Hydrocarbon Migration Systems Analysis

Developments in Petroleum Science, 35 hydrocarbon migration systems analysis DEVELOPMENTS I N PETROLEUM SCIENCE Advi...

Author: J.M. Verweij

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Developments in Petroleum Science, 35

hydrocarbon migration systems analysis

DEVELOPMENTS I N PETROLEUM SCIENCE Advisory Editor: G.V. Chilingarian Volumes I . 3 . 4 , 7 and 13 are out of print W.H. FERTL - Abnonnal Formation Pressures T.F. YEN and G.V. CHILINGARIAN (Editors) - O i l Shale D.W. PEACEMAN - Fundamentals of Numerical Resevoir Simulation L.P. Dake - Fundamentals of Resevoir Engineering K. MAGARA - Compaction and Fluid Migration M.T. SlLVIA and E.A. ROBINSON - Deconvolution of Geophysical Time Series in the Exploration for Oil and Natural Gas I I . G.V. CHILINGARIAN and P. VORABUTR - Drilling and Drilling Fluids 12. T.D. VAN GOLF-RACHT - Fundamentals of Fractured Reservoir Engeneering 14. G. MOZES (Editor) - Paraffin Products 15A. 0. SERRA - Fundamentals of Well-log Interpretation, 1. The acquisition of logging data 15B. 0. SERRA - Fundamentals of Well-log Interpretation, 1. The interpretation of logging data 16. R.E. CHAPMAN - Petroleum Geology 17A. E.C. DONALDSON, G.V. CHILINGARIAN and T.F.Yen (Editors) - Enhanced Oil Recovery, I. Fundamentals and analyses 17B. E.C. DONALDSON, G.V. CHILINGARIAN and T.F.Yen (Editors) - Enhanced Oil Recovery, 11. Processes and operations 18A. A.P. SZILAS - Production and Transport of Oil and Gas, A. Flow mechanics and production 2. 5. 6. 8. 9. 10.

(\econd cimpletely revised edition)

18B. A.P. SZILAS - Production and Transport of Oil and Gas, B. Gathering and Transport (second completely revised edilion)

19A. G.V. CHILINGARIAN. J.O. ROBERTSON Jr. and S. KUMAR - Surface Operations in Petroleum Production, I 19B. G.V. CHILINGARIAN, J.O. ROBERTSON Jr. and S. KUMAR - Surface Operations in Petroleum Production, II 20. A.J. DIKKERS - Geology in Petroleum Production 21. F. RAMIREZ - Application of Optimal Control Theory to Enhanced Oil Recovery 22. E.C. DONALDSON, G.V. CHILINGARIAN and T.F. Yen - Microbial Enhanced Oil Recovery 23. J. HAGOORT - Fundamentals of Gas Resevoir Engineering 24. W. LITTMANN - Potymer Flooding 25. N.K. BAIBAKOV and A.R. GARUSHEV - Thermal Methods of Petroleum Production 26. D. MADER - Hydraulic Proppant Farcturing and Gravel Packing 27. G. DA PRAT - Well Test Analysis for Naturally Farctured Resevoirs 28. E.B. NELSON (Editor) - Well Cementing 29. R.W. ZIMMERMAN - Compressibility of Sandstones 30. G.V. CHILINGARIAN, S.J. MAZZULLO and H.H. RlEKE - Carbonate Resevoir Characterization: A Geologic-Engineering Analysis, Part 1 31. E.C. DONALDSON (Editor) - Microbial Enhancement o f Oil Recovery - Recent Advances 33. E. F J E R , R.M. HOLT, P. HORSRUD, A.M. RAAEN and R. RISNES - Petroleum Related Rock Mechanics 34. M.J. ECONOMIDES - A Practical Companion 10 Resevoir Stimulation 36. L. DAKE -The Practice of Reservoir Engineering 37. W.H. SOMERTON - Thermal Properties and Temperature Related Behavior of Rock/Fluid Systems

Developments in Petroleum Science, 35

hydrocarbon migration systems analysis

J.M. VERWEIJ TNO Institute of Applied Geoscience, Schoemakerstr. 97, P.O. Box 6012,2600 JA Delft, The Netherlands

ELSEVIER, Amsterdam -London -New York -Tokyo

1993

ELSEVIER SCIENCE PUBLISHERS B.V. Sara Burgerhartstraat 25 P.O. Box 2 I I , 1000 A E Amsterdam, The Netherlands

ISBN: 0-444-89103-X

0 1993 Elsevier Science Publishers B.V. 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 Science Publishers B.V., Copyright & Permissions Department, P.O. Box 521, loo0 AM Amsterdam, The Netherlands. Special regulations for readers in the USA - This publication has been registered with the Copyright Clearance Center Inc. (CCC). Salem. Massachusetts. Information can be obtained from the CCC about Conditions under which photocopies of parts of this publication may be made in the USA. All other copyright questions, including photocopying outside of the USA, should be referred to the publisher. No responsibility is assumed by the publisher for any injury and/or damage to persons or property as a matter of products liability, negligence or otherwise, or from any use or operation of any methods, products, instructions or ideas contained in the material herein.

Although all advertising material is expected to confonn to ethical (medical) standards, inclusion in this publication does not constitute a guarantee or endorsement of the quality or value of such product or of the claims made of it by its manufacturer. This book is printed on acid-free paper. Printed in The Netherlands

V

ACKNOWLEDGEMENTS

I would like t o thank the TNO Institute of Applied Geoscience, Delft, The Netherlands for continued support and cooperation during the preparation of the book. The use of facilities of the TNO Institute, including library and drafting services, is also gratefully acknowledged. I thank mr Jos Rietstap for drafting most of the illustrations. Special thanks are due t o mrs Gerda Boone of Gebotekst, Zoetermeer, who typed the different versions of the manuscript and assisted in the organization of the final version. Finally, I appreciate the constructive comments of the anonymous reviewer of Elsevier, on an earlier version of this book.

This page intentionally left blank

vii

PREFACE

The distribution of oil and gas in a sedimentary basin is determined by a combination of evolutionary processes that may have taken place during the geological history of the basin. The processes that determine the hydrocarbon potential of a basin are the generation, primary and secondary migration of hydrocarbons and the accumulation and preservation of hydrocarbons in traps. The hydrocarbon migration system influences the distribution, accumulation and nature of hydrocarbons in sedimentary basins. Hydrocarbon migration systems analysis is the integrated study of the evolution of hydrocarbon migration systems in sedimentary basins. This book has been prepared t o provide geoscientists - both university students and professionals interested in or working in hydrocarbon exploration - with a comprehensive understanding of the evolution of hydrocarbon migration systems and its application for basin evaluation. For this purpose, the book treats hydrocarbon migration in a multidisciplinary and explanatory manner. In order t o present geoscientists with the necessary basic understanding of fluid flow in sedimentary basins, the whole first part of the book is dedicated t o this subject. Chapter 1 reviews the fundamentals of single-phase fluid flowlgroundwater flow and associated processes of mass, energy and chemical transport. Chapter 2 considers groundwater flow on a basin-wide scale. It provides information on the processes controlling the development of groundwater flow systems, and the characteristic physico-chemical features associated with the different types of flow system. The second part deals with the hydrocarbon system. Chapter 3 gives relevant information on the processes of hydrocarbon generation and migration in source rocks. The principles of secondary hydrocarbon migration are outlined in Chapter 4. The theoretical concepts concerning hydrocarbon migration, given in this chapter, are derived from published literature. The different migration concepts have been integrated to give a more coherent and more generally valid picture of the evolution of secondary hydrocarbon migration systems in sedimentary basins. Chapter 4 pays ample attention to the influence of hydrodynamic conditions on secondary hydrocarbon migration. Chapter 5

viii

Preface

covers the processes of hydrocarbon accumulation, entrapment and preservation, and the influence of hydrodynamic and hydrocarbon migration conditions on these processes. The third and final part of the book presents a multidisciplinary approach to identify the present secondary hydrocarbon migration systems and the geohistory of secondary hydrocarbon migration systems in a sedimentary basin, and indicates how the results can be used for hydrocarbon exploration purposes. The presented secondary hydrocarbon migration systems analysis has been developed by extending a methodology that is applied successfully by hydrogeologists for the quantitative analysis of groundwater flow systems, taking into account the principles of groundwater flow, hydrocarbon migration, accumulation, entrapment and preservation as given in the first and second part of the book.

J.M. Verweij

ix

CONTENTS PART 1

FLUID FLOW

Introduction to single-phase fluid flow Chapter 1 1.1 Driving forces 1.2 Basic equations 1.2.1 Darcy’s equation 1.2.2 Continuity equations 1.2.3 Flow equations Large scale flow of groundwater 1.3 1.3.1 Applicability Darcy’s law 1.3.2 Continuity equations 1.4 S um m a r y Groundwater flow in sedimentary basins Chapter 2 Groundwater flow in actively filling and subsiding basins 2.1 2.1.1 Driving forces 2.1.2 Permeability distribution 2.1.3 Burial-induced groundwater flow system 2.2 Tectonically-induced groundwater flow 2.3 Groundwater flow in stable subaerial basins 2.3.1 Gravity-induced groundwater flow system 2.4 Local groundwater flow systems 2.4.1 Buoyancy-induced groundwater flow system 2.4.2 Osmotically-induced groundwater flow 2.5 Interaction of groundwater flow systems 2.6 Summary PART 2

3

3 5 5 6 10 11 12 15 21

23 26 26 28 34 51

55 55 70 70

74 75 78

GENERATION, MIGRATION AND ACCUMULATION OF HYDROCARBONS

Generation and expulsion of hydrocarbons Chapter 3 3.1 Origin of natural hydrocarbons 3.1.1 Organic matter Generation of hydrocarbons from organic matter 3.1.2 Generation of hydrocarbons from coal 3.1.3 Masses of generated hydrocarbons 3.1.4 Temperature and depth of hydrocarbon generation 3.1.5 Primary hydrocarbon migration 3.2 Primary hydrocarbon migration involving active groundwater 3.2.1 flow

83

85 85 86

91 93 91

97

99

Contents

X

3.2.1.1 3.2.1.2 3.2.1.3 3.2.1.4 3.2.2

Hydrocarbons in molecular solution Hydrocarbons in micellar solution Hydrocarbons in separate phase Conclusion Primary hydrocarbon migration independent of active groundwater flow 3.2.2.1 Continuous separate phase hydrocarbon migration Diffusion-induced hydrocarbon migration 3.2.2.2 Expulsion efficiency 3.2.3 3.3 Summary

Chapter 4 Secondary hydrocarbon migration Secondary hydrocarbon migration under hydrostatic conditions 4.1 Buoyancy 4.1.1 4.1.2 Capillary pressure 4.1.3 Separate phase hydrocarbon migration Secondary hydrocarbon migration under hydrodynamic conditions 4.2 Separate phase hydrocarbon migration 4.2.1 4.2.2 Migration of hydrocarbons in aqueous solution Regional aspects of secondary hydrocarbon migration 4.3 4.3.1 Changing conditions along the migration path 4.3.2 Secondary migration efficiency 4.3.3 Hydrostatic secondary hydrocarbon migration 4.3.4 Hydrodynamic secondary hydrocarbon migration Secondary hydrocarbon migration in actively filling and 4.3.4.1 subsiding basins Secondary hydrocarbon migration in stable subaerial basins 4.3.4.2 Secondary hydrocarbon migration in tectonically affected 4.3.4.3 basins Summary 4.4 Hydrocarbon accumulation, entrapment and preservation Hydrocarbon accumulation and entrapment under hydrostatic conditions Hydrocarbon accumulation and entrapment under hydrodynamic conditions Hydrocarbon accumulation and entrapment in hydrodynamic sedimentary basins Accumulation and entrapment in actively filling and subsiding basins Accumulation and entrapment in stable subaerial basins Accumulation and entrapment in tectonically active basins Preservation of trapped hydrocarbons Hydrocarbon preservation under stable geological conditions

99 103 103

104 105 105 111 115 119 121 122

122 125

127

134 135 140 140 141 144 145 148 149

154 157 158

Chapter 5 5.1 5.2 5.3 5.3.1

5.3.2 5.3.3 5.4 5.4.1

161 162 169 178

178 181 182 183 183

Contents

5.4.2 5.5

xi

Hydrocarbon preservation under changing geological conditions 187 Summary 189

PART 3

BASIN EVALUATION FOR HYDROCARBON EXPLORATION

Application to basin evaluation Chapter 6 Hydrodynamic condition, hydrocarbon migration and basin 6.1 evaluation 6.2 Hydrocarbon migration systems analysis 6.3 Data base 6.3.1 Pressure 6.3.2 Temperature 6.3.3 Chemical composition 6.3.4 Porosity and permeability Qualitative analysis of secondary hydrocarbon migration systems 7.1 Present-day hydrocarbon migration systems 7.1.1 Identification of the depocentres Identification of the hydrostatic hydrocarbon migration patterns 7.1.2 Identification of the hydrodynamic conditions 7.1.3 Hydrodynamic conditions in subaerial regions 7.1.3.1 7.1.3.2 Hydrodynamic conditions in subsiding and filling basins 7.1.4 Identification of the hydrodynamic influence on the hydrocarbon migration system History of hydrocarbon migration systems 7.2

193 193 197

m 201

2Q5 207

208

Chapter 7

211 211 212 212 213 214 219 221

224

Chapter 8 8.1 8.2 8.2.1 8.2.2 8.2.3 8.3 8.4

Quantitative analysis of secondary hydrocarbon migration systems Present-day hydrostatic hydrocarbon migration systems Present-day hydrodynamic conditions Hydrodynamic conditions in stable subaerial regions Hydrodynamic conditions in subsiding and filling basins Hydrodynamic conditions resulting from interactions of different groundwater flow systems Present-day hydrodynamic hydrocarbon migration systems History of hydrocarbon migration systems

227 229 232 232

239 241 a43 249

References

251

Subject index

269


PART I

FLUID FLOW


3

CHAPTER 1 INTRODUCTION TO SINGLE-PHASE FLUID FLOW

The pore spaces in the subsurface are normally water-saturated. In this book, all subsurface free H20-rich fluids are referred t o as groundwater. Groundwater is present throughout the crust to depths of at least 15 t o 20 km (Bredehoeft and Norton, 1990). Oil and gas accumulations are found between the ground surface and depth levels of about 6000 t o 7000 m and deeper (Tissot and Welte, 1984). The processes of generation, migration, accumulation and preservation of natural hydrocarbons take place in a water-saturated environment. The physico-chemical characteristics of groundwater may influence each of these processes. Knowledge on groundwater characteristics is indispensable to assess this potential influence. The theory on the flow of groundwater is based on the general principles of single-phase fluid flow through porous media. The same principles also apply t o the single-phase flow of hydrocarbons. The following sections treat the driving forces for the flow of fluids in the subsurface (Section 1.1),the basic equations describing single-phase fluid flow through porous media (Section 1.2) and regional aspects of fluid flow (Section 1.3) with respect t o groundwater.

1.1 Driving forces The driving forces for the flow of groundwater are groundwater potential gradients, temperature gradients, electrical gradients and chemical gradients (De Marsily, 1986; Freeze and Cherry, 1979). Ynder groundwater flow conditions, i.e. hydrodynamic conditions, the net force E, acting on a unit mass of water can be given by e,=-grad

@,-grad T-grad E-grad C

where, @W

= groundwater potential

(1.1)

4

Chapter 1

T = temperature of the groundwater E = electrical potential of the groundwater C = chemical potential of the groundwater Under hydrostatic conditions, there is no flow of groundwater and E, = 0.

-

The groundwater potential gradient is the main driving force for groundwater flow. A t a certain point the groundwater potential Qw, i.e. the mechanical energy per unit mass of groundwater, and the corresponding total head hT,i.e. the mechanical energy per unit weight, for groundwater whose density is a function of pressure only, is given by Hubbert (1953) as

where, QW

= potential of the groundwater

(LV2)

hT g

= total head of the groundwater (L)

z

= elevation of point of measurement above datum (L) = velocity of the groundwater (LT1) = pressure of the groundwater (ML-1T-2) = density of the groundwater (ML-3)

VW

PW Pw

= acceleration due to gravity ( L T 2 )

Because velocities of groundwater are extremely low, the term v&/g is generally considered t o be negligible and Z

P

qW= g h = g J d z + j dP w

(1.3)

pw where, h = hydraulic head (L). 20

PO

The potential of a unit mass of groundwater, as given by Equation 1.3 is thus determined by its place in the earth gravitational field and by the pressure and density of the groundwater. In groundwater practice, sea level and atmospheric pressure are taken as reference state, i.e. zo = 0 and po = 1 atmosphere. If, in addition, the density of the groundwater is constant (pwisnot a function of pw,i.e. the water is incompressible), the hydraulic head becomes h = z + - P- -w

(1.4)

Pwg

The gradient of the groundwater potential, as defined in Equation 1.3, is 1

grad Qw = g grad z + -grad Pw

pw

(1.5)

5

Introduction to single-phase fluid flow

which can be rewritten as (Hubbert, 1953) 1

-grad $, = 2 --grad

p,

(1.6)

Pw

The net driving force for groundwater flow that results from the groundwater potential gradient only, can thus be expressed as

2, =-grad

$w

1

= -g grad h = g--grad Pw

p,

(1.7)

-+

The direction of this force E, is perpendicular to the equipotenti-a1 surfaces of the groundwater. The water will be driven in the direction of E,, i.e. in the direction of decreasing potential.

1.2 Basic equations

The three-dimensional flow of groundwater through the subsurface can be described by a combination of Darcy’s equation for groundwater flow with a continuity equation ( o r mass balance equation) and equations of state for the groundwater and the porous medium. A detailed theoretical overview of equations for groundwater flow is given by e.g. Barenblatt et al. (1990), De Marsily (19861, Domenico and Schwartz (1990) and Freeze and Cherry (1979). Sections 1.2.1 t o 1.2.3 present the basic equations for groundwater flow through a representative elementary volume of the porous medium (Bear, 1972) under certain restrictive assumptions, such as isothermal and isochemical subsurface conditions and absence of changes in tectonic stress.

1.2.1 Darcy’s equation Darcy’s law (Darcy, 1856) is a phenomenological law that is valid for the viscous flow of a single-phase fluid (e.g. groundwater flow) through porous media in any direction. This basic law of fluid flow is a macroscopic law providing averaged descriptions of the actual microscopic flow behaviour of the fluids over some representative elementary volume of the porous medium. For isothermal and isochemical subsurface conditions, the law can be written as (Hubbert, 1953)

cw =--grad hw P W

where,

@,

=-K grad h

(1.8)

6 +

5. K +

k PW

Chapter 1

-(pwg)

= specific discharge of the groundwater (LT-l) = k = hydraulic conductivity of the water-saturated porous PW

medium (LT-') = intrinsic permeability of the solid matrix (L? = dynamic viscosity of the groundwater (ML-lT-l)

The intrinsic permeability and the hydraulic conductivity are tensorial properties. The intrinsic permeability is a measure of the resistance to flow of a fluid in a porous medium. It depends on the physical properties of the porous medium. The hydraulic conductivity is the amount of groundwater flow per unit cross-sectional area under influence of a unit gradient of hydraulic head. It depends both on the properties of the porous medium, such as the size, shape, orientation of its interstices and the interconnection of the interstices, and on the groundwater, namely its density and viscosity. The density and viscosity of the groundwater may vary with pressure, temperature and concentration of dissolved solids, thus influencing the hydraulic conductivity. The viscosity is the property of the internal resistance of the groundwater to flow. It is most strongly influenced by the temperature: the higher the temperature the lower the viscosity of the water, resulting in higher values of the hydraulic conductivity and consequently in higher values of the specific discharge of groundwater. In a homogeneous and isotropic porous medium, where the groundwater conditions are isothermal and isochemical, the hydraulic conductivity K of the subsurface will be constant and the flow of groundwater will be in the direction of - grad qW, i.e. perpendicular to the equipotential surfaces of the groundwater. However, in for instance an anisotropic porous medium K is not constant, and the directions of Gw and - grad @w may not coincide; the flow of groundwater has a tendency t o follow the direction of highest hydraulic conductivity (Davis and De Wiest, 1967; De Marsily, 1986). By introducing the appropriate expression for the hydraulic head, i.e. Equation 1.3 or 1.4, into Darcy's equation 1.8, it describes the flow for compressible and incompressible groundwater, respectively. The flow of compressible groundwater can thus be described by a more general form of Darcy's equation, in which the gradient of the groundwater potential is given by Equation 1.5

G~ = --{pwg i; PW

grad z + grad P,)

(1.9)

1.2.2 Continuity equations A continuity equation for fluid flow through a certain representative elemental volume of porous media fixed and rigid in space is based on the law

7


of conservation of mass, i.e. mass inflow rate of fluid = mass outflow rate of fluid + time rate of change of fluid mass storage. For steady-state groundwater flow, i.e. when the magnitude and direction of the flow is constant with time, the continuity equation can be written as (e.g. Freeze and Cherry, 1979) (1.10)

The parameter pw can be neglected if pw = constant and Equation 1.10 becomes (1.11)

For unsteady-state groundwater flow, i.e. when the magnitude and direction of the flow change with time, the equation of continuity is (1.12) where,

6t

= time-rate of change of groundwater mass storage per unit of volume

Provided that the following assumptions are met, the time-rate of change of groundwater mass storage can be related to changes in the density of the groundwater, in the porosity of the porous medium and in the vertical dimensions of the representative elementary volume of the porous medium, i.e. 6AM 6t

-=--

where, n VT

V" v s

Az

6p, 6(npw)-n-+p,-+-6t 6t

6n pWn6Az 6t Az 6t

(1.13)

= porosity = V" (dimensionless) VT = V, + V, = total unit volume of porous medium (L9 = volume of the voids (L3> = volume of the solid matrix (L3>

= vertical height of elementary volume of the porous medium (L)

The assumptions are:

- The subsurface conditions are isothermal and isochemical; - There are no time-changes in temperature and chemical subsurface conditions; There are no time-changes in stress externally imposed on the representative elementary volume of porous medium; - The porous medium is compressible and behaves as a linear elastic solid;

-

Chapter I

8

- Lateral deformations of the representative elementary volume of porous medium are negligible;

- The solid grains are incompressible; - The groundwater is compressible. Forces applied to a water-saturated porous medium will cause stresses which result in strain (deformation). The stress, strain and groundwater pressure in a water-saturated porous medium are coupled, as first recognized by Biot (1941). Under the assumed stress conditions, the vertical normal component of total stress (0,) that acts downwards on a horizontal plane a t any depth is caused by the weight of the overlying water-saturated rock. This stress is born by the solid matrix of the porous medium (0,)and by the pressure of the groundwater in the pores (pJ (e.g. Hubbert and Rubey, 1959) 0,=

oe + Pw

(1.14)

and do, = do, where,

+ dp,

Ge

= vertical normal component of total stress (ML-IT-2) = vertical effective component of total stress (ML-~T-~)

PW

= pressure of the groundwater (ML-IY2)

0,

(1.15)

The weight of the overlying water-saturated rock is assumed to be constant with time, hence do, = - dp,

(1.16)

The changes in vertical effective component of total stress are thus directly coupled to changes in pressure of the groundwater. When stress is applied to a unit mass of water-saturated porous rock, a change in the total volume of the rock mass can result from: - compression of the groundwater in the pores (a change in density of the groundwater), which is controlled by the compressibility of the groundwater,

-

P,;

compression of the individual solid grains (change i n density of the solid grains), controlled by the compressibility of the grains, p,; - rearrangement of the solid grains, controlled by the compressibility of the porous rock matrix, a. In general, compressibility can be defined as the rate of change of volume V with respect to an applied stress per unit of volume (Fertl, 1976). The isothermal compressibility of water, p, is given by


dVwNw - dPw/Pw doe dPw where, = isothermal compressibility of groundwater (M-lLT? P W = volume of groundwater = nV, (L3> VW = effective component of total stress ( M L - ~ Y ~ ) Oe pw=-

9

(1.17)

The time-change in density of the groundwater can be written as

-at

p

apw

w w

at

(1.18)

The time-change in density of the solid grains has been assumed t o be negligible. The compressibility of a porous rock matrix is (1.19)

Under the assumed conditions, the vertical compressibility of the porous medium, a,,can be given by a, = - d(Az)I Az - d(Az)I Az doe dPw

(1.20)

and the time-change of the vertical dimension of the elementary volume of the porous medium with pressure is ~ ( A z )- a,(Az)- 6Pw --

6t

6t

(1.21)

An increase in effective stress results in a reduction in the total volume of the rock mass dV, = dVv + dV,. Under the assumption that the solid grains are incompressible, the reduction dV, is caused by grain rearrangements, i.e. dV, = dV,. Hence

The time-change of porosity of the porous medium with pressure is given by Sn = a,(l -n)- 6Pw 6t st

(1.23)

Chapter I

10

By introducing the equations of state for the groundwater and the porous medium (Equations 1.18, and 1.21 and 1.23, respectively) into the continuity Equation 1.13 gives (e.g. Walton, 1970) (1.24) The groundwater pressure in a representative elementary volume of the porous medium fixed in space, changes directly with the hydraulic head, i.e. dpw = pwgdhand Equation 1.24 can be written as (1.25) where, ss

= PJZ

(a, + nPw)= vertical specific storage (L-l).

