is open.

Experiment A , using montmorillonite clay

*

Both top and bottom outlets are open, simulating the presence of permeable sandstones above and below a shale sequence. The clay has been saturated with about 1.5 N NaCl solution. Results are shown in Fig. 7-17A. Large porosity decreases occur at both ends, close t o the outlets; the decrease in the middle, on the other hand, is relatively slow. Experiment B, using the same montmorillonite clay This time, the bottom outlet is closed, simulating impermeability of the underlying sandstone (Fig. 7-17B).A large porosity decrease has occurred at the uppermost part of the clay, close to the top outlet. As mentioned earlier, fluids are retained in the reservoir rock by the combined effects of pressure seal and capillary seal. The relative importance of each sealing mechanism may be influenced by the type of hydrocarbons gas or liquid in the reservoirs, and also by the timing of their segregation from water. If all the hydrocarbons are in molecular solution in water, pressure is the only effective seal. If the hydrocarbons have their own phase, most of them can be retained in the reservoir by capillary seal alone. Therefore, the timing of hydrocarbon segregation in the geological past becomes important.

* Montmorillonite No. 25, John C. Lane Tract, (Bentonite) Upton, Wyoming.

182 SHALE POROSITY

PORE PRESSURE

Fig. 7-18.Schematic diagram showing porosity and pressure distributions of shales overlying reservoir.

As most subsurface and experimental profiles suggest, the pressure seal has its maximum porosity and pressure in the middle decreasing toward the interbedded permeable zones (Fig. 7-18). In other words, the shales having the highest pressures can act as pressure seals, while the more compacted shales immediately above the reservoir act as capillary seals. In the most typical examples, therefore, these two seals are present to prevent vertical escape of fluid from the reservoir. References Berg, R.R., 1975. Capillary pressures in stratigraphic traps. Bull. A m . Assoc. Pet. Geol., 59: 939-956. Evans, C.R., McIvor, D.K. and Magara, K., 1975. Organic matter, compaction history and hydrocarbon occurrence - Mackenzie Delta, Canada. Proc. 9 t h World Pet. Congr., 3: 149-1 57 (Panel discussion). Graton, L.C. and Fraser, H.J.,1935. Systematic packing of spheres with particular relation to porosity and permeability. J. Geol., 43: 785-909. Katz, D.L. and Ibrahim, M.A., 1971. Threshold displacement pressure considerations for caprocks of abnormal-pressure reservoirs. SPE 3222, 5th Conf. Drilling and Rock Mechanics, Austin, Texas. Magara, K., 1972. Compaction and fluid migration in Cretaceous shales of western Canada. Geol. Surv. Can. Pap. 72-18: 81 pp. Schmidt, G.W., 1973.Interstitial water composition and geochemistry of deep Gulf Coast shales and sandstone. Bull. A m . Assoc. Pet. Geol., 57: 321-337.

Chapter 8 CONCEPT OF THREE-DIMENSIONAL FLUID MIGRATION

In Chapter 6,the changing values of fluid loss from a given block of shale with burial or compaction are discussed. The next question is whether we can infer the direction of fluid expulsion during burial. It is possible to infer the vertical directions of fluid flow from the fluid pressuredepth plot of a shale section discussed in Chapter 3, and the directions of horizontal fluid migration in a sedimentary basin from the potentiometric map discussed in Chapter 5. But these inferred fluid migrations are present events; whether similar fluid-flow conditions existed in the geological past is a question that must be examined separately. A case in which present fluid-flow conditions d o not necessarily reflect paleoconditions can be found in an interval where the present pressure is hydrostatic. In this interval, the present potentiometric map would suggest no horizontal fluid flow, because there is no horizontal potential difference. However, there could have been some horizontal fluid flow in the geological past. It seems likely that such ancient horizontal fluid flow would have been controlled mainly by the loading patterns of the sediment layers. Directions of horizontal and vertical fluid migration As discussed in the first section of Chapter 4 (“Compaction”), the excess fluid pressure @lo)a generated by the instant loading of a thin layer under water, whose thickness is l o , is given as (see Fig. 8-1): (Plo)a =

(Pbo

-P w k l O

(8-1)

When the thickness of the newly added sediments changes, the excesspressure increase will change accordingly (Fig. 8-2). The excess-pressure increase @ho)aat point H due to the sediments of thickness ho is similarly shown as:

If the distance between these two points is X,the horizontal excess-pressure gradient (dpa/dZ)h due t o the new loading of the wedge-shaped sediments is obtained as:

184

OLD SEDIMENTS

UTH LEVEL

SEA LEVEL

-

10.ooo'

l0,W-

20,ooo'

20,ooo'-

30,oOo'

30,000'-

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

Inner and Middle Marino Shale and Sand Focier

Fig. 8-3. Geological section through southeastern Louisiana. (From Jones, 1967.)

Carto. Sect.. Geol. Dept.. LSU

W -1

-1 W

Fig. 8-4. Geological section in western Canada. (From Gussow, 1962.)

187

188 paleofluid migration, eq. 8-1must be revised. The relationship between the thickness of a layer before and after compaction is:

where

#-,

= porosity at the time of deposition (before compaction), lo = thickness at the time of deposition (before compaction), 4 = porosity at present (after compaction), and 1 = thickness at present (after compaction).

The porosity can be expressed in terms of density as follows:

and :

where Pb = density at present (after compaction), and pm = matrix (or grain) density of the sediments.

Introducing eqs. 8-6 and 8-7into eq. 8-5,we obtain: lO(pb0-P~) = d(pb-pw)*

(8-8)

The value (pbo - pw)Zo can, therefore, be replaced by (Pb - pw)l as follows:

Suppose we have a geological section with a well or wells drilled on or around it. Data are available on the density, or on the porosity which can be converted to density, in this well. The direction of horizontal fluid migration due t o the sedimentation of the unit can be derived from this information. The value (pro)a is expressed in psi as follows:

189 C

B

A

I

t

I

,

I

I I

I I

!

I I 1

,

I

I

,

I

I I

DENSITY 11 Pdf, 1254 0 ICE1 DIRECTION OF FLUID MIGRATION

3

1 -

i

EXCESS FLUID PRESSURE

6.' PI'

5 7 p,

47

PSl

Fig. 8-5. Direction of horizontal fluid migration calculated from a schematic geological cross-section.

Fig. 8-5 shows a schematic example of the calculation. In this case, there is a layer whose density changes from 0.9 to 1.1psi/ft (or from 2.08 to 2.54 g/cc) while the thickness remains constant (10 ft). The excess pressure calculated by eq. 8-10 is shown at the bottom of Fig. 8-5 ( p , = 0.435 psi/ft). The direction of fluid movement due to sedimentation of this layer is from left (C) to right (A). This means that when deposited the layer was thicker at C than at A, although present thicknesses at C and A are the same. This schematic example suggests that using a geological section or isopach without allowing for density change could result in misinterpretation of the direction of compaction-fluid migration. Volumes of vertical and horizontal fluid movement In the previous example (Fig. 8-5), the direction of horizontal fluid movement has been discussed. The next question is how much fluid has moved horizontally and vertically? Before tackling the problem of the volumes of vertical and horizontal fluid movement, the total fluid loss from a layer due to compaction must be determined. As discussed in Chapter 6, the relationship between the volumes of rock before and after compaction is shown as follows:

190 In this equation, the rock material is assumed to have undergone no mineralogical change during compaction; porosity reduction is due to compaction alone. Eq. 8-11can also be shown in terms of density by replacing the terms lo and 1 in eq. 8-8by Vo and V as follows:

From eq. 8-12,Vo can be shown as: v o = v ( Pb - Pw P, - Pw

)

(8-13)

The volume of total fluid loss, W, due to compaction equals the difference of the rock volumes before and after compaction, or

= V(Pb-Pbo)

(8-14)

PbO - PW

This equation is similar to eq. 6-3,but is expressed in terms of density values whereas eq. 6-3is based on porosity values. The total fluid loss from the 10-ft layer in Fig. 8-5is calculated by using eq. 8-14as follows:

A : W = 2.7 cu ft/sq f t B: W = 5.5 cu ft/sq ft C: W = 8.2 cu ft/sq f t where pw = 0.435 psi (1.0g/cc) and PbO = 0.8 psi/ft (1.85g/cc). To calculate the proportion of horizontal or vertical fluid movement, Darcy’s equation can be used: (see equations 8-3,8-4) 10 - ho

(8-1 5) (8-16)

where q = volume of fluids moving through sediments per unit area and unit time, k = permeability,

191 p = viscosity, and

h and v are subscripts denoting the horizontal and vertical directions. Dividing eq. 8-15 by eq. 8-16 we obtain: (8-17) The validity of applying Darcy's equation to fluid movement in a shale sequence may be a matter for discussion. There is an opinion that Darcy's equation does not represent the fluid-flow situation in shales. However, extensive studies of undercompacted shales and abnormal pressures in the young sedimentary basins of the world suggest that the absence of permeable beds (e.g., sandstones) is probably the most important factor in causing these undercompacted shales, which have resulted from subnormal fluid expulsion (Fertl and Chilingarian, 1976). If they are interbedded with many permeable sandstones of large areal extent, the shales will lose more fluids and compact t o a near-normal level. Although we do not know the exact mechanism of fluid migration in

26

0 8.0'

52

78

1

DATA OBTAINED FROM 0

-%a

-

6.0-

----

-

*

c VI

'

VISCOUS FLOW AT DIFFERENT TEMPS. NEUTRON SCATTERING SPECTROSCOPY SELF-DIFFUSION OF 2H 'HO SELF-DIFFUSION OF 3H 1 H 0 , 2 2 N a &36Cl

I

04 0

1.0

2 .o

MJM,

3 .O

(s/s)

FROM LOW

Fig. 8-6. Relationship between viscosity of water in clay and ratio of amount of water

(M,) over amount of clay (M,) or distance ( d ) from clay surface. (From LOW,1976.)

192

qV

kv

t kh

9h

Fig. 8-7. Schematic diagram showing porportional volumes of vertical and horizontal fluid migration due to wedge-shaped sedimentary loading.

shales, such migration appears t o be influenced, or possibly even controlled, by the mechanism of fluid movement in the interbedded permeable rocks, Darcy’s equation is known t o be applicable t o such permeable rocks. However, when applying Darcy’s equation t o a sandstoneshale sequence or, possibly, a shale sequence, we would have t o vary the fluid viscosity, p , with compaction. Fig. 8-6 shows the result of estimates by Low (1976) of water viscosity in montmorillonite. The viscosity changes from about 1to 8 cP. This finding suggests that when using eqs. 8-15 or 8-16 to estimate the volume of fluid movement, one must increase viscosity as compaction progresses. Permeability will, of course, decrease at the same time. In eq. 8-17, however, which calculates the ratio of the horizontal and vertical fluid volumes, the viscosity term is not included, so that the calculation is simpler. Note that ( I o - h o ) / X in eq. 8-17 is the rate of thickness change of the new sediments, and k h and k, are the permeability values in the older and deeper sediments. The values q h and q, are, of course, volumes of fluids moving through the old sediments due to loading of the new sediments (Fig. 8-7). Because the values kh and k, vary with burial and compaction, the ratio q h / q v changes with geological time even where (lo - h o ) / X stays constant. If the kh/kv stays almost constant throughout geologic time, although kh and k, may vary, the ratio q h / q v is essentially controlled by (I0 - h o ) / X . The value (lo - h o ) / X cannot be obtained directly from a geological section or an isopach map, because the sediments have already been compacted. Eq. 8-17 can be converted t o fit the present thickness change after compaction as follows: From eq. 8-8 we obtain: lo = 1

Pb-Pw Pbg--PW

(8-18)

193

c C

Fig. 8-8. Schematic geological section of wedge-shaped sedimentary loading.

If we consider a reasonably short distance of X in Fig. 8-7,the density variation of the layer after compaction at points L and H would be negligible. Hence, we obtain: (8-19) Introducing eqs. 8-18and 8-19 into eq. 8-17we will get: (8-20) where (1 - h ) / X is the rate of thickness change with distance at present, or after compaction. Suppose we have a schematic geological section as shown in Fig. 8-8. In order t o discuss fluid movement in layers B and A during sedimentation of layer C, ( 1 - h ) / X and the density P b of layer C must be determined. Pb,, and pw may be assumed to be constant. By using the horizontal and vertical permeabilities of layer B or layer A, the ratio @&& in either layer can be calculated. The fluid movement in layer A during sedimentation of layer B can be similarly estimated by using (I - h ) / X and Pb of B, and kh/k, of A. Strictly speaking, kh and k, of A in this case are not the present permeabilities but the permeabilities during sedimentation of B. If it is, however, assumed that kh and k, may have varied with geological time but the ratio kh/k, has stayed relatively constant, the present permeability ratio may be used in eq. 8-20. Suppose there is an interbedded sandstoneshale sequence. If the fluids were t o move vertically upward through this sequence, they would have to pass through the sandstones and shales, but the vertical fluid-flow rate would be controlled by the low-permeability shales, or: kv

ksh

(8-21)

194 where ksh is the permeability of the shales (in a strict sense, in the vertical direction) thus implying that the term k , in eq. 8-20may be replaced by k s h . If the fluids were t o move horizontally, they would move through the high-permeability sandstone because it was easier. In this case horizontal movement through the shales wouId be negligible. Therefore, the sandstone permeability becomes important. The fluid flow is, however, also controlled by the thickness of the sandstones. If we consider a unit thickness of the sandstoneshale sequence, the sandstone thickness can be expressed in terms of the sandstone fraction or per cent in the unit sequence. As a result, kh may be shown as: kh

2:

s

-k,, 100

(8-22)

where S is the sandstone percent in the unit sandstoneshale sequence, and k,, is the sandstone permeability in a horizontal direction. By replacing kh and k , in eq. 8-20by k,, and hsh using eqs. 8-21and 8-22, we obtain:

(8-23) As discussed previously, in most basins, the rate of thickness change of a layer with distance (I - h ) / X would be in the order of & & or less (after compaction). The ratio k,,/k,h would usually be very large, and could easily exceed 1000. The range of S/lOO is from 0 t o 1 (the range of S is from 0 to 100%). For most sediments the value (Pb - pw)/(pb, -pw) would be roughly 2. An interesting fact about eq. 8-23is that if S or k,, is very small, the ratio qh/qv becomes very small because ( I - h ) / X for most layers is quite small. In other words, the voIume of horizontal fluid-movement relative to vertical is quite small. The horizontal movement would increase as S or k,, increased. As mentioned previously, eq. 8-23is applicable t o an interbedded sandstoneshale sequence. This model may, however, be applied to a shaly sandstoneshale sequence, by using a lower value of k,, in eq. 8-23.In massive shale zones, S is zero or qh becomes zero. Hence in the massive shales horizontal fluid movement is negligible. If, however, the horizontal permeability of a shale is much greater than the vertical, some horizontal fluid movement should occur nevertheless. If the total fluid loss from sediments is calculated by eq. 8-14,the horizontal W h and vertical W , portions of that total may be estimated by the further use of eq. 8-23as follows:

-

195

-Iw

(8-24)

1

and :

Now, let us calculate qh/qv, Wh and W, for layer A during sedimentation of layer B (Fig. 8-6) using the following values:

1 100

(1 - h ) / X of layer B = -

Pb = 2.3 g/cc

- pw of layer B Pbo - Pw Pb

= 2.0 or

pw = 1.0 g/cc

kss/ksh of layer A = 1000 (we assume that the ratio kss/ksh stays constant during burial)

S / l O O of layer A = 0.3 (or S = 30%) W of layer A = 1cu ft/sq f t

*

The answer is as follows:

1 X 2.0 = 6 _ qh - 0.3 X 1000 X 9v 100

* The total fluid loss W from layer A during sedimentation of layer B can be calculated from bulk densities of A before and after deposition of B, by using eq. 8-14.If the density of A at the end of deposition of B was 1.87 glcc, this fluid-loss volume (1 cu ftlsq ft) corresponds to a volume for a 3-ft thick shale, because W in this case can be calculated by using eq. 8-14as follows: w=3x

1.87 - 1.65 1.65 - 1

+ 1 cu ftlsq ft.

196

W, = 1- 0.86 = 0.14 cu ft/Sq ft This result shows that the volume of horizontal fluid movement in layer A during sedimentation of layer B had been 6 times that of the vertical. Therefore, in Fig. 8-8, 0.86 cu ft/sq f t of the fluids had moved horizontally from left t o right, 0.14 cu ft/sq f t have moved vertically upward. If we multiply Wh or W, by the area of distribution of layer A, the total volume of fluids that moved in either direction from this layer can be estimated. Fig. 8-9 shows the volumes of horizontal and vertical movement when the value ( I - I Z ) / X changes from & to & [ S = 30%, k s s l k s h = 1000,(Pb P,)/(pb, - p w ) = 2.0, W = 1 cu ft/sq ft]. The volume of horizontal fluid movement, Wh, increases as ( 1 - h ) / X increases. When ( 1 - h ) / X is 0.0017, Wh iS a half Of the total fluids (Wh 0.5 CU ft/Sq ft, W, = 0.5 CU ft/Sq ft). The two fluid values Wh and W,,when the ratio ks./ksh varies from 100 to 1OO,OOO are shown in Fig. 8-10 [ S = 30%, (1 - h ) / =~ (Pb - p,)/(pb, - p,) = 2.0, W = 1cu ft/sq ft]. The value Wh increases with the increase of k s s / k s h . When kss/kshequals 170, Wh iS the S a I " 8s W, (Wh = 0.5 CU ft/Sq ft, W, = 0.5 cu ft/sq ft).

A,

a Y I-

51 .6 3

4

3 g

.4

I

3 5

.2

01 71o- '1O.WO

l/lW

1000

10

*

RATE OF THICKNESS CHANGE WITH DISTANCE

X

Fig. 8-9. Volumes of vertical and horizontal fluid movement as ( I

- h ) / X changes.

197 10

8

t a e l-

.6 3

i! 3 $

z

4

3 9 .2

,--

VERTICAL MOVEMEN1

W,

c PERMEABILITY RATIO OF SAND AND SHALE

I?a k h

Fig. 8-10. Volumes of vertical and horizontal fluid movement when k s s / k s h changes.

----__ -------_____

--,/

,

M

I

a

-"y

I

40

VERTICAL M

I

w

O

1

w

V

~

~

1

70

~

~

~

I

I

m

w

100

SAND PER CENT

s

Fig. 8-11. Volumes of vertical and horizontal fluid movement when S (sandstone cent) changes.

198 Fig. 8-11 indicates the volumes Wh and W,,when the sandstone per cent S varies in the range of 0-100 [(I - h ) / =~ &,kss/ksh = 1000, (pb - p , ) / (pb,, - p,) = 2.0, W = 1 cu ft/sq ft]. w h increases as S increases. Equal volumes of horizontal and vertical fluid movement are obtained when S = 5%. This figure shows that if the sand content is extremely small (say less than 3%)most fluids move vertically. As discussed above, loading of the sedimentary column resulted in the generation of excess pressures, which controlled the directions of fluid migration during burial. However, how much of the generated excess pressure is kept in the sediments is controlled by the permeability of the sedimentary rocks. In other words, a large excess pressure generated by thick and rapid loading in the geological past does not necessarily mean the existence of high excess pressure at present. Most of these excess pressures could have been dissipated if the permeability was relatively high enough. On the other hand, even a relatively small original amount of excess pressure, if largely undissipated, might result in significantly high pressures today. Present pressure conditions can be studied by the use of a pressure-depth plot or a potentiometric map. Therefore, the combined use of pressure maps based on sediment loading patterns, (paleo events) and presentday potentiometric maps will enable us to evaluate the complete history of fluid flow in a sedimentary basin, and to predict the presence of permeable beds in the subsurface. Example of a three-dimensional fluid-flow study The concepts discussed above were applied to a structure in northern Canada in a study based on both well-log and seismic data. Fig. 8-12 shows a structure map of the main producing horizon. Several we\\8 &,iscoue~e&gars in this =and&one xesexvoix section. Fig. 8-13 is an iso-

Fig. 8-12. Structure map of the top of a sandstone reservoir in a gas field, northern Canada.

