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Lecture N0te5 1n Earth 5c1ence5 Ed1t0r5: 5.8hattacharj1, 8r00k1yn 6. M. Fr1edman, 8r00k1yn and 7r0y H. ]. Neu9e6auer, 80nn A. 5e11acher,7ue61n9en and Ya1e
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Foreword
The North African Sahara produces hydrocarbons from two major stratigraphic intervals: the Cambro-Ordovician and the Triassic. My geological speciality is the hydrocarbon setting of the Cambro-Ordovician of North America consisting of shallow-marine epeiric carbonates, especially dolostones, producing oil and gas from karst and unconformity traps. Two years ago I worked in the Algerian Sahara in the CambroOrdovician reservoirs of Hassi Messaoud. This hydrocarbon province contains one of the largest reservoirs in the world. It is drastically different from the setting in North America: the reservoirs are composed of sandstones with a porosity in the range of 15 to 2o% producing from non-marine facies, such as fluvial and take deposits and local deltaic settings. Oil production comes for the most part from sandstones and conglomerates of braided stream to alluvial facies. Reservoir quality is related to fractures and diagenetic porosity, especially through the dissolution of feldspar. I wished at the time I would have had this current manuscript of Dr. Monzer Makhous at my disposal. This extensive book presents in great detail every important aspect of the Sahara reservoirs which I studied, including regional geology, mineralogy, petrology, reservoir characteristics, especially the important secondary porosity, petrophysics, carbon isotopes, and migration and accumulation of hydrocarbons, and hydrocarbon generation and preservation. In the Sahara, as elsewhere, stratigraphic and diagenetic reservoirs are becoming the prospects of the future, and this book shows the reader how best to understand these reservoirs and explore for them. This study is most impressive. Dr. Makhows spent 17 years in the Research Center of SONATRACH evaluating zo ooo samples mineralogically and petrologically, and 5 ooo by chemical analysis. The book's ten chapters summarize the most detailed petroleum reservoir analysis I have seen in recent years. For those working in the Sahara this book is a gift. For those working elsewhere it should serve as a case history for their own areas of study. Dr. Gerald M. Friedman
Contents
Introduction
..............................................................................
1
1
Methods of Investigation
1.1 1.2
Analytical Work and Data Treated . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 6 Structure and Extent of Work . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 6
2
Characteristic Features of the Region: Geological Structure, Lithofacies, Paleogeography and Genetic Framework of Non-Structural Traps ........ 9
2.1
Geological and Structural Characteristics of the Saharan Platform and Outline of the Evolution of Its Sedimentary Basins 2.1.1 Sedimentary History and Stratigraphic Sequences . . . . . . . . . . . . . . . . . . . . . . . . Lithofacies and Sedimentary Environments of the Oil- and Gas-Bearing Formations of the Triassic Province 2.2.1 Paleozoic Sediments of the East Saharan Syndinorium (Ghadames and Illizi Basins) 2.2.2 Paleozoic Sediments of the Oued el-Mya Basin . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2.2.3 Sedimentary Genesis of the Triassic Deposits with the Northwestern Part of the Triassic Province as an Example 2.2.4 Mineral Transformations in Sandy Reservoir Rocks Resulting from the Interaction Between Interstitial Waters and Primary Components During Early Diagenis: The Diagenetic Signatures ....... 2.2.5 Evolution of Reservoirs During Progressive Subsidence (Late Diagenesis): Influence of the Sedimentary Inheritance ........... 2.2.6 Triassic Sediments of the Oued el-Mya Basin . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2.2.7 Triassic Sediments of the Ghadames Basin and the Northern Flank of the Illizi Basin . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Geological Parameters for the Formation of Non-Structural Traps and Their Mechanisms . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2.3.1 Non-structural Traps and Methods of Investigation . . . . . . . . . . . . . . . . . . . . . . 2.3.2 Lithostratigraphic Traps . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2.3.3 Lithological Traps . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2.3.4 Stratigraphic Traps . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2.3.5 Traps Related to Volcanic Deposits . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2.3.6 Traps Resulting from Differential Compaction of Sediments ............ 2.3.7 Morphological Traps . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2,3.8 Diagenetic Traps . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Brief Petroleum Geology . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
.............................................................
5
. . . . . . . . . . . . . . . . . . . . . . . . . . .
2.2
. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
. . . . . . . . . .
2.3
2.4
9
16 18 i8 25 29
37 39 42 44 46 47 49 51 54 58 58 59 61 62
X
Contents
3
Main Factors of Reservoir Compaction
3.1 3.2
Evolution of Density, Porosity and Permeability with Depth 63 Relationship Between Structure, Gravitational Compaction and Pressure Solution of Granular Reservoir Rocks 65 3.2.1 Gravitational Compaction and Pressure Solution . . . . . . . . . . . . . . . . . . . . . . . . 65 3.2.2 Structural Arrangement of Sandstones . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 74 Reservoir Compaction by Silicification from Other Sources . . . . . . . . . . . . . . . . . . . . . 75 3.3.1 from Underground Waters . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 77 3.3.2 Silicification Through Transformation of Feldspars, Detrital Illite and Other Micaceous Components . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 78 3.3-3 Silicification by Illitization of Smectite . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 82 Main Controls of the Compaction of Reservoir Rocks . . . . . . . . . . . . . . . . . . . . . . . . . . . . 85 3.4a Thickness of Sandy and Silty Reservoir Rocks . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 85 3.4.2 Argillaceous Diagenesis and Related Abnormal Formation Pressure ... 87 3.4.3 Early Development of Overgrowth Rims . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 89 3.4.4 Presence of Mesozoic Evaporites 89 3.4.5 Temperature and Pressure 90 3.4.6 Authigenic Transformation of Argillaceous Cement . . . . . . . . . . . . . . . . . . . . . . 90 3.4.7 Early Invasion of the Reservoirs by Hydrocarbons . . . . . . . . . . . . . . . . . . . . . . . . 96 Silica Solid Phase Transformation: A New Concept for Sandstone Diagenesis as Revealed in North African Basins . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 97 3.5.1 Samples . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 98 3.5.2 Methods . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 100 3.5.3 Results 101 3.5.4 Discussion . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 116 3.5.5 Solid Phase Process . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 122 Conclusions . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 129
...........................................
63
. . . . . . . . . . . . . . . . . . . .
. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
3.3
S i l i c i f i c a t i o n
3.4
. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
34
. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
3.6
4
Reservoir Decompaction and Formation of Accumulation Capacity (in Secondary Porosity) of Reservoir Rocks . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
Decompaction Due to Solution of Binding Compounds of Sandstones (Carbonates, Silicates, Aluminosilicates) 4.1.1 Relation in Time and Space (as a Function of Temperature) of Organic and Mineral Diagenesis After Formation of Reservoir Properties .... 4.1.2 Thermodynamic and Stoichiometric Regime of Formation of Carbonic and Organic Acids and Their Role in the Establishment of Secondary Porosity in Reservoir Rocks 4.1.3 Model of Differential Dissolution and Redistribution of Carbonate Cement with Compaction/Decompaction of Reservoirs in Space as Based on Carbon Isotope Data 4.2 Decompaction-Compaction by Intergranular Pressure Solution (of Quartz Grains) and Removal of SiO~ by Alkaline Solutions . . . . . . . . . . . . . . . 4.2.1 Pressure Solution and Quartz Cement 4.2.2 Factors Controlling Pressure Solution 4.2.3 Silica Budget 4.2.4 Mechanisms of Silica Transport 4.2.5 The Role of Pressure Solution in the Evolution of Porosity
131
4.1
. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
. . . . . . . . . . . .
131 131
144
149 159 161 162 166 170 173
XI
Contents
4.z.6 Conclusion . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Generalized Models of the Transformation of Oil-Bearing and Reservoir Formations . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 4.3.1 Model of Diagenesis in Space and Time . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 4.3.2 Generalized Model for the Transformation of Reservoir Rocks, Mass Transfer and the Formation of Reservoir Properties .............
178
5
General Geochemical Features of Generation, Migration and Accumulation of Hydrocarbons . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
183
5.1
Geochemical Characterization of Potential Source Rocks, Hydrocarbons and Burial Histories . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 5.1.1 Silurian Graptolitic Source Clays . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 5A.a Devonian Source Shales . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Generation and Directions of Migration . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 5.a.1 Generation in the Silurian Source Rocks . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 5.2.2 Generation in the Devonian Source Rocks . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Geochemistry of the Triassic Province . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 5.3.1 Source Rocks in the East of the Province (Ghadames and Illizi Basins) . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 5.3.2 Source Rocks in the North of the Province (Oued el-Mya and Triassic Basins) . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 5-3.3 Characterization of Petroleum in the Eastern Area of the Province ... 5.3-4 Petroleum Types and Their Variations in the North of the Province., 5-3-5 Conditions for Hydrocarbon Generation in the North of the Province .. 5.3.6 Conditions for Hydrocarbon Generation in the Eastern Area of the Paleozoic Province . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 5.3.7 Petroleum to Source Rock Correlations . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
193 197
Burial History and Kinetic Modeling for Hydrocarbon Generation
.....
207
The Models . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 6.1.1 The Program . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 6.1.2 Burial and Thermal History Modeling . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 6.1.3 Modeling Maturation History and Hydrocarbon Generation .......... 6.1.4 Additional Features of Our Burial and Thermal Modeling ............. 6.1.5 Conclusions . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 6.2 Applying the Model to Saharan Basins . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 6.2.1 Geological Framework . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 6 . 2 . 2 0 u e d el-Mya Basin . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 6.2.3 Ghadames and Illizi Basins . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 6.2.4 Southern and Western Basins . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 6.3 Summary and Conclusions . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
207 207 209 22t 224 230 230 235 240 244 251 252
4.3
5.2
5.3
6 6.1
174 175 175
183 t83 184 185 185 186 186 186 188 190 191 192
7
Degree of Preservation of Hydrocarbon Accumulation as Indicated by Carbon Isotope Analysis
7,1 7.z
Methods Employed . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 255 General Data on Carbon Isotope Composition of Sedimentary Rocks (Carbonates) and Organic Matter from Northeastern Algeria ................. 256
. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
255
xII
7-3
7.4
Contents 7.2.1 Carbon Isotope Composition of Carbonate Rocks . . . . . . . . . . . . . . . . . . . . . . . 7.2.2 Carbon Isotope Composition of Organic Matter . . . . . . . . . . . . . . . . . . . . . . . . . The M e c h a n i s m of Stable Carbon Isotope Fractionation (~2C vs. t3C) and Regularities in Their Distribution in Jurassic and Cretaceous Deposits of Northeastern Algeria . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 7.3.1 D e c o m p o s i t i o n and Oxidation of Organic Matter . . . . . . . . . . . . . . . . . . . . . . . 7.3.2 Methanogenic F e r m e n t a t i o n . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 7.3.3 Sulfate Reduction by Bacterial Activity . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Conclusions . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
256 262
263 263 266 266 267
8
Reconstruction of Temperatures from Organic and Mineral Diagenetic Criteria
8.1
8.3
Reconstruction of Temperatures from Degree of Structural Ordering in Mixed-Layer Minerals . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Crystallographic Features of Clay Minerals as T h e r m a l Indicators in Petroleum Geology . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Summary ................................................................................
271 277
9
General Conclusions . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
279
. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
8.2
10 Analytical Methods and Equipment . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Io.1 Inorganic Constituents . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . lo.1.1 X-ray Diffraction . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . lO.1.2 Scanning Electron Microscope (SEM) . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . lO.1.3 Microanalysis . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . lo.1.4 Major Element Analysis . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . lo.1.5 Trace Element Analysis . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . lo.L6 Infrared S p e c t r o m e t r y . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . lo.1.7 Physical Measurements lo.2 Organic G e o c h e m i s t r y Experiments . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . lO.2.1 Extractions and Separations 10.2.2 Vitrinite Reflectance Measurements . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . lO.2.3 Pyrolysis lo.2.4 Liquid C h r o m a t o g r a p h y . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . to.2.5 Gas C h r o m a t o g r a p h y lO.2.6 Gas Chromatography-Mass Spectrometry (GC-MS) . . . . . . . . . . . . . . . . . . . . 10.2.7 Organic Carbon Isotopic Analysis . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 10,2,8 Mineral (Carbonate) Carbon Isotopic Analysis . . . . . . . . . . . . . . . . . . . . . . . . . . lo.2.9 C a t h o d o l u m i n e s c e n c e . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
269 269
283 283 283 284 284 284 285 285 286 287 287 288 288 288 289 289 289 290 291
References .............................................................................