Jacob (1940) introduced the term specific storage. Expanding the terms on the left-hand side of Equation 1.25 by the chain rule, gives (1.26) aq are in general much greater than the Because the terms of the form pw 3

ax

terms of the form q,

h, Equation (1.26)is often simplified (e.g. Walton, 1970) ax (1.27)

1.2.3 Flow equations For steady-state flow of groundwater, a combination of the continuity Equation 1.11with Darcy's equation 1.8 yields (1.28) For flow through homogeneous isotropic media (K, = reduces to the Laplace equation d2h d2h a2h 2+--T+7=0 ax ay aZ

or V2h=0

= &), Equation 1.28

(1.29)


11

The solution h(x,y,z) describes the value of the hydraulic head at any point in a flow field. Introducing Darcy's equation 1.8 into the simplified continuity Equation 1.27 yields the general equation for unsteady-state flow through a water-saturated anisotropic medium under the assumed conditions listed in Section 1.2.2. (1.30) For flow through homogeneous and isotropic media, Equation 1.30 reduces t o the diffusion equation

s, ah V2h = -K at

(1.31)

The solution h(x,y,z,t) describes the hydraulic head at any point in a flow field at any time. The equations for steady and unsteady-state flow of groundwater (Equations 1.28, 1.29, 1.30, 1.31)can be solved for a particular hydrogeological situation by stating the appropriate initial and boundary conditions.

1.3 Large scale flow of groundwater In order to study the potential influence of groundwater flow on the distribution and characteristics of oil and gas accumulations, knowledge on the evolution of large-scale (basin-wide) groundwater flow systems is needed. The classical equations given in the previous sections describe the flow of groundwater through a representative elementary volume of porous medium under restrictive assumed conditions. On large temporal and spatial scales the actual conditions may deviate significantly from the assumed ones. On a regional scale, for example, the subsurface cannot considered t o be a homogeneous and isotropic porous medium and the subsurface conditions may not be isothermal and isochemical. When geological time scales are taken into account, changes in the direction and velocity of groundwater in large-scale systems may be influenced by time-changes in total stresses, thermal, hydrochemical and geochemical conditions. Time-changes of total stress accompany geological processes such as sedimentation, erosion and tectonic activity. The time-changes of temperature may result from changes in terrestrial heat flow, or may be caused by the displacement of a unit watersaturated porous medium along the geothermal gradient during e.g.

12

Chapter 1

sedimentation in a basin or uplift and erosion of a sedimentary basin. In turn, the chemical characteristics of both the groundwater and the rock matrix of the porous medium are changed in time as a result of changing stress and temperature conditions. Local groundwater flow conditions of importance in petroleum-related studies may also deviate from the general assumptions underlying the Laplace and diffusion equations, e.g. chemical and electrical gradients may influence the flow through low-permeable media (e.g. Neuzil, 1986). 1.3.1 Applicability Darcy’s law Darcy’s law applies to flow of groundwater relative to the solid matrix of a water-saturated porous medium (Fertl, 1976). For isothermal and isochemical subsurface conditions, Darcy’s law (Equation 1.8) is valid provided there is a linear relationship between the specific discharge of the groundwater and the gradient of the groundwater potential. A condition to be fulfilled is that the flow of groundwater should be purely laminar, i.e. not turbulent (e.g. De Marsily, 1986; Freeze and Cherry, 1979). In general, the natural flow of groundwater is laminar, except for flow through rocks with large-diameter (solution) openings. Whether or not Darcy’s law is valid for the relation between groundwater flow through a poorly permeable medium (K < 10-9m/s) and for low groundwater potential gradients 2 1is untested (Neuzil, 1986). Darcy’s law is valid for three-dimensional flow of groundwater through water-saturated porous and permeable media. On a large-scale the intrinsic permeability of the subsurface varies widely in space. The permeability is observed in nature t o vary over a t least 15 orders of magnitude (Bredehoeft and Norton, 1990; Figure 1.1).In groundwater studies, this spatial variability is generally converted t o averaged values of permeability in the two o r three principal directions of anisotropy (e.g. De Marsily, 1986). In a sedimentary basin, generally one main direction of anisotropy will be normal to the stratification and the other two parallel t o it. In fractured rocks, one principal direction of anisotropy will be in the direction of the main fractures, the other normal t o those fractures. The flow of groundwater in each of the principal directions of anisotropy is parallel to the hydraulic gradient in that direction. Consequently, by introducing the appropriate value of permeability in a principal direction of anisotropy into Darcy’s equation 1.8, the equation describes correctly the flow of groundwater in that direction. Darcy’s equation 1.8 describes the flow of groundwater as induced by hydraulic head gradients only. Darcy’s equation for groundwater flow can be generalized t o take the other driving forces for groundwater flow, i.e.: temperature gradients, electrical gradients and chemical gradients, into account (De Marsily, 1986)


-

Karst limestone Permeable basalt Fractured igneous and metamorphic rocksLimestone and dolomite -Sandstone

-

-

--

-

-

Unfractud metamorphic and igneous rocks Shale Unweatheredmarine clay Glaclal till

-

-

- -

-Silt.

loess-

-

Silty sand -Cleansand-

Gravel-

Figure 1.1 Ranges of permeabilities and hydraulic conductivities for different types of rock (modified after R. Allan Freeze and John A. Cherry, GROUNDWATER, 0 1979, p. 29. Reprinted by permission of Prentice Hall, Englewood Cliffs, New Jersey).

+

q, = -K, grad h - &grad T - K3grad E - K4grad C

(1.32)

where, K, (= hydraulic conductivity), K2, K, and K4 are phenomenological coefficients, which may be scalar o r tensorial. The flow of groundwater, electricity, solutes and heat through porous media are interdependent transport processes. The interdependency of the different processes is a reflection of the thermodynamic concept of coupled flow (Freeze and Cherry, 1979). One kind of flow influences each of the other kinds of flow, and all of the driving gradients given in Equation 1.32 are of influence on the flow of groundwater, heat, solutes as well as electricity. For example, the hydraulic head gradient does not only induce flow of groundwater, but also of heat, electricity and solutes. This coupling between groundwater flow, heat flow and flow of solutes is in part explained by the dependency of both the density and viscosity of groundwater (which appear in the Darcy equations 1.8 and 1.9) on the pressure, temperature and concentration of dissolved solids. Figure 1.2 illustrates the strong dependency of the viscosity of groundwater on the temperature. The influence of both temperature and pressure on the density of groundwater, omitting variations in chemical composition, can be given by the following equation of state (e.g. Palciauskas and Domenico, 1980)

14

Chapter 1

20

40

60

80

100

120

140

Temperature (“C)

Figure 1.2 Variation of the viscosity of water with temperature and NaCl concentration at a constant pressure of 10 MPa (modified after Garven and Freeze, 1984a, American Journal of Science, Volume 284, Figure 5. Reprinted by permission of American Journal of Science).

1

-dp,

= P,dp,

20

40

60

80

100

120

140

Temperature (“C)

Figure 1.3 Variation of the density of water with temperature and NaCl concentration at a constant pressure of 10 MPa (modified after Garven and Freeze, 1984a, American Journal of Science, Volume 284, Figure 4. Reprinted by permission of American Journal Science).

- a&dT

(1.33)

Pw

where, a, PW

= volumetric thermal expansion coefficient of the groundwater

(K-1)

= isothermal compressibility of the groundwater (M-lLT2)

The variation in density of the groundwater due to differences in concentration of dissolved solids and temperature is given in Figure 1.3. The spatial differences in groundwater density generate a buoyancy force for groundwater flow. The more general form of Darcy’s equation (Equation 1.9) may be used t o describe flow of groundwater of variable density under non isothermal conditions (see e.g. Garven and Freeze, 1984a; Hanor, 1987a; Harrison and Summa, 1991). It should be realized that for groundwater of


15

temperature- or salinity-induced variable density, the expression g grad z + 1grad pw Pw

is not equal anymore to the groundwater potential gradient, grad ,$, by Hubbert (1953)(Section 1.1).

as defined

The characteristics of the groundwater and the porous medium that appear in Darcy’s equation vary independently on a geological time scale. Equation 1.9 should be used when dealing with large temporal and spatial scales. 1.3.2 Continuity equations The conservation of groundwater mass in a large-scale groundwater flow system under steady-state flow conditions is given by continuity Equation 1.10. For unsteady-state large-scale flow of groundwater on a geological time scale, Equation 1.12 is valid. The time-rate of change of groundwater mass storage per unit of volume of saturated porous medium (6AW6t) results from the changing stress, thermal, hydrochemical and geochemical conditions in the watersaturated porous medium during its geological evolution. Hence, knowledge of the mechanical, thermal and chemical processes operating during the geological evolution of the water-saturated porous medium is indispensable for a correct understanding and assessment of the unsteady-state groundwater flow behaviour. Equations based on the concept of conservation are presented below for each of these processes to provide a general insight in the processes concerned.

Conservation of solid mass On a geological time scale the conservation of mass for the solid phase is determined by the deformation of the solid matrix and the accompanying timechanges of porosity, and the dislocation of the solid matrix in space. The deformation and dislocation of the solid matrix depend on the changing stress, groundwater pressure, temperature and chemical conditions. The changes in stress will induce both reversible elastic deformations of the porous medium as well as inelastic deformations (Houseknecht, 1987; Plumley, 1980). The stress-induced inelastic deformations may include macroscopic fracturing and dilatancy resulting from microfracturing (Palciauskas and Domenico, 1980). The pressure and temperature dependent processes, such as precipitation or dissolution of authigenic minerals, cause reversible changes of porosity, while irreversible changes of porosity are induced by pressure solution of grains at their points of contact, which occurs frequently in combination with precipitation on free surfaces of adjacent grains (Bj~rlykkeet al., 1989; Harrison, 1990; Houseknecht, 1987). On a geological time scale, the deformations of the porous medium are dominated by irreversible processes and can be considerably larger than the elastic ones for the same applied effective stress (Palciauskas and Domenico, 1989).

Chapter I

16

The general expression for the conservation of mass for the solid phase in a deforming medium may be given as (Palciauskas and Domenico, 1989) (1.34) where, Ps n vs

= density of the solid grains (ML-3) = porosity (dimensionless)

= velocity of the solid matrix with respect to fixed coordinates (LT-')

Conservation of heat The transport of heat in a water-saturated porous medium is governed by conduction in the solid matrix, transport by the groundwater (convective movement of heat) and heat exchange between the solid and the groundwater, which depends on their temperature difference (e.g. De Marsily, 1986). Assuming that, the viscous dissipation of energy is small and may be neglected, the adiabatic temperature changes are negligible, there is instantaneous thermal equilibrium between the solid and the groundwater, the thermal conductivity of the solid matrix is isotropic, a n expression for the conservation of heat in a representative elementary volume of water-saturated porous medium fixed and rigid in space is (Van der Kamp and Bachu, 1989) (1.35) where,

Fc %n

%W

%

T C

CW

T*

= thermal conductivity (MLT3K-') = thermal conductivity of the water-saturated porous medium

(MLPK-I) = n qw + (1-n) qS = thermal conductivity of the groundwater (MLY3K-') = thermal conductivity of the solid (MLPK-') = temperature (K) = specific heat (LV2K-') = specific heat of the groundwater (Lv2K-l) = specific heat of the solid (LV2K-') = heat capacity of the water-saturated porous medium (ML-'T2K-I) = npwcw+ (1-n) psc, = heat sink or source term

The heat sink o r source term represents the local sources of heat produced by radio-active decay in the porous medium and the time-change in basal heat flux.


17

Conservation of chemical mass The conservation of chemical mass in the subsurface is mainly controlled by the transport of solutes through the water-saturated porous medium, the chemical reactions in the groundwater and the reactions between the groundwater plus solutes and the solid matrix. The solutes are transported through the subsurface by convection, molecular diffusion and dynamic dispersion. Convection of solutes involves the displacement of dissolved chemical components by the flowing groundwater. Molecular diffusion is the mass flux of chemical components that takes place through both the mobile and immobile water in the porous medium as resulting from the chemical concentration gradients only. The dynamic dispersion is a physical mixing process which results from the heterogeneity of the velocity of groundwater in the porous medium (e.g. De Marsily, 1986). In the subsurface, there are variations in groundwater flow velocity inside the pores, variations in groundwater flow velocity and direction between pores and those resulting from larger-scale heterogeneity of the porous medium.

A general expression for the conservation of solute mass for a certain chemical component in a representative elementary volume of water-saturated porous medium fixed and rigid in space is (e.g. Garven, 1985; Garven and Freeze, 1984a) (1.36) where, D C

= coefficient of dispersion (a tensor) (L-1) = mass of a single chemical solute per unit volume of aqueous

C*

solution (ML-3) = net source or sink term (ML-3T-1)

The coefficient of dispersion in this advection-dispersion-reactionequation describes the combined processes of molecular diffusion and hydrodynamical dispersion (see e.g. Bear, 1972; De Marsily, 1986). The term C* indicates the disappearance or addition of a chemical component. It represents the influence of changes, exchanges and reactions of chemical components during their subsurface transport. The transport of chemical components may be stopped by physical (membrane) filtration (Bredehoeft et al., 1983; De Sitter, 1947; Graf, 1982; Neuzil, 1986). Many different types of geochemical reaction can be involved in changing the concentration of a certain solute, e.g. precipitation I dissolution; sorption via surface complexation (adsorption) o r sorption via ion exchange I desorption; oxidation I reduction; acid base reactions, complexation, radio-active decay (e.g. De Marsily 1986; Garven, 1985; Garven and Freeze, 1984a). Additional equations governing this geochemical mass transfer are

Chapter 1

la

needed to provide the proper values of C* in Equation 1.36 (see e.g. De Marsily, 1986; Garven and Freeze, 1984a; Yeh and Tripathi, 1989). The conservation of chemical mass in a multi-component groundwater flow system is given by a set of advection-dispersion-reactionequations, i.e. an appropriate equation for each chemical solute of interest. The equations describing the conservation of groundwater mass (Equation 1.12), solid mass (Equation 1.341, chemical mass (Equation 1.36) and heat (Equation 1.35) and Darcy's equation for groundwater flow (Equations 1.8 and 1.9) are coupled and nonlinear (Bredehoeft and Norton, 1990; Garven, 1985).

Conservation of groundwater mass The influence of the mechanical, thermal and chemical processes on the flow of groundwater can be introduced into the continuity equation for groundwater (Equation 1.12) assuming appropriate simplifying conditions. For example, if the assumptions listed in Section 1.2.2 are met, with the exception of the second assumption which is replaced by - There are time-changes in temperature and in total stress imposed on the elementary volume of porous medium; - There are no time-changes in chemical conditions; And, - The thermo-mechanical effects on the solid part of the water-saturated porous medium are negligible, the time-rate of change of groundwater mass per unit of volume of watersaturated porous medium can be given by (Sharp, 1983) (1.37) Taking into account the compressibility of the solid grains, Equation 1.37 may be written as (see Neuzil, 1986; Van der Kamp and Gale, 1983)

-

-

1 6pw 60 6T -V.(pwqw)= -s, -- p w ( a - a s6t) - - p w n a ~ ~ g where, = three dimensional specific storage (L-1) S S = pwg[(a - ad + n (Pw - ad] = bulk compressibility of the porous medium (LTW-1) a = bulk compressibility of the solids (LTM-1) a, = compressibility of the groundwater (LT?M-1) PW

s}

(LTW-~)

a

= az{

V

= Poisson's ratio for the porous medium (dimensionless) = mean total stress (M'PL-1) = thermal expansion coefficient of the groundwater (K-1)

ot

aTw

(1.38)


19

In order to account for additional changes in groundwater mass storage by processes not included in the terms at the right-hand side of Equation 1.38, a source or sink term (pwQ;) can be added to this equation. Such a source or sink term may account for e.g. the dehydration of gypsum and the release of interlayer water from clay minerals as a function of temperature, pressure and groundwater chemistry (e.g. dehydration of smectite, kaolinite) (e.g. Bethke, 1986b; Bjprrlykke, 1989; Buhrig, 1989; Colten-Bradley, 1987; Powers, 1967); the introduction of juvenile water from magmatic sources; the reversible change of porosity due to chemical processes and irreversible change of porosity by intergranular pressure solution; and irreversible change of porosity due to stress-induced inelastic deformation. Strictly speaking, Equation 1.38 is a continuity equation for groundwater flow through a certain representative elemental volume of porous medium fixed and rigid in space. For small elastic deformations of the porous medium, the equation can be considered t o be valid provided that the specific discharge of groundwater is taken as relative to the rock grains (Cooper, 1966 and e.g. Neuzil, 1986; Sharp, 1983). When large, possibly inelastic, deformations are involved, i.e. when the representative elementary volume of porous medium is both deformed and dislocated in space, the dislocation of the porous medium with respect t o fixed coordinates should be taken into account as well (e.g. Bayer, 1989; De Marsily, 1986; Palciauskas and Domenico, 1989; Shi and Wang, 1986). So as to take the time-change in the location of the deforming representative elementary volume of porous medium into account, the continuity equations are often given in terms of material derivatives (e.g. De Marsily, 1986). The material derivative of a property (d( Ydt) in a deforming coordinate system with velocity v, is related to the derivative in a fixed coordinate system (6( Y6t) by (1.39)

For example, the general continuity equation for groundwater (Equation 1.12; 6AM/6t = 6npJ6t) written in terms of material derivatives in a deforming coordinate system following the motion of the solids, becomes (e.g. Palciauskas and Domenico, 1989; Shi and Wang, 1986) (1.40) where, VS

= velocity of the solid matrix with respect to the fixed coordinates (LT-1)

qw

= specific discharge of groundwater relative t o the rock grains (LT-1)

Chapter 1

20

Table 1.1 Symbolic equations representing the processes of mass, energy and chemical transport related t o groundwater flow

Darcv's eauation for Proundwater flow Gw

i;

=--bwg PW

grad z + grad Pw)

Conservation of groundwater mass steady state groundwater flow condition

unsteady state groundwater flow condition

-.

6pw

-V.(pwqw)=-s,-g 6t

p w ( a - " , ) ~60 - p w n c r T w - +6T pwQ,,

6t

6t

Conservation of solid mass -V.[p, (1- n)3,1= 6[p, (1- n)l

6t

Conservation of heat V.$,

VT - cwpWGw VT + T* = (PC),

Conservation of chemical ma=

Eauations of state

pw = f(P, T, C) Pw = f(P, T, C)

fl 6t


21

1.4 Summary

The flow of groundwater in a sedimentary basin results from the combined influence of the different driving forces for groundwater flow (mechanical, thermal, chemical and electrical driving forces) and the hydraulic conductivity of the subsurface. The transport of groundwater, heat and electricity, the mass transport of chemical components and the deformation of the solid part of the subsurface are coupled processes. Table 1.1 gives an overview of the equations that symbolically represent the processes of mass, energy and chemical transport related to groundwater flow. The equations in Table 1.1are coupled and nonlinear.

22

Table 2.1 Five main types of sedimentary basins

I.

Divergent margin basins A. Rift basins 1. Rifted arch basins 2. Rim basins 3. Sag basins 4. Half-graben B. Ocean margin basins 1. Red Sea type (youthful) 2. Atlantic type (mature) C. Aulacogens and failed rifts D. Oceanic islands, seamounts, plateaus 11. Convergent margin basins A. Trenches and subduction complexes B. Forearc basins C. Interarc and backarc basins D. Retroarc (foreland) basins 111. Transform and transcurrent fault basins A. Basin setting 1. Plate boundary transform fault 2. Divergent margin transform fault 3. Convergent margin transcurrent fault 4. Suture zone transcurrent fault B. Basintype 1. Basins in braided fault systems 2. Fault termination basins 3. Pull-apart basins in en echelon fault systems 4. Transrotational basins IV. Basins developed during continental collision and suturing A. Peripheral (foreland or foredeep) basins (on underriding plate) B. Intrasuture embayment basins (remnant ocean basins) C. Hinterland foreland, strike-slip, and graben basins (on overriding plate) V. Cratonic basins From: Miall, 1990. Reprinted with permission of Springer-Verlag.