199

Fig. 8-13. Isopach map of shales overlying a sandstone reservoir (shown in Fig. 8-12), northern Canada.

Fig. 8-14. Interval velocity map of shales overlying a sandstone reservoir (shown in Fig. 8-12), northern Canada.

Fig. 8-15. Cumulative loading map of shales overlying a sandstone reservoir (shown in Fig. 8-12), northern Canada.

200 pach of the shale section overlying the reservoir section. This isopach shows that structural development probably started immediately after deposition of the reservoir sandstones. The interval velocity of this overlying shale section was evaluated from seismic and sonic-log data and is shown in Fig. 8-14.These velocity values can be converted to density values by the use of an empirical relationship. Fig. 8-15shows the final map of cumulative excess pressure caused by deposition of the overlying shale section, as based on the present thickness in Fig. 8-13and the density derived from the velocity shown in Fig. 8-14. Eq. 8-10was used for this estimate. The directions of paleocompaction fluid flow are indicated by arrows in Fig. 8-15.A t present, this reservoir section is normally pressured: that is, there is now no lateral difference in potential, so that the potentiometric map cannot be used to study lateral fluid flow. The paleoloading map shown in Fig. 8-15can indicate the directions of paleocompaction fluid movement. References Fertl, W.H. and Chilingarian, G.V., 1976.Importance of abnormal formation pressures t o the oil industry. SPE 5946, SOC.Pet. Eng. AIME. Gussow, W.C., 1962. Regional geological cross sections of the western Canada sedimentary cover. Alberta SOC.Pet. Geol., Geological Cross-Section. Jones, P.H., 1967. Hydrology of Neogene deposits in the nothern Gulf of Mexico Basin. Proc. 1st Symp. Abnormal Subsurface Pressure, Louisiana State Univ., Baton Range, La., pp. 91-207. Low, P.F., 1976. Viscosity of interlayer water in montmorillonite. Soil. Sci. SOC.A m . Proc., 40: 500-505. Magara, K.,1971. Permeability considerations in generation of abnormal pressures. SOC. Pet. Eng. J. 11: 236-242. Magara, K.,1977. A theory relating isopachs t o paleo compaction-water-movement in a sedimentary basin. Bull. Can. Pet. Geol., 25: 195-207.

Chapter 9

POROSITY-PERMEABILITY RELATIONSHIP IN SHALES

The fluid-loss calculation discussed in Chapter 6 made use of the differences in shale-porosity levels at different stages of compaction. To define the directions of fluid movement, we used Darcy’s equation (Chapter 8), in which permeability is an essential factor. An established relationship between permeability and porosity in shales would facilitate the discussion of fluid movements in shales, because shale porosity is not difficult to obtain. Very little has been published regarding the permeability-porosity relationship in shales (see Bredehoeft and Hanshaw, 1968), probably because (1)most oil companies do not bother obtaining permeability data from shales because of their low economic importance; (2) shale permeability values measured in well cores are liable to be inaccurate, since cracks or fissures due to drilling can grossly inflate the actual permeability; (3) shales do not normally produce measurable amounts of subsurface fluids from which their permeability could be estimated. In this chapter, I intend to estimate the permeability changes of Cretaceous shales in the Alberta and Saskatchewan subsurface; then to combine shale porosity with permeability in order to investigate the relationship between them. After establishing this relationship, I will discuss the amount of upward and downward water movement from the undercompacted shales at several stages of compaction. Fluid-pressure gradients and movement of fluids in shales As discussed in Chapter 3, fluid pressure at depth D in the undercompacted interval shown in Fig. 9-1 is given as follows:

P = Pwgze

+PbwdZ

-z e )

(9-1) The fluid pressure p can be divided into two parts, a normal or hydrostatic pressure, p n , and a superposed anomalous pressure, p a as follows (Hubbert and Rubey, 1959): P =Pn +Pa

Hence, pa in this case can be shown as:

(9-2)

202

SHALE POROSITY log b

-

SHALE POROSITY log 8

-

x z,

sw

0

I Fig. 9-1. Schematic porosity-depth plot of shale. Fig. 9-2. Several dh/dZ values in a typical undercompacted shale section.

(h = 2 - Ze). Then, we obtain: (9-4)

dp,/dZ is the anomalous pressure gradient and is equal to the change in anomalous pressure (above hydrostatic pressure) corresponding to a change in depth. dh/dZ is the change in h corresponding to a change in depth. Now, let us discuss the values of dh/dZ in the typical undercompacted shales, shown in Fig. 9-2. Here, dh/dZ is zero at point 0, where the tangential line on the porosity curve in the undercompacted shales is parallel to the “normal porosity (normal compaction) trend’’ line. This is expressed as follows:

(2)o

=0

At point a, at which the tangential line is vertical, dh/dZ equals 1. This is similarly shown as:

203 Above point a , dh/dZ increases upward and its value is greater than 1,or:

Below point 0, dh/dZ has negative values and decreases downward. Somewhere below point 0, there must exist a point where dh/dZ equals -1 (point b in Fig. 9-2), or:

(g)b

= -1

Below point b, dh/dZ values are given as follows:

In the equations above, 0, a and b denote the points 0, a and b, respectively, and u and d denote respective points in the upward and downward fluidmovement zones in the undercompacted shales. As explained above, dh/dZ has a positive value above point 0 and a negative below. As the value of (pbw-pw)g in eq. 9-4 is always positive, the anomalous pressure gradient dpa/dZ has a positive value above point 0 and a negative below. Fluid would move upward above point 0 and downward below it. The amount of upward fluid movement, q u , crossing a unit area normal t o the flow direction in unit time is given by Darcy’s- equation: (9-5) where k , and pu are respectively the permeability of shale and the viscosity of water at point u in the upward zone. As stated above, dh/dZ at point a equals 1 (see Fig. 9-2). Hence, the amount of fluid movement qa at point a is given as follows:

Because q u is considered t o be greater than or equal t o qa in amount, and both have negative values, we obtain:

Hence:

204

As Fig. 9-2clearly shows, the porosity value at point a is maximum in the undercompacted shale zone, and decreases upward and downward from that point. Supposing that there is a function between the porosity and permeability of shale, and that the permeability decreases as the porosity decreases, the permeability value at point a also would be maximum in this zone. According to eq. 9-8,the permeability ratio ka/k, can be calculated, if the viscosities pa and p, of the formation fluid, and (dh/dZ), are known. It is possible to read the porosity values at these points. Hence, the integration of the permeability values based on eq. 9-8with the porosity values can be used to establish a relationship between shale porosity and permeability in the subsurface. The amount of fluid movement qd in the downward zone is expressed as:

(9-9) where d denotes the downward fluid-movement zone. As the value of dh/ dZ at point b equals -1, the amount of qb is similarly shown as:

(9-10) Because qd 2 qb, the following relationship would exist in the downward zone:

(9-11) In eq. 9-11,(dh/dZ)d always has negative values. The right-hand side of eq. > 0). 9-11is, therefore, positive Eq. 9-11 as well as eq. 9-8can be used to obtain the permeability ratio of shale in the subsurface. The actual calculations in the Cretaceous shales in western Canada will be discussed in the next section. Shale porosity-permeability relationship A porosity-permeability relationship based on the method described above, for Cretaceous shales in the Alberta and Saskatchewan subsurface will now be discussed. In this area, more than 160 wells (all with sonic logs) have been selected for study (see Fig. 9-3).In fourteen of these wells,

205

U. S. A.

Fig. 9-3. Distribution of undercompacted shales of Cretaceous formations in Alberta and Saskatchewan. Solid circles: wells with sonic logs; open circles: wells with both sonic and formationdensity logs; numbers: wells shown in Figs. 9-6 and 9-7.

both sonic and formation-density logs have been run. Because the formationdensity log provides both density and porosity values, it is possible to determine a relationship between porosity and acoustic transit time in these wells. The linear relationship between shale porosity (4) and transit time (At,) in the Cretaceous shales studied is expressed as follows:

4 = 0.00466Atsh - 0.317

(9-12)

By using eq. 9-12,the transit-time data from sonic logs can easily be converted to estimated porosity values. First, I determined the “normal porosity (normal compaction) trend” in this area. The “normal porosity trend” tends to shift to smaller values from east to west in the area (shale porosity on the “normal porosity trend’’ at the same depth decreases from east to west, although the slope of the “trend” is almost constant). This decreasing trend is attributed to the existence of a greater thickness of sediments in the west in the geological past (producing more compaction) and subsequent removal of a greater thickness of the uppermost part of the sedimentary column there (see Chapter 2).

206 I

1001

-x

c

al

c

2001

w

0

6

SHALE POROSITY,

6

Fig. 9-4. Normal porosity (normal compaction) trend of Cretaceous shales in Alberta and Saskatchewan.

After the erosion thickness has been compensated for, the standard “normal porosity trend” in this area is determined, as shown in Fig. 9-4. The function of the standard “normal porosity trend” of Cretaceous shales in the area studied is expressed as follows: @

= 0.62 e-0.000588Z(ft)

(9-13)

Several values of dhld.2 based on this normal trend (see eq. 9-4) are illustrated in Fig. 9-5. Now, let us read the dh/dZ values in the undercompacted shales in the

207 SHALE POROSITY, 0 0.,05 0:l 0:3 0

.

,

, j 6

1000

-

* %%

N

i c

;2000 A

5

F: w

>

300C

400(

Fig. 9-5. Chart for the determination of dh/dZ values.

Alberta and Saskatchewan subsurface (see Fig. 9-3).k-sr this purpose, 35 wells were chosen, some of which are shown in Figs. 9-6and 9-7.These wells are in the area where Cretaceous formations contain undercompacted shales (see Fig. 9-3). The viscosity of the formation water would change mainly with temperature. I assume that the average geothermal gradient in this area is 1.So F per 100 ft of depth, and that the subsurface temperature at 2000 f t is 75"F, as based on temperature-depth data in Alberta (Oil and Gas Conservation Board, 1967) (see Fig. 9-8).The viscosity of the formation water is determined at each depth (see Pirson, 1958,p. 234,table 3-17).The possible effect of a change in water viscosity with a change in distance from the clay surfaces (Fig. 8-4) is not considered in this calculation. The ratios k a / k , and kb/ kd are calculated by eqs. 9-8and 9-11 respectively, and the porosities &,, @, , $q, and q!)d at these points are read. Fig. 9-9shows the plots of k , / k , or kb/kd (logarithmic scale) against Ga 9, or &, - @d (arithmetic scale). According to Fig. 9-9,the porosity-per-

208 6

n

1

0.6

0.1

SEA LEVEL

~-

0.6

9

Or-rq

7

II

8

2

12

'

0.6

0.1

S H A L E POROSITY 0.1

0.6

Fig. 9-6. Examples of porosity distributions in undercompacted shales in Alberta wells. Locations of wells are shown as numbers in Fig. 9-3.

meability relationship of the shale is shown by the curved solid line and shadowed area. Archie (1950) has proposed a porosity-permeability relationship for sandstones, limestones, and muddy sands. According to this relationship, an 13 15

14

17

16

:qj

18

19

#'I

2 0.1

0.1

0.6 0.1

0.6

0.1

0.6

0.1 0.6 SHALE POROSITY

0.1

SEA LEVEL

0.6

0.

0.1

0.6

Fig. 9-7. Examples of porosity distributions of undercompacted shales in Alberta and Saskatchewan wells. Locations of wells are shown as numbers in Fig. 9-3.

209

Fig. 9-8. Temperature-depth relationship in the Alberta subsurface. (Data derived from Oil and Gas Conservation Board, 1967.)

increase in porosity of about 3%produces a tenfold increase in permeability. This can also be shown as follows:

where A is constant. The values k, and k, in the'previous discussion can be shown, according to Archie's relationship, as:

k,

=A

..10@/0-03

k, = A . 10@/0.03

(9-1 5) (9-16)

210 100

100

-2

P

:: 10

10

p

s

1

1 0

0.10

6, - &, or & - 6 d

0.20

0.30

Fig. 9-9. Relationship between permeability ratio (k,/k, or k b /kd ) and porosity difference (&, - $, or $b - $d) of Cretaceous shales in Alberta and Saskatchewan subsurface.

Hence, in the upward zone:

(9-17) A similar relationship in the downward fluid-movement zones,.based on the Archie relationship, is as follows:

(9-18) Eqs. 9-17 and 9-18 are shown as straight lines (Archie) in Fig. 9-9. According t o Fig. 9-9, the increase of shale permeability with increase in porosity is less than that given by the Archie relationship based on sandstones and limestones.

Comparison of logderived and laboratoryderived porosity-permeability relationships Data on the permeability-porosity relationship in shale are rather scarce. Bredehoeft and Hanshaw (1968) compiled several sets of data. The Geological Survey of Japan measured the permeability and porosity of the mudstone cores of several stratigraphic test wells in Japan. I obtained similar data on a well in the Strathmore gas field in Alberta. Data from Japan and Alberta are listed in Table 9-1. Fig. 9-10 shows a plot of all of the above-mentioned data. Archie’s porosity-permeability relationship mentioned above is shown as a straight broken line (Archie) in Fig. 9-10. Kozeny’s relationship of poros-

211 TABLE 9-1 Mudstone porosity and permeability data of well cores in Japan and Canada Well

Depth (m)

Lithology

Permeability (mD)

Porosity

Obuchi stratigraphic well (Japan)

1454 2036 2488 3049 3053 4033

9.9 .10-3 1.2 10-2 1.0 . 10-2 9.1 .10-3 8.6 * 0

0.257 0.203 0.150 0.127 0.152 0.065

Kambara GS-1 (Japan)

1029 1609 1808 2005 2151 2295 2608 3062 3206 3503 3701

mudstone mudstone mudstone mudstone mudstone mudstone mudstone mudstone mudstone mudstone sandy mudstone mudstone mudstone mudstone mudstone mudstone mudstone siltstone mudstone sandy mudstone mudstone mudstone mudstone sandy mudstone

-

0.390 0.332 0.266 0.311 0.246 0.243 0.218 0.196 0.188 0.159 0.146

Kambara GS-2 (Japan)

Yuza GS-1 (Japan)

Strathmore 7-12-25-25W4

1000 1255 1501 1763 2508 3500 4103

7.7 -10-3 9.9 .10-3 3.0. 7.3 .10-3 1.0.10-3 2.6 .10-3 2.0 10-4 1.3 -10-3 8.0 .10-4 4.9 10-2 1.4 * 1.4 * 2.3 . 6.0 10-3 7.0 .10-3 8.0 .

-

-

0.399 0.427 0.377 0.294 0.246 0.122 0.080

807

siltstone

2.1 . 10-2

1008 1198 1398 161.4 1966 2201 2402

siltstone mudstone mudstone mudstone mudstone mudstone mudstone

2.4 .10-3 8.4 10-3 1.2 10-3 1.3 . l o + 1.0-10-3 6.0 * lo4 5.0 1 0 - 4

2586 2816

mudstone mudstone

4.0 .10-4 1.0. l o 4

0.164 0.180

shale shale

1.6 -10-4 1.0 . 1 0 4

0.195 0.205

770 ft 777 f t

Permeant

'

v

.

10,000 ppm NaCl solution

10,000ppm NaCl solution

0.427 0.382

-

-

-

0.241 0.187 0.163 0.161 0.153

1

'

22,000 mg/l NaCl solution

212

I

' 10

I

O

T ~

l

l

laboratory-measured porosity-permeability data of shales, siltstones and clays in Canada, U . S . A . , and Japan..

.......... ..................... k Montmorillonite ..............rn Bentonite ..................... b Kaolinite

Fig. 9-10. Comparison of several porosity-permeability relationships. Solid circles show laboratory-measured porosity-permeability data of shales, siltstones and clays in Canada, the U.S.A., and Japan. k = kaoljnite, rn = montmorillonite, b = bentonite.

ity and permeability based on sandstones is also shown in Fig. 9-10as a curved broken line (Kozeny). The logderived porosity-permeability relationship in western Canada when the equal signs in eqs. 9-8 and 9-11are established is illustrated as a curved solid line (Cretaceous shale). The actual relationship could be shown by the shadowed area in Fig. 9-10.The starting point for these three lines (Archie, Kozeny, Cretaceous shales) in this figure is at the point of porosity 0.2 (or 20%) and permeability 3 mD. Although the data in Fig. 9-10are derived from several different areas, this figure suggests that the logderived porosity-permeability relationship is more applicable for shales than are the Archie and Kozeny relationships.

-

213 Volume of fluids expelled downward and upward The logderived porosity-permeability relationship of shale is shown by the shadowed area in Fig. 9-9. This means that the present study could not arrive at a single function between porosity and permeability of shales, but obtained some range between them. However, the relationship between:

in eqs. 9-8 and 9-11 and (Ga - &) or ($b - & ) can be expressed as a single curved line (Fig; 9-11). By using this figure and Fig. 9-5, slopes of the shale porosity curves at each (@a - &) or (& - & ) are determined. This means that the idealized shale porosity curves in the undercompacted shale zone can be constructed by Figs. 9-11 and 9-5. Fig. 9-12 shows such examples of the .constructed porosity distributions = 0.3 or = 0.5 or 50%, C: = 0.4 or 40%, D : (A: 4, = 0.6 or 60%, B: = 0.2 or 20%). These plots are similar to many of those pressure 30%, E: seals shown in Chapter 7. The values of pa/pu and C(b/C(d can be aSSUmed to be 1 in such a short interval. (pa = p u , or p b = p d ) . Superposition of the curves in Fig. 9-12 suggest that the porosity distribution at several stages of compaction, from one stage to another, must fit the volumes of fluid expelled upward and downward between these stages. The ratio q u / q d is expressed as follows (see eqs. 9-5 and 9-9):

(9-19)

b a - 0u, Or 6b - 0 d

Fig. 9-11. Relationship between (&/pu)[(dh/dZ)], or -(&/&)[(dh/dZ)]d Or

(@b- $ d ) .

and (@a

E

214

D

E

B

C

A

TP I *1

t

I

-! 0

e!

0

0 w

0

*1

4 ' 0 0

0

e!

? 0

0

4 ' 0 0

0

-! 0

e!

0

4

0

9

0

-!

c?

0

0

p ' 4 0

0

1

4I 0 c?

4 ' 0 0

0

SHALE POROSITY

Fig. 9-12. Constructed porosity distributions of undercompacted shales ( A : $a 0.5,C:$a = 0.4, D : $a = 0.3,E : @a = 0.2).

0.6,B :

$a =

Shale porosity

log

-

Q

Fig. 9-13. Schematic diagram showing the way or reading [(dh/dZ)], and [(dh/dZ)]d for calculating the ratio of the volumes of upward and downward fluid movement (qu/qd).