293
Supplementary References
303
. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
Index .....................................................................................
309
Introduction
The Sahara contains two main hydrocarbon-bearing complexes (Fig. o.1), the older of which encompasses reservoirs in the Cambro-Ordovician and the younger one those in the Triassic. The hydrocarbon-bearing rocks have undergone secondary transformations under defined conditions to such an extent that their initial characteristics have completely disappeared. These transformations are characterized by two essential factors: the first one is the vertical inhomogeneity in the transformation of the rocks resulting from the vertical distribution of energy and matter. This is expressed by the succession of zones of compaction and decompaction which may be observed. The second factor tends to complicate the trends of the variation with depth. It is related to the particular history of development which may be subdivided into two parts. The distinction between these two stages, one during the Paleozoic and one in the Mesozoic, is highlighted by the intense Hercynian tectonic activity and orogeny followed by pronounced uplift and erosion. This has greatly lessened the heating of the rocks during the"thermal pause" in their geological history. The study presented here has made possible the elaboration and investigation of
the followingfundamental problems: t. reconstruction of the diagenetic history, unveiling of the mechanisms transforming the reservoirs which represents the key question in prospecting for and exploitation of hydrocarbon deposits; 2. outlining the new processes of decompaction of rocks at great depths by demonstrating the relative roles of the different processes at different depths. This entails the particularly important process of dissolution and transport of a tremendous amount of carbonate and silica cement under the influence of carbonic acid and in particular of organic acids like bifunctional carboxylic acids and phenols; 3. outlining the process of large-scale transport of silica under the specific conditions of the alkaline environment prevailing at great depths; 4- outlining non-structural traps formed by different mechanisms. The main aim was the characterization of the zones of decompaction and the outline of the mechanisms leading to the properties of the reservoirs. For this purpose it was necessary: - to investigate in detail the composition of the rocks with the aid of modern analytical methods and to treat the data obtained from a considerable number of samples taken from different basins in order to be representative of the different geological situations; - to design a number of models reflecting the characteristic diagenetic processes of the geological development of the basins which would allow us to establish
Introduction
z
general features of generation and migration of the hydrocarbons as well as the formation of the respective reservoirs. One of the most important rules in the formation of hydrocarbon deposits has been demonstrated, based on the fact that the primary migration of hydrocarbons follows closely on the formation of the secondary porosity, as during maturation of the organic matter the main phase of formation of the hydrocarbons takes place after the culmination of the decarboxylation which generates active acid solutions controlling the formation of the secondary porosity. This shows that the generation of hydrocar-
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Introduction
3
bons and the formation of the secondary (main) storage capacity of the reservoirs overlap in time and space. Because of this the direct links between the source of the hydrocarbons and the reservoirs in time and space favour the accumulation of hydrocarbons in reservoirs with secondary porosity. In this context the results of the present research are of great theoretical value for the study of the diagenetic transformations and of their influence on the formation of the trapping properties of the reservoirs. This leads to the establishment of new criteria for prospecting for hydrocarbon deposits, such as: 1. redistribution within the reservoirs of carbonate cement resulting from dissolution by acid solutions in diagenetically more mature reservoirs and its redeposition in overlying more mature zones. This leads to an increase in the porosity in the zones of dissolution and to a blocking of the reservoirs in the zones of precipitation; 2. correlation of the carbon isotopic composition (6'3C) of the five oil and bitumen fractions with the aim of establishing the sources of hydrocarbon formation in order to delineate which reservoirs under these conditions in the many possible sources would be the most promising targets of exploration; 3. evolution of the secondary migration of hydrocarbons by studying the mineral transformations along their routes of migration on the surface and at depth and in particular the carbon isotope composition in the carbonates and the organic matter. This approach will enable us to evaluate the degree of conservation of the hydrocarbon accumulations which underwent total or partial destruction in absence of a real cover, in particular under the conditions prevailing in certain tectonic complexes to the north of the Saharan Platform.
Practical Aspects of this Research 1. The maximum depth at which the secondary porosity resulting from mesokatagenetic decarboxylation may be preserved is more than two times larger than that at which the primary porosity may be preserved in sandstones of high mineralogical stability and more than three times larger than that in sandstones of low mineralogical stability. 2. We have tried to delineate the main zones of formation of non-structural traps which are promising, as undrilled promising anticlinal structures become rare. 3. The correlation of the carbon isotope composition (61~C) of the oil and bitumen fractions of all eventual source rock "conditions", in view of the large variety of the latter, allows us to point out the sources of the hydrocarbons which supplied the one or other reservoir and thereby to better direct the prospecting activities. 4- The chemical and kinematic simulation of the hydrocarbon-bearing basins enables us to reconstruct their burial history, the changes in the thermal regime of the sedimentary series as well as formation and migration of hydrocarbons in each oilbearing horizon. The innovation and perfecting aspect of our simulation in comparison with the other available models lies in the application of alternative methods for calculating the tectonic subsidence of the basins, the integration of the maturation of hydrocarbons as a function of burial as well as an earlier estimate of the overall potential of
4
Introduction
source rocks. For this purpose the algorithm for selecting the kinetic parameters of the reactions of hydrocarbon formation takes into account the history of maturation of the source rocks in time and space including the geological stage of maturation as well as the stage of pyrolysis of the source rock sample in the Rock-Eval system. If we omit from the calculation the geological stage, the energy spectrum of activation for the reactions is shifted abruptly- to higher values, leading to an underestimate of the hydrocarbon potential by one order of magnitude or more throughout the geological history.
Chapter 1
Methods of Investigation
The main concepts of this investigation are based on theoretical and experimental research which the author has carried out during his 17 years at the Centre of Research and Development of Messrs. SONATRACH in the fields of exploration and exploitation of hydrocarbon deposits. Particular attention was paid to the detailed study of the mineralogical and organic composition of hydrocarbon-bearing rocks, as well as of the hydrocarbons themselves and of the formation waters. For this purpose modern analytical techniques had to be employed: scanning and transmission electron microscopy (SEM and respectively TEM), SEM-image analyses, XRD, ICP, XRT, WDS and EDS microanalysis, IRS, mass spectrometry, chromatography, pyrolysis, etc. The petrographic characteristics of the reservoirs in rocks were determined in addition to the more classical methods by image analysis, permitting a characterization of geometry and structure of the pore space. The statistical treatment and graphic presentation of the enormous amount of the analytical data required a great amount of computer work. Virtually all geological problems studied here were dealt with on the basis of measured and calculated physico-chemical parameters: 1. Paleographic reconstructions were done by using, in addition to widely known geological criteria, parameters like the adsorbed complex of the clay minerals, their ion exchange capacity, their coefficient of alkalinity and the type of iron minerals present. 2. The diagenetic processes were delineated by using the structural and crystallochemical peculiarities of the clay minerals, including a calculation of the parameters of the crystal structure and the determination of the structural modifications of the various minerals. A detailed study of the mineralogy and crystallography of quartz, silica, and carbonates was carried out together with an analysis of the isotopic composition of the carbon. 3. In order to study the process of generation and migration of the hydrocarbons, one needs data on their phase composition, the mineral transformations associated with the transformation of the organic matter and the carbon isotope composition of the carbonates, bitumens and oils. 4. When carrying out the chemical and kinetic simulation of generation and migration of hydrocarbons, we have conducted experimentally an open pyrolysis at different heating regimes, taking into account the maturation of the organic matter with depth of burial of the source rock and applying other independent geothermal parameters.
6
Chapter 1 • Methods o f Investigation
1.1
Analytical Work and Data Treated The thesis is based on the results obtained by the author during his field and laboratory work while working on ten large hydrocarbon-bearing basins of the Saharan Platform. It deals with and evaluates an enormous amount oftithotogicai and geochemical data accumulated in the research centre of Messrs. SONATRACH between 1975 and 1992: the majority of this work was carried out by the author himself or under his guidance. Over lO ooo samples of sandstones and shales were subjected to integrated mineralogical analysis and analysed by XRF and SEM. Chemical analysis was carried out on more than 5 ooo samples. The petrophysical data (porosity, permeability, water saturation and density) were determined for sand- and siltstones at intervals of z5-5o m and the overall length of drii1 core studied was at least 3o ooo m. Special investigations of structure and geometry of the pore space with the aid of cathodoluminescence and SEM-image analysis were carried out on about 7oo sandstone samples. All petrographic, mineralogical and petrophysical data were compiled for each borehole in about 2o0 laboratory logs. Particular investigations like determination of the adsorbed complex, of ion exchange capacity, of coefficient of alkalinity and of type of iron-bearing minerals in combination with the mineralogy of the respective argillaceous horizons were carried out for more than 15o sections or respectively bore holes. The carbon isotope composition (613C) was performed on the carbonate cement of 18o sandstones, on the five oil and bitumen fractions in 21 source rocks and on some 50 oils. Additionally, when studying the secondary migration of hydrocarbons, the carbon isotope composition was analyzed on 15o carbonate samples and 4o bitumen samples taken on the surface. The author also made use of the results of pyrolysis of more than 2 ooo source rock samples and of the chemical and bituminological analyses of the organic matter in more than 7 ooo shales. He had at his disposal also microspectrophotometric determinations of the catagenic level of the organic matter in about 2 400 units, about 1700 chromatograms of the gaseous and liquid phases and complete chemical analyses of formation waters from about 1zoo horizons, including their mineral and organic compounds. Paleogeographic reconstructions of the sedimentary conditions were produced for many of the stratigraphic complexes of the main basins, widely using mineralogical and chemical data and in particular those on argillaceous rocks. Cathodoluminescence was greatly used for studying quartz diagenesis and for estimating the silica budget. It covered some 400 sandstone samples. 1.2
Structure and Extent of Work The first part of this book (Chapter z) deals with structures and geological history, paleogeography, the general features of development of the sedimentary basins and with the mechanisms forming non-structural traps. The second part (Chapters 3 and 4) describes the main factors of compaction-decompaction of reservoirs in the context of the diagenetic transformations and their characteristics in relation to their properties of accumulation and infiltration. The third part (Chapters 5 and 6) evaluates the geochemical peculiarities of generation, migration and accumulation of hydrocarbons
1.2. Structure and Extent of Work
7
on the basis of classical geochemical data and those resulting from chemical and kinetic modeling. The fourth part deals with specific applied aspects of isotope geochemistry for evaluation of the degree of preservation of hydrocarbon accumulation (Chapter 7) and temperature reconstruction from clay mineral and organic diagenetic criteria (Chapter 8).