CHAPTER 2 GROUNDWATER FLOW IN SEDIMENTARY BASINS

A sedimentary basin is a subaquatic or subaerial region on the earth surface in which sediments have accumulated at a greater rate and to a greater thickness than they have in adjacent areas. Sedimentary basins form by deformation of the lithosphere mainly as a result of crustal extension during divergent plate movements and by compression and crustal thickening during convergent plate movements (extension may also occur during convergent movements) (Miall, 1990). The evolution of a sedimentary basin (i.e. its subsidence history; basal heat flow history; evolution of its physical size and shape, style of deformation, bottom configuration, water depth and sedimentary fill) is related to the successive plate-tectonic settings of the basin, the geological history of the plate margin processes and the successive latitudinal (i.e. climatic) settings of the basin (see Miall, 1990). Sedimentary basins are classified on the basis of the control of plate-tectonic processes on their evolution (Allen and Allen, 1990; Miall, 1990). Table 2.1 gives a classification of sedimentary basins based on their plate-tectonic setting. Similar types of basins may show a consistent pattern in their sedimentary evolution, whilst basins of different type show correspondingly different sedimentary styles (see Allen and Allen, 1990 and Miall, 1990). Most sedimentary basins cover tens of thousands of square kilometres and may contain a thickness of over five kilometres of sedimentary fill (Selley, 1985). Numerous studies demonstrate the existence of groundwater flow on a basinal scale (Beck et al., 1989; Bethke and Marshak, 1990; Bredehoeft et al., 1988; Chiarelli, 1978; Garven, 1985, 1989; Ge and Garven, 1989; Goff and Williams, 1987; Harrison and Summa, 1991; Lloyd and Jacobson, 1987; Oliver, 1986, 1992; Torgerson, 1990; T6th, 1978, 1980; Verweij, 1990). A large-scale groundwater flow system in a sedimentary basin can be described by the threedimensional pattern of groundwater flow in combination with the physicochemical characteristics of the groundwater. Different types of groundwater flow system and different parts of a single groundwater flow system are associated with characteristic physico-chemical features. This is because the mass transport of chemical compounds, the transport of heat, the deformation of the solid part of the subsurface and the flow of groundwater are coupled processes (Chapter 1).The groundwater flow pattern, i.e. the directions and

24

Chapter 2

velocities of the groundwater in a basin, results from the combined influence of the different forces driving groundwater flow (mechanical, thermal, chemical and electrical driving forces; Section 1.3) and the subsurface distribution of hydraulic conductivities. The main classification criterion for large-scale groundwater flow systems is the dominating driving force or forces for groundwater flow. The forces inducing basin-wide groundwater flow systems are principally controlled by the following processes: - sedimentation in a subsiding sedimentary basin - introduction of heat into a basin - tectonic processes acting on a basin - infiltration of meteoric water in a subaerial basin. Each process can be associated with a particular type or types of groundwater flow system. The characteristics of a groundwater flow system are also controlled by the subsurface permeability distribution. The subsurface permeability distribution a t a certain time during the evolution of a sedimentary basin is given by the hydrogeological framework of the basin at t h a t time. This framework is characterized by the distribution, interconnectivity, thickness and dip of porous and permeable hydrogeological units (aquifers/potential carrier-reservoir rocks, e.g. sands, sandstones, carbonates, fractured rocks) and poorly permeable hydrogeological units (aquitardslpotential barrier rocks, e.g. shales, evaporites) (Figure 2. l), and by the location of geological structures and tectonic elements of importance for groundwater flow, e.g. permeable or impermeable faults, unconfomities. The hydrogeological framework of a basin is determined by the original nature of its sedimentary fill and the syn- and postdepositional mechanical, thermal and chemical deformations of the sediments. Similar types of sedimentary basin can be expected t o show corresponding hydrogeological frameworks. Processes like subsidence, heat flow and style of deformation of a basin greatly influence the main driving forces for groundwater flow. A certain type of sedimentary basin, associated with a particular history of subsidence, heat flow and deformation, may thus be related t o a particular hydrogeological framework and to a particular dominating driving force or particular evolution of different dominating driving forces for groundwater flow. The main large-scale groundwater flow systems that may develop during the different stages of evolution of a sedimentary basin are described in Sections 2.1 to 2.3.These sections also present some examples of the relation between a basin type and type of groundwater flow system. Section 2.4 gives an overview of local groundwater flow systems of interest in petroleum-related studies. When the characteristics of a groundwater flow system are adjusted to the prevailing boundary conditions and remain constant in a sedimentary basin during a certain period, the groundwater flow system is said to be in steady state. Different groundwater flow systems may coexist and interact in the same

Groundwater flow in sedimentary basins

25 East

west

upperaquifer

Deep-basin brine aauife

Shelf limestone and chert Ordovician

Lower Paleozoic carbonate aquifer

aquitad

Shelf dolomite

Shallowmarine('?) sandstone

Lower Paleozoic sandstone aquifer

Figure 2.1 Generalized lithostratigraphic units and corresponding hydrogeological units of the Palo Duro Basin, USA (modified after Bassett and Bentley, 1982. Reprinted by permission of Elsevier Science Publishers BV).

Chapter 2

26

sedimentary basin (Section 2.5). During the different stages of evolution of a sedimentary basin, different groundwater flow systems will develop. While the characteristics of the groundwater flow system change, the groundwater flow system is in unsteady state. Under unsteady state conditions relict groundwater flow systems may exist in the basin.

2.1 Groundwater flow in actively filling and subsiding basins 2.1.1 Driving forces The continued sedimentation in a subsiding basin is the overall driving force for the large-scale groundwater flow system in the basin. The hydrogeological framework of the basin strongly influences the actual physical characteristics of this basin-wide burial-induced groundwater flow system, such as the distribution of groundwater potentials and pressures and the directions and velocities of groundwater flow. Both the overall driving force for groundwater flow and the hydrogeological framework are time-dependent in an actively filling and subsiding basin. The groundwater flow at a certain time during the basin’s evolution can be described by the equations for unsteady state groundwater flow given in Section 1.3. Assuming that hydrostatic conditions prevail in a n isothermal and isochemical basin, the net force acting on a unit mass of water is zero at any point in the basin and (Section 1.1) e,=-grad

+w=Z-

grad Pw

=o

(2.1)

Pw

Hence,

The increase of groundwater pressure with depth (grad p,) follows the hydrostatic pressure gradient, which is numerically equal to p,g (Figure 2.2)) i.e. the hydrostatic pressure gradient varies with groundwater density. The vertical normal component of total stress (az)caused by the weight of the overlying water-saturated rock at a certain depth in the basin is borne by the solid matrix of the porous subsurface ( 0 , ) and by the pressure of the groundwater in the pores (p,):


27

The increase in ozwith depth is numerically equal to g { p s + ps (1- n)) and is known as the lithostatic gradient (grad o, = 23 MPa/km). Sedimentation in the basin results in an increase in the vertical component of total stress (do,) at that depth, which initially affects only the groundwater pressure (e.g. Bayer, 1989),i.e. do, = dp, creating a superhydrostatic pressure of the groundwater (also known as overpressure, excess pressure , supernormal or abnormally high pressure of the groundwater). This pressure increase, dp, is caused by the weight of both the water and the solid rock of the deposited sediments. In a permeable subsurface, the initial change in groundwater potential will induce a flow of groundwater until the groundwater pressure has returned to the hydrostatic pressure for the depth in question, i.e. until the groundwater pressure increase caused by sedimentation reflects only the weight of the water in the deposited sediments. The dissipation of the superhydrostatic pressure of the groundwater leads to an increase in effective stress and consequently to a reduction in the volume of the water-saturated rock. Hence, the process of mechanical compaction of the water-saturated sediments caused by an increase in the gravity load of the sediments is time-dependent and is influenced by the rate of increase in total stress resulting from sedimentation and by the rate of dissipation of superhydrostatic pressures of the groundwater (i.e. the rate of groundwater flow out of the sediments as controlled by the permeability of the sediments and the viscosity of the groundwater). In a sedimentary basin, both the groundwater pressure and temperature increase with depth: increasing pressure tends t o reduce the volume of a given weight of water and increasing temperature tends t o increase its volume. The effect of water expansion with increasing temperature is more pronounced, resulting in water expanding the deeper it is buried. This aquathermal effect was first identified by Barker (1972). The temperature-induced increase in depth

fresh water: p , - 1000 k g / m 3 p, g 10,000 Pa/m

-

brine:

pz

1000-

2000-

3000-

L

pressure (MPa)

Figure 2.2 Hydrostatic water gradient.

- 1200 kg/m3

28

Chapter 2

groundwater pressure resulting from this aquathermal effect is known as aquathermal pressuring and may occur at all depths. Additional mechanisms of groundwater pressure generation that are induced by the increasing pressure and temperature conditions during the burial of the sedimentary rocks, include the generation of hydrocarbons from organic matter and the resulting volume expansion, and the dehydration of minerals such as smectite, kaolinite and gypsum. The volume increase caused by the generation of hydrocarbon gases, which occurs in the temperature range of 150 - 220 "C (Chapter 3), is especially important. The stability of smectite and mixed layer clays decreases with increasing temperature and depth. Smectite may transform t o illite a t 60 - 100 "C, whereas kaolinite becomes increasingly unstable between 120 - 150 "C(Bj~rlykkeet al., 1989). The dominant mechanism of pressure generation in a filling and subsiding basin seems to be mechanical pressuring of groundwater caused by sedimentary loading (Bethke, 1985, 1986b; England et al., 1987; Harrison and Summa, 1991; Keith and Rimstidt, 1985; Shi and Wang, 1986). Shi and Wang (1986) quantitatively analysed the relative importance of mechanical pressuring versus aquathermal pressuring in the generation of groundwater pressure. They found that under normal geological conditions in sedimentary basins, the mechanical pressuring is the main mechanism in the generation of groundwater pressures. Studies by Bethke (1985, 1986b) agree with Shi and Wang's conclusion. Bethke's (1985) study showed that the aquathermal effect is probably of only limited importance in generating superhydrostatic pressures within slowly subsiding basins (c 1%of the calculated total superhydrostatic groundwater potentials). The aquathermal effect was also found t o be a less important cause of the abnormally high pressures that have developed in a rapidly subsiding shaly basin like the Gulf of Mexico Basin, USA (Bethke, 1986b; Harrison and Summa, 1991). Sharp (1983) showed that under special conditions, such as steep geothermal gradients and very poorly permeable rocks, the effect of aquathermal pressuring of the groundwater may be a significant supplementary process of pressure generation. Locally, at greater depths the groundwater pressure may be influenced by massive gas generation (e.g. Buhrig, 1989). The significance of dehydration of minerals on the generation of groundwater pressures in a sedimentary basin, or in certain parts of a basin, is disputed (e.g. Bethke, 1986b, Colten-Bradley,1987). 2.1.2 Permeability distribution The flow of groundwater in an actively filling basin is greatly influenced by the space- and time-dependent hydraulic conductivity of the subsurface. The permeability of an isotropic porous sediment is related to its porosity and grainsize distribution (Figure 2.3; Chilingarian and Wolf, 1975). Figure 2.3 shows that for the same porosity, coarser-grained well sorted sediments will have greater permeabilities than fine-grained sediments. In general, the


'0 888 Z 6 LXMJ2 4000

1

1 0

Coarse- and very coarse-grained Coarse- and medium-grained

0 0

A Fineqralned

1

1

1

1

1

1

1

1

2

4

6

8

10

12

14

16

1

1

1

1

1

1

1

1

18 20 22 porosity W)

1

1

24

26

28

30

32

34

36

Figure 2.3 Relationship between porosity and permeability of very coarse-grained, coarsegrained, medium-grained, silty and clayey sandstones (after Chilingar, 1964).

porosity of a sediment decreases with depth in a sedimentary basin (Figure 2.4). The porosity of a sediment decreases with depth of burial, leading t o an associated decrease of its permeability. However, because temperature increases with depth, the groundwater will become less viscous. Under normal geological conditions, the decrease in permeability with depth is greater than the viscosity effect (Sharp, 1983), and the hydraulic conductivity will also decrease with depth of burial. The decline in porosity of a sediment with depth of burial may result from mechanical compaction, intergranular pressure solution and cementation. Porosity loss due t o mechanical compaction or pressure solution is a function of effective stress, i.e. of total stress and groundwater pressure. Cementation by the precipitation of authigenic minerals (quartz, calcite) is influenced by the chemical composition and flow of groundwater. The groundwater flow condition in the basin thus affects all three processes that reduce the porosity of sediments with depth of burial. Mechanical compaction is the dominant process responsible for porosity reduction in argillaceous sediments (Rieke and Chilingarian, 1974). The porosity reduction in coarse-grained sediments like sands may be influenced by mechanical compaction, pressure solution and cementation (Bjorlykke et al.,

Chapter 2

30

1989; Houseknecht, 1987). The porosity of carbonate sediments and rocks is reduced by cementation, recrystallization, and mechanical and chemical compaction (Mazzullo and Chilingarian, 1992).

Compaction o f fine-grained sediments In a homogeneous isotropic basin overlying an impermeable base, the first burial of newly deposited fine-grained rocks will initially result in the compaction of the fine-grained rocks as a commensurate volume of the relatively incompressible groundwater is expelled. Under these compaction equilibrium conditions, the pressure of groundwater is near hydrostatic. The burial-induced flow of groundwater will be directed principally vertically upwards (Figure 2.5; Einsele, 1976, 1977; Magara, 1978). The rearrangement of clay mineral particles during compaction results in a preferred orientation of

10

70

30

tJOROSITY ( % )

Figure 2.4 Porosity-depth curve for sandstones, carbonates and shales (from Harrison and Summa, 1991, American Journal of Science, Vol. 291, Fig. 8. Reprinted by permission of American Journal of Science).


31

the clay particles because of their characteristic platy shape (Rieke and Chilingarian, 1974).The clay becomes vertically anisotropic during compaction (k, < kh). This vertical anisotropy in combination with overall decreasing permeability with increasing depth will eventually restrict the movement of groundwater in the fine-grained sedimentary rocks. When the expulsion of groundwater can no longer keep pace with subsidence at a certain depth, the pressure of the groundwater cannot dissipate and, consequently, the increase in effective stress slows down. As a result, the compaction of the fine-grained rocks will be delayed and the rocks will become undercompacted for the depth in question. The actual depth at which compaction disequilibrium conditions will occur is influenced by the sedimentation rate and the actual permeability of the fine-grained rocks (see e.g. Hunt, 1979; Rieke and Chilingarian, 1974; Smith, 1971). The reduction of porosity and permeability with increasing porosity 0.L

0.5

0.6

0.7

0.8

ho

height of layer 0

M,

height of mineral substance o f layer 0

r

burial induced groundwater flow as caused by sediment load o f layer ho porosity depth relation

-

relation between Zh and EM hemipelagic carbonaceous silty clay

imentation ( c m A 0 0 0 years)

0

100

200

300 EM (m)

Figure 2.5 Burial-induced vertically upward directed groundwater flow under conditions of compaction equilibrium (modified after Einsele, 1977. Reprinted by permission of Blackwell Scientific Publications Ltd.).

32

Chapter 2

effective stress caused by overburden load also depends on the loading path of the sedimentary rocks (Jones and Addis, 1985; Shi and Wang, 1986). When reburial of sedimentary rocks starts, the rocks will already be of reduced porosity and permeability. During reburial, overpressuring of the groundwater will start at shallower depths. In addition, if the effective stress during reburial is less than the maximum effective stress was in the past, the porosity will change much more slowly with the increasing effective stress during reburial than i t did as a result of mechanical compaction during first burial (Shi and Wang, 1986).

Compaction of coarse-grained sediments The evolution of porosity and permeability of coarse-grained sedimentary rocks during burial may be influenced by mechanical compaction, pressure solution and cementation. This evolution depends on the original mineralogical composition, the sorting and the packing of the sediments and the effective stress and groundwater flow conditions during burial (e.g. Bjerlykke et al., 1989; Harrison, 1990; Houseknecht, 1987). The following examples of the influence of mineralogical composition of coarse-grained sediments (sands) on the evolution of their porosity is taken from Bjerlykke et al. (1989) (Figure 2.6). Quartzarenites and subarkoses are generally less vulnerable to compaction than lithic arkoses and feldspathic litharenites. In sandstones with a high percentage of stable grains (i.e. quartzrich sandstones) the porosity decline resulting from mechanical compaction is relatively slow during shallow to moderate burial (< 3 km). A t greater depths (3 - 4 km) the porosity decline is accelerated because of increased pressure solution and stylolitization. The presence of clay laminae and micaceous layers enhance pressure solution, causing stylolites to form. The porosity of sandstones with more than 20 - 25% unstable grains will be reduced at moderate depths of burial because of the collapse of the grain framework. In contrast t o the fine-grained rocks, the initially permeable coarse-grained rocks allow enough throughflow of groundwater t o either create secondary porosity by leaching chemically unstable minerals (from rock grains or cement) (McDonald and Surdam, 1984) or to reduce porosity by cementation (silica, carbonate, authigenic clay cement). Authigenic illite, montmorillonite and kaolinite may form aggregates which reduce the permeability considerably. The remaining porosity in sandstones with early diagenetic silica or carbonate cement may be preserved relatively better during subsequent burial (Bjwlykke et al., 1989).Secondary porosity may also have a greater preservation potential than primary porosity during subsequent burial (Bj~rlykkeet al., 1989).

Groundwater flow in sedimentary basins Porosiry, 36

33 Estimated primary porosity I

0

1

2

E

Y

c;n. B

3

4

Figure 2.6 Influence of mineralogy of sandstones on porosity-depth gradient based on data from Nagtegaal, 1978 (from ELEMENTS OF PETROLEUM GEOLOGY by Robert C. Selley. Copyright (0)1985 by W.H.Freeman and Company. Reprinted by permission).

Compaction of carbonate sediments and rocks The compaction of carbonates with depth is more complex than that of sandstones, because carbonate minerals (calcite, aragonite, dolomite) are chemically less stable than silica (e.g. Chilingarian et al., 1992). At shallow burial the post-depositional evolution of porosity and permeability of the carbonate sediments is greatly affected by porosity reducing processes, such as mechanical compaction, internaI sedimentation and marine and meteoric cementation, and porosity creating processes including aqueous dissolution of carbonates (Mazzulo and Chilingarian, 1992). The processes of cementation and dissolution are influenced by the groundwater flow conditions in the basin. For example, because carbonate sediments are deposited in a shallow water environment, the sediments may come into contact with meteoric groundwater soon after their deposition. The carbonate sediments are very susceptible to leaching when invaded by acid meteoric water. The leaching creates secondary porosity. Secondary porosity may also be created by dolomitization in the mixing

34

Chapter 2

zone between meteoric and synsedimentary marine groundwater. Lithification of carbonate sediments through reprecipitation of the carbonate minerals and dolomitization may occur already during shallow burial (e.g. Chilingarian and Wolf, 1975). During deeper burial of newly deposited carbonate sediments, the primary and secondary porosity is decreased by cementation and chemical compaction. A t these deeper burial depths pressure solution causes the sedimentary grains to dissolve and cement, and stylolites to form. Stylolites may start to form at depths of 1 t o 2 km (Bj~rlykke,1989). Early formed carbonate cement may hamper later pressure solution, i.e. carbonate sediments which have been subject t o relatively early cementation may retain their remaining porosity better with depth (Bj~rlykke,1989). Aqueous dissolution of carbonates may also create secondary porosity in carbonate rocks at deeper burial. The complex evolution of porosity in carbonate sediments and rocks is reflected in the extreme lateral and vertical heterogeneity of carbonate rocks (Mazzullo and Chilingarian, 1992). 2.1.3 Burial-induced groundwater flow system The combined result of the time-dependent processes of groundwater pressure generation and dissipation determine the groundwater pressure distribution and the other associated characteristics of the burial-induced groundwater flow system at a certain moment in an actively filling and subsiding basin. The groundwater pressures in a filling sedimentary basin are generated by the combined effect of the increase in load of the water-saturated sediments and the aquathermal effect, which at greater depths is enhanced by the dehydration of clay minerals and hydrocarbon generation from organic matter. Mechanical pressuring of groundwater as a result of sedimentary loading is probably the dominant mechanism of groundwater pressure generation. In a basin consisting of homogeneous isotropic sedimentary rocks overlying an impermeable base, the burial-induced groundwater flow will be directed principally vertically upwards under compaction equilibrium conditions (Figure 2.5; Einsele, 1976, 1977; Magara, 1978). According to Bonham (1980) this groundwater flow is upward across chrono-stratigraphic units, but downward with reference to the surface of deposition.The velocity of burial-induced groundwater flow will diminish with depth. A homogeneous basin consisting of isotropic fine-grained sedimentary rocks will maintain compaction equilibrium when the sedimentation rate is slow. However, when the sedimentation rate is rapid, the expulsion of the groundwater from the sedimentary rocks at a certain depth can no longer keep pace with subsidence, because of the rapid burial of the rocks and the already reduced permeabilities of the rocks. Under compaction disequilibrium conditions the groundwater is overpressured and the compaction of the sedimentary rocks will be delayed (the


35

sedimentary rocks will be undercompacted) and consequently the basin may no longer be homogeneous.

No perfectly homogeneous isotropic sedimentary basin exists in nature, though certain parts of existing basins may be considered homogeneous, favouring burial-driven vertical upward flow of groundwater. In general, in an actively filling and subsiding inhomogeneous sedimentary basin consisting of alternating fine-grained and coarse-grained sedimentary rocks, most groundwater tends to move vertically upwards during the earliest burial stages, i.e. in the relatively shallow part of the basin ( 1 m d y e a r will induce groundwater pressures to approach lithostatic gradients. The shallow and intermediate subsystems of burial-induced groundwater flow may develop in shaly basins with moderate subsidence rates (0.1 mm - 1 mdyear). In rapidly subsiding shaly basins (burial rates > 1 mdyear), all three subsystems may occur.

40

Chapter 2

Figure 2.11 Variable depths to the geopressured zone in the Gulf of Mexico Basin, USA, modified from Wallace and others, 1979 by Harrison and Summa, 1991 (after Hamson and Summa, 1991, American Journal of Science, Vol. 291, Fig. 3. Reprinted by permission of American Journal of Science).

The actual depths at which the subsystems occur will differ for different basin types and will vary within a single basin. This can be illustrated by the variable depth t o burial-induced geopressures within, for example, the North Sea Basin (Buhrig, 1989; Cayley, 1987) and within the Gulf of Mexico Basin (Harrison and Summa, 1991; Figure 2.11). Fertl (1976) presents examples of the extremely variable depths of occurrence of geopressured zones for basins around the world. Some of these examples are in actively filling and subsiding basins. The rate of subsidence differs for different types of sedimentary basin. The rate of subsidence of cratonic basins is in the order of 10 to 20 m per million years, and basins in mobile belts (foredeep, intradeep, backdeep, marginal basins and associated troughs) subside at rates in the order of 50 or 100 m per million years or more (Tissot and Welte, 1984). The shaliness of a basin and the continuity of the shale layers depend on the original depositional environment (Figure 2.12). The original porosity, permeability and connectivity of the fine-grained and coarse-grained rocks are


41

100

\

\ Delta

Deltaic, banier(Zeito)

I

I

I fringe and delta plain

I Distr. channel (Zdto)

m e inet al) 0 560

0

I 0

1600 I

I

I

100

200 L-th

300

2600 n

1500 I

I

400

500

I

600 rn

of shale intercalation

Figure 2.12 Continuity of shale (silt) intercalations as a function of depositional environment (from Weber, 1982, Journal of Petroleum Technology, March 1982, Fig. 2. Copyright (0)1982 by Society of Petroleum Engineers. Reprinted by permission).

related to the genetic type of sedimentary basin. For example, the stratigraphy of cratonic basins is dominated by broad shallow depositional systems. The basins are characterized by a layer-cake stratigraphy with individual lithostratigraphic units sometimes traceable for thousands of kilometres and an intimate interbedding of marine and nonmarine units (Miall, 1990). The source area of the clastic sediments is continental and characterized by the presence of abundant quartz and a paucity of lithic fragments (Miall, 1990). The hydrogeological framework of a cratonic basin is thus characterized by laterally continuous hydrogeological units and a relatively small percentage of poorly permeable shales. The coarse-grained units (quartz-rich sands) probably compact slowly until depths of 3 to 4 kilometres and may provide laterally continuous escape ways for the water expelled from the compacting finegrained rocks. During the slow subsidence of cratonic basins geopressured zones will probably not be created (see also Bethke, 198613). The burial-induced flow of groundwater in a subsiding cratonic basin will belong to the shallow subsystem and possibly in the deeper parts of the basin to the intermediate system. Studies of basins underlying the North American craton confirm this general picture (e.g. Bethke et al., 1991). In contrast, aulacogens do not appear

42

Chapter 2

t o be characterized by any distinctive lithofacies assemblage or sequence. The sediments may be shallow or deep marine, carbonate or clastic, although a transition from nonmarine in the graben stage t o marine in the downwarp stage seems typical (Miall, 1990). The lateral continuity of individual hydrogeological units in an aulacogen will be slight for various reasons including the presence of abundant deep-reaching faults. The faster subsidence rates, greater percentage of fine-grained rocks and poorer lateral continuity of the hydrogeological units, potentially favour the development of all three subsystems of burial-induced groundwater flow. The present-day groundwater pressure conditions in the Viking and Central Grabens of the North Sea Basin (e.g. Buhrig, 1989; Cayley, 1987)indicate the presence of the three subsystems of burial-induced groundwater flow. The main driving force for burial-induced groundwater flow is thought to be mechanical pressuring of the groundwater caused by sedimentary loading. When the sedimentation and subsidence of a basin cease, the groundwater pressures induced by sedimentary loading and the associated groundwater flow system will eventually dissipate. This dissipation may take millions to tens of millions of years in poorly permeable basins (e.g. Lerche, 1990; Neuzil, 1986).