215 Supposing that (dh/dZ), and (dh/dZ)d values are taken at the points of the same porosity value in both upward and downward zones (Fig. 9-13),and the permeability values of the two shales are the same when their porosity values are the same, eq. 9-19is reduced to:

(9-20)

In other words, the ratio of the volumes of upward and downward fluidmovement can be determined by eq. 9-20. The values of qu/qd when the maximum porosity &, equals 0.6, 0.5, ..., 0.2 are determined from Fig. 9-12and eq. 9-20,and the average values are shown in Fig. 9-14(qu/qd+a relationship). Fig. 9-14shows a general tendency for qu/qd to decrease with any decrease in &. This means that the relative volume of downward fluid movement compared with upward increases with compaction of the shales. Fig. 9-15shows a superposition of the porosity curves, in which the porosity differences in the upward and downward fluid-movement zones between two stages of compaction fit these qu/qd ratios. The area between the two curves may show the volumes of fluids expelled from the unit shale

I

I

I

I

12

13

14

15

%/%I

Fig. 9-14. Relationship between the and average value of q u / q d , based on the constructed porosity distribution shown in Fig. 9-12.

T -rm SHALE POROSITY

c

AMOUNT OF WATER EXPULSION per tt

a

STAGES UPWARD

0

1 t: 8 0

1

DOWNWARD

30ft

I

22ft

C-D

I

I

D-E

40

31

STAGE A: (6, = 0 . 6

B: (6, = 0.5 C: Q E z 0 . 4 D: 8 , = 0 . 3 E: 6 , = 0 . 2

0

Fig. 9-15.Patterns of shale porosity distributions and amounts of water expulsion from shales at several stages of compaction.

column during these two stages of compaction. The area above line 0-0’ represents the volume of upward fluid expulsion, and the area below indicates the volume of downward expulsion. The volumes per unit vertical shale column, whose base area is 1f t 2 , are also shown in Fig. 9-15. References Archie, G.E., 1950.Introduction to petrophysics of reservoir rocks. Bull. Am. Assoc. Pet. Geol., 34: 943-961. Bredehoeft, J.D. and Hanshaw, B.B., 1968. On the maintenance of anomalous fluid pressures, I. Thick sedimentary sequences. Geol. SOC.Am. Bull., 79: 1097-1106. Hanshaw, B.B. and Bredehoeft, J.D., 1968. On the maintenance of anomalous fluid pressures, 11. Source layer at depth. Geol. SOC.Am. Bull., 79: 1107-1122. Hubbert, M.K. and Rubey, W.W., 1959.Role of fluid pressure in mechanics of overthrust faulting, I. Geol. SOC.Am. Bull., 70: 115-166. Oil and Gas Conservation Board, 1967.Pressure-depth and temperature-depth relationships, Alberta crude oil pools. OGCB-67-22, Calgary, Aka. Pirson, S J., 1958. Elements of Oil Reservoir Engineering. McGraw-Hill, New York, N.Y. Rubey, W.W. and Hubbert, M.K., 1959.Role of fluid pressure in mechanics of overthrust faulting, 11. Geol. SOC.Am. Bull., 70: 167-206.

Chapter 10

CHANGES IN SHALE PORE-WATER SALINITY DURING COMPACTION

This chapter consists of three main parts: (1)description of the salinity calculation method, (2) discussion of salinity changes during compaction, and (3) application of salinity data to exploration. The first deals with the methods of using well logs to calculate salinity in sedimentary rocks, and provides some verification of the salinity interpretation by chemical analysis. The second part discusses the effect of shale compaction on pore-water salinity (ion filtration by shale), and how to remove the compaction effect in order to identify depositional environments or post-depositional contamination. The last part describes several ways of applying salinity data to exploration. Problems associated with subsurface salinity are still in an area of scientific controversy. There are several explanations, in addition to ion filtration, for the generation of subsurface salinities greater than that of sea water: (1)evaporation of water, (2) diffusion of salt from nearby evaporite deposits, and (3) gravitional segregation (Dickey et al., 1972). The discussions in the second part of this chapter are, however, based on the premise that salinity increase, with depth or compaction, is due mainly to ion filtration by shales. Calculation of pore-water salinity from well-log data Pore-water salinity of shales

For calculating pore-water salinity in shales, the combination of a resistivity (or conductivity) log and any of the porosity logs (sonic, formation density, or neutron) may be used. The technique described in this chapter, however, uses the combination of resistivity (or conductivity) and sonic logs because they are the ones most available in wells and the least sensitive to holecaving. The relationship between the resistivity of the water-saturated rock, R , the formation-water resistivity, R , , and the formation-resistivity factor, F, is as follows (Schlumberger, 1972):

R

= FR,

(10-1)

According to Archie (1950), F is given as: (10-2)

218 where @ = porosity of the rock, a and m = coefficients, and m is called the "cementation factor". From eqs. 10-1 and 10-2, R, is given as:

R,=- mrn

(10-3)

a

The value R, can be calculated and expressed in terms of salinity (ppm NaCl) if the temperature of the rock is known. The values m and a of sandstones are usually determined by a series of laboratory measurements of the resistivity, R, of the cores whose porosity, @, POROSITY FROM SONIC LOG l p r cent) 1oD

CONDUCTIVITY FROM INDUCTIW LOG lmmho at l@F)

z

2000

2000

3000

3000

f 4000 ?

"-C'

f. 5000

5000

'I I

SALINITY -

~.

SANDSTONE

Fig. 10-1. Conductivity, porosity and salinity plots, Beaufort Basin.

r=*.+r

219 and pore-water resistivity, R,, are known. This procedure cannot, however, be applied to the shales, because saturating shale cores with water is quite difficult. Recent research work has revealed that comparison of the two normal compaction trends (conductivity and porosity) can solve the value m for shales. Fig. 10-1 shows such examples of the conductivity and porosity plots * of a well in the Beaufort Basin, Northwest Territories, Canada, on semilog paper. These plots are made by computer. From a study of the slopes of these normal trends in this well, the value m is estimated as follows: From eqs. 10-1 and 10-2, we obtain:

1R

@m

aR,

(10-4)

The normal compaction trend (porosity) of shales is usually shown as a straight line on semilog paper (porosity, logarithmic scale; depth, arithmetic scale). An example in the Beaufort is shown in the centre of Fig. 10-1. The normal porosity trend is expressed in general form as: @ = @o e-CZ

(10-5)

where @ = shale porosity at depth 2, @o = shale porosity at the surface, e = base of the Napierian logarithms, and c =constant.

The value -c is the slope of the trend line on a graphic plot. Introducing eq. 10-5 into eq. 10-4, we obtain: (10-6) where I = conductivity of the rocks. Assuming that R , is constant in the zone studied, the value @T/aR, is constant and equals the surface conductivity value extrapolated from the subsurface normal trend (conductivity) in the right-hand side of Fig. 10-1. The value -em is the slope of the normal conductivity trend in this figure.

* The shale porosity has been estimated from the sonic log by using the following empirical relationship between the sonic transit time At(ps/ft) and porosity in the Beaufort Basin : @

=

0.00374At - 0.206

220

Determining the slopes of porosity and conductivity (-c and -cm),we can calculate the m value for shales (m = -cm/-c). The average m in the Beaufort is found t o be about 2. This value has been checked in many other wells in the Beaufort. By using the result of the water salinity (which can be converted to R,) from chemical analysis of cutting samples and DST water samples, the value a can be estimated as follows (see eq. 10-3):

(10-7) because the values R, (P and m are known from the logs. As a result of comparing log data with water analyses, the value a has been found t o be about 1 for most shale samples from Gulf Coast and Beaufort wells. However, the m and a values in other areas may be different from these, and must be checked individually. Examples of the calculated cementation factor m for the Miocene mudstones in the Nagaoka Plain, Japan, are shown in Fig. 10-2. These m values are generally lower than those in the Beaufort. With the m and a values in the Beaufort, the water resistivity, R,, has

m

500

0

2000

30

w)

11

m

1.8

Fig. 10-2. Relationship between cementation factor, m, and depth of the Shiunji gas field, Japan. (From Magara, 1968.)

221 been calculated from the resistivity of the rock, R , derived from the conductivity log, and the porosity, 4, from the sonic log by eq. 10-3in this well. From the temperature-depth relationship * of the well, the value R , has been converted to water salinity in ppm NaCl. The computer plot of the salinity is shown in the left-hand side of Fig. 10-1.The pore-water salinity of several cutting samples is also plotted in this figure, along with the DST water salinity in the sandstone reservoir below about 5900 ft. It is of interest to note that the salinity in the overlying shales is comparable with that in the sandstone. Pore-water salinity of sandstone If the section contains a thick and clean sandstone saturated with formation water, its pore-water salinity can be calculated from the SP log. The SP value in this case can be approximated as

(10-8) where T = temperature (OF), R m f = resistivity of drilling-mud filtrate (ohmm), and R , = resistivity of formation water (ohm-m). Therefore R , can be calculated as follows:

R,

=

Rmf 1oSP/[0.11(460+T)]

(10-9)

Fig. 10-3shows an example of a combination plot of sandstone and shale salinities in a Gulf Coast well. The computer calculates and plots the sandstone salinity using SP when the SP value is greater than 15 mV (relatively sandy). Relatively low calculated salinity values in this sandy interval are probably due t o the effect on the log values of clays or shales in the sandstones (see the composition plot on the left). But these are not real values, so the highest salinity value in the interval would probably be the best estimate. The salinity data derived from the chemical analysis of cutting samples are also shown in this figure for comparison. When SP is equal to or less than 15 mV (relatively shaly), the computer calculates and plots the salinity based on the inductionsonic combination. Shale pore-water salinities calculated in this way have been verified by the chemical analysis of cuttings, as shown. The water salinity in the abnormal-pressure zone below 9300 ft in this

* T = 17 + 0.0152 where T = temperature (OF)and Z = depth (ft).

222 WAlER SALINIIY I VVM 1

LEGEND

aSHALE OR CLAY

FROM SP LOG

0POROSITY FROM SONIC LOG

.

4 DIREClION OF FLUID MIGRATION PRESSURE IN SHALE FROM SONIC LOG

> SALINITY FROM SP LOG ( IN SANDSTONE )

> SALINITY FROM INDUCTION k SONIC LOGS ( IN SHALE ) FROM CUlTING SAMPLES flUlD VRtSSURt FROM SONIC LOG IPS11

Fig. 10-3. Sand-shale salinity plot and composition-fluid pressure plot, Gulf Coast well A.

well (see the fluid pressure on the left) is lower than that above. The difference reveals the effect of compaction on salinity change (low salinity in the undercompacted shale zone). This problem will be discussed in the following sections. Water salinity change during compaction

Ion filtration b y clays or shales According to several experiments by Engelhardt and Gaida (1963)and Kruykhow and others (1962, see Hedberg, 1967), pore solutions exuded from lowexchange-capacity sediments (for example, kaolinite and silt) at moderate pressures show no change in concentration. Sediments of a higher exchange capacity (for example, bentonite and montmorillonite), however, exude a solution of decreasing concentration with increasing pressure. In these circumstances, the concentration of the remaining pore water will increase. Overton and Timko (1969)demonstrated the possible effect of ion filtration in the subsurface shales of the Gulf Coast area. Figs. 10-4and 10-5,

223 PORE-WATER SALINITY (1000ppm) FROM S P L O G

SHALE POROSITY 1

.2

.3

.4

Fig. 10-4. Pore-water salinity-depth and shale porosity-depth relationships in a Gulf Coast well. (Redrawn from Overton and Timko, 1969.)

P O R E - W A T E R S A L I N I T Y [lo00 ppm) FROM S P L O G

SHALE POROSITY

05

I

.2

.3

.4

.5

Fig. 10-5. Pore-water salinity-depth and shale porosity-depth relationships in a Gulf Coast well. (Redrawn from Overton and Timko, 1969.)

224 redrawn from the figures in their paper, show sandstone salinity (from the SP log)-depth and shale porosity+epth relationships on semilog paper. Overton and Timko assumed that the salinity in the sandstones is in equilibrium with that in the nearby shales *; these figures show that in normally compacted zones shale porosity and salinity are reciprocal. They stated “ . . . as the shale is compressed to one-half its original pore fraction, the water is pressed out leaving the salt behind to concentrate itself by a factor of two.” The relationship between the shale porosity, $, and salinity, C,, is given in their paper as:

$C,

= constant

(10-10)

One of their conclusions derived from the observation above is “. . .that no salt is lost from squeezed shales and clays. Only water emerges, and shales appear to be perfect ionic filters.” They tried to use this concept t o explain the presence of the “drinking water” in the shallow subsurface of their study area. (The “drinking water” in the subsurface would be the water expelled from shales.) Overton and Timko’s demonstration of a reciprocal relationship between shale porosity and salinity on the basis of the actual subsurface data is important, but the last statement, “no salt is lost . .” must be examined carefully. Fig. 10-6shows examples on semilog paper of salinity and porosity plots derived from the Beaufort wells in which the normal salinity trends are almost reciprocal to the normal porosity trends (i.e., analogous to Overton and Timko’s plots of Fig. 10-4and 10-5). As mentioned in the previous section, the general form of the normal porosity trend is expressed as:

.

$ = Go e-cz

(10-5)

Because the normal salinity trend line is reciprocal to the normal porosity line, the normal salinity can be expressed in mathematical form as:

C,

=C ,

ecz

(10-11)

* Schmidt (1973) showed that the pore-water salinity of normally compacted shales derived from chemical analysis of side-wall cores is less than that of the adjacent sandstones derived from the SP log. This may suggest that the salinities in sandstones and shales are not in equilibrium. However, because of the fact that both salinity values are obtained by different methods, we must pay special care to the interpretation of the Schmidt result; side-wall shale cores could have been swelled by the effect of low-salinity drilling fluids, and the salinity of the shales could have been reduced. Note also that the salinity data shown in Fig. 10-3suggest that the salinity in the shales is comparable with that in the nearby sandstones.

WELL NO 4

WELL NO 5 SALlN ITY (ppm .)

>0."

POROSITY(%I

0" .

!

& . oo

O .,O

SALINITY (ppm) 1p

190

POROSITY(%)

00

>0".

>oooo >ooo

1

19

190 I

3000

3000

I

I

I

I

1

I

I

-

3000

-t

I

CUTTINGS 3000 SALlN ITY

1 I

I

I

i

I

4000

4000

1

I

LOG CALCULATED SALINITY 5000

i 5000

J

5000

I

I

'I

5000

5 . LOG CALCULATED

Fig. 10-6. Salinity-porosity comparison, Beaufort Basin.

.

DST WATER SA LI NlTY SANDSTONE

SANDSTONE

226

where C, = salinity at depth (after compaction), and C,, = salinity extrapolated at the surface (before compaction). Note that the value c has a negative sign in eq. 10-5 indicating that porosity is decreasing with depth, and a positive sign in eq. 10-11, indicating that salinity is increasing with depth. Multiplying eq. 10-5 by eq. 10-11,we obtain:

$Cw = $oC,,

= constant

(10-12)

This equation is the same as eq. 10-10proposed by Overton and Timko, but may have more applications. The symbols $0 and C,, are the shale porosity and salinity at the surface, extrapolated from the subsurface normal trends. If these normal trends are drawn through a marine-shale sequence of relatively homogeneous composition, they may be considered also as the trend lines showing the history of change of the porosity and salinity of the marine shale with burial and compaction. Therefore, Go and C,, may be considered as the shale porosity and salinity when the shales were deposited (before compaction), in this case in a marine environment. In addition, if the water salinity is expressed in mg/l, $oC,, or $Cw is the weight of salt (mg) in a unit volume of shales *. Because in the low-to-intermediate salinity range, the difference between the salinity values expressed by mg/l and ppm is almost negligible, this concept can also be used when the salinity is expressed in ppm. In any case, eq. 10-12 means that the amount of salt in a unit volume of shales is the same before and after compaction. Eq. 10-12 can also be explained in the following manner. The salinity of shales increases with decreasing porosity, but the product of the salinity and porosity at depths (after compaction) is the same as their product when the shales were deposited (before compaction). The product ($C,) is thus unchanged compaction, and should hence have a unique value for the environment in which the shales were deposited. By plotting the $Cw product instead of the salinity, the effect of compaction can be removed and information about the depositional environment may be revealed. This problem will be discussed in the next section. As shown in Figs. 10-4 and 10-5, the concentration of the pore solution due to compaction is obvious. However, does this mean that the fluids expelled from the shales are completely fresh, as stated by Overton and Timko? To solve this problem, we must determine how much fluid and how much salt have been lost from the shales through compaction.

* Suppose we have 1 1 of shale with the following porosity and salinity values: q5 = 0.1 (or 10%) C, = 10,000 mg/l

In this case, q5Cw is equal to 1000 mg, or 1 1 of shale contains 1000 mg of salt.

227 As discussed in Chapter 6,the volume of fluids W that have been expelled from the shales during compaction is given as:

$0-4 w = v, - v = v---

1-@0

(10-13)

The amount of salt, S,*, in shales before compaction, Vo, is given as:

so = 4 o c w o vo

(10-14)

The amount of salt, S, in shales after compaction, V, is shown similarly:

s = 4cwv

(10-15)

The amount of salt, S1,that has been lost from the shales during compaction equals the difference in amount of salt in the shales before and after compaction, or: s1 =

so - s = doCw0Vo- @CWV #0-@

= @C,V---

1-40

(10-16)

Therefore, the salinity of the expelled fluids, Cwl, is given as:

(10-17) (see eqs. 10-13,10-16 and 10-12). Let us calculate the salinity of the expelled fluid by using Overton and Timko’s examples in the Gulf Coast (Figs. 10-4and 10-5).According to Fig. 10-4,the extrapolated value of shale porosity at the surface is about 0.37. The corresponding salinity at the surface is about 32,000ppm. Hence, we obtain: Cwl = 0.37 X 32,000+ 12,000ppm

In the same way, using the example shown in Fig. 10-5,we can calculate the salinity of the expelled fluids as follows: Cwl = 0.43 X 29,000+ 12,500ppm

* Note that the amount of salt is not the salinity (Cwo or Cw).

228

The results of the calculation above show that throughout compaction the salinity of the expelled fluids was about one-third that of original sea water. Therefore, the expelled fluids should not be fresh. Overton and Timko’s mistake in this problem results from their having neglected a very important consideration: one cubic foot of shale at depth (after compaction) was more than one cubic foot when it was deposited (before compaction). Therefore, if (as stated in their paper) no salt is lost during compaction, a cubic foot of shale at deeper depths must contain more salt than one at shallower depths. The actual observation from their plots is that the amount of salt per unit volume of shale at deeper depths is the same as that at shallower depths. Some salt, therefore, has been lost from the Gulf Coast shales during compaction. The “drinking water” in the subsurface in this area cannot be explained simply as a result of the fluids expelled from shales during compaction. Although not a perfect ion filter, shale is nevertheless a good one, and the same ion filtration observed by several compaction experiments seems to have occurred in the subsurface shales, too. The salinity of the expelled fluids discussed above has posed an interesting problem. If the expelled fluids were completely fresh, how much sal. PORE-WATER SALINITY 11000pprnJ FROM S P LOG

SHALE POROSITY

LINE A-PERFECT I O N FILTRATION LINE B - N O I O N FILTRATION

Fig. 10-7. Left-hand drawing shows imaginary salinity trends when perfect ion filtration ( A ) and no ion filtration ( B ) occur, superimposed on the actual salinity plot by Overton and Timko (1969). Shale porosity (Overton and Timko) is shown on the right.

inity would be expected in the subsurface? If the salinity of the expelled fluids were the same as that of the original sea water, what type of salinity distribution would be expected in the subsurface? In order to answer these questions, Fig. 10-7was made, in which the two salinity lines ( A and B) for the two cases mentioned above have been added as broken lines to Overton and Timko’s actual salinity plot for the Gulf Coast area. It shows that, if the expelled fluids are completely fresh, the salinity increase should be greater than was actually observed in the subsurface (perfect ion filtration, see line A in Fig. 10-7).If the expelled fluids have the same salinity as the original, the subsurface salinity is the same as the original, or no salinity change due to compaction will be observed (no ion filtration, see line B, Fig. 10-7).We may conclude that the slope of the salinity plot relative to the slope of the porosity plot is an indicator of the efficiency of ion filtration by shales; higher salt concentrations can be expected in shales of high filtration efficiency than in those of low efficiency. The efficiency of ion filtration may be related to shale composition; there would be more ion concentration for the same porosity in a montmorillonite-rich zone than in an illite- or kaolinite-rich zone, because montmorillonite has a higher exchange capacity, i.e., is a better ion filter, than other clay minerals.