Chapter 2
Characteristic Features of the Region: Geological Structure, Lithofacies, Paleogeography and Genetic Framework of Non-Structural Traps
2.1
Geological and Structural Characteristics of the Saharan a n d O u t l i n e o f t h e E v o l u t i o n o f Its S e d i m e n t a r y Basins
Platform
On the Saharan Platform, starting from the Hoggar Massif in the south, we observe a number of north-south running uplifts: (Amguid el-Biod-Hassi Messaoud, Tikhemboka-Zarza~'tine-Atrar, tdjeran-M'Zab, etc.) which are separated from each other by wide synclinoria, with the same trend underlying a number of depressions. Thus the East Saharan synclinorium may be subdivided into two depressions, the Illizi to the south and the Great Eastern Erg (or Ghadames) in the north. The Central Saharan synclinorium is also underlying two depressions, the Mouydir to the south and the Oued el-Mya to the north. On the West Saharan synclinorium we find the depressions of Ahnet, Reggane, Tindouf, Timimoune de Bechar, etc. (Figs. ~.~ and 2.2). These basins were subdivided by Guiraud et al. 0987) into North and South Saharan and a de-
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Chapter 2 • Characteristic Features of the Region I
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Fig. 2.2. Structural map of Triassic Province. I Allal High, 2 Tilrhemt-Beressof High, 30ued el-MyaMouydir Depression, 4 Hassi Messaoud-Agreb High, 5 Dahar-el Biod High, 6 Ghadames Depression, 7 Tihembouka High tailed description of their stratigraphy may be found in Legrand (1985) and Whiteman (1972). The anticlinal domes and the uplifts complicating them are generally associated with horsts of the Precambrian basement and contain Paleozoic sediments of reduced thickness with many stratigraphic gaps and unconformities covered, in the northern parts, by Mesozoic and Cenozoic deposits. The synclinoria are filled by Paleozoic and Mesozoic sediments in which the successions are thicker and more complete than on the uplifts, in particular on the western parts of the Saharan Platform (Figs. 2.3 and z.4). The Central and South Saharan basins crop out along the Ougarta belt and the
2.1 • Geological and Structural Characteristics o f the Saharan Platform
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12
Chapter 2 - Characteristic Features of the Region
Hoggar and Reguibat Massifs. Such outcrops are also observed in the monoclinal structures of South-Illizi and Tassili-Najjer in between the thickest development of Paleozoic rocks. In the formation and development of the anticlinoria and synclinoria as well as of the second- and third-order structures complicating them, large fracture zones have played an important role, amongst which some are of Precambrian age and have been reactivated during later geological phases. These fractures have led to a system of large blocks which may be observed best in the Precambrian rocks cropping out on the slopes of the Hoggar Massif and which have led in the study area to a series of buried horsts and grabens (Furon 1963). The available literature does not help in unraveling the mechanisms controling the deepening of the basins of the Saharan Platform. The basins of Algeria, Tunisia and Libya, as described above, exhibit large variations in the rate of subsidence during the various periods between the Paleozoic and the present. It is obvious that the renewed collision of the continental plates during the Phanerozoic could have created zones of reactivation and constriction (or local rifts) similar to those of northwestern Europe formed over the folded Hercynian basement and detached behind the Alpine orogenic belt (Burotlet 1967a). The sedimentary history of this region of northern Africa may be considered as having been controled by the relative movements of these blocks along fractures in the basement which have been constantly reactivated (Furon 1963). The substrate onto which the thick Phanerozoic sequences have been deposited is part of the continental platform of the western part of the African Shield. The northern boundary is formed by the fracture system limiting the southern flank of the Saharan Atlas (Burollet 1967a). This platform resulted from the collision of the East and West Saharan plates during the Pan-Africa orogeny around 55o Ma. It formed the southern coastline of the Thethys which covered a wide area of this platform during the Paleozoic and smaller areas during the Mesozoic (Dercourt et aL 199o, 1993), the only exception being the great transgression of the Upper Cretaceous (Furon 1963). The West Saharan synclinorium subsided strongly during the Paleozoic, and over it Paleozoic sedimentation took place in a vast basin with a thickness of 7-8 kin, whereas the Central and the East Saharan Synclinoria during the same time were less subsiding areas with marine deposits only 2-4 km thick (Fig. 2.5). In contrast to this, with the start of the Mesozoic, the region of the East Saharan synclinorium suffered intense burial under the accumulation of 4 km of marine and locally continental Triassic, Jurassic and Cretaceous sediments. Thus a vast Mesozoic sedimentary depositor (Triassic Basin, Oued el-Mya Basin, etc.) was formed which encompassed the present Ghadames Depression or the so-called Grand Erg Basin (Figs. 2.4 and 2.6). At the end of the Cretaceous the Eastern Erg was uplifted at the onset of the Alpine orogeny, whereafter it became covered only by thin sandy and calcareous deposits of the Mio-Pliocene. Tertiary sedimentation continued in Tunisia and further away on a large scale and with a thickness of up to 7 o00 m from the Paleocene to Pliocene to the east in the littoral zones of the Gulf of Gabes. In this area the Eocene nummulite facies is hydrocarbon-bearing. Figure 2.7 presents lithostratigraphic type sections from the Paleozoic and from the Triassic of the Triassic Province. The Caledonian and Hercynian orogenies have contributed to the main structural complexity of the Saharan Platform. Because of the Hercynian uplift and the subse-
2.1 . Geological and Structural Characteristics of the 5aharan Platform
Fig. 2.5.
13
Schematicstructure of the SaharanPlatformon the surfaceof the basement(scale 1 : 770 ooo)
quent peneplanation, this region became covered with thick layers of Meso- to Cenozoic sediments (Furon 1963). The Austrian phase of compression in the unstable zone north of the Arias Province led to the first folding, to be followed by some erosion over the eastern part of the Saharan Platform. Eventually during the Alpine orogeny of the Tertiary the unstable portions of the platform were uplifted and widely folded, leading to complex structures and thrusts. The northern province is referred to as the "Atlas Fold Zone" which resulted in oscillations and the formation of domes on the Saharan Plateau (Heybrock 1963; Burollet 1967a). Deep faults, including those of the Precambrian, and in particular those orientated north-south and east-west, as well as fold-like dislocations are widespread on the Saharan Platform. These faults have influenced the formation of structural features like synclines, grabens, etc. Most of the hydrocarbon deposits discovered so far are associated with these types of structure. The reaction of basement faults along the northern edge of the Hoggar Massif led to horsts and grabens in the Paleozoic strata. In Tunisia and Libya a series of faults created tilted blocks and grabens along the southern margin of the Thethys proto-ocean (Klitzsch 1971) as indicated by the deposition of coraliferous facies in the Carboniferous and Permian. Although later neo-tectonic movements have greatly influenced the structural features of the various basins, their main structural elements are inherited in parts from
14
Chapter 2 - Characteristic Features of the Region
STRATIGRAPHICUNITS TERT.
LITHOLOGY
MIOPLIOEOCENE
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2.1 - Geological and Structural Characteristics of the Saharan Platform CENTRALPAler
15
EASTAERNPART
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Fig. 2.7.
o~ o"
Correlationof typical Paleozoicand Triassic cross sections for the TriassicProvince
tectonic trends in the Precambrian basement (Burollet 1967b). The post-Panafrican picture of the Saharan Platform and the structural impact of the Caledonian and Hercynian orogenies may be deduced from the Phanerozoic structural evolution of northern Africa as presented by Klitzsch (1971, 1981,1986), Klitzsch and Wycisk (1987) and Schandelmeier et al. (1987).
16
Chapter 2, Characteristic Features of the Region
2.1.1 Sedimentary History and Stratigraphic Sequences Sedimentary history of the Saharan Platform started in the Cambrian and has continued until the present. It is generally held that the platform existed as a cratonic plateau throughout this period (Burollet 1967a). Epirogenic folding and faults have controled sedimentation in the various basins. The oldest rocks in the area are the folded Precambrian igneous and metamorphic rocks found in the south and in the Hoggar which constitute the basement to the sediments deposited from the Paleozoic up to the present. The Cambrian rocks cover a nearly peneplained basement and include thick cross-bedded non-fossiliferous fluvial sandstones. At the start of the Paleozoic the South Pole was located just north of Africa in the Japetus Ocean (Petters 1991). Quartzitic sandstones were deposited during the Cambrian throughout northern Africa in a progressing marine transgression lasting until the Ordovician (Buroltet 1989; Klitzsch 199o). The Hoggar and Reguibat massifs were not yet developed during this period, having appeared only after the Ordovician when the intercratonic basins of Tindouf, Reggane, Ahnet, Mouydir, Ghadames, Illizi, Murzuk and Kufra subsided (Petters 1991 ). At the onset of the Ordovician the Caledonian movements caused marine transgressions leading to alternations of shallow-marine sandstones and mudstones. During the Ordovician a far-reaching marine transgression was underway, depositing thick marine sandstones. Melting of the glacial cover of the Sahara which took place at this time rather than a more direct tectonism could have caused this large-scale transgression depositing trilobite- and graptolite-bearing black shales (Beuf et al. 1966;Rognon 1971). The Cambro-Ordovician sandstones represent important reservoirs in the oil fields of Hassi Messaoud and Rhourde el-Baguel. There is no major sedimentary hiatus between the Ordovician and the Silurian. A renewed transgression took place at the beginning of the Silurian with the deposition of grey and black graptolitic mudstones, siltstones and fine-grained sandstones in the centre of the depository which coincided with Eastern Erg Basin. This interval contains source rocks rich in organic matter. Rocks dating from Early Devonian are made up of sandstones with plant remains resting discordantly on the Silurian deposits. Throughout the Devonian, sedimentation is characterized by a repetition of transgressive/regressive cycles with an alternation of mudstones and sandstones. The reservoirs in Devonian sandstones represent important productive levels in the Illizi Basin. Uplift occurred during the upper Devonian and even part of the upper Silurian in the southern part of the Eastern Erg Basin. The first signs of the Hercynian orogeny at the onset of the Carboniferous put an end to the stable conditions on the platform (Bishop 1975).All the marine transgressions of the Lower Paleozoic had come from the west. The Carboniferous transgressions, however, now come from the Thethys to the north (Hecht et al. 1964; Hoffmann-Roth 1966 ). The Carboniferous argillaceous schists resulting from a major eustatic transgression are also considered to be potential source rocks. Shallow-marine sediments, i.e. evaporites, dolomitic marls and oolitic limestones, were deposited throughout the Illizi and Ghadames Basins. The rocks of the uppermost Carboniferous were deposited in a regressive shallow-marine environment. The Hercynian uplift took place along pre-existing basement axes such as the Anguid-E1-Biod axis and its extension the E1-Agreb-Hassi Messaoud axis (Fig. 2.z). Deposits from the Precambrian and the Carboniferous were eroded
2.1 . Geological and Structural Characteristics of the Saharan Platform
17
in those areas which had been subjected to this uplift (Burollet 1989). This erosion improved the reservoir characteristics of the Cambrian sandstones of Hassi Messaoud. The main feature of the post-Hercynian erosion on the Saharan Platform is that of a "T'-shaped anticlinorium extending from Algeria to Tunisia. To the east, a continuation of the Hercynian uplift at Nefusa in Libya extends towards the east to join up with this antidinorium. The absence of Permian deposits in Algeria suggests that this region remained a continental high during this period. The marine transgressions of this period were terminated by the thick Permian deposits of Tunisia and Libya. The Permian argillites form a cover for the Silurian reservoirs referred to as "Acacus" in the Libyan oil fields. The restriction of the West Thethys basin and the post-Hercynian subsidence along the edge of the African continent led to a new sedimentary cycle comprising thick Triassic and Liassic evaporite sequences. During the Triassic the waning Hercynian tectonic cycle permitted a renewed subsidence of the Saharan shield, marking the start of a second important sedimentary cycle. The Triassic sediments are developed over the northeastern portion of the Saharan Platform, i.e. the basins of Oued el-Mya and Ghadames, the Triassic Basin and on the northern flank of the Illizi Basin. This area of Triassic rocks is referred to as the "Triassic Province". In areas of the Hercynian discordance trenches were filled by sandstones of fluvial and shallow-marine origin. These sandstones contain one of the most important reservoir rocks of the Triassic region. Andesitic and basaltic lava flows are present in certain zones of the Triassic, forming the cover over part of the Cambrian oil field of Hassi Messaoud and over the Triassic oilfields in the Ouargla area. This volcanic episode suggests a thinning of the continental crust and the underlying thermal anomaly responsible for the Permo-Carboniferous uplift. The basal Triassic sandstones are overlain by calcite- and dolomite-rich mudstones of the so-called Carbonate, the "Lower Argillaceous" and the "upper Argillaceous". The latter result from marine to hypersaline environments and are themselves covered by the "Salif~re", a sequence of mudstones and anhydrites grading into halite. These evaporites belong to the most important cover rocks of the hydrocarbon deposits of the Triassic Province. They are not developed in the Illizi Basin because of a facies change towards clastic rocks derived from the neighbouring Hoggar Massif. The post-Hercynian subsidence was much less pronounced in the Illizi Basin than in those farther north. With the Devonian reservoirs at rather shallow depths (8oo-1 ooo m) in the Iltizi, the lurassic deposits follow- concordantly on those of the Triassic (Busson 1967). The Liassic is made up of evaporites, the overlying Dogger becoming increasingly argillaceous upwards, and the Malta comprises essentially mudstones. The Cretaceous sediments consist of evaporites, limestones, dolomites and thin sandstone beds. The calcareous littoral facies of the Apt and Alb are oil-bearing in Tunisia and sandstones of the same age represent the main aquifer of the northern Sahara. The subsequent Tertiary sedimentation in Tunisia led to great thicknesses of up to 7 ooo m. These sediments represent the succession from the Paleocene to Pliocene in the east and in the littoral parts of the Gulf of Gabes in Tunisia. The Alpine orogeny of the Tertiary raised the unstable parts of the platform and resulted in a multitude of faults and other complex structures. At the end of the Cretaceous the northern basins were uplifted simultaneously at the onset of the Alpine orogeny, to become covered by thin deposits of sands and limestones dating from the Mio-Pliocene.