Pattern of burial-induced groundwater flow Different groundwater modelling studies have provided a general insight in the pattern of burial-induced groundwater flow in sedimentary basins (Bethke, 1985; Bethke et al., 1991; Bethke and Marshak, 1990;Harrison and Summa, 1991; Magara, 1978). Magara (1978) presented a fluid-flow model for a compacting sandstone-shale sequence. This model showed that the groundwater moves principally through the more permeable sandstone layers from an area of more loading (thicker deposition) to one of less. In a sedimentary basin consisting of alternating laterally continuous fine-grained and coarse-grained sedimentary rocks, the groundwater below a certain depth will flow through the individual coarse-grained rocks and will be directed from the depocentre(s1of the basin to its edges. The results of later modelling studies of burial-driven flow in inhomogeneous basins confirmed this tendency for groundwater to flow laterally towards the edge of the basin as previously predicted by Magara (Figure 2.13). Bethke’s (1985) studies further showed that the groundwater fluxes along lateral flow paths increase toward the basin edges. According to Bethke, this is because compaction-driven flow is cumulative along a flow path, with the volume flux at a point along the path nearly equal to the rate of pore volume collapse for the entire path up to that point. Bredehoeft et al. (1988) presented field evidence for the existence of burial-induced lateral groundwater flow in the rapidly subsiding South Caspian Basin. The distribution of superhydrostatic pressures in part of the South Caspian Basin, which consists predominantly of sands and shales, was interpreted by Bredehoeft et al. (1988) to indicate the occurrence of a lateral flow of groundwater focussed through the sands, from the basin centre towards its edges (Figure 2.14). The lateral

43


l

!

I

r

~ (

l r

r

l r

l r

( 1

l

l r

~ '

( ~

'

~

~

~

l

~

max. v,

-0.1997 cm/yr

rnax. vZ =0.0033 cm/yr t=50 M.Y. max. vzm =0.0050 cm/yr

0

0

-

50 kin

max. v,

=0.2060 cm/yr

max. v,

=0.0056 cm/yr

max.

-0.0050 cm/yr

VZm

groundwater velocity vector groundwater velocities a r e shown relative t o the subsiding medium equipotential o f 0.106 MPa

-0.106-

+

subsidence velocity o f the basement relative to fixed elevation

Figure 2.13 Calculated directions and velocities of burial-induced groundwater flow in a subsiding inhomogeneous basin, after 50 and 100 million years of subsidence (after Bethke, 1985, Journal of Geophysical Research, Vol. 9, no. B8, Fig. 5, p. 6822. Copyright by the American Geophysical Union).

groundwater flow pattern in a compacting basin can be related to the variations in thickness of its sedimentary fill, i.e. t o the geometry of the base of the sedimentary fill of the basin. Figure 2.15 shows the groundwater flow patterns for different geometries of simple compacting basins. The classification of the basin geometry types is according to Pratsch (1982). Pratsch's classification is based on the assumptions that the fine-grained and coarse-grained rock units are continuous throughout the basin, the geometry of the interface between the fine-grained and coarse-grained rock units follows the geometry of the basin and that the permeabilities of the individual rock units are constant.

l

Chapter 2

44

4 Stmctures 0Oil and gas fields

-

0

50

lOOkm

A. Location of area 1. I1 and 111 near the Baku Archipelago

6 . Hydraulic head of groundwater in sand units in areas I, I1 and 111

Area II

Area I

Area 111

0 2 4 1

g6

50

8 2 4 6

0

2

4

60

2 4 6 Pressure (km water)

0

2

4

6

0

I

b

k

m

C. Pressure-depth curves for sand and shale units in areas I, I1 and 111 Figure 2.14 Distribution of superhydrostatic pressures of the groundwater in part of the South Caspian Basin near t h e Baku Archipelago, indicating lateral flow of groundwater through sand units from area I11 towards area I (from Bredehoeft e t al., 1988. Reprinted by permission of the American Association of Petroleum Geologists).

The velocities of groundwater flow in actively filling and subsiding basins depend, among other things, on the sedimentation rate in the basin. The burialinduced velocity of groundwater flow has been calculated to be less than 2 m d y e a r during gradual subsidence of the Illinois Basin, USA, in Permian


4-5

A

a

a p l e circular symmetrical basin No concentration of groundwater flow

b

a p l e circular asymmetrical basin Concentration of groundwater flow towards narrow concave side B o f the basin

E E

c

c

c

C

A

c

Simple elonqate symmetrical basin Concentration o f groundwater flow towards the long flanks A and B o f the basin

d

Simple elonqate asymmetrical basin Concentration o f groundwater flow towards t h e long flanks A and B o f t h e basin

e

Simple elonqate symmetrical curved basin Concentration o f groundwater flow towards concave long flank A o f t h e basin

f

Simple elonqate asymmetrical curved basin Concentration o f groundwater flow towards concave long flank A o f the basin

-

isopach sedimentary f i l l

T

I depocentre of

t h e sedimentary basin

______

*

basin axis burial induced groundwater flow direction

Figure 2.15 Lateral groundwater flow patterns for different geometries of simple compacting basins (classification of basin geometry types according to Pratsch, 1982).

46

Chapter 2

times (average rate of sedimentation near the depocentre = 0.03 m d y e a r ; velocity of groundwater, v, = qdn, is taken relative to the medium) (Bethke et al., 1991). In the Gulf of Mexico Basin, USA, with recent sediment accumulation rates of up t o 2.2 mdyear, the velocities of groundwater flow in large parts of the basin are in the range 0.1 to 0.001 m d y e a r (Harrison and Summa, 1991). In the shallow subsystem of burial-induced groundwater flow in the Gulf of Mexico Basin, the velocities of groundwater flow are a few tens of centimetres per year, while minimum velocities in the order of 5 x m d y e a r are found in the geopressured zone near the basement (Harrison and Summa, 1991). The velocities of groundwater flow i n actively filling and subsiding basins possibly range from fractions of millimetres per year in the deep geopressured subsystem of groundwater flow t o centimetres per year in the shallow subsystem. In theory, there may be no vertical groundwater flow between different coarse-grained rocks below a certain depth in a simple compacting basin consisting of laterally continuous fine-grained and coarse-grained rocks. In reality, however, facies changes will occur and permeable faults, fracture systems, salt diapirs etc will disrupt the lateral continuity of the individual fine-grained and coarse-grained rocks, thus allowing vertical communication between different coarse-grained rocks, which may result in vertical upward flow of groundwater. In the deep subsystem of burial-induced groundwater flow, geopressured conditions prevail in both the fine-grained and the coarse-grained rocks. Several authors have discussed the formation of fractures (microfractures; hydraulic fractures) in the geopressured zone (Cathles and Smith, 1983; Domenico and Palciauskas, 1979; Du Rouchet, 1981; Etheridge et al., 1984; Nur and Walder, 1990; Palciauskas and Domenico, 1980). When the groundwater pressure reaches its upper limit, i.e. the sum of the least principal stress plus the tensile strength of the medium, the water-saturated rock may fracture and faults may open or reopen (e.g. Du Rouchet, 1981; Price, 1980a). The created fractures have a preferred orientation normal or nearly normal to the least principal stress (Palciauskas and Domenico, 1980). According to Nur and Walder (1990) the natural hydrofracturing leads t o a rapid release of water together with a sudden drop in groundwater pressure. As a result of the decline in groundwater pressure, the hydrofractures will close again. Subsequently the groundwater pressure will increase again in an actively subsiding basin leading t o another cycle of hydrofracturing, release of groundwater and sealing (see Nur and Walder, 1990). In addition to this discontinuous flow of groundwater related to hydrofracturing, episodic flow of groundwater may also occur along faults opened by high superhydrostatic pressures (Price, 1980a) o r along tectonically active faults (Hooper, 1991). According t o Hooper (1991) the hydraulic conductivity along growth faults increases as these faults become more active. This is considered to be caused by


47

dilation-increased permeability, refracturing of mineral zones and seismic pumping (Hooper, 1991). Since activity of growth faults is related to sediment accumulation rates, the amount of groundwater flowing up faults should be greatest when accumulation rates are fast (Hooper, 1991). During periods when the faults are inactive, probably no vertical upward flow of groundwater occurs and lateral flow of groundwater across the faults may also be restricted. During continued burial, large lateral differences in groundwater potential may build up across the closed faults reflecting restricted flow conditions. The vertical upward escape of groundwater from a deep geopressured subsystem of groundwater flow may occur as a periodic flow of groundwater focussed along distinct vertical pathways (faults, along salt diapirs) or directed through a hydrofractured zone in the upper poorly permeable part of the geopressured system (i.e. a zone of seal failure). In a simple compacting basin, local vertically upward groundwater flow will not disturb the general picture of the basin-wide lateral groundwater flow pattern. In a simple inhomogeneous compacting basin, the sedimentary rocks dip towards the centre of the basin: the moving groundwater following the dip directions of the individual coarse-grained rocks therefore also has a vertical component of flow. While flowing from the depocentre to the basin edges, the groundwater will reach ever shallower parts of the basin. Indicators of burial-induced groundwater flow In addition to the groundwater pressure distribution, the burial-induced flow of groundwater is associated with several physical and chemical characteristics of the sedimentary basin. These include the distribution of temperature, salinity and chemical composition of the groundwater, and the distribution of diagenetic minerals in the basin. In the subsurface, heat moves from the centre of the earth outwards through the sedimentary crust into the ocean or atmosphere, where it is lost as radiant energy. The subsurface temperature distribution is influenced by conductive heat transport, which depends on the thermal conductivity and the specific heat properties of the rock matrix and the pore fillers (e.g. water), and by convective heat transport by groundwater flow (Section 1.3.2). The general direction of burial-induced groundwater flow is from the deeper and hotter parts of a sedimentary basin to its shallower and cooler parts. This upwarddirected component of groundwater flow may induce positive temperature anomalies, when the groundwater flow velocities are sufficiently high (Bethke, 19851,i.e. in areas where upward directed groundwater flow is concentrated, such as along the basin edges, along permeable faults and other permeable escape routes. Such thermal effects caused by burial-driven groundwater flow in slowly subsiding basins are probably negligible because of the slow

48

Chapter 2

groundwater flow velocities perpendicular to the hydrostatic isotherms (Bethke, 1985; Cathles and Smith, 1983). In rapidly subsiding basins, the deeper subsystem of burial-induced flow of groundwater may be associated with relatively high temperatures if the rocks are undercompacted. Because pore fillers, such as groundwater, have much smaller thermal conductivities than the rock minerals, the relatively large porosity of undercompacted sedimentary rocks in the geopressured zone reduces the thermal conductivity and therefore raises the geothermal gradient of most rocks (e.g. Hunt, 1979). Concentrated vertical upward flow of groundwater escaping from the high-temperature geopressured zone along e.g. faults or salt diapirs, will induce positive anomalies of temperature, thermal gradient and heat flow at shallower depths in the basin (Bodner and Sharp, 1988; Hooper, 1991; Jensenius and Munksgaard, 1989). The salinity, i.e. the amount of total dissolved solids, of the groundwater in the permeable coarse-grained rocks of both actively filling and subsiding sedimentary basins and stable subaerial basins generally increases with depth at rates ranging from 50-80 t o 300 r n m m (Hunt, 1979; Selley, 1985). Salinities as high as 400,000 mg/l have been observed (Kreitler, 1989). Although, the salinity of groundwater generally increases with depth in a sedimentary basin, there is a wide variation in the actual total content of dissolved solids and chemical composition of the groundwater with depth. The gradual change in chemical characteristics of the groundwater in a clastic basin is influenced by the original chemical composition of the rock and the synsedimentary water, and the increase in solubility of most rock minerals with depth. A t relatively shallow depths ion exchange with clay minerals may play a role, and membrane filtration and the expulsion of ions plus water from compacting fine-grained rocks may increase the salinity of the adjacent coarse-grained rock. A t greater depths, i.e. at increased temperatures (> 80 "C) and assuming near hydrostatic groundwater flow conditions, the groundwater in the coarsegrained rocks will tend t o be in equilibrium with most minerals present (Bj~irlykke, 1989). There is no general agreement as to the origin of the great salinity of the deeper groundwaters, i.e. of the deep basin brines (Na-Cl brines and Na-Ca-C1 brines) (Hanor, 1987b; Kreitler, 1989). Different processes have been proposed. The main ones are (Hanor, 1987b): membrane filtration or reverse osmosis; subsurface infiltration of subaerially produced brines (residual bittern brines); dissolution of evaporite minerals. Several authors have postulated that finegrained rocks (shales) may act as semi-permeable membranes that retard the passage of ions while allowing relatively unrestricted passage of water (Bredehoeft et al. 1963,1982;De Sitter 1947; Graf, 1982).De Sitter (1947)proposed a filtration mechanism t o account for the subsurface increase in groundwater salinity. The fine-grained sedimentary rocks are thought to acquire the property of a semi-permeable membrane by compaction. The process of

Groundwater flow in sedimentay basins

49

membrane filtration requires large volumes of water to pass through the poorly permeable layers in order to create brine concentrations (Kreitler, 1989). In an actively filling and subsiding basin, active cross-formational flow through poorly permeable layers probably only occurs in the shallow subsystem of burial-induced flow of groundwater. In slowly subsiding basins, burial-induced flow of groundwater is unable to cause concentration of brines by membrane filtration (Bethke, 1985). The other two brine-forming processes are restricted to basins in which evaporites have been deposited. The large influence that evaporites may have on the chemical composition of groundwater in sedimentary basins is indicated by Bjorlykke (1989). He states that sedimentary basins along the Atlantic Coast where Mesozoic evaporites occur have chemical groundwater compositions which are essentially different from groundwater compositions in basins north of this area of occurrence of evaporites. The infiltration of bittern brines in the basin is a syngenetic process of brine formation. The location of the brine in the sedimentary basin may change during the hydrodynamic evolution of the basin. Dissolution of the extremely soluble evaporites (e.g. halite) will result in an increase of the groundwater salinity near salt layers and diapirs (e.g. Morton and Land, 1987; Stoessel and Moore, 1983). The subsequent movement of the groundwater of increased salinity through relatively permeable layers or along vertical routes in the shallow and intermediate subsystem of burial-induced groundwater flow may spread the saline water through the basin. In the deep subsystem with restricted groundwater flow conditions, the movement of saline groundwater by diffusion may become important ( B j ~ l y k k eet al. 1988; Ranganathan and Hanor, 1987). The present-day occurrence of high salinities of the groundwater in sedimentary basins is the combined result of the water-rock interactions and groundwater flow conditions that occurred during the evolution of these basins. In addition to the periods of burial-induced flow of groundwater, the periods of meteoric flow may have greatly influenced the present-day salinity distribution in a basin because of their great potential for changing the chemical composition and salinity of the groundwater. The reasons for this include the relatively large groundwater flow velocities and fluxes in a meteoric system (Section 2.3)and the associated greater potential for membrane filtration. In addition to the generally observed increase in salinity of the groundwater in the permeable coarse-grained layers in an actively filling and subsiding basin, the difference between the salinity of groundwater in fine-grained rocks and that of groundwater in the adjacent coarse-grained rocks increases with depth (Dickey, 1988; Hunt, 1979). The groundwater salinities in the coarsegrained rocks exceed those in fine-grained rocks (e.g. Fertl, 1976). In the shallow subsystem of burial-induced groundwater flow, the process of membrane filtration may cause a relative increase in the groundwater salinity of the coarse-grained rocks. In the shallow and intermediate subsystems of burial-induced groundwater flow, both water and dissolved ions may move from the compacting fine-grained rocks into the adjacent coarse-grained rocks.

50

Chapter 2

However, the movement of some water molecules is inhibited by their adsorption on mineral surfaces of the fine-grained rocks (structured water), while the dissolved ions are relatively free to move into the coarse-grained rock, increasing the salinity of its groundwater. As compaction proceeds all the larger pores in the fine-grained rock lose their salts to the coarse-grained rocks, until finally only small pores filled with fresh (structured) water remain (Hinch, 1980). Dehydration of clay minerals at greater depths, i.e. at higher temperatures and pressures causes the groundwater salinity in the finegrained rocks t o decline (e.g. Bjprrlykke, 1989). Many researchers have shown that the salinity of the water expelled from fine-grained rocks decreases with increasing overburden pressure (Fertl, 1976). In the intermediate subsystem of burial-induced groundwater flow, there is no cross-formational flow through the poorly permeable fine-grained rocks, while the permeable nature of the coarse-grained rocks still allows throughflow of groundwater. Groundwater of different origins and chemical compositions may be introduced into the coarsegrained rocks; this in turn also influences the interactions between water and rock. The chemical compositions of groundwater in the coarse-grained rocks in the intermediate subsystem may be extremely variable. In the deep subsystem of burial-induced groundwater flow, the flow of groundwater through both the fine-grained and coarse-grained rocks is restricted. As a consequence, the groundwater in the geopressured zone supplies o r removes hardly any solids in solution except by diffusion. This probably restricts the water-rock interactions to isochemical reactions (Bjprrlykke et al., 1989). Possibly, the increase with depth of the salinity of groundwater in the coarse-grained rocks in the geopressured zone is slowed down, because dissolved solids are no longer expelled by the fine-grained rocks, dissolved solids are not supplied by flowing groundwaters and there is less chemical activity as a result of the restricted flow of groundwater through the coarse-grained rocks. Deviations from these general variations of the groundwater salinity in the different subsystems of burial-induced groundwater flow of an actively filling and subsiding basin can be expected in areas with concentrated flow of groundwater. In a basin-wide burial-induced groundwater flow system, appreciable concentrated upward-directed flow of groundwater, for instance along the basin edges, permeable faults or other vertical escape ways, will induce positive groundwater salinity anomalies in the shallower parts of the basin. The concentrated vertical upward flow of relatively warm, saline and high pressured groundwater may cause diagenetic reactions when this groundwater enters shallower parts of a basin (e.g. Glasmann et al., 1989a, 1989b; Hooper, 1991; Mathieu and Velde, 1989). For instance, a drop in pressure of the upward-flowing groundwater may induce diagenetic reactions, such as calcite and silica cementation (e.g. Harrison, 1990). Illite formation in coarsegrained rocks may be associated with the introduction of groundwater of deeper


51

origin. Lee et al. (1989) studied the timing and conditions of diagenesis of the Permian Rotliegende sandstone in the southern part of the North Sea Basin using burial histories, petrographic data, WAr ages and 180/160ratios. They interpreted the data from a sample in the Groningen area to indicate a late intense period of illite formation related to a period of groundwater flow along a major fault nearby. Hence, zones with appreciable vertical upward flow of groundwater in a burial-induced groundwater flow system can be characterized by positive pressure, temperature and salinity anomalies and associated diagenetic mineral assemblages.

2.2 Tectonically-induced groundwater flow

Plate-tectonic processes control the evolution of sedimentary basins and consequently influence the evolution of groundwater flow within the basins. In this section tectonically-induced flow of groundwater is restricted to groundwater flow which is related directly to tectonic activity. The influence of the increase in stresses of tectonic origin on the development of superhydrostatic groundwater pressures in a sedimentary basin was recognized more than 30 years ago by Hubbert and Rubey (1959) and Rubey and Hubbert (1959). More recently, Oliver (1986,1990,1992) suggested a causal relationship between plate-tectonic interaction in zones of convergence and large-scale flow of groundwater. The tectonic origin of ancient groundwater flow conditions in zones of plate convergence and continental collision has been demonstrated by several, mainly geochemical, studies (e.g. Dorobek, 1989; Losh, 1989; Roberts, 1991; Vrolijk and Myers, 1990). From geochemical, thermal and structural geological data the present-day occurrence of tectonically-induced flow of groundwater has been inferred to occur at the Barbados Ridge Complex along the detachment zone separating the underthrusting Atlantic Ocean crust and the overthrusting Caribbean plate (Moore et al., 1987, 1988). The occurrence of tectonically-induced flow of groundwater has been proven, but as yet no detailed tectonically-induced groundwater flow system has been defined. In sedimentary basins within the area of influence of tectonic activity, the directions and velocities of groundwater flow will be affected by the magnitude and directions of the tectonic driving force for groundwater flow in combination with the tectonically changed hydrogeological framework of the basin. The tectonic driving force for groundwater flow may interact with other driving forces acting on the groundwater in a particular basin, such as mechanical driving forces caused by sedimentary loading in a filling and subsiding basin (Section 2.1) and the relief of the water table in subaerial basins (Section 2.3).