Time of the first occurrence o f ion filtration Figs. 10-4and 10-5show that the surface salinity extrapolated from the subsurface-salinity trend in marine shales is about 32,000 or 29,000ppm, or close to present sea-water salinity. But such is not always the case. The extrapolated salinity from rocks older than Pliocene age is usually lower than present sea-water salinity. Overton and Timko have already noted this problem: Fig. 10-8shows their plot of the extrapolated surface salinity * from various formations of different ages versus geological times, in which the older rocks show a lower extrapolated surface salinity. They explained this by ion diffusion; longer geological time has caused more diffusion of salt ions to reduce pore-water salinity. This diffusion would probably shift the entire salinity profile toward a lower value, so that the extrapolated surfacesalinity values would also be lower. This explanation seems quite possible. However, I can suggest another possible solution. If the sediments were deposited under marine conditions and filtration occurred immediately afterward, a Gulf Coast-type salinity plot might be expected (extrapolated surface salinity close to that of sea water). If filtration did not occur until the shales reached a certain depth of burial or compaction level, the salinity profile shown in Fig. 10-9might be observed.

* Expressed in this figure as “ocean water salinity, ppm NaCl after diffusion”.

230

\

PIioceneO\

Recent I

lo6 10

5

10’

5

Id

5

,

w4

iL.5

lo5

Ocean woter salinity, ppm NoCl after diffusion

Fig. 10-8. Extrapolated surface salinity (= ocean-water salinity after diffusion) and geological age relationship. (After Overton and Timko, 1969.)

Before the shale reached depth D, no filtration occurred; i.e., the porewater salinity was the same as the original (see also line B in Fig. 10-7). Actually we might not be able t o see such constant salinity in the relatively shallow interval of a real salinity plot, because the salinity data for such shallow intervals are not always available, or there are sometimes no marine sediments at such shallow depths. Below depth D, filtration would start and the salinity would increase with compaction or depth. If we extrapolate from the deeper portion of the salinity trend, we will obtain a lower surface-salinity value than that of sea water (Fig. 10-9). The critical depth or compaction for the beginning of ion filtration may also depend on shale composition; montmorillonite may start filtration earlier, i.e., at a shallower burial depth, than illite and kaolinite. In summary, the fact that the surface salinity extrapolated from subsurface marine sediments older than Pliocene is lower than present sea-water salinity, may be explained by (1)diffusion in the geological past, as proposed by Overton and Timko, or (2)the combined result of diffusion and delayed onset of ion filtration, possibly an effect of shale composition. In line with the above reasoning, it might be noted that the surface salinity extrapolated from Cretaceous sediments (probably younger than Tuscaloosa) in the Beaufort wells shown in Fig. 10-5is about 12,000ppm (see also Fig. 10-8).

231 SALINITY EXTRAPOLATED SALINITY AT SURF ACE

\

-SEA

+

DEPTH

~

WATER SALINITY

ION.FILTRATION BEGINS

I

Fig. 10-9. Schematic salinity profile.

“Salinity X porosity ” p l o t

The relationship between the salinity and porosity of shales (eq. 10-10, = constant) arrived at by Overton and Timko is based on the logderived values of salinity and porosity in the Gulf Coast. The same concept seems to apply also to most of the Beaufort sediments. However, we do not know whether it can be applied to shales from all parts of the world. In order to solve this problem, the salinity and porosity of shales must be studied on a worldwide basis - a large-scale research project. While I will not attempt to solve that problem in this chapter, I will try to show some examples of the salinity-porosity relationship from other basins. Hedberg (1967) studied the pore-water chlorinities of shales on the basis of chemical analysis of cores and samples, and measured pore space on the basis of total bulk volume and grain volume. Assuming that the pore spaces had been occupied by water in the subsurface, he then calculated the salinity of the formation water. The shale porosity can also be determined from the ratio of the pore volume over the total bulk volume. Fig. 10-10 shows several examples of the chloride content (ppm) versus porosity in the Burgan field in Kuwait, and for several oil fields in Texas. The relationship between the chlorinity * and porosity in Fig. 10-10may be approximated

4C,

* Salinity (ppm NaCI) may be calculated by multiplying the chlorinity by 1.65.

232 a 0

BURGAN (Kuwait)

0

0 THOMPSON

ITexasl

X ANAHUAC ITexal

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

V V

v

V

x X

b

0

HEBERT II 1 ( T s x a l

.

0 .

01

a

20

10

a

00

50 CHLORIDE

ppm

10

en

w

.

im

110

0

Ilhouundl

Fig. 10-10. Shale porosity-chlorinity relationship, Kuwait and Texas. (Data from Hedberg, 1967.)

x

0

.

X

... " .

0 0

0

0

0

0

0

x x

0 0.

0

x 1 . x

. *

0

1

.

.

WELL A

o WELL B

GULF COASl

I WELL C I

10

20

30

a

sa

60

CHLORIDE ppm

10

80

IThouundl

Fig. 10-11. Shale porosity-chlorinity relationship, Gulf Coast.

90

rm

110

233 by a hyperbola. The chlorinity increases as the porosity decreases (or $Cw = constant). Therefore, eq. 10-10 seems to be applicable to these examples as well. Fig. 10-11 shows other examples from Gulf Coast wells, based on chemical analysis of cutting samples using the same technique as Hedberg. The chlorinity is plotted against porosity and a similar relationship is indicated, although in this case more scattering of the points was observed. This result could be explained by the presence of nonmarine sediments in the section or, perhaps, by fresh-water contamination through permeable beds before any significant compaction and ion filtration started. Such contamination could dilute the water in the sediments before compaction. If this is the case, the chlorinity (or salinity) after compaction could also be reduced. Application of water salinity data to exploration Figs. 10-1 and 10-3 show good agreement of logderived salinity and chemical analysis. Further documentation is provided in Fig. 10-12 (Beaufort), and Figs. 10-13 and 10-14 (Gulf Coast). These figures show that reliable pore-water salinity data for clastic sections can be obtained from the interpretation of well logs (SP, resistivity or conductivity, and sonic). In this section are described several uses of the calculated salinity data in exploration. Proximity t o bedded salt deposits

Fig. 10-15 shows a salinity plot and a combination “composition-fluid pressure” plot * of a Canadian Arctic Island well. This well was drilled close to one of several diapiric structures as revealed by seismic surveys. It was not known whether the structures were salt domes or shale diapirs. Water recovered from a test of the sandstone below 7500 f t had a salinity of 220,000 ppm NaCl, which is confirmed by the SP log in the sandstone (see “log calculated sand salinity” in Fig. 10-15).By using the concept of “salinity X porosity” mentioned previously, the maximum possible salinity from shale compaction and ion filtration was calculated to be about 150,000 ppm NaC1. This shale salinity value was confirmed in the overlying shales (see “log calculated shale salinity” in Fig. 10-15). This result suggests that water in that aquifer is abnormally saline. The high probability that structures in the area are saltcored further suggests that this abnormal salinity is due to salt solution and consequent charging of the aquifers with concentrated brine **. The shales, which are almost impervious, appear to have a

* The composition (quartz and clay) is calculated from gamma-ray log. ** Recently, a second well, drilled high on a nearby structure, has confirmed the presence of salt core.

WELL N0.3

WELL N0.2

WELL NO.l

SALINITY

*P &P i I

...e LOG :.., CALCULATED

300(

i.:'

.:.( : y LOG CALCULATED . .(. f.. SALINITY

.i

1'.

LOG CALCULATED .:*SALINITY

:.

SALINITY

.

I.'..

400(

'i.

I

t

I ;.

x

..

'

!

CUTTINGS

, . ' .I/SALINITY KWM

'.. I:

I

\CUTTINGS

1 .

*-DST

SALINITY 1 I

-I

6400

CUTTINGS SALINITY

I

SANDSTONE

+- DST WATER SALINITY Fig. 10-12. Salinity plots, Beaufort Basin.

DOLOMITE

DEVONIAN SHALE

WATER SALINITY

I--.

WAIEP SALINllV IWMl

VOLUME FRACTION FROM SP

r

l

.

yt ;In r

'

. . . :.

r

v

9200-

--

9400 I

9600-

Y

n 9800 -

LEGEND 10,000-

10,200-

p I:

.

I. 0

2wo

6ooo

1o.m

QUARTZ

2

4 10,400 -

10,600

>

1

a SHALE OR CLAY

FROM SP LOG

0POROSITY FROM SONIC LOG t DIRECTION OF FLUID MIGRATION

.

PRESSURE IN SHALE FROM SONIC LOG

> SALINITY FROM SP LOG ( IN SANDSTONE ) > SALINITY FROM INDUCTION & SONIC LOGS ( I N SHALE )

-

I SALINITY FROM CUTTING SAMPLES

10,800I

FLUID PRESSURE FROM SONIC LOG (PSI)

Fig. 10-13. Sand-shale salinity plot and composition-fluid pressure plot, Gulf Coast well B.

LEGEND QUARTZ SHALE OR CLAY

1

FROM SP LOG

POROSITY FROM SONIC LOG DIRECTION OF FLUID MIGRATION PRESSURE IN SHALE FROM SONIC LOG SALINITY FROM SP LOG ( IN SANDSTONE ) SALINITY FROM INDUCTION 6 SONIC LOGS ( I N SHALE ) SALINITY FROM CUTTING SAMPLES

-

FLUID PRESSURE FROM SONIC LOG IPS11

Fig. 10-14. Sandsh ale salinity plot and composition-fluid pressure plot, Gulf Coast well C.

236 FLUID PRESSURE (PSI) 0

0

0

0

,3,MUD

...,.. . ......

.- .

Fig. 10-15. Sand-shale salinity plot and composition-fluid tic Island well.

WEIGHT (LB/GAL)

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

pressure plot, Canadian Arc-

normal salinity, and have not been charged by salt diapirs. Therefore, the combination plot of the sand and shale salinities may indicate whether a well is in an area of salt diapirism or not. Although it is not so in this example, the salinity in shales adjacent to the high-salinity permeable rocks could also be affected. A b normal pressure

Since presentday shale salinity depends on porosity, as shown in Figs. 10-10 and 10-11, salinity is a useful tool for indicating undercompacted shales with which abnormal pressures are usually associated; shale salinity is abnormally low in the undercompacted shales or the abnormal-pressure zone. This problem is discussed by Overton and Timko (Figs. 10-4and 10-5). Fig. 10-3 shows an example of the association of abnormal pressure with abnormally low salinity in shales below 9300 ft. Also, in the well shown in Fig. 10-14, there is an abnormal-pressure zone below about 9000 f t and the shale salinity decreases gradually with depth.

237

Fault zone Overton and Timko stated that a fault zone can also be detected by subnormal salinity. By way of explanation, they suggested stress relief along the fault zone. Another explanation might be fresh-water contamination through the fault planes. I have not been able to obtain such an example of salinity decrease at fault zones in Canadian wells.

Permeable sandstone o f significant extent At about 7500 f t in the well shown in Fig. 10-14, the fluid pressure in shales drops toward sandstones, suggesting that the fluids in the shales have been drained into the sandstones. As the fluids are expelled, the shales are compacted and the pore-fluid pressure in them decreases. A t the same time, the salinity of the shale pore water increases. If the sandstones are permeable and of large areal extent, the fluids in adjacent shales easily drain into them and the amount of fluid loss increases toward the sandstones. In other words, the salinity increases gradually in the direction of the sandstones. Such increases are seen at about 7550 f t in Fig. 10-14, and at about 9300 ft and 9800 f t of the well in Fig. 10-13. If the sandstones have low permeability or are of small areal extent, the fluids from the shales will not be expelled into them properly, and little or no increase in compaction or salinity toward the sandstones will be observed. Therefore, in the mixed and undercompacted shale zones, a gradual increase in salinity toward the sandstones may suggest that the latter have high permeability and large areal extent.

Depositional environments Fig. 10-16 shows a comparison of a calculated salinity plot ( A ) and a palynology log of a Beaufort well. Below about 1800 ft, the salinity is conspicuously higher than it is at shallow depths. The palynology log shows that marine species are correspondingly abundant below the 1800-ft level. As discussed in the previous section, within sediments of the same depositional environment, salinity will increase with compaction as a result of ion filtration. The increasing salinity trend in Fig. 10-16A might, therefore, be attributed to ion filtration. To remove the compaction effect from the salinity, a “porosity X salinity plot” can be made, as shown in Fig. 10-16B. Because the ‘‘4 X C,” product has a relatively constant value for the same depositional environment, it is possible to detect gross environmental differences from this plot; the marine sediments have a significantly higher product value than the nonmarine. Fig. 10-17 shows a similar plot from another well. Increases in salinity ( A ) and the product of porosity and salinity ( B ) below about 3400 f t seem to

Fig. 10-16. Comparison of salinity plot and palynology log of a Beaufort well.

239

WELL N o . 2 PALYNOLOGY (A) WELLN0.2 SALlNlTV WM

-Y ?. 200c

3000

3000

t

E x

400C 4000

. . DOLOMITE

500C

SI

Fig. 10-17. Comparison of salinity plot and palynology log of a Beaufort well.

correspond with the abundance of marine species in the palynology log in the same zone. A salinity profile of seven Beaufort wells is shown in Fig. 10-18.Note that the log-calculated salinity in shales is very close to the salinity of the water samples from sandstones by DST. Zones containing predominantly marine species taken from the palynology logs are also shown. In summary, the salinity plot, or the compaction-corrected salinity (#C,) plot may provide useful information about depositional environments. Fresh-water contamination Suppose we made a shale salinity plot, and observed abnormally low salinity values in a generally high-salinity zone. From the salinity plot alone, we

WELL NO. 5 SALINITY, ppm

.

_-_-._

4

EEL

I

o o o

SAND-

P P P

WELL NO. 3 SALINITY, pprn

WELL NO. 6

.

-.-

METAMORPHIC ROCKS I

=

SAND STONE SALINITY OF WATER FROM DST

M A R I N E SPECIES PALYNOLOGY L O G

Fig. 10-18. Water salinity in shales, Beaufort Basin.

SAND STONE UK

=

SAND STONE

UPPER CRETACEOUS FOR MAT IONS

SAND STONE

DOLOMITE

LK

=

IMPERIAL FORMATION 4-

LOWER CRETACEOUS FORMATIONS

241 do not know whether these low salinity values are due to depositional environment, undercompaction, or fresh-water contamination after deposition. If we convert the salinity plot t o a “salinity X porosity” plot (9 X Cw), thereby removing the compaction effect, and the low values persist, they may be explained by depositional environment or fresh-water contamination. If the paleontological data show no variation in environment (or show it to be all marine), the low-salinity zone must be explained by fresh-watercontamination at a later stage of burial. For example, in Fig. 10-16A, low salinity is indicated in the 3000-3300-ft zone *. According t o the palynology log, marine species are abundant in this interval. However, the “porosity X salinity’’ plot (Fig. 10-16B) also shows low values in this zone (i.e.y after compaction effect has been removed). Therefore, this low salinity is most likely due t o fresh-water contamination, probably related to the presence of an unconformity at about 3300 ft. A similar interpretation may be made in the low-salinity zone of 43004600 f t in Fig. 10-17. There is an unconformity in this well at about 4500

ft. If the contamination is through permeable sandstone, the abnormally low salinity may be observed in the sandstones. This is the reverse of the case shown in Fig. 10-15, where the sandstone is charged by concentrated brines from nearby salt deposits and consequently has an abnormally high salinity. References Archie, G.E., 1950. Introduction to petrophysics of reservoir rocks. Bull. A m . Assoc. Pet. Geol., 34: 943-961. Burst, J.F., 1969. Diagenesis of Gulf Coast clayey sediments and its possible relation to petroleum migration. Bull A m . Assoc. Pet. Geol., 53: 73-93. Dickey, P.A., Collins, A.G. and Fajardo, M.I., 1972. Chemical composition of deep formation waters in southwestern Louisiana. Bull. A m . Assoc. Pet. Geol., 56: 1530-

1570. Engelhardt, W.V. and Gaida, K.H., 1963.Concentration changes of pore solutions during compaction of clay sediments. J. Sediment. Petrol., 33: 919-930. Hedberg, W.H., 1967. Pore-Water Chlorinities o f Subsurface Shales. Univ. Microfilms, Ann Arbor, Mich. (Thesis, Univ. Wisconsin). Magara, K., 1968. Compaction and migration of fluids in Miocene mudstone, Nagaoka Plain, Japan. Bull. A m . Assoc. Pet. Geol., 52: 2466-2501. Magara, K.,1974. Compaction, ion-filtration and osmosis in shales and their significance in primary migration. Bull. A m . Assoc. Pet. Geol., 58: 283-290. Overton, H.L. and Timko, D.J., 1969. The salinity principle - a tectonic stress indicator in marine sands. Log Anal., 10: 34-43.

* An important freshening mechanism, in addition to those mentioned above, could be montmorillonite dehydration (Powers, 1967 and Burst, 1969). This effect in the Gulf Coast area is discussed by Schmidt (1973)and Magara (1974).However, in the wells shown in Figs. 10-16and 10-17,montmorillonite dehydration has apparently not taken place (there is montmorillonite still remaining in the relatively deep sections).

Powers, M.C., 1967. Fluid-release mechanisms in compacting marine mudrocks and their importance in oil exploration. Bull. Am. Assoc. Pet. Geol., 51: 1240-1254. Schlumberger, 197%.Log Interpretation 1. Principles. Schlumberger, New York, N.Y., 113 pp. Schmidt, G.W., 1973. Interstitial water composition and geochemistry of deep Gulf Coast shales and sandstones. Bull. Am. Assoc. Pet. Geol., 57: 321-337.