18
Chapter 2 - Characteristic Features of the Region
2.2 Lithofacies and Sedimentary Environments of the Oil- and Gas-Bearing Formations of the Triassic Province 2.2.1 Paleozoic Sedimentsof the EastSaharan Synclinorium(Ghadamesand Illizi Basins) The study of the complex rocks in the Illizi and Ghadames Basins permitted a detailed reconstruction of the paleogeographic conditions of the sedimentation during the Paleozoic. The paleogeographic reconstructions and the respective conclusions drawn are based on: • • • •
the study ofthe adsorbed complex and the alkalinity coefficient in the argillaceous rocks; the lithological descriptions; the study of the main features of the clay minerals byXRD and etectron microscopy; analysis of the modifications of the iron-bearing minerals.
These investigations led to the identification of a number of series and at the same time facilitated an evaluation of the physico-geographic conditions under which the respective rocks were formed.
2.2.1.1 Cambrian In the study region the paleogeographic history may be subdivided into two stages, viz. the Lower Cambrian characterized by predominantly continental environments and the Middle Cambrian with transitional and marine sedimentation. The Lower Cambrian is represented by alluvial and proluvial deposits, conglomerates, pebble beds and coarse angular and poorly sorted sandstones. The transitional environments are located in the central part of the Illizi Basin (IRLq), the higher-lying parts of the Tiguentourine - La Recul6e area (Tg-38) and in the southern part of the Ghadames Depression (SOHq and SEDq). These deposits are composed of coarse sandstones and microconglomerates with intercalations of argiltites and tigillites as well as oolites and chamosites. The depositional area was probably situated at the foot of a delta slope. The features of the adsorbed complex of the Cambrian argillites were intensely transformed by diagenetic reactions, making them unsuitable for any interpretations. The complex composition of the clays which consist of kaolinite, illite, chlorite and mixedlayer minerals, together with the absence of organic matter leads to the conclusion that the areas of denudation must have been of a local nature. At the same time, we may reconstruct a temperate to humid climate interrupted at times by periods of relative aridity. This explains the lack of laterites on the eruptive and metamorphic rocks.
2.2.1.2 Ordovicion After a phase of uplift affecting the entire region and accompanied by intense erosion lasting throughout the entire Cambrian we observe a burial of the East Saharan
2.2 • Lithofacies and Sedimentary Environments of the Oil- and Gas-Bearing Formations
19
synclinorium and a marine transgression over a large part of the region. The deposits of Unit III in the southern border areas of the Ghadames Depression and the Illizi Basin are represented by: 1. coastal argillaceous sandstones deposited in freshwater (alkalinity coefficient
In contrast to what we have noted in the Ordovician, the influences of deep structural features on the differentiation of the various basins and on the sedimentation are less clearly expressed here. Some of these, the high lands of Oued Namous, the trenches of the Ougarta chains, the Gourara ridge and the Azzel Matti Depression, are still there during the Silurian, but at the same time tectonic movements make their appearance over wide areas of the shields. Their preferred directions at present, at least in parts, are inexplicable (Legrand 1981). The most pronounced marine facies, argillaceous or calcareous, is of Lower Silurian age. This is evidence of the vast marine transgression following the end of the glaciation of the Upper Ordovician. Higher up, these marine influences become progressively less pronounced. The marine facies was preserved probably" only around the edge of the Ghadames Depression and its southern border where the influence of the glaciers had been considerably smaller. The Silurian is characterized by a widening of the basin resulting from a warmer climate and the melting of the glaciers. Over a large part of the study region its sediments are made up of argillaceous-sandy material deposited in marine environments
20
Chapter 2 • Characteristic Features of the Region
(coastal, infralittoral, sublittoral). The Silurian pelites of the Sahara serve as indicators of an extensive eustatic transgression. The alkalinity coefficient is
Devonian After the maximum of the Silurian transgression a partial regression took place which resulted from the orogenic processes of the Ardennian phase, leading to uplift in parts of the Illizi Basin and the Ghadames Depression. These uplifts resulted in an intensification of the erosion in these regions.All this led at the outset of the Lower Devonian to the deposition of thick sandstone sequences of Siegenian age when the region was subjected to a new transgression. The respective rocks rest with an unconformity on pronounced erosive surfaces of Silurian rocks. At the same time, to the south and southeast of the Itlizi Basin, coastal sandstones were deposited which contain pyrite, siderite and pebbles of ferruginous rocks. In the central part of the Illizi Basin and over the Ghadames Depression (WT-1, HD% RYB-1, AKF-1) we observe sediments characteristic of deeper marine environments, i.e. fine-grained sandstones with intercalations of clays and silts. The characteristic feature of these Siegenian sandstones is the chloritic composition of the clay fraction of their cements and the chloritic-illitic nature of the argillaceous intercalations. Because of this situation, these deposits could have been derived from a hard substrate (effusive or metamorphic rocks) inundated by the Siegenian sea and open into the direction of the present Libyan coast. 2.2.1.5
Emsian The Emsian period is characterized by a deepening of the basin and, as a direct consequence of this, by the increasing role played by argillaceous strata. At the same time
2.2
• Lithofacies and Sedimentary Environments of the Oil- and Gas-Bearing Formations
21
we observe a shrinkage of the area underlain by the basin. To the east of the Ghadames Depression information gleaned from drill sites RYB-1 and AKF-t together with the presence of brachiopods suggest that sedimentation here took place in a shallowmarine environment. A vast number of mineralogical analyses on samples of this region show that the clays consist of illite and chlorite together with some siderite and calcite. In chemical analyses one finds that they contain a notable amount (7-12%) of ferrous iron (Fe ~+) and that Fe3+ is lacking. The alkalinity coefficient ranges from 2-4 with a mode around 3. The deepest part of the depository was located in the central part of the depression (HAD-l) because of the following observations: • • • •
thick deposits of sandy-argillaceous and argillaceous sediments; clay minerals composed of chlorite, illite and mixed-layer clays; alkalinity coefficient between 6 and lo; presence of p o t a s s i u m and sodium ions in the a d s o r b e d complex reaching 6-8 mg eq -I per loo g of rock, corresponding to a salinity of 2o-35 g 1-~.
The Fe 2+ content (present as pyrite and siderite) of the argillaceous rocks and of the cements in the arenaceous stringers amounts to up to 18% and that of Fe 3+ is about 2%. These parameters correspond to shallow-marine to even oceanic environments. Nearly all argillaceous rocks of the Lower Devonian section contain a mixed-layer swelling mineral with layers of smectite. In the diffractograms this mineral exhibits a widening of the lo X on the side of smaller angles, the widening assuming the shape of a peak at lo X in the case of a higher content of swelling phases. Saturation with ethylene-glycoI changes the prone of the lO ~ peak which becomes more symmetrical and centred at 9.8 A. Mineralogical analysis furnishes morphological data on the clay minerals in the Lower Devonian section. Regarding the origin of the mixed-layers, we observe, e.g. in the Tiguentourine (Tg-38), that there was a degradation of detrital illite by a leaching under the influence of freshwater. In the Zarzaitine (Zt-1), however, we identified a mixedlayer phase with 3o% swelling layers in which the particles are elongate platy, possibly resulting from the recrystallization of smectite. The presence of a small quantity of smectite particles in the rocks of the Lower Devonian section might be due to the transformation of felsic volcanic components introduced into the sediments. One reason for the preservation of this smectite could be the potassium deficit in the enclosing rocks. The southern boundary of the Ghadames Depression and the Illizi Basin during the Lower Devonian was characterized by shallow-marine domains in which wellsorted arenaceous sediments with a high content of organic matter were deposited. Within these rocks there is also a high content of pyrite, siderite and apatite as well as high concentrations of Fe 2+ (up to 8%). On the basis of these indications we may consider these formations as potentially hydrocarbon-bearing. 2.2.1.6 Giver
After a period of relative stability at the end of the Lower Devonian we observe another regional transgression with a m a x i m u m during the Givet. Over most of the
22
Chapter 2 • Characteristic Features of the Region
Ghadames Depression and the Illizi Basin we find argillaceous sediments indicative of a deep open-marine environment with an alkalinity coefficient of 4-6. To the east (RYB-1) and south (EAL-1) along uplifted zones sandy-argillaceous sediments suggest a shallow coastal environment (alkalinity coefficient - 1). Towards the west the basin became increasingly shallower with the deposition of argillites and carbonates. The Middle Devonian formations differ from those of the Lower Devonian in the composition of the clay minerals and the sandstone cements which consist essentially of kaolinite whereas illite appears to be dominant in those of the Lower Devonian. The presence of kaolinite, the high content of organic carbon and the presence of brachiopods and corals favour a warm humid climate. 2.2.1.7
Upper Devonian During the Upper Deonian period, Frasnes and Famenne, the marine environment continued. Over the continent a thick kaolinitic weathering crust was formed, resulting in an enrichment of the argillaceous rocks in kaolin. During the same period, to the west and in the central part of the depression (WT-I, HAD-l) sedimentation took place under normal salinity with alkalinity coefficients of 3-6. However, in the Illizi Basin (WIA-1, FRG-1, IRL-1) we observe a decrease of the salinity in the uppermost Devonian. This may be ascribed to the existence of the Amgnid-E1-Biod barrier which led to a relative isolation of the Illizi Basin from the rest of the Sahara. The Upper Devonian formations are characterized by chemical peculiarities, like, over a large portion of the basin, iron oxide concentrations sometimes amounting to 17-18% (SED-1) and Fetota 1 attaining values of 22% (SED-1 between 2 645.2 m and z 876.0 m). However, the vertical distribution of the iron content is by no means regular. Actually, a sample taken in the same area but at a depth of 2 649 m (Reservoir F-2) contained 20% Fe2+ compared to 24.5% Fetotal. Beds with higher Fe2+ concentrations usually contain chamosite and siderite as diagenetic minerals. The formation of diagenetic chamosite took place through metamorphism of detrital kaolinite or of fine-grained clay minerals like illites or mixed-layer clays. This formation could be noted under the electron microscope and by XRD-analysis. Peaks at 7.05 A, 3.02 A and 2.5 ~. indicate the presence of chamosite whereas a peak at 3.56 alongside the 3.5z A peak of chamosite may be ascribed to kaolinite. In the Ghadames Depression the sediments deposited farther away from the coastline generally contain much less Fetotal (3-6%), with Fe2+ predominating (WT-1, RYB-1). Part of this Fea+ may be fixed in the chamosite. The origin of ferruginous sediments in a cold environment is difficult to explain as it is generally held that areas in which the mobilization and transportation of iron takes place are more climatically extreme, like hot and humid regions. Without entering into a discussion of this problem we would like to point out that, in addition to humidity and temperature, the oxygen content of the atmosphere and the density of the plant cover over the emerged lands also play an important role in the dissolution of iron. During the periods prior to the appearance of photosynthetic land plants the atmosphere is generally thought to have been oxygen-deficient. Under such reducing condition the iron could have been dissolved and transported independently of the
2.2 • Lithofacies and Sedimentary Environments of the Oil- and Gas-Bearing Formations
~
"
~
~
/
J
23
I
:..3
Fig. 2.8. Modelof typicalpaleographicand depositionatenvironmentsof Devonianclasticsediments
position of the basin in relation to the climatic belts. The formation of goethitic oolites is generally considered as an indication of oxidizing conditions, but the Devonian oolites are made up mainly of chamosite. As a result of the Hercynian orogeny at the end of the Upper Devonian there is a slight uplift and a partial regression of the sea, a situation best developed in the eastern parts. The sediments characteristic of the transgression are ctastic materials deposited under shallow water with a sparse fauna. A schematic reconstruction of the various types of the Devonian paleoenvironments is shown in Fig. 2.8. 2.2.1.8
Carboniferous The onset of the Carboniferous is characterized by a maximum regression of the sea as indicated by the absence of the Lower Tournai and only minor thickness of the Upper Tournai beds in the Illizi Basin. The latter have also been observed in the central parts of the Ghadames Depression (HADq, SEDq). In contrast to this, Carboniferous deposits are totally lacking in the west (WTq) and south (SOH-~, ZM-0, indicating that the sea essentially covered the central portion and was open towards the Libyan coast. The analysis of the adsorbed complex in the argillaceous rocks of the Tournai in the Illizi Basin (IRL-1, EAL-0 indicates that the salinity of the Tournai sea must have been normal (alkalinity coefficient 3.5-4.5). The Visd is characterized by further subsidence of the Carboniferous Basin. In the central portion of the Ghadames Depression (HADq) sandy-arenaceous and calcareous sediments were deposited with a thickness of 7oo m, showing that the subsidence
z4
Chapter 2 . Characteristic Features of the Region
was very strong and of an eustatic nature. The lithological features and structures of the rocks such as argillaceous sandstones with contorted bedding and the abundant faunal elements show that over the East Saharan synclinorium during the Vis6 a warm shallow sea existed, wide open to the east and with a salinity below that of the Tournai (alkalinity coefficient 1.9-2.6). The clay mineral spectrum is made up essentially of illite, kaolinite, chlorite and mixed-layer swelling clays of the illite-smectite type. The qualitative and quantitative composition of the clay minerals depends on their depositional environment. In littoral zones where the influence of the continents is pronounced, the kaolinite content may reach 30% whereas in the deeper parts of the sea it will be only lo%. For the mixedlayer clays the opposite trend may be observed. The widespread presence of these minerals, the genetic conditions of which are similar to those of smectite, suggests an arid climate over the East Saharan synclinorium during the Vis& This also explains the formation of evaporites in isolated highly saline basins in the western Saharan region. 2.2.1.9