52

Chapter 2

A direct tectonic driving force for groundwater flow in sedimentary basins at o r near zones of plate convergence is the increase in both vertical and lateral compressive stresses. The increase in lateral compressive stresses at a zone of continental collision may propagate far into plate interiors and may change the intra-plate stress fields at distances of more than 1000 km from the collisional margin (Klein and Barr, 1986; Ziegler, 1987). Kooi et al. (1989) presented a synthetic paleo-stress curve inferred from the foreward modelling of the North Sea Basin stratigraphy. This curve shows an increase in compressive stress of 4 kbar (400 MPa) over a period of approximately 10 million years in the Early Tertiary. The compressive stresses are probably primarily controlled by European-African plate convergence and continental collision. The strong phase of Early Tertiary compression in the North Sea Basin corresponds in timing with the occurrence of strong folding phases in the Alpine domain (Kooi et al., 1989). The Alpine deformation front is more than 800 km away from the North Sea Basin. The increase in lateral compressive stress of 400 MPa/lO million year is significantly larger than the minimal increase in vertical stress necessary to induce geopressured conditions in a sedimentary basin. For instance, Harrison and Summa (1991) calculated for the Gulf of Mexico Basin, USA, that sedimentation rates of 1 mndyear, i.e. 1000 dmillion year corresponding to an increase in vertical stress of approximately 23 MPdmillion year, will induce groundwater pressures to approach lithostatic gradients. From this example, it seems likely that changes in tectonic stresses acting on a sedimentary basin a t a considerable distance from a zone of convergence, may be large enough to change significantly the groundwater pressure condition in the basin. Changing groundwater pressure conditions affect directly the system of groundwater flow in the basin. In addition, the directions of groundwater flow may also be influenced indirectly by the tectonically increasing groundwater pressures because under high superhydrostatic groundwater pressure conditions, the water-saturated rock may fracture and faults may (refopen (Section 2.1) and displacement of the rocks along faults may be facilitated (Hubbert and Rubey, 1959; Rubey and Hubbert, 1959). The fractures and active faults may become the preferred escape routes for the groundwater in a tectonically-induced geopressured zone. A sedimentary basin at o r near a convergent plate margin (e.g. a foreland basin) may be affected by both vertical and lateral compressive stresses. An initial increase in lateral tectonic compression may amplify the subsidence of a foreland basin (Allen and Allen, 19901, affecting the rate of sedimentation in the basin. A burial-induced groundwater flow system and the associated distribution of superhydrostatic pressure can be expected to develop in such an actively filling and subsiding basin. The continued lateral tectonic compression may also directly affect the magnitude and distribution of groundwater pressures. Lateral tectonic compression may readily raise the already superhydrostatic groundwater pressures to lithostatic pressures. The increased groundwater pressures, i.e. the decreased effective stresses may


53

facilitate faultinglthrusting (Hubbert and Rubey, 1959; Rubey and Hubbert, 1959). The vertical load of thrust sheets emplaced over the foreland basin will further increase the groundwater pressures below the thrust sheets and induce groundwater flow away from the zone of convergence. Several studies suggest that the tectonically-induced flow of groundwater is an episodic and focussed flow of groundwater which is directed away from the area of convergence (Bradbury and Woodwell, 1987;Dorobek, 1989;Duane and De Wit, 1988; Losh, 1989; Moore et al., 1987, 1988; Vrolijk and Myers, 1990). The tectonically-induced flux of groundwater focussed along active faults or available aquifers may be large enough to create geochemical and temperature anomalies in the subsurface (e.g. Moore et al., 1987, 1988;Deming et al., 1990). The episodicity of groundwater flow along faults or basal aquifers in thrust sheets has been related to episodic movement of thrust sheets (e.g. Bradbury and Woodwell, 1987;Moore et al., 1987,1988;Roberts, 1991). Ge and Garven (1989)used numerical modelling to evaluate the role of compressional tectonics in driving regional groundwater flow in a foreland basin in a late period of thrusting and uplift (Figure 2.16). Their model assumed that emplacement of the tectonic loads on the foreland basin platform is instantaneous. Ge and Garven (1989)found that lateral compression and tectonic loading can create significant large-scale flow of groundwater soon after compression of the foreland basin, with flow rates in the order of 0.001 0.01 m per year for thrust sheet loads 1 to 10 km thick. They found that the groundwater flow velocities increase more significantly in the vicinity of the loading area. The flow of groundwater is directed away from the area of loading and most groundwater is focussed into a basal aquifer. The study suggests that cycles of tectonically-induced flow could exist repeatedly during the history of foreland basin thrusting (Ge and Garven, 1989).

-

-

Compression and v8rtlcal loading

Figure 2.16 Conceptual model of a foreland basin in a late period of thrusting and uplift (from Ge and Garven, 1989, Geophysical Monograph 48, Fig. 5, p. 148. Copyright by the American Geophysical Union).

54

Chapter 2

B a s e Tene Marie Galante Basin (Guadeloum)

Barbados Ridge

Atlantic Plain

i: CEPM line CRV 128

Inset 1

(no vertical exaggeration)

A = Accreted sequences D = D&dlement

U = Underthrust sequences

0 = Top Of Oceanic Crust

4

L B 6

Inset 2

Figure 2.17 Generalized cross-section of the northern Barbados Ridge and adjacent Lesser Antilles arc, showing t h e decollement zone separating the underthrusting Atlantic ocean crust and the overthrusting Caribbean plate. Inset 2 shows the upward directions of fluid flow along faults in the Barbados Ridge accretionary prism, along the dbcollement and in the underthrust sediments (modified after Moore e t al., 1987. Reprinted by permission from NATURE vol. 326, p. 786. Copyright (0)1987 Macmillan Magazines Ltd.;and after Moore e t al., 1988 in Geological Society of America Bulletin, Vol. 100, pp 1580 and 1590).

In subduction zones, subduction of water-saturated sedimentary rocks, i.e. continued burial of the rocks and the associated increase in vertical stress and temperature, will induce superhydrostatic pressures in the subducting rocks if groundwater cannot escape fast enough to keep up with the rate of subduction. Near lithostatic pressures have been inferred t o occur along the d6collement zone, i.e. the detachment zone, between subducting and overriding plates (Moore et al., 1987, 1988; Vrolijk and Myers, 1990).The detachment zone appears t o constitute a preferred lateral escape way for the overpressured groundwater in the underthrust sedimentary rocks (Moore et al., 1987, 1988). Vrolijk and Myers (1990) propose the term tectonic aquifer t o characterize the watertransmitting properties of a detachment zone. The adjective ‘tectonic’ signifies the role that deformation along the detachment zone plays in increasing the permeability (Vrolijk and Myers, 1990). At the Barbados Ridge Complex, separate groundwater flow systems appear t o be present in the accretionary wedge of the overthrusting Caribbean plate and the subducting Atlantic plate.


55

Lateral flow of ‘Atlantic’ groundwater through the detachment zone is preferred to vertical upward flow through the accretionary wedge (Figure 2.17). According to Moore et al. 1987,this may in part be explained by the expected change in stress orientation across the top of the detachment zone (Figure 2.17). The low plunge of the maximum principal stress in the accretionary wedge would discourage the opening of steeply dipping fractures by high groundwater pressures while encouraging flow along gently dipping surfaces (Moore et al., 1987). Flows along both the detachment zone and the faults within the accretionary wedge of the Barbados Ridge Complex are apparently episodic (Moore et al., 1988).

2.3 Groundwater flow in stable subaerial basins

When a sedimentary basin emerges and stabilizes above sea level, infiltrating meteoric water will become the main driving force for the development and maintenance of a regional groundwater flow system, i.e. the gravity-induced cross-formational groundwater flow system (also called topography-induced groundwater flow system). Buoyancy forces may play a role by modifying gravity-induced flow systems. Under certain conditions local flow systems may form within the gravity-induced flow system (Section 2.4). 2.3.1 Gravity-induced groundwater flow system In various papers (T6th, 1962,1963,1970,1978,1979,1980) T6th outlined the theoretical and observed characteristics of gravity-induced groundwater flow in geologically mature drainage basins. According to T6th, the groundwater flow pattern in a geologically mature sedimentary basin (i.e. a tectonically stable and non-compacting basin) is induced by the relief of the water table, which in most cases follows its ground surface topography. The water table represents the surface at which groundwater pressure equals that of the free atmosphere. Given adequate rainfall the water table closely follows the topographic relief and this relief can be considered to generate the differences in the potential energy of the groundwater. The rock framework of a geologically mature basin is hydraulically continuous, i.e. there is no absolute permeability barrier to water flow, and the gravity-induced cross-formational groundwater flow pattern will only be modified by the differences in subsurface permeability (T6th, 1980). In a geologically mature basin, the possible existence of driving forces for groundwater flow other than gravity are considered to be temporary or local (T6th, 1980).

T6th describes the gravity-induced groundwater flow pattern in terms of interdependent flow systems as existing in drainage basins. A drainage basin is a depression of the ground surface, partly or entirely surrounded by

Chapter 2

56

relatively high areas and underlain a t some depth by an effectively impermeable base. A flow system is defined by T6th (1963) as a set of flow lines in which any two flow lines adjacent at one point of the flow region remain adjacent through the whole region; they can be intersected anywhere by an uninterrupted surface across which flow takes place in one direction only. Under steady-state conditions, the groundwater flow system in a homogeneous and isotropic drainage basin with simple ground surface geometry consists of three parts; the area of recharge, mid-line and discharge (Figure 2.18). Meteoric water infiltrates in the upland recharge area and subsequently descends in that area, flows laterally through the area of midline, and finally ascends towards the ground surface in the topographically low discharge area. In the mid-line area, there is no vertical flow. As a consequence, the groundwater potential does not change with depth in the midline area, and the corresponding gradient of groundwater pressure is hydrostatic (Figure 2.19). Below recharge areas the groundwater potential decreases, reflecting the downward flow of groundwater. The corresponding groundwater pressure gradient is subhydrostatic. In discharge areas, the vertical upward flow of groundwater is reflected in increasing groundwater topographic e l e v a t i o n above s t a n d a r d datum

-

mid line

I

line o f equal hydraulic h e a d

20,000 f e e t (6096 m l

standard datumJ

g r o u n d w a t e r f l o w line

Figure 2.18 Cross-section showing theoretical distribution of hydraulic head and gravityinduced groundwater flow pattern in a homogeneous and isotropic drainage basin with simple ground surface geometry (modified after T6th, 1962, Journal of Geophysical Research, Vol. 67, no. 11, Fig. 3, p. 4380. Copyright by the American Geophysical Union).


57

depth Im)

soperhydrostatic pressures (discharge area) hydrostatic pressures (midline area) subhydrostatic pressures (recharge areal

2000

3000

.

\'

pressure

Figure 2.19 Change of groundwater pressure with depth in different parts of a flow system of a hypothetical homogeneous drainage basin with simple ground surface geometry (modified after Tbth, 1980. Reprinted by permission of the American Association of Petroleum Geologists).

potentials with depth and in superhydrostatic pressure gradients (Figure 2.19). When the geometry of the basin's ground surface is not simple but complex, i.e. local relief is superimposed on the regional slope of the basin, flow systems of three different orders can be distinguished: local, intermediate and regional (Figure 2.20). Each flow system consists of the same three parts as mentioned above with the corresponding descending, lateral and ascending directions of groundwater flow. In a local flow system, the recharge and discharge areas are contiguous. In intermediate and regional flow systems, the recharge and discharge areas are separated by one or more local systems. In regional systems, the groundwater flow path is between the highest upland recharge area and the lowest discharge area of the basin (Figure 2.21). The depth of penetration of the various flow systems in a homogeneous isotropic basin is a function of the ratio of local relief and regional slope (T6th, 1963), and may reach several thousand feet in a homogeneous basin under the effect of local topography of a few tens of feet (T6th, 1980).The intensity of flow in a single flow system decreases with increasing depth and away from the area of mid-line, and near-stagnant conditions exist in the lower corners of the flow field (Figure 2.18). Where two or more flow systems meet or part, the groundwater flow directions may change significantly over the interface between the different flow systems. In the areas between these different flow systems, the groundwater movement will be relatively slow to zero, and the areas are referred to as near stagnant to stagnant (Figure 2.20).

58

Chapter 2

topographic elevation and head o f w a t e r above s t a n d a r d datum in f e e t i m l ,valley b o t t o m

"-3 02,

potential distribution on t h e s u r f a c e o f t h e ,theoretical f l o w reqion

20,000 f e e t (6096 m l C'

--

-_*--

v ... .....

regional slope o f ground s u r f a c e

water divide

s t a n d a r d datum

-.-.-.

boundary b e t w e e n groundwater flow systems

line o f equal hydraulic head I

i n t e r s y s t e m s t a g n a n t zone

2

b o t t o m s t a g n a n t zone

g r o u n d w a t e r f l o w line local groundwater flow system

intermediate groundwater f l o w s y s t e m I I regional g r o u n d w a t e r f l o w s y s t e m

Figure 2.20 Cross-section showing the theoretical distribution of hydraulic head and gravityinduced groundwater flow pattern in a homogeneous and isotropic drainage basin with complex ground surface geometry (modified after T6th, 1970. Reprinted by permission of the National Research Council of Canada).

In nature, the subsurface of a drainage basin will not be isothermal and isochemical, nor will it be isotropic or homogeneous. Temperature and salinity variations in a basin influence the density and viscosity of the groundwater and consequently the direction and velocity of groundwater flow (Section 1.3).The temperature and salinity variations and the anisotropy and inhomogeneity of the subsurface may change the gravity-induced groundwater flow pattern as illustrated in Figures 2.18 and 2.20, but will not change its general character (e.g. Garven and Freeze, 198413;T&h, 1978, 1980).

59


a

S h p l e circular symmetrical basin No concentration o f groundwater flow

b

S h p l e circular asymmetrical basin Concentration o f groundwater flow towards n a r r o w concave side B o f the basin

I

B

B

c

c

c

c A

A

c

Simple elongate symmetrical basin Concentration o f groundwater flow t o w a r d s t h e long f l a n k s A and B o f t h e basin

d

Simple elonqate asymmetrical basin Concentration o f groundwater flow t o w a r d s t h e long f l a n k s A and B o f t h e basin

I

I

c

c c

___

J,,I,p,E e , " , , y ~ , , I I , I I E , , I ~ . . L L " ,

1s"

Y--III

Concentration o f groundwater flow towards concave long flank A o f t h e basin

1

I

area o f recharge

,

-",,,p,': _ _

_

~ , " " , j ~ , C

~

'-',""""'ILOL

L",

.LY

- 1 - . . 1

Concentration of groundwater flow towards concave long flank A o f the basin

g r a v i t y induced g r o u n d w a t e r f l o w direction

ground surface geometry (ground surface geometry types patterned after basin geometry types given by Pratsch, 1982).

Chapter 2 elevation above datum in f e e t Im)

feet m)

(-1219) *

c'

-

regional slope of ground surface

___-----line K , ,K,,Kj

30 miles (68 km)

o f equal hydraulic head

groundwater flow line relative hydraulic conductivities

(1219)

Discharge area

6000

(1829) (2438)

8ooo{ Pressure

...-.--.. Calculated pressures in low permeabilitylayer Figure 2.22 Distribution of hydraulic head and gravity-induced groundwater flow pattern, and change of groundwater pressure with depth in different parts of the flow system in a hypothetical inhomogeneous drainage basin with a n areally extensive poorly permeable Iayer and simple ground surface geometry (modified after T6th, 1980. Reprinted by permission of the American Association of Petroleum Geologists).


61

The modifying influence of the subsurface permeability distribution, as studied by theoretical models (Freeze and Witherspoon, 1967; Garven, 1989; Garven and Freeze, 1984a,b;T6th, 1962,1963,1978)indicate that Layers of poor permeability may cause large losses of energy (Figure 2.22). Figure 2.22 shows the pressure-depth curves for an inhomogeneous drainage basin with an areally extensive poorly permeable layer and simple ground surface topography. The groundwater in the poorly permeable layer and basal aquifer is relatively more underpressured in the recharge area and more overpressured in the discharge area. I n a n inhomogeneous drainage basin with laterally extensive hydrogeological units, the flow of groundwater is essentially lateral in the aquifers (Figure 2.22). Vertical flow of groundwater is more widespread across the poorly permeable layer than in the aquifer. In a n inhomogeneous drainage basin with laterally extensive hydrogeological units, the regional flow tends to be focussed into units of relatively good permeability (Figure 2.23). The degree of focussing is controlled by the permeability contrast between the very permeable units and the surrounding hydrogeological units (Garven and Freeze, 1984b). Garven and Freeze (1984b) calculated that specific discharge rates on the order of 1to 10 m3/m2/yearare possible in basin aquifers.

0.1 s

0 0

0.1s

0.2s

0.3s

0.4s

0.5s

0.6s

0.7s

0.8s

0.9s

S

0

0.1s

0.2s

0.3s

0.4s

0.5s

0.6s

0.7s

0.8s

0.9s

S

0.1 s

0

Figure 2.23 Cross-sections showing the theoretical distribution of hydraulic head and gravity-induced groundwater flow pattern in two inhomogeneous drainage basins with laterally extensive hydrogeological units and a complex ground surface geometry (after Freeze and Witherspoon, 1967, Water Resources Research, Vol. 3, no. 2, Fig. 3, p. 628. Copyright by the American Geophysical Union).

M

Chapter 2

0.1

s 0 0

0:l S

0:2 S

0:3S

014 S

0:5 S

016 S

0:7S

OhS

0:9 S

S

0

0.1s

0.2s

0.3s

0.4s

0.5s

0.6s

0.7s

0.6s

0.9s

S

0.2 s

Figure 2.24 Cross-sections showing the theoretical effect of a partial aquifer of relatively good permeability on the distribution of hydraulic head and gravity-induced groundwater flow pattern in two drainage basins (after Freeze and Witherspoon, 1967, Water Resources Research, Vol. 3, no. 2, Fig. 4, p. 629. Copyright by the American Geophysical Union).

-

Lateral variations in permeability, caused by e.g. facies changes unconformities, faults, can profoundly affect the gravity-induced groundwater flow system (Figure 2.241, which in some cases can lead t o the creation of local flow systems (Garven and Freeze, 1984b).

Numerous field examples of regional gravity-induced groundwater flow systems as compiled for example from literature by T6th (1980) confirm the existence of hydraulic continuity i.e. of cross-formational flow, in geologically mature drainage basins. Field evidence from actual cross-formational flow through poorly permeable layers is rare. Alexander et al. (1987) present an example of measurements of head and chemical composition of groundwater in argillaceous formations in the middle Thames Valley area, UK, which are consistent with the concept of cross-formational flow. The theoretically determined influence of laterally extensive hydrogeological units of different permeability on the gravity-induced groundwater flow and associated groundwater potential and pressure distributions in sedimentary basins is confirmed by field observations. For example, regional underpressuring of groundwater in Mesozoic and Paleozoic rocks in the Denver Basin, USA (Figure 2.25, Belitz and Bredehoeft, 1988) and in the Deep Basin Brine aquifer in the Palo Duro Basin, Texas, USA (Senger and Fogg, 1987) were shown to result from steady-state gravity-induced groundwater flow


0

Pressure (MPa) 10 20

30

-z 5

i

.

+

Figure 2.25 Subhydrostatic groundwater pressures in Cretaceous D and J sandstones in northeastern Colorado and Nebraska panhandle USA (after Belitz and Bredehoeft, 1988. Reprinted by permission of the American Association of Petroleum Geologists).

in the basins and the presence of poorly permeable units (Cretaceous shales and Permian evaporites, respectively), which hydraulically isolate these underpressured deep aquifers from the recharge zones (compare with Figure 2.22). Although poorly permeable units cause a large downward loss of groundwater potential, it has been shown that vertical leakage, i.e. crossformational flow, through such units on a basin-wide scale can contribute significantly to the total groundwater flow in the system (Bredehoeft et al., 1983; Senger and Fogg, 1987). The focussing of gravity-induced groundwater flow into regionally extensive aquifers has been observed in e.g. the Denver (Belitz and BredehoeR, 1988), Palo Duro (Senger and Fogg, 1987) and Kennedy Basins, USA, (Gosnold, 1990) and in sedimentary basins in the UK (Downing et al., 1987).

Indicators of gravity-induced groundwater flow Different physical and chemical phenomena are considered to be associated with regional gravity-induced groundwater flow (e.g. T6th, 1972, 1980). These phenomena may be apparent in the subsurface (e.g. the distribution of pressure, temperature, salinity] geochemical characteristics of groundwater; the distribution of diagenetic minerals) or at the ground surface (e.g. moisture conditions). Under isochemical and isothermal conditions in a stable drainage basin, groundwater moves in the direction of decreasing groundwater

64

Chapter 2

A.

r

Temperature estimate (contours“c)

s.Ll

2000 m

-2000

m

0

B.

10

20

30

40

Distance (km) Temperature anomaly (contours“c)

50

60

70

80

DR Duchesne River Formation

Uinta Formation

Green River Formation

Wasatch Formation

Figure 2.26 Temperatures and temperature anomalies along a characteristic cross-section in the Uinta Basin, USA (after Willet and Chapman, 1987. Reprinted by permission of Editions Technip; and Willet and Chapman, 1989, Geophysical Monograph 47, Fig. 2, p. 31. Copyright by the American Geophysical Union). 2.26a Temperatures estimated from local formation gradients. 2.26b Temperature anomalies calculated from t h e estimated temperatures in Figure 2.28a by removing a constant regional temperature gradient of 25 ‘Ckm. 2 . 2 6 ~ Temperature anomalies calculated by modelling conductive heat transport and advective h e a t transport by gravity-induced groundwater flow through the Duchesne River Formation and the Upper Uinta Formation.


Rocky Mountains

Interior plains Foothills

major regional drainage basin

regional

regional lateral fiow

dsChafQ8

’

r -

local

Recharge area

local

local

Discharge area

Figure 2.27 Schematic model of the relation between cross-formational gravity-induced groundwater flow and changes in temperature and heat flow with depth in Alberta, Canada (after Majorowicz et al., 1985. Reprinted with permission from Journal of Geodynamics 4, Fig. 10, p. 280; Copyright, 1985, Pergamon Press Ltd.).

potentials. The groundwater potential distribution and the pressure distribution are direct indicators of the groundwater flow directions. The temperature, salinity and chemical composition (including isotopic composition) of the groundwater change systematically along its flow path from the recharge areas, where meteoric waters infiltrate, to the groundwater discharge areas. Forced convection of heat by flowing groundwater (Section 1.3) can significantIy change the purely conductive subsurface thermal regime, i.e. the temperature distribution and heat flow pattern, in a basin. Theoretical studies show that the main effects of a gravity-induced groundwater flow system on the temperature distribution in a basin occur in the recharge and discharge areas

66

Chapter 2

of the system, i.e. in areas of vertical groundwater flow (e.g. Smith and Chapman, 1983; Woodbury and Smith, 1985). The downward flow of infiltrated meteoric water decreases the subsurface temperatures in recharge areas, while the upward flow of groundwater increases the temperatures in discharge areas. Recharge areas are characterized by negative temperature anomalies, relatively low geothermal gradients and low heat flows, and discharge areas by positive temperature anomalies, relatively high geothermal gradients and high heat flows (Figure 2.26). The geothermal gradient and heat flow will increase with depth in recharge areas and decrease with depth in discharge areas (Figure 2.27). These thermal characteristics associated with large-scale gravity-induced groundwater flow have been recognized in numerous basins worldwide, e.g. Western Canada Sedimentary Basin (Garven, 1989; Hitchon, 1984; Jones and Majorowicz, 1987; Majorowicz et al., 1984, 1985); Kennedy, Denver and Williston Basins, USA (Gosnold, 1985,1990; Gosnold and Fischer, 1986), Illinois Basin, USA (Bethke, 1986a1, Michigan Basin, USA (Vugrinovich, 1989), Uinta Basin, USA (Willet and Chapman, 1987, 1989), Hungarian Basin (Afoldi et al., 1985; ErdBlyi, 1985), Rhine Graben, Germany (Person and Garven, 1989, 1992). Alps and Northern Foreland, Switzerland (Bodmer and Rybach, 1985),Liaohe Basin, China (Wang and Xiong, 1989).