Chapter 11 IMPORTANCE OF ABNORMAL PRESSURING IN SHALE DIAPIRISM

In explaining the mechanism of salt or shale diapirism, the so-called buoyancy concept is often used. Salt and diapiric shales both have lower densities than the rocks that surround them, so that they tend t o float in the heavier rocks. In the case of salt diapirism, this buoyancy effect associated with certain differential loading of overlying sedimentary rocks can be sufficient t o cause the salt t o move upward, because salt can flow or yield relatively easily. In the case of shale diapirism the mobility of the diapiric material (shales) is less easily explained. For one thing, the density contrast between the diapiric shales and the surrounding sedimentary rocks is usually less than in the case of salt diapirism. If clays or shales s t a r t to move at near-surface depth almost immediately after deposition, the mobility presents no problem. This type of clay or shale movement must be distinguished from typical shale diapirism, where deep-seated shales move upward for a long distance t o form a diapiric structure. In the case of the near-surface clay or shale movement, these mobile materials may have stayed at relatively shallow depths all the time, while surrounding sediments continued to be buried deeper. Therefore, this type of structure may be called a pseudodiapir. The density of pseudodiapiric materials would be quite low, because they have never been deeply buried; they are also relatively soft and mobile. On the other hand, in the real shale diapirs, the shales were once relatively deeply buried. Most of these shales were undercompacted, so that their densities tend to be subnormal; usually they are in the range from 2.1 t o 2.3 g/cc, as compared with other normally compacted shales at similar depths of 2.2-2.5 g/cc. Fig. 11-1shows typical diapiric shales and normal sediments in the Gulf Coast. The typical diapiric shales are compacted t o the extent that normal sediments would be at 7000-8000 f t burial depth. In other words, they are relatively well compacted. If some differential stress is applied t o these shales for a long geological period, they may be deformed by “yield” or “flow” to some extent, but the degree of deformation in this case would be much less than in the previous case of unconsolidated pseudodiapiric material. In this chapter, I intend to offer another explanation of the upward movement of relatively deep-seated and well-compacted shales to the one based on yield or flow under differential stress. It depends on the extremely high

Dens ty - g/cc 3

5

L

u-

87J

5 a 12-

S

c

B 16C

201

I

\

Fig. 11-1. Assumed density contrast for salt, undercompacted shale and normal section. (From Musgrave and Hicks, 1968.)

fluid pressure that can be generated by the aquathermal effect. Hedberg (1974)recently suggested the importance of methane generation in shale diapirism. I too believe it can contribute t o overpressuring and diapirism. However, quantifying this effect is not an easy task. The important factors in the generation and growth of diapirs in the subsurface can be summarized as follows: (1)The underlying layer must have a certain degree of mobility. (2) There must be at least a break or a fracture in the overlying sediments to enable deep-seated diapiric movement. Pressure in the diapiric shales should overcome the internal friction of the overlying sediments. (3) Differential loading of the overlying sediments may also be a contributing factor. (4)Buoyancy too may contribute, but this effect becomes relatively important only when the diapir has reached a sufficient height. This chapter will discuss the importance of the four points mentioned above in interpreting the mechanism of shale diapriism. For generating mobility in shales, restricted fluid expulsion from the shales under the influence of an aquathermal-pressuring effect is very important, as will be shown below. Abnormal pore pressures and their significance to shale mobility Shale compaction model without the aquathermal-pressuring effect Fig. 11-2 shows a schematic diagram of the compaction history of a shale at depth D. When the shale was deposited under water, it had an initial

245 A SHALE POROSITY

B

-

SHALE PORE-PRESSURE

IPS11

I

k

C Lu

1

c COMPACTION TERMINATION DEPTH

De

Dl

D

D

Fig. 11-2. Schematic diagrams showing compaction and pore-pressure histories of a shale now at depth D where there is no aquathermal-pressuringeffect.

porosity &,. During the early stages of compaction, the fluids in the shale were expelled normally, so that it reached the equilibrium of compaction after each small increment of loading. Fluid pressures during these stages were near hydrostatic. The porosity at each of these compaction stages can be approximated by an exponential function which relates it to burial depth (see Rubey and Hubbert, 1959). This “normal compaction trend” is shown as a straight line on semilog paper (Fig. 11-2A). Hubbert and Rubey’s (1959)equation for the effective stress a in a fluidsaturated porous rock under overburden is:

a=s-p

(11-1)

where S is the total overburden load and p is the fluid pressure. During these early stages of normal compaction, the fluid pressure is almost hydrostatic. In the Gulf Coast area the hydrostatic-pressure gradient is commonly taken as 0.465 psi/ft. The overburden-pressure gradient for the same area, on the other hand, may be approximated as 1 psi/ft (Dickinson, 1953).Therefore, during these early stages the effective stress increased at a rate of about 0.535 psi/ft (1-0.465). The compressional strength or grain-to-grain bearing strength of the shale also increased at approximately this rate, because the shale had almost zero strength at the time of deposition, and the compres-

246 sional strength increases as the effective stress increases. Suppose that fluid expulsion from this shale stopped at depth D,, so that during subsequent burial there was no further compaction or, in other words, no change of the shale strength afterward. The pore pressure in this shale when it was buried t o D, was near hydrostatic. The rate of increase in pore pressure since this burial stage would be about 1 psi/ft (or overburden-pressure gradient), because the effective stress or compressional strength in this shale is unchanged during successive burial, but the overburden load increases at about 1 psi/ft. In this case the increased overburden load after burial to D , is entirely supported by the pore fluids. The history of porepressure change is schematically shown by the heavy line with an arrow in Fig. ll-2B.The fluid pressure line between D, and D is parallel to the overburden-pressure line. The effective stress or compressional strength of the shale increased t o its maximum value of 0.535 X D , at depth D,. The same stress or strength was maintained during burial from D , t o D, because there was no further compaction in this model. The possible effect of thermal expansion of water on pore pressure is here ignored. An important outcome of the above discussions is that with this first model the pore pressure in the undercompacted shale would never reach or exceed the overburden load, because sediments usually compacted normally at early stages of burial, and the restricted fluid expulsion began at a later stage. The value of the effective stress or compressional strength remained the same throughout burial from D , t o D. Hubbert and Rubey (1959)stated that slippage along any internal plane in a rock should occur when the shear stress along the plane reaches a critical value. Using the Mohr-Coulomb law, they expressed this as: 7 = 70 + u

tan 4

(11-2)

where u is the normal stress (which is the same as that given in eq. 11-1), ro the shear strength of the rock, when u is zero, and 4 the angle of internal friction, which has an average value of about 30” for a wide variety of rocks (Hubbert and Rubey, 1959). Hubbert and Rubey further stated that “once a fracture is started, ro is eliminated” and therefore a simpler equation can be used: 7=

u tan4 = ( S - p ) tan4

(11-3)

If the compaction termination depth shown as D , on this model (Fig. 11-2)is, for example, 8000 f t (about an average figure for the Gulf Coast overpressured shales according to Magara, 1975), the effective stress is given as:

+

uEoOoft = 0.535 X 8000 4300 psi

247 Then, the critical shear stress for sliding after fracturing would be as follows (see eq. 11-3): 7

+ 4300 X tan 30” = 2500 psi

This simple calculation shows that, even in the relatively deep sections in the overpressured shales, a stress of about 2500 psi is necessary to cause slippage in the rocks. This means that under the conditions of this first model the overpressured shales would not be completely mobile even after generation of an initial fracture. The shales can, however, move very slowly by yielding even in this case. Shale compaction model with the aquathermal-pressuringeffect Barker (1972) discussed the aquathermal-pressuring effect in shales, using the temperature-pressure-density diagram for water. Magara (1975) applied this concept to the overpressured shales in the Gulf Coast, using Hottman and Johnson’s (1965) shale compaction-pressure data. He stated that “if complete isolation of pore water developed in the subsurface, pore pressure thereafter would increase as a result of aquathermal effect, at a gradient of 1.8 psi/ft, provided the average geothermal gradient is 25”C/km (or 1.37”F/ 100 ft). This gradient is almost twice the average overburden-pressure gradient in the Gulf Coast area.” The actual subsurface data in this area, however, leads to the conclusion that “the average net pressure gradient since isolation is about 1.4 psi/ft; this lower gradient suggests some leak of pore water from the undercompacted shale (i.e., isolation was not perfect)” (Magara, 1975). In Fig. 11-3, the left-hand diagram (A) of the schematic shale porositydepth relationship is the same as that in Fig. ll-2A, but the rate of porepressure increase since isolation at D , in the right-hand diagram (Fig. 11-3B) is higher than that in the previous figure; by taking into account the actual aquathermal effect in the Gulf Coast shales, the rate becomes 1.4 psi/ft. So, in this case, the rate of increase of the pore pressure from D , on is greater than that of the overburden pressure. Therefore, if burial continues, the pore pressure in this shale should eventually equal or even exceed the overburden pressure. If the isolation D , (compaction termination depth) is 8000 ft, the pore pressure could reach the overburden-pressure level at a depth of about 20,000 feet (Magara, 1975). A t this depth, the effective stress becomes zero (eq. 11-1).Hence, the critical shear stress for sliding after initiation of a fracture is also zero (eq. 11-3).In other words, once a fracture is generated, the shales at 20,000 ft or deeper could become completely mobile. The depth at ‘which shales become mobile would vary with their termination depth (D,),the geothermal gradient, and the effectiveness of the retention of the generated presssure. Once generated, this unstable condition does not

248 A SHALE POROSITY

B

-

SHALE PORE-PRESSURE

-

LOGARITHMIC SCALE

I

r

L YI 0

D

-COMPACTION TERMINATION DEPTH

De

D

Fig.11-3. Schematic diagrams showing compaction and pore-pressure histories of a shale now at depth D,where there is an aquathermal-pressuring effect.

terminate until the shales lose their fluids or pressure to a certain critical level. Another interesting point related t o Fig. 11-3 is that the effective stress and compressional strength of the shale increase together at a rate of 0.535 psi/ft to burial depth D,, but below D , the effective stress decreases with burial eventually to zero, while the compressional strength stays at the same level as that at D , (maximum strength). In other words, stress and strength are not identical in the interval D,-D of the second model. We may conclude that complete elimination of internal friction in the deep-seated shales is possible if pore fluids are isolated t o the degree in the undercompacted shales of the Gulf Coast, and if the temperature increases afterward. If, under these circumstances, an initial fracture is generated, the shales would move relatively easily. Movement of rocks overlying diapiric shales Suppose that the deep-seated shales are mobile to the degree at which a diapir could start. Their pore pressures are near lithostatic. These shales are mobile at least within this deep and super-pressured zone. In order to have a shale diapir, however, the overlying rocks would have to move t o some

249

Fig. 11-4. Seismic illustration showing differences between fault systems formed by differential compaction and gravity slide. Dashed white line shows the configuration of shale masses. (From Bruce, 1973.)

extent with the mobile shales. Such movement would be controlled by the internal friction of the overlying rocks. The generation and progress of contemporaneous faults were recently discussed by Bruce (1973).He stated that “these contemporaneous fault systems may be formed either by differential compaction or gravitational sliding.” This suggests that these fractures and faults are closely related to sedimentation and compaction processes themselves. In other words, there would be no need for strong, regional tectonic forces to explain these faults and fractures in this area. Examples of the faults are taken from Bruce’s paper and shown in Fig. 11-4. Assume that fractures or faults due to differential compaction and gravitational sliding occur in the sediments above mobile super-pressured shales. If the force in these deep-seated mobile shales exceeds the frictional resistance in the overlying sediments, a piercement could occur. The left-hand diagram in Fig. 11-5 shows a schematic porosity-depth relationship of shales at present. Note that in Figs. 11-2 and 11-3 the compaction history of the shale at D (present) is illustrated, while in Fig. 11-5 the present porositydepth relationship of the entire shale sequence is shown. The shales in the shallower interval (above A ) are normally compacted, having the hydrostatic pore pressure indicated in the right-hand diagram. The shale at B is under-

250 S H A I F PORF-PRFSSLIRF SHALE PORE-PRESSURE

IPS1 I

-

\

\

Fig. 11-5. Schematic shale porosity-depth and pore-pressuredepth relationships at present in an area where there is an abrupt lithology change from underlying shales to overlying interbedded sandstones and shales.

compacted and overpressured. The normal compaction and normal fluid expulsion of this shale stopped at a relatively early stage of burial, while the overlying shales continued to compact normally. Because of the aquatherma1 pressuring effect after the termination of compaction, the shale at B now has pore pressures equal to the overburden pressure. This means that at depth B pore fluids are carrying the entire overburden load. To produce a diapir in this situation, the pressure in the shale at B must become high enough to overcome the frictional resistance in the overlying sediments. The maximum internal friction in the sediments is at the maximum compaction or maximum effective stress horizon (A). If the depth of this horizon is about 20,000 f t , the critical shear stress to cause slippage can be calculated by using eq. 11-3as: 720.000ft = 0.535 X

20,000 X tan 30"

+ 6200 psi So, in this model a pressure of about 6200 psi in excess of the overburden pressure is necessary to create a diapir. In other words, it is rather unlikely. This type of relatively thick normal compaction interval can develop in an

SHALE POROSITY

I

A

-

\

SHALE PORE-PRESSURE IPS11

-

251

A

&

0

I B

Fig. 11-6. Schematic shale porosity-depth and pore-pressure-depth relationships at present in an area where massive shales continued to be deposited.

area where the zone of sandstoneshale interbed is relatively thick. Below the interbedded zone, pore pressure increases abruptly. Fig. 11-6 shows another shale porosity-depth model in which all the shale beds below depth A have terminated normal compaction and normal fluid expulsion at A. Pore pressures below depth A increase continuously to reach the overburden pressure at B. Note again that Fig. 11-6shows the present shale porosity-depth relationship of the shale sequence. It is different from Figs. 11-2and 11-3,which show the compaction history of a specific shale. A shale porosity-depth relationship such as shown in Fig. 11-6 might develop if there was continuous deposition of shales. There would have been restricted fluid expulsion in all shales below the compaction level at depth A. The fluid pressure in this deep zone would increase at a rate of about 1.4 psi/ft because of the aquathermal-pressuringeffect. Suppose that the pore pressure at depth B is the same as the overburden pressure. As stated previously, if depth A is 8000 ft, depth B would be approximately 20,000 f t for the average undercompacted shale in the Gulf Coast (Magma, 1975).The shale at depth B is mobile. If this mobile shale moves upward only a minute distance, as indicated by A 2 in Fig. 11-6,and the fluid pressure gradient inside the shale is near hy-

252 drostatic *, a small excessive pressure would exist over the overburden pressure at that new level. Generation of pressures slightly more than the overburden pressure is possible, because rocks usually have some tensile strength. The excessive pressure Ap in this case is given as:

Ap = (1 - 0.465)X A 2 = 0.535 X AZ psi The critical stress AT for slippage after a fracture initiation at this point can be calculated by using eq. 11-3as:

AT = A D t a n 4 = (1.4- 1) X AZ X tan

30"

+ 0.231 X AZ psi This model shows that, once a shale reaches the depth at which the pore pressure is equal to the overburden and the shale becomes mobile, the shale immediately above could also move upward quite easily. Note that in the model shown in Fig. 11-5 the frictional resistance of the normally compacted shales overlying the super-pressured shale is very high, so that movement of the overlying rocks is not easy. If the mobile shale continues to rise to higher levels, the frictional resistance increases (Fig. 11-6).As mentioned above, however, the excess of pore pressure over overburden pressure increases at an even higher rate, provided that the pore-pressure gradient inside the mobile shale is near hydrostatic. It might not be. If the diapir were already high, but if the diapir were growing, the shale would tend to expand with a decrease of confining pressure, and its temperature would decrease. These two factors would reduce the pore pressure in the diapiric shale. The rate of temperature decrease may be related to the rate of diapiric growth; if the growth is fast, the temperature may not drop very much. In this case, the shale would retain most of its high pressure, facilitating more growth. Because the excessive pore pressure Ap in the rising shale is more than twice the frictional resistance in the second model (Fig. 11-6), it seems plausible that, even with a certain pore-pressure drop due to some temperature decline and shale volume expansion, such mobile shale would most likely be able to pierce through overlying sediments relatively easily. The most resistant sediments to movement would exist at A. because at

* Pressure in this mobile shale is much higher than that of surrounding rocks, but pressure gradient within the mobile shale is assumed to be near hydrostatic. This assumption is not too unrealistic, because near-vertical pressure charging is possible under these lowinternal-friction or slippery conditions within the shale.

253 this level the internal friction is greatest, and this point is remote from level B where the shale diapir started (Fig. 11-6). Moreover, by the time the top of the moving shale has reached depth A, the cooling effect may have become significant. Once depth A is passed, however, the diapirism may progress relatively easily, because above this depth the internal friction of the strata declines. At those shallow depths, the buoyancy effect may also contribute to movement, provided the shale is still mobile. Diapiric movement would terminate as soon as the shale loses its mobility. So far I have discussed the effect of internal friction of the sediments overlying a diapir: the sharp contact of normal- and abnormal-pressure zones (Fig. 11-5), and the continuous development of abnormal pressure below a certain depth (Fig. 11-6). The first condition may develop where there is an abrupt change in lithology, e.g., upper sandstoneshale interbed and lower massive shale. The second may exist where shale deposition is relatively thick, and any sandstone interbeds in the shallower zone are relatively thin. The actual subsurface shales may not be the same as is assumed in the two situations discussed. It is more likely they will fall somewhere between these two extremes *. Nevertheless, the examples show the importance of differing internal friction in the overlying sediments for the occurrence of diapiric movement. Bruce (1973) recently proposed an explanation for the mechanism of the development of contemporaneous faults under an environment of progradational sedimentation, using his fig. 2. The same figure is adopted and shown in Fig. 11-7 of this chapter **. In the centre of Fig. 11-7D a thick shale mass is shown associated with several faults in the layers above it. On both sides of this shale mass, sandstoneshale interbeds extend from the surface to relatively great depths. Below these interbeds are more shales. The possible shale porosity profiles for these environments are shown in Fig. 11-7E. The porosity-depth profile shown in Fig. 11-6 may be the one to fit in the central part of the diagram, where the shales are thick and massive. The profile shown in Fig. 11-5 may be used at both sides, where the sandstoneshale interbeds reach t o depth. By the same logic as applied in Figs. 11-5and 11-6, chances of shale diapirism are greater in the central zone. The development of contemporaneous faults over such an area could be an important factor in initiating diapirism - a possibility also suggested by Bruce (1973). The presence of sandstones, which are usually more rigid than the average shale, would also prevent the occurrence of diapirism off the shale mass.

* In the examples in Figs. 11-5 and 11-6, it is assumed that there are shales at depth and sandstoneshale interbeds at shallower depths, simulating the typical Gulf Coast examples. If, however, there are thick and permeable zones at the deep subsurface, the profile might be completely different. ** See Bruce's (1973) paper for an explanation of contemporaneous fault generation using this schematic diagram.

254

Fig. 11-7. Diagrammatic illustration showing four stages in the development of a residual shaie mass. (From Bruce, 1973.) Possible porosity-depth relationships are added at the bottom ( E ) of the diagram.