Upper Carboniferous The Upper Carboniferous is characterized by a regression of the sea and the deposition of sandy-argillaceous sediments impoverished in fauna. During the Namur we observe over a short period the appearance of a continental environment throughout the area of the present Sahara. The arrival of a renewed transgression coming from the Tunisian coast to the northeast initiated the formation of an epicontinental sea over a large portion of the itlizi and Ghadames Basins up to Tassili des Adjers during the Westphal. At the end of the latter there was another retreat of the sea and lagoonal and continental environments became established, accompanied by the deposition of clays and gypsum interrupted by rare sandy ingressions. 2.2.1.10
Conclusion The data presented above confirm that over a large part of the Paleozoic the paleogeographic conditions over the East Saharan synclinorium were such that the respective sediments can be considered as potentially hydrocarbon-bearing: " The tectonic regime was characterized by constant subsidence of the basin with few phases of uplift. This led to the deposition of thick sequences of argillaceous rocks with a notable content of organic matter, which could be considered as potential source rocks, and intercalated with them sandy beds were deposited as the eventual reservoirs. • An increase in the sedimentation rate as well as the preservation of organic matter in the respective sediments led to a reducing environment isolated from the influence of phenomena which could induce oxidative processes. • The predominance of a warm humid climate led to the development of a rich fauna and flora on the continent and to an intense erosion of the surface layers. • We also observe, in time and space, the existence of a shallow-marine (littoral and lagoonal) environment.
2.2 • Lithofacies and Sedimentary Environments of the Oil- and Gas-Bearing Formations
25
2.2.2 Paleozoic Sediments of the Oued eI-Mya Basin Throughout the entire Paleozoic the Oued el-Mya region was subjected to generally marine sedimentation. As the Paleozoic sea was unstable, it experienced several cycles of transgression and regression. During the Lower PaIeozoic (Cambro-Ordovician), periodical marine transgressions came from the northwest and covered the entire study area. At the base of a transgression sandstones and siltstones accumulated which contain thin conglomerate bands. During the extension of the transgression and the deepening of the sea, fine-grained sediments of a homogeneous facies continued and gradually covered the entire region. During the maximum of the marine transgression the facies conditions were most favourable for the formation of regional cover sequences for the Cambrian and Ordovician reservoirs. This was the period of the "EI-Gassi-", the "Azzel-" and the "Micro-
congtomeratic-Argillites". The regressive cycle of the Lower Paleozoic sea generally led to the deposition of sandy sediments with rounded grains of coarse to medium size which could be potential reservoirs for hydrocarbons. These are the "E1-Atchane-", "Ouargla-", "Oued Saret-" and "Ramade-Sandstones" (M'Kratta). The arenaceous formations "Zone Ra" and"Hamra Quartzite" occupy a special position in the history of the Lower Paleozoic. Their sedimentation is related to the marine transgression which took place during continuous subsidence, the latter being compensated by the regular supply of detrital material from the continent by a dense fluvial network. Thick arenaceous beds were formed by reworking and redeposition of the detrital material in the marine environment. These are widely distributed over the region and with their homogeneous lithology they are considered as potential reservoirs for the Cambrian and the Ordovician. In the dynamics of facies development during the Lower Paleozoic there is a certain similarity throughout virtually the entire Ordovician. The western, central and northwestern parts of the region were subjected to the stable accumulation of finely banded argillaceous deposits. This facies is unfavourable for the formation of good reservoirs in contrast to the predominantly siky-sandy deposits of the rest of the region which are favourable reservoirs. During the Lower Paleozoic the continent was under a temperate humid, sometimes even cold climate. This conclusion is based on: • • • •
the pronounced predominance of clastic formations, a large illite content in the argillaceous deposits and in the matrix of the sandstones, a low presence of carbonates and phosphates, and a tow iron content.
The transport of the detrital material to the zone of accumulation took a long time. The submarine currents and the oscillatory movements of the sea level played a rather important role in transporting and reworking of the terrigenous materials. The glaciation which covered a large portion of the areas to the south exercised considerable influence on sedimentation towards the end of the Ordovician. Through glacial rivers and icebergs the melting ice served as the main source of detrital material for the
26
Chapter 2 • Characteristic Features of the Region
marine basin (Beuf et al. 1966; Rognon 1971). During the Lower Paleozoic the main supply of detrital material for the marine basin came from sources located to the north, northeast and east of the Oued el-Mya Basin. A considerable part of the Lower Paleozoic sediments and in particular of the Ordovician ones resulted from the erosion and redeposition of earlier sequences. In the Silurian the rapid warming of the climate, leading to a massive melting of ice, resulted in a marine transgression of glacial-eustatic nature. The abrupt transgression limited the sedimentation of detrital material and is characterized by the accumulation of argillaceous-calcareous sediments throughout. In the Oued el-Mya area the end of the Silurian is marked by a gradual marine regression leading to active detrital sedimentation. From the Silurian onwards during the rest of the Paleozoic the centre of subsidence of the Oued el-Mya region was shifted gradually towards the southwest when the sea periodically came back. At the same time the uplift of the regions to the north and northeast increased their importance as a source of detrital material. The south and southeast always represented an essential source province. Under these conditions thick clastic sediments were deposited as well as carbonates during the short periods of a dry climate, in particular during the Middle and Upper Devonian. Over the sourcd regions of the sedimentary materials periodically a weathering crust was established. Over most of the Oued el-Mya Basin the rapid transgression was accompanied by the accumulation of mostly dark-coloured argillaceous-calcareous sediments rich in organic matter and frequently fossiliferous (brachiopods, graptolites, lamelli-branchs, coinoids). Silty-sandy zones are always present as are pyrites and micas, generally muscovite and rarely biotite. The clays, accounting for the majority of the section, are illitic and illito-kaolinitic with a small content (5-15%) of mixed-layer minerals. The carbonates are prominently developed in the lower part of the sequence. They are crystalline organo-detritic limestones with a variable clay content, as well as marls and argillaceous dolomites. The carbonate content increases in a regular manner towards the north and northwest. In the southeast of the region we observe in the sediments a higher percentage of sandstones and siltstones, sometimes accounting for up to about 40% of a section. The sandstones are fine-grained, well-sorted and subrounded with an argillaceous-calcareous cement. The thickness of Silurian sediments varies from 40 to 15o0 m, increasing towards the south and southeast of the region. The sedimentary conditions probably were not really calm. The lithofacies features of the Silurian appear to represent shallow- to moderately deeper marine conditions in the study area. The deepest part of the basin was located in the central region of the Oued el-Mya which is essentially argillaceous (9o-loo%), It extends over about one third of the region and would form an excellent cover for the Ordovician reservoirs. Towards the east and southeast of the region we note a regular change in sediment composition which indicates a progressive shallowing of the basin. The shallow portions of the sea are found in the north and northwest where silty-sandy formations assume an important role in the lithological composition of the succession, accounting for an average of 30% of the Silurian. The thickness of the sediments investigated indicates either a continuous subsidence of the basin or the development of grabens. Considering the data on the Fe2+/Fe3+we may visualize a reducing environment throughout the study area.
2.2 - Lithofacies and Sedimentary Environments of the Oil- and Gas-Bearing Formations
27
Towards the northwest and in the extreme north the proportion of sandstones and siltstones in the succession decreases and intercalations of carbonates, mostly argillaceous sandstones, make their appearance. These sediments formed in a shallow sea. They accumulated under deeper water than those in the north. According to the adsorbed complex the salinity of the Silurian sea must have been low, probably because of the melting of the glaciers which ted to a reduction of the salinity of the basin waters. The basin was not deeper than about 2oo m. Judging from the regional facies changes and the thickness of the Silurian sediments we can conclude that the main supply of terrigenous material came from the southeast where massive melting of glaciers took place. Another source of supply of detrital material was to the north of the basin. The predominance of leukoxene or zircon, of tourmaline and anatase indicates that argillaceous, crystalline and metamorphic schists as well as felsic intrusions formed the main source rocks. The prominent development of kaolinite in the clays may be used as evidence of the presence of a weathering crust in the respective source region. 2.2.2.1
Devonian In the Oued el-Mya region the Devonian sediments are only of local extent. In the Gedinne the regression started probably as a result of the Ardennian phase. During this regression the sea became increasingly shallower and the deposition of siltysandy sediments increased. In the southwestern part of the area the sediments are silty argillites with thin intercalations of marls and argillaceous limestones containing debris of thin-walled shells of brachiopods and lamellibranchs. Towards the end of the Gedinne this silty-argillaceous material was gradually replaced by sandy deposits. The dark-coloured Gedinnian argillites consist of illite (lo-9o%), kaolinite (3o-5o%), chlorite (lO-lOO%) and lo-2o% mixed-layer clays. Pyrite is also present and the bitumen contents are elevated. Based on the composition, stratification and grain size statistics we may conclude that these sediments probably are of shallow-marine origin, the original sea becoming gradually deeper towards the west and southwest. The frequent changes in Fe2+/Fe3+ are evidence of an oxidizing environment. During the Siegenian the regression continued with the accumulation of fine- to medium-grained sandstones which are well to moderately sorted, sometimes quartzofeldspathic with an argillaceous to calcareous matrix and containing frequently silty micaceous argillites and rarely layers of dark limestones with fossil debris. The heavy minerals are generally of a stable nature such as tourmaline, zircon, leukoxene, rutile and pyrite. In the extreme southwest the Siegenian sedimentation is represented by illitic argillites with low concentration of chlorite (lo%) and illite-smectite (lo-2o%). From grain size statistics the Siegenian deposits appear to have been deposited in a shallow-marine environment. During the Emsian the sea probably became wider and deeper because of a transgression, argillites representing the predominant rock type. During the Lower Devonian the deepest portion of the sea was located to the southwest of the region, where we observe the greatest thickness of the succession with 5oo-8oo m. Towards the north, northeast and southeast the sea became gradually shal-
28
Chapter 2 - Characteristic Features of the Region
lower. The increasing content of silty-sandy sediments as well as the occurrence of phosphate concretions are evidence of this situation. 2.2.2.2
Middle and Upper Devonian The Middle Devonian is characterized by a transgression accompanied by the accumulation of argillaceous-calcareous sediments. During the Upper Devonian the region was under a stable marine environment in which argillaceous-sandy sediments with a thickness of 9o m were deposited. In general the sandstones are fine- to medium-grained and moderately sorted with subrounded grains, in an argillaceous-calcareous matrix, rarely exhibiting cross-bedding and containing argillaceous laminae. The lithological features as well as grain size statistics show that the sandstones were deposited in a shallow-marine environment. The intercalated argiltites are mica-rich (>1o%) and contain 3o-6o% itlite, 2o-3o% kaolinite, lo-2o% chlorite and lo-3o% mixed-layer clays. In the Upper Devonian the argillites with the highest day content are found in the extreme southwest where we also find intercalations of limestones and marls. The sediments mentioned were probably deposited in a marine environment of moderate depth as indicated by the complete absence of Fe3+. The relatively high san&tone content in the Strunian sediments indicates a regression of the sea at the end of the Devonian. The Devonian sea is characterized by a tow salinity (alkalinity coefficient 2-3), whereas the widespread occurrence of carbonates and kaolinite indicates a warm humid climate over the continent. 2.2.2.3
Carboniferous Like the sediments of the Devonian those of the Carboniferous are only locally developed and because of this only poorly studied in the area in question. We therefore present here only a few general remarks. 2.2.2.4
Lower Carboniferous During the Tournai the Oued el-Mya region experienced a short regression characterized by the accumulation of fine-grained, poorly to moderately sorted sandstones containing micaceous stringers with a slight carbonate content and illitic to kaolinitic in composition with up to 2o% mixed-layer clays and chlorite. Grain size statistics of the sandstones indicate that sedimentation took place in a shallow-marine environment. From the Upper Tournai onwards throughout nearly the entire Upper Carboniferous, stable marine conditions prevailed with the deposition of up to 90o m of finely laminated argillaceous sediments. Highly metamorphed argillaceous schists as well as felsic and intermediate intrusions together with hydrothermal deposits probably represented the source rocks of the Lower Carboniferous sediments. According to the adsorbed complex the salinity of the Lower Carboniferous sea was similar to that of the Devonian.