Meteoric water has an extremely small content of total dissolved solids, is slightly t o moderately acidic and has a large oxygen content. After infiltration in the recharge area, the change in chemical composition of the groundwater of meteoric origin involves many different types of geochemical reaction, e.g. dissolution and precipitation; oxidation and reduction; adsorption anddesorption; acid-base reactions and complexation (e.g. Hem, 1985). At shallow depths, i.e. at low temperatures, the reaction rate will be slow, and the groundwater may be oversaturated or undersaturated with respect to several minerals. The composition of groundwater is initially influenced by the chemical characteristics of the infiltrating water and the climatic conditions prevailing during infiltration. Subsequently, the concentration and type of ions in the groundwater will be controlled by - the initial chemical composition of the groundwater; - the rock framework and its mineralogy: the distribution, solubility and sorption capacity (ion exchange) of the rock minerals. The soluble components of rocks (halite, gypsum, anhydrite, calcite, dolomite) are of special influence on the composition and salinity of groundwater; - the groundwater flow system, i.e. the groundwater flow velocity and flow path of the groundwater, which determine the residence time of the groundwater, the depth of the flow system and hence the temperature and pressure conditions. The chemical characteristics of groundwater of meteoric origin change systematically in the direction of flow in a gravity-induced groundwater flow system (Bredehoeft e t al., 1982; Collins, 1975; Herczeg et d.,1991; T&h, 1980).The


67

recently infiltrated meteoric waters at shallow depths in recharge areas have a small content of total dissolved solids. The salinity of the groundwater increases in the direction of flow, i.e. with increasing depth (increasing temperature) and with increasing residence time of the groundwater in the subsurface. The salinity increases with depth and along the flow path, mainly because the solubility of most minerals increases with increasing temperature, the water reacts with the more readily soluble minerals present along its flow path and because of salt filtering by semi-permeable membranes. Hence, the salinity of shallow groundwater in a recharge area of a certain gravity-induced groundwater flow system is less than that of the groundwater in the deeper parts of the system and less than that of the groundwater in the discharge area. In recharge areas, the groundwater wiIl initially have a high oxygen content and will oxidize organic matter and minerals. As a result, the oxygen content of the groundwater will decrease in the direction of flow. Chebotarev (1955) observed that in general the chemical evolution of groundwater of meteoric origin is related t o a distinctive sequence of dominating anions. The dominating anion changes from bicarbonate to sulphate to chlorine in the direction of groundwater flow and with increasing residence time of the groundwater. In addition to changes in the physico-chemical characteristics of the groundwater itself, the infiltrating meteoric water causes chemical and physical changes in the solid rock matrix, e.g. by diagenesis of the sediments, by changing its porosity and permeability and by forming ore deposits (Table 2.2). Diagenetic reactions that occur in a gravity-induced groundwater flow system include the leaching of quartz and chert, alteration of K-feldspar and plagioclase to kaolinite, the dissolution of high Mg-calcite and aragonite, the precipitation of calcite and smectite (Harrison, 1990). Metals leached in trace amounts from rocks near the surface, which are more soluble in an oxidized than in a reduced state, will precipitate at the boundary between oxidizing and reducing groundwater (e.g. uranium, vanadium, Bethke et al., 1988; Bj~rlykke, 1989). Several studies suggest that basin-wide gravity-induced groundwater flow may lead to the genesis of Mississippi Valley-type lead-zinc deposits in groundwater discharge areas (Bethke, 1986a; Bethke and Marshak, 1990; Garven, 1985; Garven and Freeze, 1984a,b; Deming and Nunn, 1991). In a recently uplifted sedimentary basin, or after an increase in water table relief by e.g. a fall in sea level, groundwater of meteoric origin will come into contact with original, probably more saline, groundwater. The meteoric water may dilute the original water in the rocks and dissolve minerals that were stable under originally more saline conditions, changing the porosity and permeability of the rocks. In the zone where nonsaline meteoric waters mix with more saline groundwater, dolomite may be formed in both carbonate and siliciclastic rocks (Harrison, 1990). Domenico and Robbins (1985) indicate that groundwater of meteoric origin may never completely replace the original connate groundwater in a sedimentary basin. Hence, the chemical

Table 2.2 Some of the different kinds of ore deposits whose origin depends in part on flowing groundwater Type of deposit

Example

Type of flow system

Factors contributing to precipitation

Nickel laterite

New Caledonia

Shallow, water table

Weathering and changing E H - ~ Hat the water table

Laterite bauxite

Jamaica

Shallow, water table drainage helped by karst

Accumulation as residual deposit accompanying weathering

Supergene sulfide

Chuquicamate, Chile

Shallow, water table

Weathering and changing EH-PH at the water table

Calcrete uranium

Yeelirrie, Australia

Discharge end of shallow groundwater flow system

Dissolution from source rock, transport, and precipitation due to evaporation and decomplexation

Roll-front uranium

Texas coastal plain

Shallow groundwater

Leaching of ash, transport, and precipitation at redox front

Unconformity-related uranium

Athabasca district Saskatchewan, Canada

Deep groundwater flow related to faulting

Mixing of oxidizing uraniferous and reducing waters

Mississippi Valley type Lead Zinc deposits

Pine Point, Northwest Gravity o r compaction flow Territories, Canada of brines from deep sedimentary basins

Porphyry copper

San Manuel, Kalamazoo, Arizona

Convection in response to intrusion of a stock or dike

Mixing of meteoric and magmatic fluids and cooling

Lode gold deposits

Carlin, Nevada

Fluid convection of meteoric water deep in the crust

Leaching of source rocks, transport and deposition in fractured rocks due to declining temperature

Leaching from sedimentary source rocks, transport, and deposition due to declining temperatures and and possibly changing EH-PH

From: Domenico and Schwartz, 1990; Copyright 0 1990 by John Wiley & Sons, Inc. Reprinted by permission.


89

characteristics of groundwater and solid rock matrix in part of a gravityinduced groundwater flow system may still reflect the connate groundwater conditions in certain basins, even a long time after steady-state groundwater flow has been reached. At shallow depths in recharge and discharge areas, there is a difference in the ground moisture supply because of the difference in groundwater flow direction. In recharge areas with vertical downward groundwater flow, relatively unmoist conditions can be expected, whereas in discharge areas with vertical upward groundwater flow very moist conditions will prevail. In turn, these differences in moisture supply may produce observable contrasts in the vegetation, erosional features, soil types and surface accumulations of salts (T6th, 1980).

Unsteady-state groundwater flow conditions A present-day regional groundwater flow pattern may not be in steady state. For example, the observed groundwater flow pattern may not be entirely the result of the present-day relief of the water table. The relief of a water table in a tectonically stable subaerial basin may change in time as a result of climatological changes, eustatic sea level changes o r erosion of the ground surface topography. The rate at which changes in the relief of the water table will be reflected entirely in the regional groundwater flow pattern depends on the magnitude of the change in water table relief, the total depth of the groundwater flow system and the hydraulic properties of the subsurface (especially the thickness, lateral continuity and permeability of poorly permeable layers). Considering a geological time-scale, very poorly permeable layers will delay rather than inhibit the transmission of water table changes to the underlying parts of the basin. It may take millions to tens of millions of years for hydraulic heads at depth i n a multi-aquifer system to adjust to changes in head at the ground surface (England and Freeze, 1988;Senger et al., 1987;Tbth, 1978;T6th and Corbet, 1987;T6th and Millar, 1983). Under unsteady-state conditions, the regional groundwater flow pattern in all or part of a sedimentary basin is not entirely the result of the present-day relief of the water table. Different relict groundwater flow patterns reflecting former water table configurations may exist at different depths (e.g. T6th and Corbet, 1987). Besides changing the hydraulic head at the ground surface, erosion may also influence the groundwater pressure conditions by changing the subsurface temperature conditions and the stress conditions caused by erosional unloading. Erosional unloading can induce dilation of poorly permeable units (decompaction). When the unloading is rapid, groundwater flow into the poorly permeable unit may be too slow to accommodate the pore

Chapter 2

70

volume increase and groundwater pressures in the poorly permeable unit may decrease to subhydrostatic values (Neuzil and Pollock, 1983;Neuzil, 1986).T6th and Corbet’s (1987)study in part of the Western Canada Sedimentary Basin shows that erosional unloading may result in extensive regions of subhydrostatic pressures and associated unsteady-state groundwater flow conditions.

2.4 Local groundwater flow systems Two types of local groundwater flow systems may develop in sedimentary basins at depth ranges of interest for studies of hydrocarbon migration and accumulation: flow systems driven by buoyancy (Section 2.4.1) and those driven by osmosis (Section 2.4.2).

2.4.1 Buoyancy-induced groundwater flow system Buoyancy-induced flow of groundwater is driven by density gradients of the groundwater. The density of the groundwater depends on the pressure, temperature and chemical characteristics of the groundwater (Section 1.3.1). Darcy’s equation 1.9 gives the specific discharge for groundwater of varying densities. Equation 1.9 may be rewritten as (e.g. Garven and Freeze, 1984a).

where, Po CLr

PO

h

k

= reference density of the groundwater (ML-3) = pdp = relative dynamic viscosity of the groundwater = reference dynamic viscosity of the groundwater which is defined for

the same pressure, temperature and chemical characteristics as po (ML-IT-’) =&+z POg =-GPog PO

The first term represents the buoyancy force. The buoyancy force and the hydraulic gradient may act simultaneously on the groundwater. However, under certain conditions, the spatial differences in groundwater densities alone may induce a recirculating flow of groundwater (free convection of groundwater). The development of buoyancy-driven convection cells in sedimentary basins with normal geothermal gradients has been inferred to account for the volumes of groundwater necessary t o explain observed diagenetic characteristics of sedimentary rocks (e.g. Wood and Hewett, 1982;


71

Davis et al., 1985; Rabinowicz et al., 1985; Haszeldine et al., 1984; Gerretsen et al., 1991) and to explain observed chemical characteristics of groundwater around salt domes (e.g. Hanor, 1987a; Ranganathan and Hanor, 1988).

Free thermal convection The increasing temperatures with depth in a sedimentary basin cause a thermal expansion and hence a decrease in density of the groundwater with depth. This vertical density stratification may induce free convection of the groundwater if the critical Rayleigh number (R, = 40) is exceeded (e.g. Wood and Hewett, 1982).

where, k g aTw (PC),

AT P W

KTm

= permeability of the isotropic porous medium = acceleration due to gravity = volumetric thermal expansion coefficient of the groundwater = heat capacity of the groundwater = temperature difference over a vertical distance Az of the porous medium = dynamic viscosity of the groundwater = thermal conductivity of the water-saturated porous medium

Bethke (1989) calculated that a 100-m thick isotropic water-saturated permeable unit under a normal geothermal gradient requires a permeability of about 2 x 10-12 m2 (= 2D) t o satisfy the Rayleigh criterion; and a kilometre thick unit requires a permeability of about 2 x 10-14 m2 (20mD). Vertical anisotropy of a hydrogeological unit, caused by e.g. intercalations of poorly permeable layers (shales) in an otherwise homogeneous isotropic unit, increases the stability of a thermal stratification as compared with a corresponding isotropic unit of the same horizontal permeability (Bjorlykke et al., 1988; Bethke, 1989). The critical Rayleigh number is considered t o be a rather restrictive criterion in sedimentary basins, because of the common characteristics of the hydrogeological framework of basins (i.e. the wide range and heterogeneous distribution of permeabilities and thicknesses of hydrogeological units) (Bjgrlykke et al., 1988, 1989; Bethke, 1989). According to Bj~rlykkeet al. (1988, 1989) thermal convection is unlikely t o occur under normal geothermal conditions in sedimentary basins with horizontal hydrogeological units and horizontal isotherms. In basins, or parts of basins, with high geothermal gradients, free convection of groundwater may create an important groundwater flow system (Chapman and Rybach, 1985; Person and Garven, 1989). Figure 2.28 gives an overview of typical heat flow values associated with different types of sedimentary basin as given by Allen and Allen (1990).

Chapter 2

72

EXTENSIONAL BASINS

1 -

Active Ocean ridges and volcanoes

I;:;:-

Approximate global average heat flow

';...: : ! ] T I

:.:.:.I

120

Active (syn-rift) back-arc basins

85

Active (syn-rift) rift or passive margin

(post-rift) rift or passive margin

.. COMPRESSIONAL BASINS

..

1

I

1

Collisionalfoldbelt

70

I AI Ocean trench foreland basin (foothills-margin) 80

40

unrelated to arc magmatism

35

..... STRIKE-SLIP BASiNS Active strike-slip, deep lithosphere involvement 100 Active strike-slip, shallow thin-skinned (crustal) extension only

60 1. BASEMENT

... ..

Precambrian Shield

Oceanic crust (200Myr)

"""

(..A_..

20

40

60

80

I

I

1

I

I

1

mWm-2q

I

I

2

100

120

140

160

I

I

I

1

1

3

180 I

I

4

Heat Flow Units

F i g u r e 2.28 Typical h e a t flows associated w i t h different types o f sedimentary basins (from A l l e n and Allen, 1990. Reprinted by permission o f Blackwell Scientific Publications Ltd.).

73


c

Gravitational potentiel surface \

Ti

Figure 2.29 Schematic illustration of free thermal convection in sloping parts of a hydrogeological unit with isothermal and impermeable boundaries (after Wood and Hewett, 1982. Reprinted with permission from Geochimica et Cosmochimica Acta, Vol. 46, Copyright 1982, Pergamon Press Ltd.).

For sloping isotherms, and corresponding lateral differences in groundwater densities, free convection of groundwater may occur at any value of the Rayleigh number (Wood and Hewett, 1982)(Figure 2.29). The velocity of groundwater flow in the convection cells is affected by the lateral temperature gradient, the permeability of the hydrogeological unit and the geometry of the unit. Bjerlykke et al. (1988)calculated that for a normal geothermal gradient of 30 ' C h , the velocity of groundwater will be 10 d y e a r in a 10 metre-thick permeable unit with a permeability of 1D for a 15' slope of both the permeable unit and the isotherms. For a 1' slope, the velocity of the groundwater was calculated to diminish to 0.1 d y e a r . Large lateral changes in temperature can be expected to occur around intrusions in sedimentary basins e.g. igneous intrusions and salt diapirs. The free thermal convection of groundwater, as described here, is attributable to temperature-induced density differences of the groundwater only if it is assumed that there are no other driving forces for groundwater flow. This assumption is unlikely t o be met in sedimentary basins. According to Bethke (1989)there has been little work to determine the extent to which free convection persists in the presence of other groundwater flow systems.

Free thermohaline convection In sedimentary basins, the gradient of groundwater density resulting from the change in solubility of minerals with temperature, is very small (Bjerlykke

74

Chapter 2

et al. 1988). The density of groundwater is strongly influenced by the concentration of very soluble salts. Steep density gradients can be expected to occur near evaporites. Differences in groundwater salinity may cause a density flow of groundwater (e.g. salt-water intrusions in coastal aquifers, density flow around salt domes where active halite dissolution is taking place). Free convection of groundwater induced by density differences resulting from the combined influence of salinity and temperature gradients (thermohaline convection) has been studied near salt diapirs (e.g. Hanor, 1987a; Ranganathan and Hanor, 1988; Evans and Nunn, 1989). The calculations of Evans and Nunn (1989) indicate that upward flow of groundwater near the diapir edge is present over most of the vertical extent of the salt dome. This groundwater flow is considered to be caused by viscous drag of brines flowing off the top of the dome and upward buoyant forces caused by the increasing temperature with depth and a salinity inversion caused by preferential dissolution of salt a t the dome crest. The last effect is enhanced by salt diapirism and diminishes after diapirism ceases (Evans and NUM, 1989).

Methanogenic convection Park et al. (1990) suggested an alternative mechanism that may induce free convection in the subsurface: the methanogenesis-driven convection system. By numerically simulating reaction-transport equations, Park et al. (1990) calculated that the generation of methane from kerogen maturation can induce kilometre scale flow in a porous medium because of the density dependence of groundwater on methane concentration. 2.4.2 Osmotically-inducedgroundwater flow The flow of water through a semi-permeable membrane (clay, shale) from water with a small concentration of dissolved solids to water with a greater concentration is called osmosis (e.g. Bredehoeft et al., 1982; Neuzil, 1986). The osmotically-induced flow of water occurs because of a difference in vapour pressure across the membrane (Hinch, 1980). The aqueous activity will be relatively small in water with a relatively large concentration of dissolved solids, because more water molecules are bonded on the dissolved ions (Hinch, 1980). In a sandstone-shale sequence with water of equal chemical concentration, the aqueous activity of the shale water will be less than that of the sandstone-water, because water molecules are adsorped on the large mineral surfaces of the shale (Hinch, 1980). As a consequence, the water salinity differences that may exist in sandstone-shale sequences in the intermediate and deep subsystems of burial-induced groundwater flow may actually be in osmotic equilibrium.

The osmotically-induced flow of water across a semi-permeable membrane can cause superhydrostatic o r subhydrostatic pressures. When the osmotic pressure difference induced by the groundwater at opposite sides of the


75

membrane is exceeded by an externally applied force, groundwater may flow through the semi-permeable membrane in a direction opposite to the direction of osmotic flow. This reverse osmosis may cause membrane filtration (salt sieving or ultrafiltration). The role of semi-permeable membranes and the associated importance of osmosis in sedimentary basins is disputed (Neuzil, 1986).

2.5 Interaction of groundwater flow systems The four processes controlling the development of groundwater flow systems in sedimentary basins are sedimentation in a subsiding basin, introduction of heat into a basin, tectonic deformation of all or part of a basin and infiltration of meteoric water in a subaerial basin. During subsequent stages of evolution of a sedimentary basin, different processes or combination of processes may dominate the development of a groundwater flow system in the basin. A t a certain time during the evolution of a basin different groundwater flow systems may co-exist and interact. During active sedimentation in a subsiding basin, the burial-induced system of groundwater flow will initially be the main flow system. Local buoyancyinduced groundwater flow systems may develop in certain basins, e.g. in rift basins with high heat flow and in basins with magmatic intrusions, while in e.g. foreland basins tectonic driving forces may influence the burial-induced groundwater flow. In general, when the subsiding basin is surrounded by areas of high topographic relief, o r when part of the basin has emerged and stabilized above sea level, the regional groundwater flow system for the entire basin is determined principally by the interaction of the burial- and gravity-induced groundwater flow systems. In a simple subsiding and filling sedimentary basin that is surrounded by areas of high topographic relief, the burial-induced flow of groundwater is directed radially outwards from the basin’s depocentre towards its edges, where it will meet the oppositely directed gravity-induced groundwater flow. The position of the interface between the two groundwater flow systems is determined by the permeability distribution in the subsurface and the magnitude and directions of the driving forces for the burial- and gravity-induced flows. Bethke et al. (1988) and Harrison and Summa (1991) modelled the interaction of burial- and gravity-induced flow systems during the evolution of the Gulf of Mexico Basin, USA. They showed that the present-day gravity-induced groundwater flow system reaches depths of 2 km and extends tens of kilometres into offshore strata in the rapidly subsiding and filling Gulf of Mexico Basin. Changing rates of sedimentation and eustatic rises and falls

Chapter 2

76

in sea level were found t o be important in shifting the interface between the gravity- and burial-induced groundwater flow systems. Figure 2.30 shows the distribution of the gravity-induced system of groundwater flow in the Gulf of Mexico Basin during Miocene times and the present-day, as calculated by Harrison and Summa (1991). In the Miocene, meteoric water may have penetrated deeper into the basin in comparison with the present-day situation, because the opposing burial-induced forces were smaller during the Miocene (Harrison and Summa, 1991). Figure 2.31 shows the effect of a fall in sea level on the basinward extension of the gravity-induced groundwater flow system. N

S

Present day 01-

2-

3-

!i 2:

1 ;09-

10

Miocene

0

so

lWkm

Gravity-induced groundwater flow system Burial-induced groundwater flow system Vmax Maximum velocity of groundwater flow Figure 2.30 Calculated distribution of gravity- and burial-induced groundwater flow systems in the Gulf of Mexico Basin for present-day and Miocene times (after Harrison and Summa, 1991, American Journal of Science, Vol. 291, Fig. 25. Reprinted by permission of American Journal of Science).

77


N

8

200 rn fall in sea level

0

elo-

0.0

I 0.001

0.01

c

average initial hydrocarbon potential ( k g / k g o f rock)

Figure 3.14 Variation of average bulk oil expulsion efficiency with average initial hydrocarbon potential (after Cooles et al., 1986. Reprinted with permission from Organic Geochemistry, Vol. 10, Copyright 1990, Pergamon Press Ltd.).

facilitates the escape of liquid hydrocarbons. Cooles et al. (1986) found very high expulsion efficiencies associated with high maturities of the source rock irrespective of the initial petroleum potential of the source rock. They believe that these high efficiencies are the result of oil cracking to gas at these high maturity levels.

For a certain geological heating rate, the petroleum expulsion efficiency, in combination with the type and richness of a source rock, determine whether oil, gas condensate o r gas will be expelled over a certain temperature range (Mackenzie and Quigley, 1988; Figure 3.15). For the same source rock, the expulsion efficiencies increase with increasing hydrocarbon generation rates. The highest expulsion efficiencies of oil occur in rich source rocks containing type 11kerogen during the peak phase of oil generation, when oil is expelled in a separate phase (Figure 3.16). The expulsion efficiencies in a single source rock may vary due to capillary effects (Leythaeuser et al. 1987,1988; Mackenzie et al. 1987, 1988). The capillary forces enhance the efficiency of expulsion close t o the interfaces between source rock and carrier rock. Thin source rock layers may exhibit higher expulsion eficiencies than thick source rock layers because of the capillary effects near the edges of the source rocks.

Generation and expulsion of hydrocarbons

117

CLASS1 1

PGI 0 1

PEE 0

Petmkum exp4ed

1

PGI 0 1

PEE 0

Petrokum expelled CLASS3 1

PGI 0 1

PEE 0

Petmkum expelled

80

120

160

200

240

PGI = Petroleum Generation Index PEE = Petroleurn Expulsion Efficiency

Figure 3.15 Principal petroleum phases expelled from three classes of source rock over relevant temperature ranges assuming a mean heating rate of 5 'C per million years (after Mackenzie and Quigley, 1988. Reprinted by permission of the American Association of Petroleum Geologists).

Figure 3.16 shows that some expulsion does occur prior to the peak phase of oil generation. During this early expulsion stage the edges of thick source rock layers and the thin source rock layers are depleted (Leythaeuser et al., 1987). The rate of hydrocarbon expulsion from mature rich oil-prone source rocks is about 8 x 10-15 to 8 x 10-14 mh-%-I, according to a rough estimate made by England et al. (1987). England et al.'s calculations are based on the subsurface conditions given in Table 3.4.

Chapter 3

118

1

I

I

I

I

0

-1

0

n 0.5

-

0.8

-

n

early expulsion

x

-

L

L

3

+ m

E

peak generation

0-

Figure 3.16 Schematic representation of amount of hydrocarbons generated in, and expelled from a type I1 kerogen-bearing source rock as a function of organic matter maturity for the initial part of the oil window (after Leythaeuser e t al., 1987. Reprinted with permission from the Proceedings 12th World Petroleum Congress, Houston, Vol. 2, Fig. 2a, p. 229).