Significance of differential loading If, under the geological and physical environments as illustrated in Figs. 11-7D and l l - 7 E , more pressure is applied in the off shale-mass areas than in the shale-mass area, the chances of generating diapirism will be increased. Such would be the case in the actual subsurface, because there are thicker and more compacted sediments outside the shale-mass area. Fig. 11-8 shows the overburden-pressuredepth relationships corresponding to the two typical cases shown in Figs. 11-5 and 11-6, based on Dickinson’s (1953) density-depth relationship for the Gulf Coast area. The solid curve in Fig. 11-8

255 OVERBURDEN PRESSURE, PSI 1ow0 2woo

I

I

I

--

I

30000

I

I

PRESSURE BASE0 ON DICKINSON'S I19531 DENSITY-DEPTH RELATIONSHIP IN GULF COAST PRESSURE WHERE COMPACTION HAS TERMINATED AT B O W FT

Fig. 11-8. Overburden pressure-depth relationships in the Gulf Coast area.

shows the overburden pressure versus depth relationship where Dickinson's densitydepth trend is applied. This may simulate the case in the off shalemass areas of Fig. 11-7. Where shale compaction has terminated at a depth of about 8000 ft, the rate of increase of overburden pressure below 8000 f t would be lower than in the previous case; this is shown as a broken straight line in Fig. 11-8.This second case may be applicable for the shale-mass area shown in Fig. 11-7. The overburden-pressure difference at 20,000 f t between off-shale and onshale-mass areas is about 500 psi. The difference increases with depth. Forces by differential loading would push the mobile shale from the outside t o the shale-mass area. Buoyancy effect

Once a diapir grows upward, the density difference between the inside and the outside of the diapir could generate a buoyancy effect. Gussow (1968)

256 estimated such pressure differences due to salt diapirism, based on Dickinson’s (1953) density-depth relationship and an average salt density of 2.2 g/cc. According t o GUSSOW, the maximum pressure difference of 2500 psi is attained at a depth of about 3000 f t , t o which salt grew from 28,000 ft. The density contrast in the case of shale diapirism would, however, not be as much as that for salt, because of a lower density contrast in the‘ case of shale, so that the buoyancy effect for shale diapirs should be smaller. As stated previously, buoyancy becomes effective in the later stages of diapirism, but is not important in the earlier stages. References Barker, C., 1972.Aquathermal pressuring - role of temperature in development of abnormal-pressure zones. Bull. Am. Assoc. Pet. Geol., 56: 2068-2071. Bruce, C.H., 1973. Pressured shale and related sediment deformation: mechanism for development of regional contemporaneous faults. Bull. Am. Assoc. Pet. Geol., 57 :

878-886. Dickinson, G., 1953. Geological aspects of abnormal reservoir pressures in Gulf Coast Louisiana. Bull. Am. Assoc. Pet. Geol., 37: 410-432. Gussow, W.C., 1968. Salt diapirism: importance of temperature, and energy source of emplacement. Mem. Am. Assoc. Pet. Geol., 8: Diapirism and Diapirs, pp. 16-52. Hedberg, H.D., 1974. Relation of methane generation to undercompacted shales, shale diapirs, and mud volcanoes. Bull. Am. Assoc. Pet. Geol., 58: 661-673. Hottman, C.E. and Johnson, R.K., 1965. Estimation of formation pressures from logderived shale properties. J. Pet. Technol., 17: 717-722. Hubbert, M.K. and Rubey, W.W., 1959. Role of fluid pressure in mechanics of overthrust faulting. I. Geol. SOC.Am. Bull., 70: 115-166. Magara, K., 1975. Importance of aquathermal pressuring effect in the Gulf Coast. Bull. Am. Assoc. Pet. Geol., 59: 2037-2045. Musgrave, A.W. and Hicks, W.G., 1968. Outlining shale masses by geophysical methods. Mem. Am. Assoc. Pet. Geol., 8. Diapirism and Diapirs, pp. 122-136. Rubey, W.W. and Hubbert, M.K., 1959. Role of fluid pressure in mechanics of overthrust faulting, 11. Geol. SOC.Am. Bull., 7 0 : 167-206.

Chapter 12 ESTIMATION OF OIL-GENESIS STAGE

In the preceding chapters, the important factors for hydrocarbon accumulation, such as structural timing, drainage (which is largely controlled by reservoir development), sealing and fluid migration have been discussed from the standpoint of a shale compaction study. One important factor that has not so far been touched on is the source and the state of maturation of the hydrocarbons. The generation of hydrocarbons is a matter for a geochemical text book and I have no intention of discussing these details here. However, the maturation of petroleum, which is controlled by a time-temperature function (Connan, 1974), will be covered in this book. We learned in Chapter 2 how t o estimate the maximum burial depth of a given bed; this information will help us evaluate more realistically the maximum temperature the bed reached. Time-temperature relationship for hydrocarbon generation Most petroleum hydrocarbons have been generated by a subsurface thermal process from the organic matter deposited along with fine-grained clastic sediments. Connan (1974) documented a time-tempetarture function by

CODE 1. A M A Z O N B A S I N , B R A Z I L 2. P A R I S B A S I N , F R A N C E 3. A O U I T A I N E B A S I N , F R A N C E 4. EL A A l U N A R E A , RIO D E O R O , W . A F R I C A

5. D O U A L A B A S I N , C A M E R O O N 6. O F F S H O R E T A R A N A K I B A S I N , N E W Z E A L A N D

7. C A M A R G U E B A S I N , F R A N C E

8. O F F S H O R E T A R A N A K I B A S I N , N E W Z E A L A N D 9 . LOS A N G E L E S B A S I N , C A L I F O R N I A

10. V E N T U R A B A S I N . C A L I F O R N I A 11. A O U I T A I N E B A S I N , F R A N C E

TEMPERATURE

Fig. 12-1. Time-temperature relationship of petroleum genesis. (From Hunt, 1974.)

258

00

!OO

Y

t’ Y

I

Y

y Y

$00

L

>

v)

VI 3

00

00

Fig. 12-2. Oil-genesis chart 1 .

examining accumulated oils of known geologic ages and the maximum (in this case, present) temperatures. Solid circles in Fig. 12-1 show the data he obtained. The plot in this figure suggests that the younger the rocks, the higher the temperature needed to generate hydrocarbons in them. Using Connan’s data, Hunt (1974) proposed possible oil-generation and gas-generation zones as shown in Fig. 12-1. To evaluate the status of oil genesis from Fig. 12-1, one must know the present temperatures of the formations involved, or calculate them from the depths, geothermal gradient and surface temperature. Such calculations, if many formations or depth points are being evaluated, are time-consuming. To simplify the process, Hunt’s data from Fig. 12-1 have been replotted as a function of geological time and, depth on five oil-genesis charts, each with a different geothermal gradient and incorporating a selection of surface

259

100 O

4 1

' I

200

300

I

I

500

400

"I

600

700

M V

' l J L 1 J . l . L r - - . 100

200

300

Y

c 3

e P Y

5 400

5 i 3

Y)

m

Y) 3

500

600

20

40

----

60'F

SURFACE TEMPERATURE

OIL GENESIS CHART GEOTHERMAL GRADIENT

2

I.S*F/IOO FT

I

60

I

700

SURF4CE TEMPERATURE

Fig. 12-3.Oil-genesis chart 2.

temperatures (Figs. 12-2,12-3,12-4,12-5 and 12-6).These oil-genesis charts are valid only for areas where the burial rate was relatively uniform.

Chart description Five oil-genesis charts (Figs. 12-2,12-3,12-4,12-5 and 12-6)show the relationships of geological time and depth for five geothermal gradients ranging from 1"F/100 ft to 3"F/100f t in 0.5"F/100f t increments. A stratigraphic scale is included, based on Holmes' geological time scale (1960 version) taken from Wetherill (1966,p. 518). Each chart was constructed for a surface temperature of 60°F;however, additional scales are provided t o accommodate surface temperatures ranging from 20°F t o 80°F. The subsurface temperature at a given depth (corre-

26 0

- 100 - 200 - 300 w

-400

f

w I w

-500

2 237

2

-600

*

-700

- 800

20

40

60°F

SURFACE T E M P E R 4 l U R E

OIL GENESIS C H A R T GEOTHERMAL GRADIENT

3

2.F1100FT

7 60 BO'F

&SO0

SURFICE T E M P E R 4 l U R E

Fig. 12-4.Oil-genesis chart 3.

sponding to a specified geothermal gradient and surface temperature) is provided by the scale at the extreme right-hand side of each chart. Use of charts Select the appropriate chart based on the known or predicted geothermal gradient. Plot the geological times (or horizons) versus the depths, using the depth scale corresponding to the known or estimated surface temperature. By connecting these points, the time-depth relationship for the area in question is obtained. An example is shown by points A to C on oil-genesis chart 1 (Fig. 12-2, for a geothermal gradient of 1"F/100 f t and a surface temperature of 60°F).

261

0

MY

\ I 100

>I L

TlON 200

0

300

\ 0

400 Y

-L

Y

100

20

5 5 Y

\

I

g

600

" Y

i (L

700 0 '3

30

800

\ 30

900

1000

10

20

40

60'F

OIL GENESIS CHART

SURFACE TEMPERATURE

GEOTHERMAL GRADIENT

4

2.5'F/IOOFT

\ 40 I 60 I F SURFACE TEMPE TURE

1100

Fig. 12-5. Oil-genesis chart 4.

Points A , B and C represent the following horizons and depths: Point

Stratigraphic horizon

Depth (ft)

A

Top of Cretaceous Top of Jurassic Top of Triassic

7000 17,500 25,000

B C

~

____ ~

_

_

_

_

For this example, the main oil and gas zone is the Cretaceous, with the Jurassic and older horizons containing gas. These oil-genesis charts should be used only for areas where the burial rate was relatively uniform. If there have been significant variations in the rate of burial and/or erosion, the charts may result in erroneous estimates of oil genesis.

262

20

40

6O'F

SURFACE TEMPER4TURE

OIL GENESIS CHART GEOTHERMAL GRADIENT

5

J*F/IOOFT

. -

60

80-F

SURF4CE TEMPERLTURE

Fig. 12-6. Oil-genesis chart 5.

Comparison of oil-genesis chart and world oil and gas reserves The above-mentioned oil-genesis charts can be documented by the occurrences of oil and gas in the world. For this comparison the data used were the reserve figures (ultimate recoverable oil and gas) of the world giant oil (>500 million bbl) and gas (>3.5 trillion cubic feet) fields listed in tables 1and 2 of Geology of Giant Petroleum Fields (Halbouty, 1970). The reserve figures were grouped on the basis of geological age and summed within each age group. The original data are derived from about 270 giant fields throughout the world. The subsurface temperatures or geothermal gradients for these fields are not readily available, but a gradient of 1.5"F/100 f t may be assumed as an average (see, for example, fig. 6 of Some Fundamentals of Petroleum Geology, Hobson, 1954).

263

MILLION Y E A R S 100

200

300

400

500

600

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0

10.000

I-

w

W LL

=- 20,ooc

I-

n

w

a

’ ?,

-4-7 GAS

/)

0/

PHASE OUT

Q.0’

/’

30,000

NO GAS

1’

I I 40,000

SURFACE TEMPERATURE 60’ F GEOTHERMAL GRADIENT l . 5 ° F / 1 0 0 FEET

LEGEND

]FROM

&GAS

GIANT FIELDS

-

0

50 B I L L I O N BBLS

0

3 M ) l R l L L l O N BBLS

Fig. 12-7. Comparison of Hunt’s oilgenesis stages with the reserves of the world’s giant oil and gas fields.

With this assumption the reserves are plotted on the oil-genesis chart of this geothermal gradient (1.5”F/100ft). Fig. 12-7 shows such a plot; it is interesting to note that the oil reserves drop sharply below the “oil phase-out”1ine.

264

Even below this line there is the chance of a large gas reserve. As mentioned earlier, the assumption of 1.5"F/100f t for all of these giant fields may not be reasonable. The burial histories of many of these fields may have been relatively complicated and, therefore, plotting geological time versus depth on a selected chart may produce an oversimplified result. Some of the relatively shallow occurrences of oil and gas in the Tertiary rocks (see Fig. 12-7) may, for example, be due to the effects of late-stage uplift and erosion, but most of the detailed geological information for these fields is not readily available. References Connan, J., 1974. Time-temperature relations in oil genesis. Bull. A m . Assoc. Pet. Geol., 58: 2516-2521. Halbouty, M.T., 1970. Geology of a Giant PetroIeum Field. AAPG, Tulsa, Okla., 575 pp. Hobson, G.D., 1954. Some Fundamentals of Petroleum Geology Oxford Univ. Press, London, 139 pp. Hunt, J.M., 1974. How deep can we find economic oil and gas accumulations? SPE 51 77, 1974 Deep Drilling and Production Symp., Prepr., pp. 103-110. Wetherill, G.W., 1966. Radioactive decay constants and energies: Section 23 in Handbook of Physical Constants. Geol. SOC.A m . Mem., 97: 513-519.

Chapter 13

ESTIMATION OF PALEOPORE PRESSURE AND PALEOTEMPERATURE

Chapter 12 discussed the importance of subsurface temperature in the generation of petroleum. If deposition and burial were relatively uniform and continuous, the oil-genesis stage can be estimated from the geological time and present temperature of a given bed by using one of the charts included in that chapter. However, if the area being studied underwent significant erosion in the geological past, a bed’s present temperature will not be the maximum temperature that bed ever attained. In such a case one easy way to obtain the possible maximum temperature is to estimate the temperature that would have been reached at the time of maximum burial. This estimate can be made if we know the amount of erosion (see Chapter 2) and if we assume the geothermal gradient and surface temperature when maximum burial was reached. The validity of such an estimate can be examined by a technique based on inclusion thermometry, which this chapter will discuss. In applying this technique, it is necessary first to estimate the paleopore pressure. Estimation of paleopore pressure For this discussion, let us use an example in the western Canada Foothills. Fig. 13-1,which is the same as Fig. 2-11,is a transit-time-depth plot of the 5 well, which shows undercompaction Pacific Amoco Ricinus 16-29-34-8-W below about 3000 ft. Within this general area, the Cardium sandstones are known to be overpressured at many locations. However, these abnormal pressures are probably lower now than they were in the geological past. The significant uplift and erosion that took place in the area would have caused a marked decline in pore pressure, as a result of the decrease in temperature and slight expansion of the pore spaces. The pore pressure in these undercompacted zones before erosion, can be estimated from the known amount of erosion, as discussed in Chapter 2,and the subsurface transit-time data. In the well shown in Fig. 13-1,for example, the Cardium sandstone is at a depth of about 6200 ft. The estimated thickness of erosion is about 4600 f t (see Chapter 2). In other words, at one time, before the area was eroded, the Cardium was buried to a depth of 10,800ft. The shales above and below the Cardium are undercompacted, having a transit-time value of about 76 &ft. This level of compaction is equivalent t o that in the normal compaction zone at the present depth of about 3700 ft. The maximum burial depth of this equivalent level, therefore, is about 8300 ft.

266 SHALE TRANSIT TIME (ps/FT) 50

100

116

200

300

400

0

2500

4000

-+ u-

I

k W

0000

0

8000

10000

Fig. 13-1. Shale transit-time-depth plot of the Pacific Amoco Ricinus 16-29-34-8-W5 well. (From Magara, 1976.)

From these observations it is possible to make the following burial-compaction model. During the initial burial, from the surface to approximately 8300 ft, compaction was normal. Fluid expulsion during this stage was also normal. At 8300 f t the pore fluids were locked in completely, so that from there to the maximum burial depth of 10,800 f t there was no further compaction. At a later stage the area was uplifted and the uppermost sections removed by erosion, leaving the depths of the Cardium at about 6200 ft. One may ask, what would happen if the generation of these undercompacted shales were “gradual,” rather than “abrupt” as such a model of normal compaction-po compaction assumes? In that case fluid expulsion and shale compaction might have been continuous during burial, at rates that decreased gradually with time. In many cases the “gradual” model would be quite realistic. However, the “abrupt” model is simpler and more convenient t o use for estimating paleopore pressures. Yet the results obtained by the two methods are not significantly different. Therefore, the “abrupt” model will be used in this chapter. The pore pressures in the Cardium during the early stages of burial to

267 8300 f t would have been near normal or hydrostatic, because the compaction was almost normal. This means that the pore pressure at a depth of 8300 f t would have been about 3600 psi, assuming a hydrostatic-pressure gradient of 0.44psi/ft. The rate of increase of pore pressure after this critical stage (or the onset of the restricted fluid expulsion) was discussed by Magara (1975).If there is no aquathermal-pressuring effect during burial, the pore pressure would increase at the same rate as that of the overburden pressure, or at about 1 psi/ft. However, this is not the case in most sedimentary basins because the temperature also should increase during burial, resulting in a much higher rate of pressure increase. Magara (1975),in a study of the Gulf Coast shales, estimated the rate of pore-pressure increase since isolation of pore fluids to be about 1.4 psi/ft. Applying this result to the shale in the well being studied, an increased pressure of about 3500 psi [ 1.4 X (10,8008300) psi] above the hydrostatic pressure at the isolation depth (8300 ft) can be calculated. This means that the Cardium would have been overpressured to about 7100 psi (3600 + 3500 psi) when the maximum burial depth of 10,800f t was reached. The paleopore pressure, the maximum burial depth, and the pressuredepth relationship of the Cardium in four wells in western Canada was calculated using the preceding method. The results (Table 13-1)indicate the possible existence of significantly high pressures in the Cardium in the geological past. The previously mentioned calculation is based on actual shale compaction data in wesMrn Canada and knowledge of the aquathermalpressuring effect obtained in the Gulf Coast area. The validity of applying the aquathermal-pressuring effect observed in the Gulf Coast to western Canada may be a matter of discussion. It is not possible to study such a phenomenon directly in western Canada because of the effect of late-stage disturbance on pore pressure in this area. We can, however, examine the possibility of such a phenomenon on the basis of what is known of the area’s thermal and compaction histories. The pressure generated by the aquatherTABLE 13-1 Paleopore pressure, maximum burial depth, and paleopressure/depth of Cardium sandstone at four locations in western Canada Paleopore pressure (Psi)

Well

Location

A

16-29-34-8-W5 (well shown 7100 in Fig. 13-1) 3-5-34-8-W5 (Fig. 2-12) 10,000 12-8-36-5-W5 (Fig. 2-13) 5200 10-26-36-6-W5 (Fig. 2-14) 5400 _. ___ - - - - - - --

-_ _

B C

D ____

~

-

~

__

__~--

Maximum burial depth (ft)

_~__.

_ _._ -

-_

10,800 13,000 10,200 8900 -_

Paleopressure/depth (psi/ft) --

-- --

0.66

-~

0.77 0.51 0.61 __~_

__

268 mal effect is a function of at least two important factors: geothermal gradient, and the retention of generated pressure.

Geothermal gradient The average geothermal gradient in the Gulf Coast is about 25"C/km (or 1.37"F/100 f t ) according to Barker (1972), although it varies widely within this area. The average gradient in the western Canada basin is about 1.7 to 1.S0F/100 f t , based on data given by Magara (1972, fig. 11).This suggests that the aquathermal effect could have been more pronounced in western Canada than in the Gulf Coast because of the higher gradient in western Canada. In other words, applying the results obtained in the Gulf Coast to western Canada would not, at least, cause an overestimation of pore pressures. The difference of the actual subsurface temperature (not the gradient) in these two areas should not cause a significant difference in generation of aquathermal pressures. This is because the aquathermal-pressuring mechanism is related t o relative change in temperature since isolation of pore fluids, rather than to actual temperature.

Retention of generated pressure The composition of shales and the level of compaction would be important factors in retaining the generated pressure. Because of a lack of data and because both factors probably vary widely within each area, it is not easy to evaluate them. It is assumed that there was no significant compaction of undercompacted shales during and after erosion. Therefore, the paleopore pressure can be calculated from the present compaction data. However, this may not be true always, because the pore pressures in these shales probably dropped during and after erosion mainly because of a cooling effect. The decreasing pore pressure would have caused the late-stage compaction of shales. In other words, the undercompacted shales at present are more compacted than those in the geological past (especially before erosion). It follows that the calculated paleopore pressure using the present compaction data and the estimated amount of erosion may be the lowest possible estimate. Therefore the calculated paleopore pressures in the Cardium may be lower than the pressures that actually existed in the geological past, but may be quite reasonable estimates with our current knowledge of shale compaction. These estimated paleopressures based on shale compaction data are at least more realistic than those by Currie and Nwachukwu (1974). They simply assumed hydrostatic pressures throughout geological time.