2.2 - Lithofacies and Sedimentary Environments of the Oil- and Gas-Bearing Formations
29
2.2.2.5
Upper Carboniferous During the Namurian the sea was completely absent from the Oued el-Mya region, During the Moscovian a renewed transgression of the sea took place. As this went on, the silty-sandy sediments were gradually replaced by fossil-rich argillaceous-calcareous sediments. The sandstones contain subrounded grains of variable granulometry, they exhibit cross-bedding and contain up to 4% iron oxides and up to 2% glauconite. Their cement is illitic-argillaceous. The finely banded argillaceous sediments are illitic or kaolinitic with a constant content of about lo% mixed-layer minerals and <5% chlorite. Carbonates are generally represented by micro- to macro-crystalline dolomites and organogenic limestones. The thickness of those shallow-marine Upper Carboniferous sediments is 25-9o m. According to the adsorbed complex the salinity of the sea must have been normal. Frequent variations of Fe~+/Fe3+ are evidence of an oxidizing/reducing environment. From the end of the Carboniferous onwards the Oued el-Mya region, like the rest of the Sahara, was subjected to continental conditions probably as a result of the Hercynian orogeny. Permian sediments are entirely lacking from the study area. 2.2.2.6
Conclusion Over a long time the marine Paleozoic basin possessed only a limited connection with the ocean and was regularly supplied by freshwaters from the continent as indicated by the low values of the adsorbed complex (alkalinity coefficient) and by the marine fauna. Judging from the degree of chemical activity, the seawater was essentially weakly alkaline. During nearly the entire Paleozoic the environment was largely reducing and only rarely oxidizing/reducing. This is indicated by: • • • •
the the the the
constant predominance of Fea+ over Fe3+; constant presence of pyrites; periodic appearance of chamosite and glauconite; development of siderite in the cement of the sandstones.
For the Cambrian, Ordovician and Silurian we have outlined a degree of lithologicaI heterogeneity. This heterogeneity allowed us to better appreciate the dynamics of the sedimentary environments and of the tectonic framework of the respective source areas. The sediments with the greater lithological heterogeneity are those of intermediate lithofacies and of the transgressive cycles as opposed to those of the regressive cycles during which the sedimentary conditions were the least stable. 2.2.3 Sedimentary Genesis of the Triassic Deposits with the Northwestern Part of the Triassic Province as an Example
This investigation was carried out :in a portion of the Triassic Province in which diagenesis was rather diverse. Because of geological peculiarities like low general sub-
Chapter 2 • Characteristic Features of the Region
30
sidence, predominance of evaporites and especially of halite, etc. we observe zones of low diagenetic intensity and widespread sediments preserving their initial sedimentary features. This allows us to reconstruct their sedimentary history and that of the Triassic diagenesis without much extrapolation and uncertainty. 2.2.3.1
General Data of the Lithofacies Complexes of the TriassicDeposits The complex petrographical, mineralogical and chemical study of the Triassic deposits of the region (see Table 2.1) and the resulting lithological and facies parameters allow us to distinguish, from bottom to top, three terrigenous facies types: fresh- or brackish water deposits, calcareous evaporitic deposits and halite-bearing evaporitic formations. The main criteria for this identification of Iithofacies were drawn from the data on the mineralogical and chemical composition of the succession and especially from typomorphic and genetic peculiarities of the clay minerals. The association and the crystalline characteristics of the latter furnish indications on the genetic conditions of the rocks enclosing them and their significance for obtaining geological information has been amply described (Weaver 196o; Kiibler 1963, 1964,1973; Tardy 1971). In the present study the clay minerals were not investigated soMy by classical phase analysis with the aid of x-rays which, as is general knoMedge, would not yield any information on the genetic conditions of these minerals, but also by structuro-crystallochemical and typomorpho-genetic procedures. Typomorphic and crystallochemical
Table 2.1.
Subdivisionof the argillaceous-arenaceousTriassicdeposits
Northern Triassic province Main evaporite sequence
MASSIVESALT ARGILLITE D2 EVAPORITE D2
Southern Triassic province No clear correlation
MASSIVE SALT
Upper ArgillaceousArenaceous (TAGS) Arenaceous Triassic
ARGILLACEOUS CARBONATE
T2
Group 3:Sandstone A
T1
Group 2:Argillite between A and B
Group 1: Sandstone B Argillite between B and C Sandstone C Lower series
Argillaceous-calcareous
Lower ArgillaceousArenaceous (TAGt)
2.2 • Lithofacies and
Sedimentary Environments of the Oil- and Gas-Bearing Formations
31
features were not determined for the same minerals formed in different geological environments. In a saline basin there is a whole array of successive steps from a freshwater environments to moderately or highly saline conditions and we can study nature and dynamics of the crystallochemical transformations of the various clay minerals during the interactions of these minerals under changing conditions. Because of this we could follow the compositional modifications of the clay minerals in rocks characterizing different stages in the salinification of a basin starting from initially homogeneous terrigenous material. Because of their crystallochemical features the clay minerals are indicators of the physico-chemical parameters of the various facies of the sedimentary basin investigated and also for the post-sedimentary history of the enclosing rocks under conditions ofepigenesis and regional metamorphism (Kfibler 1964;Dunoyer De Segonzac 197o; Hayes 197o; Foscolos and Powel11979). Furthermore, an investigation of a vertical sequence of clay minerals from saline formations may furnish information on specific features and degree of secondary transformation of the enclosing salt-bearing strata. Kaolinite as a mineral characteristic of continental deposits is present virtually throughout all facies studied. However, variations in its composition during the early stages of diagenesis, as compared to those of chlorite, illite and mixed-layer clays, is of no special diagenetic significance. Because of this, it will not be discussed at this stage. 2.2.3.2
Facies in Terrigenous Argillaceous-Arenaceous Deposits of Fresh and Brackish Water 2.2.3.2.1
General Situation This complex of deposits formed during a short interval of sedimentation which took place discordantly on eroded Paleozoic rocks. Halite and sulfates here are virtually completely absent. Dolomite is also lacking (HRSq, OEN-1, SMZ-1) or only weakly developed (DJB-4, EHA-1) (see Fig. 2.9). The large quantities of dolomite observed in the Lower Triassic in drill hole 0MZ-1 are present in the form of large pebbles derived from higher-lying deposits. In this complex there are frequently conglomerates, microconglomerates and large rock clasts. As far as the rock chemistry is concerned, we note low concentrations of Na, Ca, Mg and K, not taking into account the presence of dolomite pebbles derived from higher areas. 2.2.3.2.2
Clay Minerals In all types of rocks, i.e. in the cement of sandstones, in argillites and in carbonates, we note essentially dioctahedral clay minerals (see Fig. 2.1o). The day mineral of sedimentary and early diagenetic origin was montmorillonite. Data on the chemical composition of the sandstones and argillites indicate a high Fe3+-content in the form of hydrogoethite [(FeO(OH)) x nHaO] entering the structure of the chlorite (vermiculite) and also, as we shall see later, that of illite, a process which we define for the time being as the end of the successive diagenetic transformations of the primary clayey material.