Table 3.4 Assumed average subsurface conditions for rich oil-prone source rocks

TOC Hydrogen index Petroleum generated Expulsion efficiency Temperature range of hydrocarbon generation Geological heating rate Source rock thickness Source rock density Petroleum fluid density Data from England e t al., 1987

= 4% weight = 0.7 k&g TOC = 0.02 kgkg rock

= 100% = 120- 150 "C = 1- 10 'Clmillion year = 100 m = 2400 kg m a = 650 kg m a

Generation and expulsion of hydrocarbons

119

3.3 Summary Hydrocarbons are generated from finely disseminated organic matter in fine-grained sedimentary rocks, such as shales, mudstones and fine-grained carbonates. When sediments rich in organic matter are buried as caused by continued sedimentation in a subsiding sedimentary basin, successive steps of evolution of organic matter occur, i.e. diagenesis, catagenesis and metagenesis. Kerogen, the insoluble part of organic matter in sedimentary rocks that is formed during diagenesis, is the main precursor of hydrocarbon compounds. Temperature increase is of primary importance for hydrocarbon generation. During burial, the kerogens (type I, 11, and 111) are altered, with trends to lower hydrogen to carbon and oxygen t o carbon ratios due to the generation and liberation of hydrocarbons, carbon dioxide and water. Some organic matters (kerogens) are less able to generate oil than others, but any organic matter may generate gas, provided it is buried sufficiently deeply for long enough. Kerogen type 111, the insoluble organic matter derived from terrestrial plants, generates comparatively less oil than the kerogen types I or 11. At greater depths type 111 kerogen, and coal, may be a good source of hydrocarbon gas. For average geothermal gradients and geological heating rates, the principal zone of oil generation occurs at burial depths between ca 2500 and 5000 m, corresponding to temperatures between 100 and 150 'C. Hydrocarbon gases (methane) are generated simultaneously at these depths. Significant gas generation occurs at burial depths beyond 4000 m until ca 7000 m, corresponding to temperatures of 150 - 220 "C. Thermal cracking of oil to gas takes place in the temperature range 150 - 190 'C. Different primary migration mechanisms are probably responsible for the transport of hydrocarbons through the hydrocarbon-generating source rocks in sedimentary basins. The three major mechanisms of primary hydrocarbon migration seem to be: - primary migration of continuous separate phase hydrocarbons (oil and gases and mutual solutions of these) driven by hydrocarbon potential gradients, including capillary-pressure driven primary migration of separate phase hydrocarbons (oil and gases) - groundwater-driven primary migration of hydrocarbons in aqueous solution (mainly gases and t o a less extent the most water-soluble liquid hydrocarbons) - diffusion-driven primary migration of hydrocarbons (principally gases) in aqueous solution, and diffusion-driven primary migration of hydrocarbons through organic matter network.

120

Chapter 3

The effectiveness of the different primary migration mechanisms is determined by the subsurface characteristics of both the hydrocarbons to be moved and the medium through which the movement takes place. During the peak phase of hydrocarbon expulsion, the hydrocarbons are transported through the already low permeable source rock mainly as a continuous single phase fluid in a vertically upward o r downward direction. The driving forces for the separate phase hydrocarbon migration are large hydrocarbon potential gradients which are related t o large groundwater potential gradients generated in the source rock by e.g. sedimentary or tectonic loading, volume expansion of organic matter during hydrocarbon generation, thermal expansion of groundwater, and infiltration of meteoric water. Under average geological conditions, the active groundwater-driven transport of both liquid and gaseous hydrocarbons in aqueous solution is probably not important a s a primary migration mechanism at depths corresponding t o the peak phase of oil and gas expulsion from good source rocks. A t relatively shallow depths active groundwater flow may be an important migration mechanism for (biogenic) gaseous hydrocarbons. Possibly primary solution migration of the more water-soluble constituents of liquid hydrocarbons takes place in good source rocks at shallow depths corresponding to early phases of oil generation and in leaner source rocks at depths corresponding to the entire phase of oil generation. Diffusion of light hydrocarbons through water-saturated source rock may be active at all depth ranges. It is considered to be a relatively minor transport process in comparison with separate phase hydrocarbon migration. It may influence, however, the composition of the hydrocarbons expelled from the source rock. In rich source rocks with practically negligible matrix permeability, primary migration may possibly also occur by activated diffusion of hydrocarbons through organic matter network. The primary migration mechanism or combination of mechanisms by which the hydrocarbons are moved through and expelled from the source rock, is of influence on the geochemical composition of the expelled hydrocarbons. Oil expulsion from good oil-prone source rocks is very efficient, whereas oil expulsion from leaner source rocks is relatively inefficient. Probably, most of the oil generated in leaner oil-prone source rocks will remain in the source rock and be cracked t o gas a t higher temperatures and expelled as gas condensate followed by dry gas. Gaseous solution can be an effective migration mechanism for oil generated from mature type I11 kerogen. The expulsion of gas is very efficient .

I21

CHAPTER 4

SECONDARY HYDROCARBON MIGRATION

Secondary hydrocarbon migration is the movement of hydrocarbons after expulsion from a source rock through carrier and reservoir rocks or fault and fracture systems. During the peak phase of hydrocarbon expulsion, the hydrocarbons are probably mainly expelled as a single phase fluid (Chapter 3). Expulsion of hydrocarbons in aqueous solution may be of importance before the start of the peak phase of hydrocarbon expulsion from good source rocks, and possibly during the entire phase of hydrocarbon generation in leaner source rocks. After expulsion from the fine-grained source rock the hydrocarbons enter the water-saturated relatively coarse-grained carrier rock or reservoir rock. The magnitude and direction of the driving forces for hydrocarbon migration through the carrier rock are influenced by, amongst other things, the hydrodynamic conditions in the rock and the form in which the hydrocarbons move. Sections 4.1 and 4.2 give information on the driving forces for secondary hydrocarbon migration under hydrostatic and hydrodynamic conditions, respectively. The characteristics of a hydrocarbon migration system in a sedimentary basin at a certain moment during its evolution are defined by the masses and composition of hydrocarbons available for migration and the starting point of hydrocarbon migration (i.e. by the location, type, richness and maturity of the hydrocarbon-generating source rocks); the hydrogeological framework of the sedimentary basin; and the hydrodynamic condition of the basin. Different hydrocarbon migration systems may coexist and interact in the same sedimentary basin. During different stages of evolution of a basin, different hydrocarbon migration systems may develop. The secondary hydrocarbon migration system distributes the hydrocarbons in a sedimentary basin in a way that may lead to either concentration of hydrocarbons into economic accumulations or loss of hydrocarbons due to dispersion, destruction or escape into the atmosphere (e.g. Demaison and Huizinga, 1991). Secondary hydrocarbon migration systems are treated in Section 4.3. In Section 4.3 attention is focussed on the regional pattern of hydrocarbon migration.

Chapter 4

122

4.1 Secondary hydrocarbon migration under hydrostatic conditions Generally speaking, hydrocarbons may leave the fine-grained source rocks as separate phase hydrocarbons, or as hydrocarbons in aqueous solution. The primary migration and expulsion of dissolved hydrocarbons occurs principally under hydrodynamic circumstances, although light hydrocarbons may also be expelled by molecular diffusion under hydrostatic conditions (Chapter 3). Molecular diffusion of light hydrocarbons through water-saturated carrierreservoir rocks is not thought t o be significant as a secondary migration process (Leythaeuser et al., 1982, England et al., 1987). In this section, description of the secondary hydrocarbon migration under hydrostatic conditions will be restricted to separate phase hydrocarbon migration. After expulsion from the source rock, the separate phase hydrocarbons may initially form small oil globules or gas bubbles in the larger pores of the coarsegrained carrier rocks. With continued supply of hydrocarbons from the source rock, the oil globules or gas bubbles grow in size and become subject to buoyancy forces (Section 4.1.1).Upward secondary migration of oil or gas in the watersaturated carrier rock will occur whenever the buoyancy force is larger than the capillary force that resists this migration (Sections 4.1.2 and 4.1.3).The nature of the multiphase flow through the carrier rocks is laminar (nonturbulent) and is dominated by capillary forces in the pore network of the carrier rocks (England et al., 1987). 4.1.1 Buoyancy

In an isothermal and isochemical body of water under hydrostatic conditions, there is no flow of water. The potential of the water is constant, at any point in the water, and assuming the density of the water is constant, the potential can be given by (Chapter l), +w

= gz + PW = constant

(4.1)

Pw

where, g

= potential of the water (L?F2) = acceleration due to gravity ( L P 2 )

Z

= elevation (L)

Pw

= pressure of the water (ML-lY2) = density of water (ML-3)

+W

Pw

Hence, the net force acting on a unit mass of water is zero (Figure 4.1) (Section 2.1.1).

-

Ew=-grad +,=O

(4.2)

Secondary hydrocarbon migration

t 7

-grad p,

pw

E,=O

Figure 4.1 The magnitude and direction of the net driving forces for separate phase oil and gas migration under hydrostatic conditions.

The net forces acting on a unit mass of separate phase oil or gas completely immersed in water under isothermal and isochemical conditions, as given by Hubbert (1953),is: (4.3) PO

where, = g z + -Pw $0

Po $0

Po

= potential of the oil ( m 2 ) = density of the oil ( m a ) (4.4)

where,

%J

p,

= potential of the gas (LV2) = density of the gas (ML-3)

These net forces acting on oil and gas can also be expressed in terms of gravity, and the net force acting on water (Hubbert, 1953):

Chapter 4

124

(4.5)

(4.6) From Equations 4.5 and 4.6 it follows that since g is fixed, the direction and magnitude of the net force acting on a unit mass of oil or gas immersed in water under isothermal and isochemical conditions depend upon the density of acting on the the oil o r gas and the direction and magnitude of the net force water.

ew

eo eg

and will be directed Under hydrostatic conditions Rw = 0 and therefore vertically upwards under these circumstances and Equations 4.5 and 4.6 will reduce to (4.7) and

These upward directed forces are generally called buoyancy forces and are the driving forces for separate phase oil o r gas migration through water under hydrostatic conditions. The magnitude of the buoyancy force for a vertical length zo of a body of oil or gas immersed in water can be expressed by

or

E, = zog

(Pw - P g )

(4.10)

Equations 4.9 and 4.10 show that the greater the density difference between the immiscible phases oil and water, or gas and water, the greater the buoyancy force for a given vertical length of the body of oil or gas. Small changes in the density of either hydrocarbon o r water will be of great influence on the magnitude of the buoyancy force (Davis, 1987). For a subsurface density of oil po= 600 kg m-3, a subsurface density of gas pg = 200 kg m-3, a groundwater density of pw = 1000 kg m-3, the buoyancy force per unit vertical length of the immersed hydrocarbon body is 4000 Pa m-1 and 8000 Pa m-1, respectively. For the same densities of oil and gas, and a groundwater density of pw = 1200 kg m3 the buoyancy forces increase to 6000 and 10,000 Pa m-1, respectively.


125

The buoyancy force is also considered a driving force for secondary separate phase hydrocarbon migration through water-saturated rocks (Schowalter, 1979). However, in the subsurface the hydrocarbons must migrate through the rock pores and they encounter a resistant force to movement when the diameter of the oil globule or gas bubble is larger than any throat connecting the pores of the rock. This resistant force is indicated as capillary pressure. The net force acting on a unit mass of hydrocarbon in water-saturated rock under isothermal and isochemical conditions, as given by Hubbert (1953),is

(4.11) where, PC subscript hc

= capillary pressure (ML-'T-~)

= hydrocarbon

4.1.2 Capillary pressure When two immiscible fluids (or a fluid and a gas) are in contact, molecular attractions between similar molecules in each fluid are greater than the attractions between the different molecules of the two fluids and a clearly defined interface exists between them. The force that acts on this interface is called interfacial tension (or surface tension in case of a gas-fluid contact). As a result of this force, a pressure difference exists across the interface. This pressure difference is known as capillary pressure and is given by the following equation (Dake, 1978): (4.12) where, PC

Y-+1 1'

1 2'

= capillary pressure (ML-lY2) = interfacial tension (MY2) = curvature of the interface; rl and r2 are the two principal radii measured in planes at right angles to each other and normal to the interface (L) = pressure in the hydrocarbon (ML-lT2) = pressure in the water (ML-"F2)

Phc Pw When an immiscible fluid or a gas is completely immersed in another fluid it assumes a spherical shape of minimum surface area. The curvature of the interface is spherical and (l/rl + l/r2)in Equation 4.12can be replaced by 2/r:

2Y Pc =I-

(4.13)

When two immiscible fluids are in contact with a rock surface, the capillary pressure is also influenced by the wettability of the rock. The wettability of a

Chapter 4

elevation1

Pv'=Po

capillary tube'

pressure

Figure 4.2 Capillary tube experiment for two immiscible fluids (after Dake, 1978).

rock is expressed by the contact angle 8 of the hydrocarbon and water against the solid pore walls as measured through the water phase (Figure 4.2). For rock-fluid systems with contact angles between 0" and go", the rocks are generally considered water-wet; for contact angles greater than go", the rocks are considered oil-wet. For a hydrocarbon-water-rock system, the capillary pressure is given by (e.g. Dake, 1978).

(4.14) where, pd

= hydrocarbon-water displacement pressure (ML-'T2)

R

as measured through the water phase = radius of largest connected pore throats in the rock; R defines the radius of curvature of the hydrocarbon-water interface R = r cos 8

e

= contact angle of hydrocarbon and water against the solid pore walls

The displacement pressure is a rock property and is defined as the force required t o replace water from a cylindrical pore with oil or gas. Hence the displacement pressure determines the minimum buoyancy pressure needed for migration. Secondary hydrocarbon migration generally occurs through water-saturated sedimentary rocks, i.e. through rocks that are water-wet. As water is generally considered a perfect wetting fluid (Schowalter, 1979),the contact angle 8 in Equation 4.14 for hydrocarbon-water-rock systems can be taken to be zero. If, in addition, the hydrocarbon-water interface is assumed to be spherical, then Equation 4.14becomes identical with Equation 4.13.


127

The magnitudes of the displacement pressure, and capillary pressure, are inversely proportional to the largest connected pore throats in the rock, R (R < radius of rock pore). The average pore diameters in shales are around 5 . 10-9 10 . 10-9 m a t depths of about 2000 m and will be smaller at greater depths (Tissot ant Welte, 1984). Assuming an interfacial tension of y = 0.03 N m-1, the capillary pressures in such shales may reach values of 10 MPa or more. The largest available pores are generally one or two orders of magnitude larger than the average pore sizes (Mann, 1989). For example, England et al. (1987) give a magnitude cd 10-8 m for the mean radius of the larger interconnected pores in shales at depths of approximately 3 km, and a mean radius of 10-6 m for sandstones at the same depths. The capillary pressure in the interconnected pores of sandstones with a pore radius of 10-6 m and an interfacial tension of y = 0.03 N m-1 is 0.06 ma. 4.1.3 Separate phase hydrocarbon migration Berg (1975) amongst others, has described separate phase hydrocarbon migration from a rock pore through an adjacent pore throat in a water-wet rock under hydrostatic conditions. The following outline of the migration process is largely based on his work. The capillary pressure of a spherical globule of oil (or gas bubble) that is in equilibrium with the surrounding pore water is given by Equation 4.13 (Figure 4.3a):

where, r 5 rp rP

= radius of the rock pore

A

0

Pc = 2Y

%>% rt

D

C

rp

2Y r

-

2Y r

a a

-<rp

rt

Figure 4.3 Migration of an oil globule through pore throats in a water-wet rock (after Berg, 1975. Reprinted by permission of the American Association of Petroleum Geologists).

Chapter 4

128

When the buoyancy force is large enough, the globule is forced upwards through an adjacent pore throat and the globule will be distorted (Figure 4.3b). The capillary pressure at the upper end of the distorted globule in the pore throat is: pt = 2y/rt. The capillary pressure is greater in the throat than in the pore: (2y/rt) > (2yh ). In this situation, the capillary pressure gradient is P directed downwards in opposition to buoyancy (Berg, 1975): (4.15) ZO

where, ZO

= vertical height of the globule

When the buoyancy force is greater than the capillary pressure gradient, the globule can rise further. The next situation encountered by the globule is depicted in Figure 4 . 3 ~ .The globule has moved halfway through the throat where the radii are equal a t the upper and lower end of the globule and the capillary pressure gradient has become zero. The globule can easily continue to rise by buoyancy . As soon as the globule attains a position more than halfway up the pore throat, the radius at its upper end is larger than at the lower end, and the capillary pressure gradient and the buoyancy force both act in the same upward direction. Secondary migration of an oil globule or gas bubble in a water-wet rock will occur if the buoyancy force acting on an oil globule or gas bubble is large enough t o overcome the capillary pressure gradient within the globule or bubble caused by the greater pressure in the pore throat than in the pores, i.e. (Berg, 19751,

In an equilibrium situation (Berg, 1975): (4.16) Under a given subsurface condition (i.e. a given hydrogeological framework and pressure, temperature, chemical conditions), the only variable in Equation 4.16 is the vertical height of the oil globule or gas bubble (i.e. the hydrocarbon column). The value of zo may increase by the accretion of the oil globule/gas bubble, e.g. because of a continued supply of oil or gas from the source rock. The maximum height of an oil globule or gas bubble that can be held in place, is called the critical height z, and is given by (Berg, 1975):


129

(4.17) For migration to occur a continuous oil stringer must extend through the interconnected pores of a water-saturated carrier rock (e.g. Berg, 1975;England et al., 1987; Schowalter, 1979). England et al. (1987) calculated that the hydrocarbons must fill c. 50% of all available carrier rock pore volume (which probably corresponds to 1 t o 10% of the rock’s cross-sectional area which represents the more coarse-grained parts) in order to create an interconnected pathway allowing hydrocarbon movement to occur. Secondary vertical upward migration of hydrocarbons through a carrier rock can continue as long as the buoyancy force of the hydrocarbon column is greater than the resistant force of the carrier rock. The hydrocarbons will migrate in a tortuous manner, focussed into the path of lowest resistance by moving through part of the rocks with the largest connected pore throats or lowest displacement pressure, i.e. through the larger pores and the coarsegrained parts of the carrier reservoir rock (Dembicki and Anderson, 1989; England et al., 1987; Schowalter, 1979). The smaller pores will remain waterfilled. A t the base of a migrating hydrocarbon stringer small isolated droplets or bubbles will be left behind as the stringer migrates upwards. The residual hydrocarbon droplets or gas bubbles are permanently trapped by capillary forces. The amount of hydrocarbons left behind in a carrier rock will be approximately equal to the amount necessary to form an interconnected pathway initially (England et al., 1987). The capillary-trapped gas and the soluble portion of the residual oil may dissolve in the pore water and dissipate by diffusion. As the migrating stringer loses oil at its base, the vertical height of the hydrocarbon column diminishes and consequently the buoyancy force will be reduced. Eventually, migration will stop until additional hydrocarbons migrate upwards along the same pathway to the stalled stringer, thus increasing the vertical height of the hydrocarbon column, i.e. the buoyancy force, and migration will continue. This continuous pulsating migration process, as described by Schowalter (19791, will continue as long as hydrocarbons are being added up. When the thus migrating hydrocarbons encounter an overlying rock with small pore diameters that exert capillary pressures that are too high to be overcome by the buoyancy force of the hydrocarbons, initially, the hydrocarbons will spread out along this seal boundary (barrier rock) (Figure 4.4). If the barrier rock is inclined, the hydrocarbons will start to migrate laterally updip along the seal boundary perpendicular to its strike as soon as the critical height of the hydrocarbon column is attained again by addition of hydrocarbons. For a dip of the carrier rock - barrier rock interface (a),the length of the hydrocarbon stringer (L) needed to obtain the critical vertical hydrocarbon column height (z,) is given by

130

Chapter 4 a

. . . _ . ., . , . ;. . . . . . . . :.. .' . .. . . .. _ . . . .. . . . .; .. .. .

U g r a t i o n of finely dispersed hydrocarbons Initially. the hydrocarbons entering a t the base o f a horizontal carrier rock are very finely dispersed and the buoyancy forces are s t i l l too small t o initiate hydrocarbon migration.

b

Vertical upward hydrocarbon migration through carrier

rock

Continued supply of hydrocarbons from the source rock increases the vertical height of the hydrocarbon column lz,) As soon as z, is large enough, i e as soon as the buoyancy force of the hydrocarbon column is greater than the resistant force of the carrier rock, vertical upward migration through the carrier rock w i l l start

c

... .. ..... .... . . . . '

. .. ..

Capillary pressures excerted by the barrier rock resist vertical upward hydrocarbon migration Hydrocarbons accumulate and spread out along the horizontal barrier rock - carrier rock interface.

. . . ., .

d

-..

..

. q T 1 )

= constant

- Obsewed temmratures Figure 7.1 Variations of temperature with depth for measured temperatures and for temperatures calculated for two assumed conductive thermal equilibrium conditions.

Qualitative analysis of secondary hydrocarbon migration systems

219

plotted on maps for appropriate depth levels. Areas of high heat flow and low heat flow for a certain depth can be delineated on the map and compared with the previously identified pattern of regional groundwater flow. Salinities of groundwater are relatively low in recharge areas and relatively high in discharge areas. The chemical evolution of the groundwater in the direction of flow is related t o a distinctive sequence of water types. Chebotarev (1955)showed that the water type changes from bicarbonate to sulphate to chlorine from areas of recharge to areas of discharge. The interpretative significance of the hydrochemical composition of the groundwater for evaluating gravity-induced groundwater flow systems is widely recognized by hydrogeologists (e.g. Back, 1960;BredehoeR et al., 1982;Engelen and Jones, 1986; Herczeg et al., 1991;T6th, 1978,1980). The interpretation of water analysis data should include the identification of hydrochemical water types. Graphical representations of water analysis data provide means for identifying hydrochemical water types. Well known techniques for graphical representation of analysis data or for representing the absolute or relative contents of single ions or their ratios as used in hydrochemical groundwater and/or petroleum studies may be applied (e.g. Collins, 1975; Hem, 1985; Domenico and Schwartz, 1990). Widely used graphical representations include the bar graphs introduced by Collins (1923),the Stiff pattern diagram (Stiff, 19511, the pie-diagram and the trilinear Piper diagram (Piper, 1944) (Hem, 1985). The spatial changes in groundwater composition at shallow depths, along an inferred groundwater flow path, or in a regional aquifer can be studied on a map or cross-section by plotting the graphical representations for the different sample points or by contouring salinities and concentrations of single ions a t their ratios. Herzceg et al. (1991) illustrate the change in concentration in a single ion from an inferred recharge to discharge area with concentrations of single ions plotted against distance from the recharge zone. In addition to these graphical methods, statistical methods can be helpful in evaluating the available hydrochemical data (e.g. Davis, 1986). The procedure to combine readily available information on the driving forces for groundwater flow (i.e. the relief of the water table) and the hydrogeological framework of the basin with the various indicators of gravity-induced groundwater flow leads t o the identification of the main recharge and discharge areas and the general pattern of gravity-induced groundwater flow, including preferred paths of groundwater Sow, in all or part of a subaerial basin.