269 Estimation of paleotemperature Currie and Nwachukwu (1974) discussed the results of their research on the homogenization temperature of liquid-gas (or vapour) inclusions in the mineral filling that now occupies some of the fracture openings in the Cardium sandstone. Their objective was t o determine the temperature at which fluid inclusions were formed in fracture-filling material (mainly quartz). They first made thin-sections of those quartz fillings that contain fluid inclusions. The thin-sections were heated under a petrographic microscope by using a Leitz heating stage, until the bubble in each inclusion disappeared. The temperature when the bubble disappeared was measured and was called the "homogenization temperature." The same measurement was made for many samples from a single reservoir t o obtain a range of temperatures. Then this range of homogenization temperature was introduced into the pressure-temperature-specific volume diagram for water t o obtain the range of temperature under subsurface conditions (or under high pressure). The maximum temperature in this subsurface range may be assumed t o be close t o the temperature when the sedimentary deposits reached the maximum burial depth. In other words, this temperature seems to indicate the temperature of the deposits immediately before significant erosion took place. The minimum temperature in this range would be close to the present subsurface temperature. In this discussion, the fractures and infillings are assumed to have been caused mainly by the changes of the subsurface stress field associated with uplift and unloading. The ranges of the homogenization temperatures of the Cardium section determined by Currie and Nwachukwu (1974) are shown in Table 13-11. Wells A, B, and C in Table 13-11,are the same as those in Table 13-1.Well E in Table 13-11 is only a few miles away from well D shown in Table 13-1. Well F in Table 13-11is not far from well C. Because of their proximity, the latter two sets of wells may be combined t o study the subsurface temperature at the time of maximum burial. TABLE13-I1 Ranges of homogenization temperature of Cardium sandstone at five locations in western Canada (Determined by Currie and Nwachukwu (1974)) Well

A B C E F

Location

Homogenization temperature

16-29-34-8-W5 4 5-108" C 3-5-35 -8-W 5 46-100" C 1 2-8 -36-5-W5 50-85°C 12-24-36-6-W5 5144°C 10-5-36-5-W5 49-88°C _ _ _ _ ~ _ _ . _ _ _ _ _

( 113-226" F )

(113-21 2"F) (122-185°F) (124-183°F) (120-190°F)

- -.-.__-

270 1WW

’

10000

I>B?$ 500

K w 5000

5000

0 LT

250

0

50

100

0

I 150

I 200

I

250

0

I 300°F

TLMPERATURE

Fig. 13-2. Graphs showing ranges of homogenization temperature of Cardium sandstone and interpreted ranges of temperature and pressure under subsurface conditions at five locations in western Canada. Alphabets refer to wells shown in Tables 13-1, -11 and 111. (From Magara, 1976.)

As mentioned previously, estimation of the paleotemperatures may be made by using the pressure-temperaturespecific volume diagram for water (Fig. 13-2). The vertical axis shows the pressure in bars and psi, and the horizontal axis shows the temperature in both Celsius and Fahrenheit. The thin diagonal lines indicate the specific volumes (cc/g) which are reciprocals of the densities (g/cc) of water. In well A , for example, the maximum homogenization temperature was 226°F or 108°C (Table 13-11).This value is shown as A on the bottom horizontal axis of Fig. 13-2. It is assumed in this case that the pressure in the inclusion when the homogenization temperature is determined is zero psi. The estimated paleopore pressure in the Cardium is about 7100 psi (Table 13-1).Therefore, point A on the horizontal axis is moved parallel with the equal-specific-volume lines to a new point shown as A ’ in Fig. 13-2 corresponding to a pressure of 7100 psi. The reason for moving point A parallel with the isospecific volume lines is that the volume of the fluid inclusion should be unchanged under sub-

271 TABLE 13-111 Estimated temperature and geothermal gradient o f Cardium sandstone at time of maximum burial and at near present *** ___

Well

__

At maximum burial depth

-

___--__.__

A B C ED FC

--

_.

Near present _____-

Temperature ("F)

Gradient * ("F/100 ft)

Temperature (OF)

Gradient ** ("F/100 ft)

280 293 230 230 234

2.1 1.9 1.8 2 .o 1.8

142 153 154 160 151

1.8 1.4 1.8 1.8 1.7

* Based on the surface temperature of 50"F. ** Based on the present average surface temperature of 32°F. *** For locations of wells A , B , ..., F, refer to Tables 13-1 and 13-11. surface conditions and under the homogenization stage under the microscope. The temperature of point A' is about 280°F (139"C), which is considered to be the paleotemperature when the Cardium sandstone was buried to the maximum burial depth (10,800 ft). Assuming that the surface temperature was about 50°F (10°C) at that time, a geothermal gradient of about 2.1" F/100 f t is calculated. The minimum homogenization temperature for this well is marked a on the bottom axis of Fig. 13-2. The point corresponding to the subsurface condition after erosion is labelled a ' , which has a temperature of about 142°F (61°C). This temperature is slightly higher than the present temperature of this section. The geothermal gradient is calculated t o be about 1.8"F/100 ft as based on the average surface temperature of 32°F (OOC), which seems t o be a reasonable figure. This gradient is in agreement with the general geothermal gradient observed in this general area (Magara, 1972). The points for the other wells are also shown in Fig. 13-2, as indicated by B-B', b-b' and so on. In the case of each of the last two wells in Table 13-11,a combination of the two wells, as mentioned previously, is used for the study, shown as ED-ED', ed-ed', etc., in Fig. 13-2. The calculated temperatures and geothermal gradients are summarized in Table 13-111. The results in Table 13-111 suggest that the Cardium sandstone reached temperatures as high as 300°F (149°C) at the time of maximum burial. Its present temperatures range from 140 to 160°F (60 to 71"C), which is about one-half of the maximum. These estimated paleotemperatures are higher than those given by Currie and Nwachukwu (1974). The reason for the difference is that in their estimate they ignored the importance of the presence of undercompacted shales in this area. They assumed that the pore pressure

27 2 was hydrostatic at the time of maximum burial. Their method of estimating the thickness of erosion is also quite interpretive, based on an average denudation rate (Currie and Nwachukwu, 1974). The values given in this chapter are believed to be more realistic. The estimated geothermal gradient in Table 13-111shows that the gradient generally was higher at the time of maximum burial than it is at present. In this calculation an average surface temperature of 50°F (10°C) was used for the time of maximum burial. (This suggests that the average temperature of the sediments at the time of sedimentation was higher than the present surface temperature.) If, however, the surface temperature was higher than 50°F (lO"C), the geothermal gradient must have been lower. The other important factors affecting the interpretation of the geothermal gradient are the estimates of eroded thickness and the paleopore pressure, discussed previously. If the thickness of erosion was greater than estimated, the paleogeothermal gradient would have been lower than estimated. This is because the burial depth would have been greater but yet the paleotemperature would not have been much higher. However, if the paleopore pressure was higher than estimated, because of the possible late-stage compaction of undercompacted shales and the possible underestimation of the aquathermal-pressuring effect as discussed previously, the actual paleotemperature and, therefore, the paleogeothermal gradient would have been higher than the writer's estimate (see Fig. 13-2). Although estimates of thickness of erosion and of paleopressure made in this chapter are believed to be more accurate than those made by Currie and Nwachukwu, the estimated geothermal gradients at the time of maximum burial shown in Table 13-111may include both positive and negative errors. In any case, there seem to be more factors that might cause positive errors on the paleogeothermal gradient than factors that might cause negative errors. Thus, it appears that the actual paleogeothermal gradient would have been lower than the preceding current estimates. From these observations and interpretations, it appears that the geothermal gradient probably did not change very much in the geological past. Currie and Nwachukwu's estimate showed the changes of the gradient from 1.9 (maximum burial) to 1.1 (near present) "F/100 ft in some wells (see their fig. 6). These differences are probably caused by their lower estimation of erosion and pore pressure. Their interpreted near-present geothermal gradient of 1.l0F/100 ft is considerably different from the present geothermal gradient of 1.7 to 1.8"F/100 f t in this area (Magara, 1972). This lower value (1.lo F/100 ft) seems t o have resulted from use of the high surface-temperature value of about 59°F (15°C). The present surface temperature extrapolated from the subsurface data in this area is approximately 32°F or 0°C (see Magara, 1972, fig. 11).In their fig. 6, Currie and Nwachukwu indicated the geothermal gradient lines to be straight, on the assumption that the pore pressure has been hydrostatic (or that pore pressure is a linear function

273 of depth). If the pore pressure was abnormal and its abnormality varied in the geological past, the geothermal gradient lines must be curved rather than straight. In summary, the maximum temperature estimated by the method mentioned above will in many places provide a more realistic estimate of the oilgenesis stage (Chapter 12) than the present subsurface temperature. References Barker, C., 1972. Aquathermal pressuring - role of temperature in development of abnormal-pressure zones. Bull. Am. Assoc. Pet. Geol., 56: 2068-2071. Currie, J.B. and Nwachukwu, S.O., 1974.Evidence of incipient fracture porosity in reservoir rocks at depth. Bull. Can. Pet. Geol., 22: 42-58. Magara, K., 1972. Compaction and fluid migration in Cretaceous shales of western Canada. Geol. Surv. Can. Pap,, 72-18: 81 pp. Magara, K., 1975. Importance of aquathermal pressuring effect in Gulf Coast. Bull. Am. Assoc. Pet. Geol., 59: 2037-2045. Magara, K., 1976. Thickness of removed sediments, paleopore pressure, and paleotemperature, southwestern part of Western Canada Basin. Bull. Am. Assoc. Pet. Geol.,

60: 554-565.


Chapter 14 PRIMARY HYDROCARBON MIGRATION

Hydrocarbons generated in fine-grained sedimentary rocks are probably disseminated at first, but eventually they must move from their host rocks into more permeable. and porous sedimentary rocks t o form an accumulation. The movement of hydrocarbons from nonreservoir rocks to reservoir rocks is called primary migration, and is distinguished from their concentration and accumulation within the reservoir rocks known as secondary migration (Levorsen, 1967). A discussion of the primary migration of hydrocarbons includes three different kinds of problems: (1)The form in which they migrate, such as molecular solution, micellar solution and separate hydrocarbon phase. (2) The cause of migration. (3) The water source and the cause of its movement. Whether the hydrocarbons being discussed are mainly gas or mainly liquid will affect the proportions of them that are moving in solution and in separation, because the solubility of gaseous hydrocarbons is generally higher. The migration mechanism for hydrocarbons in separate phase may differ from that for hydrocarbons dissolved in water. If water movement is important in moving hydrocarbons, the source of the water and the cause of its movement must be examined carefully. The molecular solubility of liquid hydrocarbons in water at relatively high temperatures was recently discussed by Price (1976), who showed that the solubility increased with increasing temperature (Fig. 14-1). The solubility of the Farmer's oil at 160°C is approximately 150 ppm, and the curve shows that it tends t o increase with further increase in temperature. However, these temperature values are much higher than the known temperature range of 60" to 150°C for active oil generation. Dickey (1975) on the other hand, suggested that the flowing stream would have to contain at least 10,000 ppm of hydrocarbons at the time of primary migration. Therefore, it may be very difficult t o believe that most oil migrates as a molecular solution in water. In the case of migration of gas, the situation can be completely different. According t o Dodson and Standing's chart (1944, see Fig. 14-2), the solubility of natural gas in water ranges from 4 cu f t per barrel of water at 500 psi, up to 30 cu f t at 10,000 psi. In other words, most of the gas might migrate in solution in water at the primary migration stage. Another solution mechanism, called micellar solution, was proposed by

27 6

-E

l

FARMERS WHOLE OIL

ALASKA REEDY CREEK

L. A. STATE

TEMPERATURE IN OC

Fig. 14-1. Solubilities of two whole oils (Wyoming Farmers and Louisiana State) and four topped oils (Amoseas Lake, Reedy Creek, Alaska and Union Moonie) as functions of temperature in water. Topping temperature is 2OO0C (392OF). (From Price, 1976.)

Baker (1962). He suggested that hydrocarbon solubility is substantially high if the water contains micelles formed by soaps of organic acids. However, there are several reasons why Baker’s proposal is not plausible as the princi-

-

EXPERIMENTAL D A T A

---- E X T R A P O L A T E D

DATA

Fig. 14-2. Solubility of natural gas in formation water in cu ft/bbl. (From Dodson and Standing, 1944.)

277 pal mechanism of hydrocarbon migration in the subsurface, First, there is no good evidence that such solubilizing micelles exist in substantial quantity in shales. And even if they d o exist in shales, they would not be easily moved because they are not small. Then, the micelles would increase the solubility of the heavier hydrocarbons in water only to a few parts per million - nowhere near the 10,000 ppm o r more that now appears to be necessary (Dickey, 1975). Another difficult point in believing micellar solution to be important in primary migration is that the process of unloading the hydrocarbons carried by the fluid (water, micelles and hydrocarbons) at the final trapping position in the reservoir cannot be thoroughly explained. The preceding discussion may lead us t o conclude that the larger proportion of liquid hydrocarbons must migrate in a separate phase, although the rest can migrate in solution in water. In the case of gas migration, the proportion moving in solution in water can be relatively large, because of its greater solubility in water. Although the form in which hydrocarbons moved at the time of primary migration, and the mechanism of that migration, are not completely understood, the movement of water in fine-grained source rocks must be one of the most important factors. The amounts of organic matter and generated hydrocarbons in the source rocks are quite small in comparison with the amount of water. The movement of the large quantity of water must have influenced and may have controlled the direction and effectiveness of hydrocarbon migration. If the water concerned is meteoric water, the direction in which it moves is controlled by excess pressures generated by the difference in elevations of the water-intake areas of the aquifers. If, however, the moving fluid originated in the sediments, the loading of the sediment layers would be the principal cause of the excess fluid pressure that determines the direction of fluid movement. Compaction fluid movement This problem was discussed in Chapter 8 ; the fluids move from an area of more loading (thicker deposition) to one of less. The volume of horizontal fluid movement relative to vertical increases as the permeability and thickness of interbedded permeable rocks increase, and as the rate of thickness change of newly deposited sediments increases. If the shales are thick and homogeneous, most fluid will move vertically. The presence of some contiguous or lenticular sandstones in a thick shale sequence may not drastically change this basic direction of fluid flow. Price (1976) recently proposed the importance of growth faults through massive .and undercompacted shales in the Gulf Coast as the main fluidmigration pathways. He suggested that hydrocarbons generated in the deep and hot undercompacted shale section have migrated upward along these

278 FLUID PRESSURE 1000 p r i 0 2 4 6

Fig. 14-3. Fluid-pressure profile in the Beaufort Basin.

faults, in the form of molecular solution in water. The fluid-flow model discussed in Chapter 8 , however, indicates that the principal direction of flow through massive shales is vertically upward, whether the shales are faulted or not; in other words, such upward movement of generated hydrocarbons through these shales is always possible. Its importance in the total petroleum accumulations in this area, however, may not be so great because the total volume of vertical fluid flow through these undercompacted shales may not have been large. The model discussed in Chapter 8 is applicable if the sediments reached the compactionequilibrium condition after each increment of instantaneous loading. If, however, some shales were t o stay (slightly) undercompacted while other shales attained almost compaction equilibrium, significant pressure differences and barriers within the shale zones would be developed. This type of facies was named “mixed compaction facies” (Evans et al., 1975). It occurs in an intermediate depth range below the normal compaction facies. Examples of the calculated fluid-pressure profiles of the mixed compaction facies in the Beaufort Basin, Canada, are shown in Fig. 14-3.Fluid moves from a higher excess-pressure point to a lower, and the inferred directions of fluid flow are shown by arrows. Similar pressure or compaction patterns were reported in other sedimentary basins (see Chapter 5 ) . Once an interbedded sequence has reached such an intermediate compaction stage, essentially all the compaction fluids may have to move laterally

279 Ol

O l

$65

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10000 12500 FEET

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c 20052

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u

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30000

30000

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5000 10000 CUMULATIVE WATER VOLUME

CU FT I S 0 FT

15000 CU FT is(1 FT CUMULATIVE WATER VOLUME

Fig. 14-4.Cumulative water-loss volumes from shales in the Gulf Coast (combined vertical and horizontal migration model). (From Magara, 1976.)

through the interbedded sandstones. There is some vertical fluid flow in the shales, too, but the flow is only local. In summary, the development of the mixed compaction facies could facilitate lateral fluid flow from syncline areas, and this flow would take place after the sandstoneshale sequence has reached an intermediate depth range where petroleum may have been generated by the thermal process. The discussion in Chapter 8 introduced the method of calculating the horizontal and vertical fluid volumes that have moved from a given block of rock. In the subsurface, however, fluids expelled from the other blocks below and beside a particular block will also influence the fluid-flow condition within that block. In other words, the cumulative effect of fluid migration will be three-dimensional. Although estimating such a fluid-migration condition is extremely complicated, it may be worthwhile in that the migration of hydrocarbons may be affected by the cumulative fluid migration after the hydrocarbons have been generated. Magara (1976)recently estimated cumulative compaction fluid volumes, using Dickinson's (1953)porosity-depth curve and a simplified Gulf Coast model. In this model, the upper geological sequence is composed of sandstoneshale interbeds in which fluids have moved horizontally, and the lower sequence consists of massive and homogeneous shales where compaction fluid 'has moved vertically upward. The horizontal migration distance in the upper sequence is assumed to be 10 miles, and the total thickness of the sedimentary column 33,000 f t (10km). Fig. 14-4shows the cumulative vol-

280 umes of fluid loss since burial t o 2000 f t from a shale column whose base area is 1 sq f t (the respective depths to the boundaries of the upper and lower sequences are assumed t o be 9500 f t and 12,500 ft). It is interesting t o note that the cumulative fluid-volume plot based on the model that simulates the Gulf Coast sedimentary basin resembles the oilproduction frequency plot for the same area (Burst, 1969). This similarity suggests that fluid movement due to sediment compaction is one of the controlling factors in hydrocarbon occurrence in that area. A fact that could affect the importance of mechanical shale compaction and fluid expulsion in petroleum migration is that the rate of compaction decreases continuously as the shales become more deeply buried. In other words, by the time the source rocks had reached deep burial where the ternperature was high enough to generate hydrocarbons, the movement of cornpaction fluid might have become too slow and insignificant. If the fluids expand at such depths, the expansion might facilitate latestage fluid movement. Subsurface temperature increase with burial depth might cause such fluid expansion in most sedimentary basins. Aquathermal fluid movement Fig. 14-5 is a pressure-temperature diagram for water with selected isodensity lines, adapted from Barker (1972). The vertical scale is pressure in psi, and the horizontal scales are temperature in both Centigrade and Fahrenheit. Density values in g/cc (and specific volume values in cc/g) of water are shown along the isodensity lines. The original data for constructing this diagram were obtained by Kennedy and Holser (1966). The three geothermalgradient lines of 25"CFm (1.37"F/100 ft), 18"C/km (1"F/100 f t ) and 36"C/ km (2"F/100 ft) for hydrostatically pressured water (the system is not closed) are superimposed; the lines intercept water isodensity lines whose values decrease as the pressure (or burial depth) increases. A hydrostaticpressure gradient of 0.47 psi/ft was used. This progression t o lower densities and higher specific volumes means that a given weight of water expands with burial: the reason is that the increase of pressure associated with the 0.47 psi/ft hydrostatic gradient is inadequate t o hold the water volume constant. The amount of expansion can be derived easily from the specific volume values (cc/g), shown in brackets. When the geothermal gradient is 25"C/km (1.37"F/100 ft), for example, the specific volume increases from 1 cc/g at 0 psi pressure to 1.10 cc/g at 11,600 psi, which corresponds to a burial depth of about 25,000 ft. Thus, a 10%water expansion results from about 25,000 f t of burial; this is a significant amount. Continuous expansion of water for the three geothermal gradients is depicted in Fig. 14-6, where specific volume of water (cc/g) is shown on the vertical scale and depth (ft) on the horizontal scale. At 20,000 ft, for example, about 3% expansion has occurred for the geothermal gradient of 1"F/

281

TEMPERATURE b 32

O C

4 100

ZOO

300 TEMPERATURE

400

SO0

O F

Fig. 14-5. F'ressure-temperature-density (or specific volume) diagram for water. Three geothermal lines of 25O, 18O and 36OC/km for hydrostatically pressured fluid are superimposed on a basic diagram derived from Barker (1972).