3z
Chapter 2 • Characteristic Features of the Region I
4° O
Be-1 oORA-I
oHR-13 o HR-6bis o HR-82
DJB-4 o
Berriane Ghardaia
°HIR-831 --J oGHA-1
/
oOEH-I
143; DJF-3 I 2
oHRS_1 | o HST-1 o Ae-2 1435
o~
OO
o OMZ'I
DJF-lO oDFA_I ODF-1 o MtA~3
I
o Af-2 O
1436
I Af-1 oAf. 4 o [
EHA-I
ATR-3 SDA-2 o o
3ooS~Z-I o Af-3
I
o Ak-1
32°
OSB-I o[
32°4o
Fig. 2.9. Locality plan of drill holes (scale 1 : 1ooo ooo)
The composition of the clay minerals in the cement of the arenaceous rocks is highly complex. The main constituents are minerals of the chlorite and illite group. The complex composition of the cement results from the variety of clays present, and mainly those of the chlorite group. An important role is played by swelling chlorite (or swelling vermiculite) which is found frequently together with dioctahedraI chlorite and nonswelling illite. In diffractograms of oriented sample preparations we observe a d(om) reflex at 14 A and a large, weakly crenulated peak of illite around d(om) values of lO-ll A. Saturation with ethylene-glycol leads to a peak at 17.7 -~. The intensity ratio of the peaks at 13.8 A and lo A on the diffractograms after heating to 35o °C suggests that part of the 14-A layers belong to a swelling chlorite (vermiculite). In addition to swelling chlorite (vermiculite) there is also in association with mixedlayer minerals of the illite-montmorillonite type (sometimes Fe-bearing) a chlorite that is unstable on thermal treatment. The diffractogram of oriented samples is char-
2.2 - Lithofacies and Sedimentary Environments of the Oil- and Gas~Bearing Formations
Ill Salt-bearing(haltic) and terrigenousfacies l
It Lagoonalterrigenous andcarbonatefades
Mg chlorite Fe-Mg-montmorillonite ~ (trioctahedral) ~%~ Fe-illite (dioctahedral)
Mg-montmoriilonite (trioctahedral}
Montmodtlonite"-'*
chlorite (swelling)
----*
veryfew swellinglayersand strong tendancyfor orderedcrystalstructure heterogeneousstructure;the amount of mixed-layerclaysof illitemontmorillonite-typeis reduced Mg chlorite Fe-Mg-chloJite~ " (tdoctahedral) + Illite-rnont (defective) morillonite
T
I
33
\
(tricctahedral) Fe-rnontmorilfonite- _~ Fe-illite illite (dioctahedral) Terrigenousdetdtal Fe-montrnorillonite / (dioctahedral) facieswith weaksalinity (dioctahedral) ~ . . . ~ Chloritewermiculite __~ (Fe-Mg)-chlorite (swelling) (dioctahedrat) (dioctahedraD
Feqllite
(dioctahedrat) + lllite-montmoriltonite
Fig. 2.10. Model of early diagenesis of clay minerals in the Triassic sediments of the Sahara
acterized by an intense d(ool) peak at 14 A. On saturation with ethylene-glycol d(ool) reflexes at 17.6 and 14 A make their appearance. Gradual calcination of the sample up to 55o °C leads to the gradual disappearance of the 14-A peak and to a widening of the lo-X peak. This may be explained by a progressive compression of the chlorite layers. The diffractograms of non-oriented samples exhibit only a highly asymmetrical peak on the side of smaller angles with d(om) = 1.49o A. In some samples a chlorite is predominant which may be distinguished by a series of peaks at 14, 7, 4.7 ~, etc. which remain stable even after saturation with ethylene-glycol and calcination to 55o °C. Illite appears systematically in all argillaceous fractions. In diffractograms of oriented samples there is generally a large, frequently crenulated peak in the range lO-12 ~. Saturation with ethylene-glycol leads to the appearance of an asymmetrical reflex on the side of smaller angles with d(ool) = lO-9.8 ,L 2.2.3.2.3
Partial Conclusion The clay mineral spectrum is characterized by the paragenesis of dioctahedral chlorites exhibiting primary crystallochemical features with dioctahedral illites sometimes containing swelling layers. This association is quite typical for continental deposits or those from brackish-water basins. In addition to chlorites with a stable structure we also observe swelling chlorites which are unstable on heat treatment and which occupy a sort of intermediate position between chlorite and vermiculite. It appears to represent a process of incomplete crystallization of minerals of the chlorite-vermiculite group in continental basins or under freshwater to brackish conditions. The minerals may thus be considered as metastable intermediate forms. 2.2.3.3
Carbonate and Terrigeneous Facies of the Saline Lagoons 2.2.3.3.1
General Situation Rocks of this complex make up the majority of the Triassic sediments T I, T 2 and the argillaceous carbonate of the Triassic Basin and of the Hassi R'Mel area (HR). They
34
Chapter 2 . Characteristic Features of the Region
form an argillaceous-arenaceous association with a notable dolomite fraction, either in the form of intercalations Of the terrigenous rocks or as cement to the respective sandstones. The quantity of dolomite varies widely between 2 and 80%. Its distribution on surface is by no means more homogeneous: large quantities are observed in the areas HRS-1, 0EH-1, SMZ-1 and DJB-4 (see Fig. 2.9). In the region of EHA-1 it is less frequent and still less around OMZ-z. However, around DJB-4 and OEH-~ calcite appears together with dolomite, whereas around OMZ-1 and SMZ-1 anhydrite makes its appearance. Spectrochemical analysis of the samples invariably shows a notable increase of Ca, Mg, SOl- and CO3- in the stratigraphic interval considered here. On the whole, the mineralogical and chemical data indicate clearly the initial stages of the salting-up of the basin, viz. the dolomite-sulfate stage with a salinity of 7-20%. In other words, the terrigenous-chemical complex delineated here was deposited in a basin that became increasingly salty. The carbonates (dolomite and calcite) as well as the anhydrite still exhibit primary features, a conclusion drawn from the study of numerous thin sections. Complementary proof for this is furnished by the dolomite pebbles within the underlying sequence especially in the area of OMZ-z. The conclusions on the sedimentary conditions of the terrigenous-chemical complex are supported more clearly by the typomorphic and crystallochemical observations on clay minerals described below. 2.2.3.3.2
Clay Mineral Paragenesis Throughout the areas studied we observe in rocks of different granulometry, from sandstones to rocks of the terrigenous-chemical complex, primary clay mineral associations. The main components of the associations are: a number of varieties of Mgchlorites, primary hydromicas grouped together under the term ferruginous illites, trioctahedral swelling minerals and mixed-layer clays. Trioctahedral Swelling Mineral. This mineral, possessing properties intermediate between montmorillonite and chlorite, is found in the siltstones and the sandy intercalations with dolomitic cement. A particular feature of this mineral is its swelling after saturation with ethylene-glycol to d(om) = 17.8 A and the removal of its Io.8 ~ peak on heating to 550 °C. Chlorite with Structural Defects. The mineral may be considered as a phase following the formation of a brucite layer in the series trioctahedral montmorillonite --~ trioctahedral chlorite. It is observed especially in the cement of the sandstones and in the siltstones of the Upper Lagoonal Complex. This chlorite with structural defects is characterized by a lower thermal stabilitiy, as on calcination to 600 °C the peak passes to lo.2 -A.
According to the XRD-data this mineral is very close to the dioctahedral Fe-Mg chlorites present in the cement of sandstones from the desalinized complex. Trioctahedral M g - C h l o r i t e . These phases are observed mostly in the dolomitized rocks and in the argillites in the form of trioctahedral chlorite. They are essentially Mg-rich. In the dolomites and argillites they are associated with large quantities of iron-bearing itlite.
2.2 • Lithofacies and Sedimentary Environments of the Oil- and Gas-Bearing Formations
35
Iron-Bearing Illite. This hydromica mineral of the type "green mica" or Fe-illite is present frequently in variable quantities in the clayey fraction of the different rock types of the saline complex where it is part of the paragenetic association Mg-chlorite + mixed-layer chlorite + mixed-layer mineral of illite-montmorillonite type. The Fe-itlites were noted in dolomitized sandstones and dolomites as well as in the argillaceous intercalations and argillaceous siltstones. Under the electron microscope the hydromica particles appear as only slightly characteristic laminar-elongate platelets. In the diffractograms, basal reflexes d(om) are present at lO-lOa ~., values that change neither on saturation with ethylene-glycol nor on heating to 550 °C. For d(o6o) a value of 1.5Ol A was noted, 2.2.3.3.3
Crystallochemicai Characteristics of the Clay Minerals of the DolomiteSulfate Stage of Solinization and Conditions of Formation At first sight, the association of clay minerals of this facies is rather common in terrigenous rocks. Chlorite and illite group minerals are widely distributed with only low contents of kaolinite and mixed-layer clays. However, detailed investigation of these minerals reveals crystallochemical features typical for an authigenic origin. The chlorite is trioctahedral and magnesian. The micaceous minerals are distinguished by a high Fe-content, belonging to the Fe-bearing illite group (Buatier et al. 1989). The predominant formation of the one or other mineral, i.e. Mg-chlorite or Fe-illite, in the saline facies depends mainly on the physicochemical conditions of sedimentation and in particular on the salinity and the cation/anion-content in the bottom water and the muddy water. The deposits are characterized by the absence of any clear connection between the type of detrital material and their clay mineral association. The latter is clearly controlled by the physico-chemical conditions of the different depositional environments. For identical detrital material the resulting clay mineral composition wilt differ completely between, e.g., tagoonal and marine complexes. Crystallochemical and typomorphic data enable us to propose a scheme for the formation of the main compound of this complex on the following path: Mg-montmorillonite --~ trioctahedral swelling mineral --~ chlorite with structural defects --~ Mgchlorite (see model in Fig. 2.1o). The factor responsible for the diversity of the intermediate phases of this series is the degree of establishment and perfection of the brucite layers and not the ordering in distribution of the interfoliar intervals between montmorillonite and swelling chlorite. As far as the hydromicas are concerned we may say that the formation of Fe-illite is typical of the early stages of sedimentogenesis and diagenesis. This conclusion is based on the presence of this mineral in numerous recent lakes in association with dolomite. A factor favouring the formation of minerals close to Fe-illite appears to be a high content of K+ in the bottom waters. Hydromicas are authigenic minerals rather easily forming in a sedimentary environment. Fe-illites would be the analogs of glauconites in gradually salting-up basins. The out-flow of bottom waters and the fixation of the bivalent cations Ca 2+ and Mg2÷ (in dolomite) in case of a relatively low concentration of the latter (sodic-calcic lakes) lead to the relative enrichment of these in potassium. The latter will then be fixed by the silicate compounds in solution leading to the synthesis of Fe-illites. In case
36
Chapter 2 • Characteristic Features of the
Region
the solution becomes enriched in Mg (sodic-magnesian lakes) which is incompletely fixed in dolomite we observe the concurrent formation of chlorite or of a swelling trioctahedral magnesian mineral intermediate between montmorillonite and chlorite. 2.2.3.3.4
Partial Conclusion The results obtained on the association of the clay minerals as well as on the crystallochemical pecularities of the Triassic deposits in the Triassic Province lead us to conclude that the association Mg-chlorite + swelling trioctahedral mineral + Fe-illite may be interpreted as an indication of the dolomite-sulfate stage of the salinization of a sedimentary basin of the terrigenous-chemical type. 2.2.3.4
Facies of Terrigenous-Hofitic Deposits 2.2.3.4.1
General Situation This facies is represented by an argillaceous-arenaceous complex with a high halite content ranging from z-3% to 15-2o% in certain regions. In the stratigraphy the complex occupies the upper part of the Triassic above the carbonate-terrigenous complex. There is a gradual replacement of the (dolomitic) carbonate constituents by halitic compounds. This follows from the appearance of a certain amount of rock salt downstream of the carbonate-terrigenous complex whereas the dolomites are strongly reduced upstream of the terrigenous-halitic complex up to complete disappearance. 2.2.3.4.2
Composition and Crystallochemical Characteristics of the Clay Minerals The clay mineral spectrum is notably less differentiated than in the other facies, the dominant minerals being trioctahedral chlorites and dioctahedral illites. In the chlorite structure, non-swelling layers predominate whereas the alternation of layers of different types exhibits a trend towards ordering. The proportion of mixed-layer minerals of the illite-montmorillonite type decreases especially as one approaches the massive layers of rock salt. Trioctahedral Illite. This is observed in all argillaceous fractions. The characteristic
feature of this mineral in the terrigenous-halitic complex is its heterogeneity. In diffractograms of oriented samples the peak occurs in the range lo-lo.3 A. Based on spectral analyses the hydromicas may be classed as Fe-illites. Trioctahedral Chlorite. This is encountered in virtually all fine-grained fractions of
the sections studied. There are no swelling layers in the structure of this mineral which exhibits a high degree of crystallinity. The d(o6o) reflex at 1.532 ,~ indicates that the chlorite is trioctahedral. There are no mixed-layer clays of the chlorite-montmorfllonite type, nor swelling chlorites nor, in view of the tack of any other structure, any minerals containing hydrate-sheets between their layers.
2,2 . Lithofacies and Sedimentary Environments of the Oil- and Gas-Bearing Formations
37
2.2.3.4.3
Partial Conclusion The components making up the fine-grained fractions of the terrigenous-halitic complex are essentially Mg-rich trioctahedral well-crystallized chlorites without swelling layers as well as Fe-illites, the structure of which does not contain swelling layers (see model in Fig. 2.1o). At the same time there are no mixed-layer species of the chloritemontmoritlonite type which were so characteristic of the carbonate-terrigenous complex (dolomite-sulfate facies). Mixed-layer clays of the illite-montmorillonite type also diminish throughout the complex. Thus, under conditions of higher salinity (-28%) in a hydrodynamically stable environment the composition of the clay mineral spectrum may vary considerably. There is a synthesis (or transformation) of clay minerals with a stable lattice when halite is being deposited. This is evidently tied to the physico-chemical conditions leading to the start of crystallization of halite (rock salt). A basin saturated in NaC1 is characterized by a low degree of hydratization. The lithification of halite during which any excess water is being eliminated takes place rapidly and favours the dehydration of the interfoliar space of clay minerals which prove to be most apt for transformations. Based on the above observations the well-crystallized Mg-chlorites and Fe-illites may quite rightly be considered as index minerals for the salinization stage in the Triassic Province as well as in other similar provinces.