7.1.3.2 Hydrodynamic conditions in subsiding and filling basins Each of the three subsystems of burial-induced groundwater flow that may develop in a subsiding and filling basin modifies the migration, accumulation and entrapment of hydrocarbons in a specific way in comparison with theoretical hydrostatic conditions for that basin (Table 6.1). Large groundwater

220

Chapter 7

potential gradients and concentrated flow of groundwater in the intermediate and deep subsystems of burial-induced groundwater flow are of major influence on the migration, accumulation and entrapment of the hydrocarbons. Whether all three subsystems of groundwater flow are present in a certain basin depends on the basin’s subsidence rate and permeability characteristics. In rapidly subsiding shaly basins with burial rates of more than 1 m d y e a r all three subsystems may occur (Chapter 2). Present-day systems of hydrocarbon migration are being considered here, and therefore the present conditions of burial-induced groundwater flow that result from geologically recent deposition of sediments should be established. An appropriate geological time interval should be chosen for the analysis of burial-driven flow in all or part of the basin. Which time-interval is chosen depends on the geological evolution of the basin, the rate of sedimentation during different geological time intervals and the resulting total thickness and lateral thickness variations of the sediments in each time interval. Subsequently, a depth-contour map of the base of the sedimentary fill of the basin as deposited during the chosen time-interval, should be selected from published literature or constructed from geological and geophysical data. The thickness of the sedimentary fill of the basin gives a first indication of the subsidence rates in the basin during the chosen time-interval. The depocentres of a sedimentary basin contain the maximum thickness of the geologically recent sedimentary fill and are the most likely parts of the basin for the occurrence of all three subsystems of burial-induced groundwater flow. In the depocentres geopressured conditions may occur in the geologicalIy recent sedimentary fill and in the underlying rocks. The geometry of the base of the sedimentary fill can be used to identify qualitatively the potential directions of lateral burial-induced groundwater flow in the intermediate and deep subsystems. The lateral component of the burial-driven groundwater flow in the sedimentary fill and in the underlying rocks will be directed from the depocentreb) of the appropriate part of the basin to its edges (Figure 2.15). The actual occurrence of the three different subsystems of burial-induced groundwater flow can be verified from published groundwater pressure data and from pressure indicators (drilling mud weight, and other drilling and well logging, surface and borehole seismic data; Chapter 6). The combined interpretations of both drilling and logging indicators is outlined in detail by Fertl(197f3, Sahay and Fertl(1989) and Magara (1978) for quantitative pressure evaluations. As outlined in Chapter 6, the application of these procedures is restricted t o superhydrostatic conditions associated with undercompaction of siliciclastic sedimentary rocks. Provided these conditions occur in the studied basin, the depth of the interfaces between the different subsystems of burialinduced groundwater flow and the areal distribution of the subsystems can be estimated from the analysis of the pressure indicators in combination with the groundwater potential characteristics of the subsystem given in Section 2.1.3.


221

Hydrocarbon migration can be greatly affected by concentrated flow of groundwater. In a burial-induced groundwater flow system, appreciable concentrated upward directed flow of groundwater occurs, for example, along the basin edges, salt diapirs, permeable faults or other vertical escape ways. This concentrated flow of groundwater induces positive pressure, temperature and salinity anomalies and associated mineral assemblages in the shallower parts of the basin (Section 2.1.3). Hence, areas where upward-directed groundwater flow is concentrated can be recognized from the analysis of temperature, salinity and geochemical information (see also Section 7.1.3.1); e.g. by identifying the location of temperature anomalies on geothermal gradient maps, heat flow maps or isotherm maps, or the location of positive salinity anomalies from isosalinity maps. The qualitative assessment of the hydrodynamic conditions in a filling and subsiding basin results in the identification of the areal distribution of the different subsystems of burial-induced groundwater flow, the general pattern of lateral burial-induced groundwater flow and the identification of zones of concentrated upward directed flow of groundwater.

Identification of the hydrodynamic influence on the hydrocarbon migration system The application of the above described stepwise procedure of analysis results, amongst other things, in four basin-wide maps. The first map shows the currently active and/or nonactive depocentres of the sedimentary basin. The locations of the depocentres indicate the most probable locations of mature and/or overmature source rocks and the probable starting points of secondary hydrocarbon migration. Each of the three remaining maps shows the potential area of influence of a single driving force or combination of driving forces for secondary hydrocarbon migration: the second map shows the area of influence of buoyancy and capillary forces: and the third and fourth map show the area of influence of the two main driving forces for groundwater flow as controlled by sedimentation in a subsiding basin and infiltration of meteoric water in a subaerial basin, respectively.

7.1.4

The four maps in combination with additional information obtained in the analysis on e.g. depths of influence and intensity of each of the driving forces, permit the identification of the main characteristics of the present-day migration pattern (Table 7.1). Whether or not the identified hydrodynamic conditions in the basin will modify the migration, accumulation and trapping conditions for hydrocarbons in a certain basin in comparison with theoretical hydrostatic conditions for that basin depends on the type or types of groundwater flow system in the basin and on the location of the groundwater flow systems in the basin with respect to the location of the hydrocarbon generating source rocks (Chapters 4 and 5). For

222

Chapter 7

Table 7.1 Identification of hydrocarbon migration patterns

Hydrodynamic influence

Hydrodynamic condition

no

Yes

Characteristics migration pattern Hydrostatic migration pattern; preferred lateral migration towards parts of basin showing largest dips

Burial-induced shallow subsystem intermediate subsystem

deep subsystem

k Hydrostatic migration pattern Lateral migration pattern; length migration path differs from hydrostatic path (longer or shorter) Restricted lateral migration; focussed vertical migration

Gravity-induced

Lateral migration pattern; enhanced migration towards discharge areas; length migration path differs from hydrostatic path (longer or shorter)

Tectonically-induced

Enhanced focussed migration

example, very shallow gravity-induced groundwater flow systems probably do not reach depths where hydrocarbon migration takes place, and burial-induced groundwater flow systems in very slowly subsiding basins probably do not change the theoretical hydrostatic hydrocarbon migration and accumulation system. By comparing the map showing the pattern of theoretical hydrostatic hydrocarbon migration with the maps showing the different patterns of groundwater flow associated with the different identified groundwater flow systems, it can be established whether the location and directions of preferred flow paths for hydrocarbon and groundwater coincide, and whether under hydrodynamic conditions the potential distribution of migrated hydrocarbons valid for hydrostatic conditions is maintained, reinforced or changed. Figure 7.2 gives an example of such a comparison of maps. Figure 7.2 presents maps of the theoretical patterns of hydrostatic hydrocarbon migration and of the


hydrocarbon migration directions-

223

structure contours o f basin f l o o r

basin axis

'ective depocentre

a

"pLe

clrcular symmetrical basin No p r e f e r r e d hydrocarbon migration directions

a

5 A p i e circular s y m e t r l c a l basin No concentration o f groundwater flow

B

B

c

C

c

A

C

A

c

Simple elonqate symmetrical basin P r e f e r r e d hydrocarbon migration t o w a r d s t h e long f l a n k s A and B o f the basin migration preference A and B over C

c

Simple elonqate symmetrical basin Cencentration o f groundwater flow towards the long f l a n k s A and B o f the basin

e

Simple elonqate symmetrical curvedP r e f e r r e d hydrocarbon rnigratlon t o w a r d s concave l o n g flank A migration preference A over B over C

e

Simple elongate symmetrical curved basin Concentration o f groundwater flow t o w a r d s concave long flank A o f the basin

Figure 7.2 Comparison of maps showing patterns of theoretical hydrocarbon migration with maps showing groundwater flow patterns in intermediate subsystems of burial-induced flow for three hypothetical hydrodynamically and geometrically simple basins.

groundwater flow patterns in intermediate subsystems of burial-induced flow for three hypothetical hydrodynamically and geometrically simple basins. In such hydrodynamically and geometrically simple subsiding basins, the lateral flow of groundwater in an intermediate subsystem of burial-induced flow, is directed from the depocentre to the basin's edges and will be parallel to the directions of secondary hydrocarbon migration as induced by buoyancy forces

224

Chapter 7

alone (Figure 7.2). The lateral hydrocarbon migration distance increases because of the hydrodynamic condition in such a basin and the hydrocarbons will accumulate below fine-grained rocks of hydrodynamically increased sealing capacity, at greater lateral distances from the depocentre than in a comparable hydrostatic basin. Changing the hydrodynamic conditions in these geometrically simple basin to those associated with a deep burial-induced subsystem of groundwater flow will result in a completely different hydrocarbon migration system. The influence of the deep burial-induced flow system is t o restrict the lateral hydrocarbon migration and enhance focussed vertical migration. In addition t o the importance of hydrodynamic conditions for basin evaluation that results from their modifying influence on secondary hydrocarbon migration and entrapment in a basin, certain hydrodynamic conditions can also be considered as direct indicators of potential hydrocarbon migration paths. For example, preferred paths of groundwater flow may coincide with preferred paths of secondary hydrocarbon migration (e.g. zones of concentrated vertical upward flow of groundwater from a deep burial-induced subsystem of groundwater flow may indicate vertical escape ways for hydrocarbons; a pattern of lateral groundwater flow focussed through a regional aquifer directly indicates a potential lateral migration pattern for light hydrocarbons in aqueous solution andlor hydrocarbons very finely suspended in water). After having established the present-day hydrocarbon migration characteristics in a basin, evaluation of these characteristics leads to the selection of areas where preferred migration of hydrocarbons takes place, and, consequently, provides insight into the potential basin-wide distribution of hydrocarbons.

7.2 History of hydrocarbon migration systems

Hydrocarbon generation from mature source rocks may continue over a long geological time interval. When the secondary hydrocarbon migration pattern changes over such a time interval, the ultimate distribution of the hydrocarbon accumulations will be the common result of different patterns of hydrocarbon migration. Knowing when hydrocarbon formation started, the history of secondary hydrocarbon migration patterns can be reconstructed qualitatively from the geological evolution of the sedimentary basin. Secondary hydrocarbon migration is induced by hydrocarbon potential gradients, which in turn result from buoyancy forces, capillary forces and net


225

driving forces for groundwater flow. As a first step in assessing the history of hydrocarbon migration patterns, the areal influence of the net driving forces for groundwater flow should be established for subsequent time periods. The hydrogeohistory of a basin, i.e. the evolution of its hydrodynamic conditions and associated groundwater flow systems, can be reconstructed qualitatively from the basin’s geological evolution by identifying the periods of general subsidence and sedimentation, periods of non-sedimentation, uplift or erosion or periods of a low stand of sea level, and periods of increased tectonic activity. For each time period the theoretical hydrostatic hydrocarbon migration pattern should be reconstructed first. A map showing the hydrostatic hydrocarbon migration pattern can be constructed with the help of a paleodepth contour map of the basement of the sedimentary basin for the appropriate time interval.

For a period of general subsidence and sedimentation, the main characteristics of the burial-induced hydrodynamic condition should be determined. Ideally, an isopach map of the reconstructed original thickness of the sediments accumulated during that selected period is available and will reveal the location of the depocentres i n the basin and will give the sedimentation rates in the different depocentres during that period. From the estimated sedimentation rates in combination with knowledge on the general lithostratigraphic buildup of the basin, it can be deduced whether or not all or part of the burial-induced subsystems of groundwater flow could have existed in the selected period. The presence of deep subsystems of burial-induced flow will probably have been restricted to the depocentres. The general pattern of groundwater flow in identified intermediate subsystems of burial-induced flow can be reconstructed from the isopach map of the sedimentary fill in the same way as outlined in Section 7.1.3.2. During a period when all or part of the basin is subaerial, meteoric water will have infiltrated into the subsurface and initiated the development of a gravity-induced groundwater flow system. The areal extent of such a flow system can be inferred from the areal extent of the continental parts of the basin during a certain period, as can be deduced from paleogeographic information. The depth of penetration of a paleo gravity-induced flow system cannot readily be estimated, because detailed information on paleotopographic relief of the ground surface is generally lacking. If possible, regional recharge and discharge areas should be indicated on paleogeographic maps. The general picture of the evolution of hydrodynamic conditions thus obtained can be verified with the help of published information on indicators of paleogroundwater flow conditions. Such indicators are paleopressures, temperatures and -chemical composition of groundwater. For example, geochemical studies on reservoir diagenesis or genesis of ore deposits, may

226

Chapter 7

provide valuable information for evaluating paleogroundwater flow conditions (e.g. information on paleotemperatures, -pressures and -hydrochemistry as derived from fluid inclusions; and on dating of formation of diagenetic minerals). Information on paleotemperatures may also be derived from biomarkers and vitrinite reflectance data and from fission track analysis (Section 6.3). In addition, physico-chemical characteristics of known hydrocarbon accumulations will reflect former hydrodynamic conditions (e.g. a waterwashed and biodegraded hydrocarbon accumulation presently located in a burial-induced groundwater flow system reflects the former existence of a gravity-induced groundwater flow system). By combining the appropriate maps the main characteristics of secondary migration for separate phase hydrocarbons, hydrocarbons in aqueous solution or in very fine suspension, can be identified for each time period. In addition, the influence of periods of increased tectonic activity should be taken into account in the evaluation, because of the direct and indirect effects of tectonic activity on migration, accumulation and entrapment of hydrocarbons (Sections 4.3.4.3and 5.3.3). For each time period the resulting information on directions and lengths of preferred paths of hydrocarbon migration can be used to evaluate the potential distribution of hydrocarbons that have migrated from known or inferred source rock positions and the possibility of the previously accumulated hydrocarbons being destroyed or remigrated.

227

CHAPTER a QUANTITATIVE ANALYSIS OF SECONDARY HYDROCARBON MIGRATION SYSTEMS

A quantitative analysis of secondary hydrocarbon migration systems should result in figures for the volumes and compositions of hydrocarbons migrating in a sedimentary basin as a function of time and space. Ideally, all aspects of a secondary hydrocarbon migration system in a sedimentary basin at a certain time during the basin’s evolution, should be quantified in the analysis, i.e. the masses and initial composition of hydrocarbons available for secondary migration, the three-dimensional migration pattern, the flux of migrating hydrocarbons and the migration losses. Different procedures are available for the quantitative determination of the masses and initial compositions of hydrocarbons available for secondary hydrocarbon migration (e.g. Duppenbecker et al., 1991;Mackenzie and Quigley, 1988;Tissot and Welte, 1984). The reader is referred to the published literature for a detailed outline of these procedures. This chapter presents approaches for the quantitative determination of the remaining characteristics of present and past secondary hydrocarbon migration systems. Emphasis is placed on the identification and integrated analysis of a wide variety of observable present-day physico-chemical characteristics of fluids and rocks in a sedimentary basin in order to quantify present-day migration characteristics and t o place constraints on the reconstruction of the history of hydrocarbon migration systems in the basin.

A quantitative analysis of present-day secondary hydrocarbon migration for basin evaluation can be restricted to the prospective parts of a sedimentary basin as selected on the basis of the previously described qualitative study (Chapter 7). The quantitative assessment of present-day hydrocarbon migration systems is described separately for hydrostatic and hydrodynamic conditions of the prospective parts of the basin (Sections 8.1,8.2and 8.3). Section 8.4 briefly describes the available approaches for a quantitative analysis of the evolution of secondary hydrocarbon migration systems.

Chapter 8

MaD view

A'

m .-..-. *

; -.*

:

Structurecontours top carrier rock Hydrocarbon expelling source rocks Catchmentarea

Cross section

A'

a

Hydrocarbon expelling source rock Drainagevolume

Figure 8.1 Map view and cross-section of hypothetical prospective area.

Quantitative analysis of secondary hydrocarbon migration systems

229

8.1 Present-day hydrostatic hydrocarbon migration systems In order t o quantify the separate phase hydrocarbon migration under true or assumed hydrostatic conditions, the location of hydrocarbon expelling source rocks, the amount and characteristics of expelled hydrocarbons and the basin's hydrogeological framework should be known. The hydrostatic separate phase hydrocarbon migration starts in the porous and permeable hydrogeological units (i.e. carrier-reservoir rocks) adjacent to the expelling source rocks. As outlined in Chapter 4, secondary hydrocarbon migration in hydrostatic basins is a preponderantly lateral migration through carrier-reservoir rocks. After expulsion from the source rock, hydrocarbons will move updip along the upper boundary of the adjacent carrier-reservoir rocks until traps are encountered. The migrating hydrocarbons will seek the shortest possible migration paths. In order to establish the hydrocarbon migration pattern, the geometry of the upper boundary of these carrier reservoir rocks and the location of permeable and impermeable zones along this boundary should be derived from the known hydrogeological framework of the basin (Section 6.3.4). The hydrocarbon migration pattern from source rock to potential trapping positions can be constructed with the help of a depth-contour map of the geometry of the upper part of the appropriate carrier-reservoir rocks (Section 7.1.2). The geometry of the upper boundary of the carrier-reservoir rock and the location of vertical seals along the migration path determine the possible trapping locations. After having identified the potential hydrocarbon migration paths from expelling source rock to a trapping position, the total volume of hydrocarbons lost WL) along the migration pathways can be estimated from the total volume of rock through which the hydrocarbons migrate (the drainage volume VD) and the mean porosity of that rock (n) (Section 4.3.2, Equation 4.25: V, = n&VD; S, = apparent residual saturation, estimated at 1 - 3%) as proposed by Mackenzie and Quigley (1988; Figure 8.1). Knowing the volume of petroleum expelled from that part of the source rock that provides a drainage area for the trap being evaluated, the total hydrocarbon charge that potentially is available for the trapping location can then be estimated from the difference between the volume of hydrocarbons expelled from the source rock and the volume of hydrocarbons lost during secondary migration (Mackenzie and Quigley, 1988). Repeating this procedure for all potential trapping positions along the identified migration pathways in the studied part of the basin, the potential trapping locations can be ranked according to hydrocarbon charge. The volume of hydrocarbons that actually have reached a certain trapping position in a presently stable hydrostatic basin is, in theory, also influenced by the time required for the hydrocarbons to reach the trap in relation t o the time

Chapter 8

230

elapsed since hydrocarbon expulsion from the source rock started to the present-day. The updip lateral migration of oily hydrocarbons proceeds at a specific discharge rate in the order of millimetres per year (Section 4.1) corresponding to velocities of tens of centimetres per year. Hence, the specific discharge rate will probably not be a limiting factor on the hydrocarbon charge for the generally encountered migration distances of less than 30 km (Section 4.3). Quantitative information on migrating hydrocarbons can be obtained by using the specific discharge equations given in Section 4.1. For example, the specific discharge for hydrostatic updip migration can be calculated from the equations

-2

r

khc = T n S h c

8Z

To solve these equations, requires knowledge on geometrical and hydraulic properties of the rocks (Section 6.3.4), densities of groundwater, and densities and viscosities of hydrocarbons. The density of the groundwater can be estimated from Figure 1.3 for different temperatures and salinities. As a first approach, the densities and viscosities of hydrocarbons estimated from, respectively, Figure 4.7 and Table 4.2 can be used in the calculations. The reader is referred to England et al. (1987) and Mackenzie and Quigley (1988) for more precise techniques to determine the density and viscosity of hydrocarbons for various subsurface conditions and for different compositions of the hydrocarbons. Different modelling approaches simulating hydrostatic hydrocarbon migration have been developed (Lehner et al., 1987; Sylta, 1987, 1991a, 1991b). Figure 8.2 shows the result of an example calculation, as given by Lehner et al. (19871, for hydrostatic separate phase hydrocarbon migration through a hypothetical carrier rock of indefinite thickness at different times during subsidence. It has been assumed in the calculation that the location of hydrocarbon input into the secondary migration system shifts with continued subsidence. Sylta (1991a) incorporated modelling of phase behaviour in the modelling of secondary hydrocarbon migration by using an equation of state of a multicomponent hydrocarbon mixture. Sylta’s modelling approaches also account for migration losses (Sylta, 1987, 1991).

-.

+ r: _, l i i i ! l i i l x

a

a

,+

2

v)

0

w

al

m

m

h

h

h

h

\D 0

-

v

h

a

0

al h

w

.r

M

In

3 .

0

In

3

9


--

I

l

P

h

v

231

Y

sC !i

7

8

2

U

Y ul

Figure 8.2 Calculated hydrocarbon column heights below hypothetical cap rock structure at different times before present (from Lehner et al., 1987. Reprinted by permission of Editions Technip).

232

Chapter 8

8.2 Present-day hydrodynamic conditions

The quantitative analysis of hydrodynamic conditions in a prospective part of a basin includes the identification of the different types of groundwater flow system (gravity-induced, burial-induced, tectonically-induced), the quantitative assessment of the characteristics of the identified groundwater flow systems, the interaction and genetic explanation of the flow systems. For this purpose, a quantitative integrated analysis of direct and indirect indicators of regional groundwater flow is used in combination with groundwater flow modelling techniques. Section 6.3 provides information on the data sources for the principal direct and indirect indicators of flow (groundwater pressure; and temperature, salinity and chemical composition of groundwater). Irrespective of the type of sedimentary basin under consideration, i.e. the type of groundwater flow system under consideration, the current direction and intensity of groundwater flow at each point in a basin under isothermal and isochemical conditions are directly related to the groundwater potential gradient, the density and viscosity of groundwater and the permeability of the subsurface (Chapter 1). On a regional scale, the combination of groundwater pressure, density and viscosity data and permeability data directly indicate e.g. the lateral groundwater flow pattern, areas of concentrated horizontal or vertical flow and the location of geopressured zones. Some characteristics of indirect indicators of flow may be associated with a particular flow condition independent of the type of flow system under consideration (e.g. positive anomalies of groundwater temperatures and salinities at shallow depths may indicate focussed vertical upward flow of groundwater). A correct genetic interpretation of the direct and indirect indicators of groundwater flow and the associated groundwater flow directions and velocities requires knowledge on the type and evolution of the groundwater flow systems involved. The previously described qualitative study is a first approach to differentiate between the different groundwater flow systems (Chapter 7). 8.2.1 Hydrodynamic conditions in stable subaerial regions The hydrodynamic condition in a stable subaerial region is given by the characteristics of the prevailing gravity-induced groundwater flow systems.

The present-day groundwater flow systems in the selected study area may not be in accordance with the relief of the present-day water table, and as a consequence the characteristics of the flow system cannot be reliably inferred solely from water table relief and subsurface permeability distribution. A direct determination of the quantitative characteristics of the present-day gravityinduced groundwater flow system, requires data on pressure, density and viscosity of groundwater and data on the permeability distribution


233

supplemented with data on the relief of the present-day water table, o r eventually on the topographic relief of the ground surface. The interpretation of these data should be integrated with a n evaluation of the different characteristics associated with gravity-induced groundwater flow, as apparent at the ground surface (moisture conditions) or in the subsurface (temperature, salinity, hydrochemical composition, isotopic composition of groundwater) (Chapter 2). T6th (1978) proposed an analysing technique for the quantitative assessment of cross-formational gravity-induced groundwater flow systems and the genetic explanation of the identified systems, based on the integrated interpretation of the following five pressure-related parameters: - potentiometric surface - pressure-depth relation - dynamic pressure increment - hypsographic distribution - water table elevation.

Potentiometric surface The groundwater potentials of a certain hydrogeological unit can be calculated from groundwater pressure and density data (Chapter 1). Groundwater potentials are often expressed in equivalent fresh-water heads. Equipotential lines, or potential contours, connect points of equal groundwater potential for a single hydrogeological unit. The potential contours for a hydrogeological unit add up to a relief map of the potentiometric surface for that unit. A potentiometric surface map of a hydrogeological unit provides a regional picture of the magnitudes and directions of the groundwater potential gradients, which is also of direct importance in analysing hydrocarbon migration systems. For analysing groundwater flow conditions, the direction and spacing of the potential contours of the potentiometric surface of a hydrogeological unit can be used to identify the lateral component of the groundwater flow pattern, the location of lateral barriers to flow andor the location of vertical escape ways from the unit. Pressure -depth relation A regional picture of groundwater pressure-depth relations is indicative of the following groundwater flow characteristics: groundwater flow condition (hydrostatic or hydrodynamic), groundwater flow direction (descending, horizontal or ascending), absence or presence of pressure barriers. Dynamic pressure increment The dynamic pressure increment, Apw, at a certain depth is the difference between the hydrostatic and the hydrodynamic pressure-depth relation at that depth (T6th, 1978). T6th (1978) showed that the dynamic pressure increment is a function of both ground surface elevation and depth of measurement. He proposed the use of a two-dimensional presentation of this relation (the elevation-depth pattern: Apw-z-d)as an analysing tool. The elevation-depth

Chapter 8

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