100 ft, about 7% expansion for 1.37°F/100 ft, and 15%for 2"F/100 ft. Fig. 14-6 shows that rates of increase in specific volume, or rates of expansion, increase with burial depth. This fact is interesting because the amount of water expelled by compaction decreases with burial depth, but the subsurface temperature tends to expand water volume. This expansion could facilitate fluid migration at depth and hence could favour hydrocarbon migration. Expansion of rock grains also may be considered in the discussion of fluid migration. The grain expansion would create more intergrain spaces, thus more spaces for water. Its effect, however, is much less significant: the thermal expansion of quartz, for example, is only about & that of water (see Skinner, 1966). Thermal expansion data for dry clay matrix are not readily available; the value for quartz may be the closest approximation. In other

282

2 0 )OO

00

DEPTH [ F T

3 c I00

I

Fig. 14-6. Specific volume (of waterkdepth relationships in normally pressured zones for three geothermal gradients of 25', 18' and 36'C/km. (From Magara, 1974b.)

words, if the ratio of volume of water to that of rock grains is more than about 1:15 (porosity is more than about 6%),the effect of water expansion overrides that of grain expansion, resulting in water movement. In the Gulf Coast, a shale porosity of 6% would not be attained above 24,000 f t (Dickinson, 1953). Note that the above-mentioned aquathermal model is valid when pore water is not completely isolated. Such a relatively open system is developed in the normal and mixed compaction facies, which usually occur in the shallow t o intermediate depth range in many sedimentary basins. If the pore fluids are more isolated, as in the case of undercompacted facies, the fluid cannot expand freely and the fluid pressure will increase (Magara, 1974b). The directions of fluid migration due to the aquathermal effect are from a hot place to a cold, from a deep section t o a shallow, and from a basin's centre t o its edges. These directions are essentially the same as those of fluid movement caused by sediment compaction. Therefore, the significance of the aquathermal effect in the subsurface may simply be t o increase the effectiveness of compaction fluid flow at deep burial.

283 B

A SHALE

C

FLUID PRESSURE WATER SALINITY

POROSITY

IN SHALE

I N SHALE

SHALE

SHALE

+

DEPTH

DEPTH

4 ~

SHALE

SHALE

R

\+

E

T I O ON F OSMOTIC FLUID FLOW

C

7

LHYDROSTATIC PRESSURE

Fig. 14-7. Schematic diagram showing shale porosity, fluid pressure and pore-water salinity distributions in interbedded s a n d s h a l e sequences (From Magara, 1974a.)

Now let us assume a geological model at intermediate depths in which sandstones are interbedded with shales. A shale porosity profile such as is shown in Fig. 14-7A may be developed. If the interbedded sandstones are permeable, the maximum fluid expulsion or the maximum shale porosity reduction will occur in the shales directly above and below the sandstones. The porosity in the middle of a shale bed may remain relatively high. The corresponding fluid-pressure plot is shown in Fig. 14-7B,in which arrows depict the inferred directions of compaction fluid flow. If water expands from the thermal effect, water will move within the shale bed from the centre to the upper and lower edges, because more expansion can be expected at the point of higher porosity (more water). The directions of the small-scale fluid migration due to the aquathermal effect are essentially the same as those of compaction fluid migration. Osmotic fluid movement In many sedimentary basins the salinity of the formation water increases with depth or compaction. These salinity values are usually higher than that of sea water (about 35,000ppm). In the undercompacted zones, the salinity is lower than those of normal and mixed compaction zones. The principal cause of these salinity variations in sedimentary rocks may be ion filtration by shales (see Chapter 10). Ion filtration by clays or shales has also been documented by laboratory methods (McKelvey and Milne, 1962;Engelhardt and Gaida, 1963),which showed that clays and shales filter salt from a solution. Therefore, the fluids

284 moving through the shales must be fresher than the original solution that saturated the shales. As mentioned in Chapter 10, Hedberg (1967) studied pore-water chlorinities and porosities in shales, using cores from several areas in the world. Fig. 10-10 shows the chloride content (ppm) versus porosity plots from the Burgan field in Kuwait and several oil fields in Texas. The relation between the chlorinity * and porosity in the Burgan data may be approximated by a hyperbola: the chlorinity increases as the porosity decreases. The data from the three Texas fields are too scattered and insufficient to prove or disprove the hyperbola relation. It is, however, interesting that most of the plotted data from Texas fall within the extension of the general Burgan trend. Combining the the concept of ion filtration and the shale porosity profile as shown in Fig. 14-7A enables a possible water-salinity profile for the shales to be drawn (Fig. 14-7C). Salinity is the reciprocal of shale porosity; i.e., it increases as the porosity decreases. Salinity, therefore, would increase from the centre to the edges of each shale bed. Because osmosis tends to move water from a fresher to a more concentrated side, the fluid-flow direction due to osmosis can be inferred as shown by the arrows in Fig. 14-7C. The osmotic-pressure difference due to salinity change is probably not very large as compared with that due to compaction. According to the chart shown by Jones (1967), the osmotic-pressure difference caused by a salinity difference of 50,000 mg/l is only about 600 psi (see Fig. 4-18 in Chapter 4). Because osmotic fluid flow is in the same direction as compaction fluid flow, however, osmotic flow could facilitate the primary hydrocarbon migration from shales to permeable sandstones. This combined fluid flow due to compaction and osmosis may continue until the shales reach equilibrium, at which time no fluids can be expelled from them by compaction, and salinity also may reach equilibrium. If any of the freshening mechanisms alter the salinity distribution at the later stages, the osmotic fluid flow may be changed also. The most important point in this combined mechanism of flow, however, is that the salinity contrast resulting from ion filtration seems to start at relatively early stages of compaction, and the resultant osmotic-pressure difference seems to support fluid flow from the shales at the early stages of water expulsion. Another important effect of the combined fluid flow on hydrocarbon migration is that the fluids moved both by compaction and by osmosis are relatively fresh, so that hydrocarbons would be more soluble in them than in more saline water. High hydrocarbon solubility will favour hydrocarbon migration.

* Salinity (NaCl) may be calculated by multiplying the chlorinity by 1.65.

285 Fluid movement due to clay dehydration Powers (1967)showed that alteration of montmorillonite to illite in the Gulf Coast area begins at a depth of about 6000 ft and continues at an increasing rate to a depth, usually about 9000-12,000 ft, where there is no montmorillonite left. This alteration offers a mechanism for desorbing the last few layers of adsorbed or bound water in clay and transferring it into interparticle locations as free water. If the last few layers of bound water have a greater density than free water, this released water tends to increase its volume as it is desorbed from between unit layers. If water expansion is restricted, the pore-water pressure will increase to abnormally high levels. According to Burst (1969),clay dehydration depends mainly on subsurface temperature, the average dehydration temperature being 221°F in the Gulf Coast. Certain chemical conditions for potassium fixation also are required for this conversion. Phase change and possible expansion of bound water at the time of dehydration may, as proposed by Burst (1969),be important agents for flushing hydrocarbons, at least from clay-interlayer locations t o interparticle locations (shale pore space). Martin (1962) summarized data on adsorbed water density in montmorillonite analysed by several different investigators. This summary is shown in Fig. 14-8,which plots the calculated and measured water density versus amount of water in the clay (g HzO/g clay) Fig. 14-8appears to support Powers’ (1967)and Burst’s (1969)proposals of the higher-than-normal (greater than 1 g/cc) density of the adsorbed water. However, the validity of

0

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Fig. 14-8. Adsorbed water density on Na-montmorillonite. (From Martin, 1962, for which see also the references in this figure.)

286 the higher-than-normal density shown in Fig. 14-8 is not very great, because most of these higher values were derived from calculations rather than direct measurements. Martin (1962) stated that “the only unambiguous adsorbed waterdensity data are (those) of Anderson and Low (1959)”, which show values less than 1g/cc. Therefore, from the data existing at present, it is difficult t o prove or disprove the water expansion and flushing effect associated with clay dehydration. However, we may be able t o say that clay dehydration could be an additional source of liquid water at relatively deep burial where hydrocarbons may have been generated. Van Olphen (1963) demonstrated that at 25°C the pressure needed to remove the last interlayer of water is 65,000-70,000 psi, and that needed for the second-to-last water interlayer is 30,000 psi. These values are considerably higher than the pressure at depths less than 20,000 ft. Overburden pressure alone, then, may not suffice to release at least the last two layers of bound water. This is the main reason why Burst and Powers developed their concepts of the temperaturedependent water-release mechanism associated with clay-mineral conversion. If, however, the interlayer water is released by clay dehydration in response to temperature, and subsequently remains in the pore spaces as free water, the same overburden pressure could be enough to push it out of the shales, provided drainage is available. This type of water movement is essentially the same as that caused by sediment compaction discussed above (Magara, 1975a). The validity of the average dehydration and mineralconversion temperature of 221°F proposed by Burst must also be examined. Schmidt (1973) studied the proportions of expandable clay (mostly montmorillonite) and nonexpandable clay in a well drilled in the Gulf Coast (Fig. 14-9). This figure shows that the rate of mineral conversion increases at about 10,500 ft, which corresponds t o a subsurface temperature of about 200°F (Fig. 14-10). However, the geothermal gradient in this well also increases at that depth (10,500 ft), which is the top of the undercompacted (abnormally pressured) section. Because water has a thermal conductivity significantly lower than that of most rock matrix, the undercompacted section, which contains an excessive amount of water, tends to have a thermal conductivity lower than that in the normally compacted section (Lewis and Rose, 1970). If heat flows upward at a given rate, the geothermal gradient in the undercompacted section would become greater than that in the normal compaction section. An important point shown in Fig. 14-9 is that clay-mineral conversion is not a drastic event. The conversion temperature of 221”F, as suggested by Burst, may not be required. Rather, conversion seems t o begin almost immediately after deposition, and continues to depth. The higher the geothermal gradient, the faster the rate of conversion. Because in essentially all the world’s sedimentary basins temperature tends

287

0

LESS THAN 2 . 0 MICRON SIZE C L A Y

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L E S S THAN 0.15 MICRON SIZE C L A Y

I-

Y Y)

8-

I II -

-

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a

12-

14 I 100

40

60

80

20

% EXPANDABLE CLAY

Fig. 14-9. Plot of per cent expandable clay versus depth showing accelerated increase in diagenesis of montmorillonite to mixed-layer mon tmorilIoniteillite. (From Schmidt, 1973.)

.

TEMPERATURES FROM BOTTOM HOLE PRESSURt SURVEYS I N SHUT-IN PRODUCING WELLS

aTfMPERATURES OF SHALLOW GROUND WATERS MEASURED B Y THE ffiGS

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to increase with depth, the bound water will be released in any case. Claymineral conversion could create an additional source of liquid water at depth. Its significance in primary migration, however, cannot be understood clearly, because whether such conversion causes fluid expansion and migration is not known. Other possible causes of primary migration There are several other possible causes of primary migration, such as capillary pressure, buoyancy, diffusion, generation of hydrocarbons - especially gas, etc. These causes are mostly unassociated with the movement of water. Although there is no solid reason to deny their importance, I personally feel that water movement of some kind must be important at the primary-migration stage, and that therefore a mechanism unrelated to water movement may be of secondary importance. As pointed out previously, we are dealing with a large amount of water and a relatively small amount of hydrocarbons in the sediments, which have a very fine network at the time of primary migration.

Form of hydrocarbons at primary migration If all the hydrocarbons are in molecular solution in water at the primarymigration stage, estimating the volume and direction of sediment-source water as discussed is of prime importance in understanding hydrocarbon migration. The water volume may be tied directly t o the amount of hydrocarbons. However, the solubility of liquid hydrocarbons in water is relatively low even at elevated temperatures (Fig. 14-1).Micellar solution as proposed by Baker (1962)cannot be very important in the subsurface for the several reasons mentioned above. According to Dickey’s (1975)estimate, there must be at least 10,000 ppm hydrocarbons in the flowing stream at the time of primary migration. Another approach to estimating the required concentration of oil in the flowing stream is as follows: Tissot and Pelet (1971)analysed the amounts of hydrocarbons, resins and asphaltenes in shales adjacent to a reservoir in Algeria. Fig. 14-11shows the results of their analyses in mg/g organic carbon. Although the amounts of resins and asphaltenes in the shale remain relatively constant, the amount of hydrocarbons decreases toward the reservoir, suggesting primary hydrocarbon migration. The difference in hydrocarbon contents at the 14-m point and at the near-reservoir point is about 40 mg/g organic carbon. If the level of total hydrocarbon generation per gram of organic carbon is constant throughout the shale section, this 40 mg represents the lowest possible amount of hydrocarbons expelled per gram of organic carbon from the shale closest to the reservoir. If the shale has a density of 2 g/cc and 1 weight per

289 mdg ORGANIC CARBON

20

40

M

HYDROCARBONS

80

1W

120

ALGERIA

RESERVOIR

Fig. 14-11. Plot of amounts of hydrocarbons, resins and asphaltenes versus organic carbon (g) of Devonian shales adjacent to reservoir in Algeria. Original data derived from Tissot and Pelet (1971).

cent of organic carbon, 1 cc of this shale lost 0.8 mg of hydrocarbons *. If the porosity difference between these two points is lo%, which seems to be the largest porosity difference possible under these conditions, the amount of hydrocarbons in the flowing stream can be estimated to be about 8000 PPm. As mentioned above, this estimate is based on the lowest possible estimate of hydrocarbons in the compaction fluid; the true value could be higher. In any case, this figure is at least one order higher than the highest molecularsolubility figure in the temperature range for oil generation, and is surprisingly close to the >10,000 ppm given by Dickey (1975).Note that the density and porosity data for the shales studied by Tissot and Pelet (1971) are not readily available, so that they have had to be assumed for this estimate. Vyshemirsky et al. (1973)experimented with squeezing the mixture of clay, liquid hydrocarbons and water up to 300 atoms. They found that the

* The figures used for this estimate would be the lowest possible values to produce the lowest possible hydrocarbon yield.

290 amount of hydrocarbons squeezed with the water was more than could be accounted for by the solution mechanism alone. From the above estimate and other observations, it is clear that the greater proportion of liquid hydrocarbons must move in a separate phase. Gas, however, can migrate in aqueous solution because of its higher solubility. The question then arises: Why is the movement of water important if most of the liquid hydrocarbons move in their separate phases? The next section will suggest an answer. Migration of oil in oil phase

A comprehensive discussion of the mechanisms associated with oil-phase migration was published in 1954 by Hobson. Recently Dickey (1975) rediscussed this possibility from a slightly different angle. In compacted shales, the larger proportion of water is electrically charged at the clay surfaces, and has a relatively high viscosity (Fig. 8-6), which means that some water is semisolid. The amount of liquid (or free) water in the compacted shales is probably not great. In these circumstances, if the shales compact further, the oil as well as the liquid water will migrate provided the oil saturation in the liquid phase is higher than the critical value for oil migration. Basing his argument on the concept of relative permeability in sandstone, Dickey (1975) stated that, “oil, will move along with the water only if it occupies about 20 per cent or more of the pore volume” (p. 341). If the sandstone is partly water-wet and partly oil-wet, the critical residual-oil saturation can be as low as 10%.Dickey also suggested that the residual-oil saturation in shales may be less than 10% and possibly as low as 1%, because a considerable fraction of the internal surfaces of shales can be oil-wet (p. 342). A schematic diagram of the relative-permeability and oil- (or water-) saturation relationship is shown in Fig. 14-12. The critical residual-oil saturation is marked by an X and for shales may be assumed t o be a value between 10 and 1%.For oil t o migrate along with water, this critical oil saturation must be exceeded; that is, if the oil saturation is at X ’ in Fig. 14-12, there will be some oil migration. If, for example, the oil saturation in the total water (solid and liquid) is 100 ppm (0.01 wt%), and if only 1%of the water is in liquid phase (and 99% is solid), the oil saturation in the liquid water will be 10,000 ppm (1 wt%). Assuming that the density of oil is 0.8 g/cc and that of water 1g/cc, this figure will correspond to about 1.2 ~01%.This is the concept suggested by Dickey. If some of the liquid water is expelled, possibily with a small amount of oil, then the oil saturation of the liquid phase in the shale pores will increase, ensuring more oil migration. However, as the liquid water is further expelled as compaction proceeds, the permeability will be reduced to

291

X'Y

x

0

OIL SATURATION

0

WATER SATURATION

100%

Fig. 14-12. Schematic diagram showing relative permeabilitywater (or oil) saturation relationship for sandstone.

an extremely low level and the movement of the total fluids (water and oil) eventually may become difficult. Possible changes in oil saturation with the gradual removal of liquid water are depicted in Fig. 14-13. The two diagonal lines refer to the original oil concentrations of 1 0 ppm and 100 ppm, respectively, when the liquid water occupied 10% of the bulk shale volume. If the liquid is expelled from these shales until 0.01% liquid water remains (there would still be a lot of solid water left at this stage), the respective oil saturations in the liquid phase will become 10,000 ppm ( 1 wt%) and 100,000 ppm (10 wt%) *. The corresponding volume percentages are about 1.2 and 12%, respectively. At this stage, oil may move along with the water (Fig. 14-12). The boundary between the solid and liquid water phases in shales would not be clear-cut; the change is probably quite gradual. In other words, it would be difficult to define how much solid water and how much liquid water are in a shale at any given compaction stage. However, for further discussions in this chapter, it would be wise to get some approximate figures on the amounts of solid water in shales. Let us assume that there is an illite clay sample whose bulk density is 2 g/cc. The specific surface area of illite clay is about 100 m2/g (Grim, 1953, p. 311). If illite has one solid-water

* Movement of oil up to the state of 0.01%liquid water is ignored in this case.

292

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.

10

100

1.m

1o.m

1M.MO

Wl

01

1

1

10

W1

01

1

1

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lOppm OIL CONCENTRATION WHEN LIOUID WATER OCCUPIES 10% OF BULK VOLUME

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1.m.MO ppm

loow% I W VOL %

CONCENTRATION OF OIL IN LIQUID PHASE

Fig. 14-13.Chart showing increasing tendency of concentration of oil in liquid phase a liquid water per cent in shale decreases, assuming no oil migration.

layer whose thickness is about 2.5 A (or 0.26 nanometer), then 1 cc of this illite contains about 0.06 cc of solid water. In other words, about 5 vol% of illite clay would be solid water. Shales usually contain other nonclay minerals (quartz, feldspar, etc.), so the actual per cent of solid water in an illitic shale may be slightly lower than this estimate. Now let us estimate the amount of solid water in montmorillonite clay, Theoretically speaking, montmorillonite has a specific surface area of about 800 m2/g (Grim, 1953, p. 311); therefore, if only one water layer is considered to be relatively solid, the amount of solid water might be about 40%. Because montmorillonite should have more than one water layer, this estimate probably is too low. However, we also have t o allow for the effect of nonclay minerals in actual shales, which will reduce the proportion of solid water in bulk volume as discussed above. The other factor we might have to consider is that numerous water layers attached t o montmorillonite clay surfaces will reduce the bulk density of the sample significantly, so that to assume a bulk density of 2 g/cc may not be warranted; it could be lower. This factor in turn would reduce the surface area within 1 cc of the montmorillonite clay sample, and hence the percentage of solid or structured water. Powers' (1967) estimate is that about 60% of montmorillonite clay is (structured) water. In the light of the above considerations, let us assume that 4 6 5 0 % of montmorillonitic shale is relatively solid or structured water. This reasoning suggests that montmorillonitic shale contains approximately 8 t o 10 times as much solid water as illitic shale.

293

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Transnational Migration and Human Security: The Migration-Development-Security Nexus

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Modelling of powder die compaction

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The Sand Compaction Pile Method

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Global Migration and Development

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Fluid and Electrolyte Balance

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