2.2.4 Mineral Transformations in Sandy Reservoir Rocks Resulting from the Interaction Between Interstitial Waters and Primary Components During Early Diagenesis: The Diagenetic Signatures During the first stage of evolution of the interstitial waters, prior to their large-scale expulsion from the argillites, the amount of fluids in the deposits was low (ferralitic environment). The formation of cements was characterized by reactions associated with the bacterial reduction of sulfates. Fine-crystalline and micritic carbonates are formed then. Argillaceous sandstones represent the most favourable environment for the reduction of sulfates. The interstitial waters become acidic when hydrogen sulfide and carbonic acid are formed according to the following formula: 2(CH2OH) + S042- = H2S + 3HCO3 The transformation of micas and feldspars buffers the formation of acidic solutions by increasing the concentration of K+, Fe2+,Mg2+,AP ~ and Si4+ in the pore waters. These pore solutions are not displaced and the mineral authigenesis is controlled by a local equilibrium. Under these conditions the primary mineral composition of the deposit controls the distribution of the cement. In the early diagenetic environment the unstable silicates are dissolved and replaced by clay minerals leading to the liberation of potassium, aluminium, silica, iron and magnesia which then go into solution. As a consequence, many authigenic minerals form depending on the physico-chemical conditions prevailing in the eogenic pore waters. The appearance of mixed-layer minerals of the illite-montmorillonite type and subsequent ion exchange lead to the assimilation of potassium and magnesium from
38
Chapter 2 Characteristic Features of the Region .
the solutions. What is more, a notable quantity of potassium is fLxed in intermediate potassic sanidines (KA1Si3Os), a polymorphic variety of orthoclase stable in the diagenetic environment in question. The authigenetic formation of feldspars which leads to a high K+/H÷ is controlled by the reactions described above. When this formation continues it may be considered as a first indication for the arrival in this environment of components which maintain the K+/H+. The formation of poikilitic carbonate could be explained only by the advent of methane in the reservoirs when the interstitial fluids had started to become displaced. The methane originated largely outside the Triassic reservoirs in the Paleozoic argillites rich in organic matter. A certain amount of silica was deposited from the pore waters as overgrowths on quartz. The deposition of iron was controlled by the oxidation/reduction potential (Eh) of the pre-surface conditions during the early stages of diagenesis. Under alkaline-neutral conditions of oxidation in the pore waters the dissolved iron is rapidly oxidized and deposited on site in the form of an intermediate iron oxide which eventually is transformed into hematite. Other metals highly soluble in the oxidized state are fixed in the same way in authigenic or primary clay minerals. The chemistry of the two types of clay mineral formation (substitution or precipitation within the pores) is extremely variable. Under analogous physico-chemical conditions, illite, montmorillonite and chlorite may form within the same sandstone depending only on the chemistry around the grains upon which these authigenic clay minerals are deposited (Roberson and Lahann 1981). Actually, all the irons necessary to form the authigenic day minerals observed in the Triassic reservoirs of the Sahara during the eogenetic processes may be derived from the dissolution and clayey replacement of detrital silicate grains with the exception of Caz+ and HCO3 of the carbonate cement which have to be accounted for by outside sources. Glauconite and berthierite form in interstitial waters of marine origin or directly below the sediment water interface under conditions of oxygen deficiency or a lack of it when it changes between the surrounding environment and the interstitial waters of the substrate. The distribution of glauconite is controlled by the availability of iron and potassium as well as by the equilibrium between the supply of sedimentary material and its dissemination (Stille and Clauer 1994). There is also some berthierite observed in the redeposited sandstones located to the north of the Ghadames. Glauconite forms under open-marine conditions whereas berthierite is indicative of freshwater. The latter is also observed in the form of large scales replacing detrital micas. The presence of detrital berthierite may be used as evidence for marine sedimentation under some influence of fresh (meteoric) water, i.e. on the littoral shelf close to the continent, an environment clearly developed in the southwest of Hassi R'Mel. Lateral variations in the eogenetic association may be related to changes in sedimentary conditions or to the starting mineral composition. The development of nitrates was probably controlled in part by sedimentary rhythms. In the border zones of the Triassic Basin of the Sahara where the source of supply (Hoggar Massif) was probably closer by, calcretes were formed and the deposits were cemented by calcite with a heterogeneous structure. In the case of farther removed source areas and a more stable sedimentation the sandstones are characterized by a poikilitic nodular dolomite cement. Variations in the composition of the sandstone cements between calcite and dolomite may be explained by the ionic composition of the solutions in the
2.2 - Lithofacies and Sedimentary Environments of the Oil- and Gas-Bearing Formations
39
diagenetic environment resulting from the percolation of water from the edge of the basin towards the "depo-centre". At the same time, the contrast between the conglomeratic arkosic sandstones with predominantly illitic cement and the shelf sandstones with mainly mixed-layer minerals of the montmoriltonite-iltite type might be ascribed to differences in the mineralogy of the original detrital material. The feldspar-rich arkoses widely distributed north of Oued el-Mya after alteration contained in their interstitial waters a large proportion of potassium and here illite should be the most likely clay mineral. The quartz-rich sandstones (arenites) with conglomerates characteristic of the southern Triassic Province in which quartz dominates contain less potassium in their pore waters and here montmorillonite clays will be predominant as their ions are supplied from the dissolution of ferromagnesium minerals.
2.2.5 Evolution of Reservoirs During Progressive Subsidence (Late Diagenesis): Influence of the Sedimentary Inheritance During early diagenesis the detrital materials react with the interstitial waters, thereby creating conditions of equilibrium between the latter and the sediment. Although there will be no real chemical equilibrium achieved, the highly unstable diagenetic minerals will undergo decomposition and in their place stable authigenic phases will form. As the diagenetic mineral complex undergoes important later changes, it is necessary that the factors controlling the system, notably pressure, temperature and chemistry of the pore waters, are changing (Benbaccar and Fritz 1993).Thus, during later burial the mineral associations will reflect the modifications characteristic of the depth as well as the increase of temperature and pressure and additionally the evolution of the pore waters of the sediments themselves and of the waters migrating in from the surrounding sediments. These variations take place throughout the entire buried complex and thus lead to uniform transformations on a large scale and not just on that of the micropores. Towards the end of the Mesozoic, the Triassic basins were controlled by fault tectonics and achieved depths of locally over 3.5 km with paleotemperatures above 11o °C. The mixed-layer minerals montmorillonite-illite and montmorillonite-chlorite, typical of shallow depths, became replaced with increasing depth by the higher-ordered phases of illite and respectively chlorite passing through ordered mixed-layer phases or by allevardite and kalkbergite in the case of the montmorillonite-illite mixed-layer and corrensite in the case of the montmorillonite-chlorite mixed-layer. The latter phases (illite and chlorite) form in environments in which, depending on the facies, the original detrital material did not contain abnormally high concentrations of iron and magnesium. However, over most of the basins of the Saharan Platform and in particular in the Triassic Province, the Ghadames and the Illizi Basins, the terrigenous facies, depending on the source of supply, are distinguished by an elevated Fe- or Mg-content in the montmorillonitic phases and in some cases of both these elements concurrently. Under these conditions the catagenetic final mineral phases illite and chlorite to some extent will inherit the Fe-rich and/or Mg-rich nature of the detrital terrigenous phases. Fixation of iron and magnesium in the authigenic phases is in perfect agreement with the observed types of crystal structure, i.e. trioctahedral vs. dioctahedral. The corresponding conditions with depth are registered by the intercalated argillaceoussilty rocks. Furthermore, especially in the delta'fc complexes occurring throughout the
40
Chapter 2 . CharacteristicFeatures of the Region
Fig. 2.11. Flow chart of clay Initial sediment transformation in late diagenesis in sandstones (TriassicProvI DETRITAL ince) ILLITE
Solution
Intermediate
Final
'
[ DETRITAL I ILLITE
CHLORITE ~ _ ~ MONTMOR.
AOT"'GEN'CI
MONTMORLLONTE
I KAOLINITE
CORRENSITE|Il CHLORITE I (ordered ch/M) I I KALKBERG I 'I ,
K+
,,L,TE Ir MONTMOR.~
¢
ALLEVAR- II ~ , i(nonordered)i I DITE m
ILLITE ]
Triassic Province, the early non-ferriferous carbonates everywhere with depth are gradually replaced by carbonates with a higher Fe-content. Under these conditions it is also possible that in the mixed-layer clay structures the reaction between detrital potassic feldspars and montmorillonite will lead to the illitization of montmorillonite. The silica, iron, calcium and magnesia ions liberated in this process will be incorporated in the cements or enter into new authigenic phases (see Fig. ~.n). It is at this stage of the invasion of the reservoirs by active (acid) solutions that secondary porosity is formed on a large scale. Bacterial fermentation and thermal decarboxylation of organic matter in the shales introduce agressive waters rich in carbon dioxide into the reservoirs which selectively dissolve the silicates and the initial carbonate cement. At this stage also K+, Fe ~+,AP + and SiO~ penetrate into the reservoirs which facilitate the subsequent formation of kaolinite, feldspars and iron-bearing carbonates in the sandstone cement. The late, and possibly final, phase of the invasion of fluids into the sandstones is the eventual generation of hydrocarbons from the source rocks rich in organic matter. The advent of the hydrocarbons in several cases is associated with a second stage of formation of secondary porosity. Evidence for a considerable transfer of ions by the solutions comes from the observation that with depth, especially in the Triassic sandstones, there is a considerable development of secondary porosity because of a strong elimination of carbonate and siliceous cement. A mineralogical precursor of great value for a development of secondary porosity is the presence of a type of carbonate cement filling the portions overlying sandy bodies. However, really irrefutable proof for the dissolution and transfer of carbonate cement and the subsequent increase of the porosity in the horizons in which carbonates and feldspars had been dissolved, together with a drop in porosity in the overlying zones in which the carbonate became deposited under the appropriate physico-chemical and geological conditions, is found in the distribution of the carbon isotopes (~13C). This mechanism will be described in detail later in the section devoted especially to it.
2.2 - Lithofacies and Sedimentary Environments of the Oil- and Gas-Bearing Formations
41
The dissolution of the carbonate cement by carbonic acid leads to a progressive increase of the pH in mineralogically more "mature" sandstones, Under these conditions the presence, in the solution, of aluminium ions resulting from the decomposition of feldspars leads to the oversatnration of the solutions against kandites as confirmed by the deposition of kaolinites in the newly formed pore space. Under deeper burial the silica liberated during the destruction of the feldspars and illitization of montmoritlonite will be deposited, as the temperature drops during the displacement of the interstitial waters, above the more "mature" rocks in the form of oriented quartz aggregates filling the pore space of the less "mature" sandstones. The increase of temperature with depth leads to a drop of hydrogen ion concentration of the solutions so that Fe2+ and Mg2+ liberated during illitization easily enter into the structure of carbonates to form ferriferous dolomite, siderite and ankerite. A late structural inversion defined in particular in slightly subsided marginal basins, especially on the northern flank of the Illizi Basin, caused meteoric waters to penetrate into the respective sandstones. Under the conditions of the original distribution of temperature, pressure and pore water chemistry the eogenetic and mesogenetic mineral complexes became unstable and started to react with the actual subterranean waters in an attempt to achieve a new equilibrium. These weakly mineralized waters with a near-neutral pH became able to leach the carbonate, sulfate and h/tlite cements and generated a secondary porosity.
Summaryfor Sections
2.2.3-2.2.5
1. The Triassic sandstones encompass a complex of continental sediments formed essentially under fluvial and littoral conditions in a number of tectonically active basins with different maximum burial depths ranging from z to 4.8 km. Despite differences in the transport conditions of the detrital materials we observe in the different sedimentary basins a similarity between the respective diagenetic phenomena resulting mainly from the sedimentary environment. The above-described diagenetic processes are best developed in moderately subsided basins. The detrital grains underwent transformations leading to a well-defined equilibrium with the chemical conditions prevailing, at the paleosurface. The unstable silicates became dissolved or replaced and the ions set free under these circumstances contributed to the deposition of authigenic minerals like generally mixed-layers of the illite-montmorillonite type, potassic feldspars, iron-free carbonates and hematite. Lateral variations of the early diagenetic associations reflect changes in chemistry and temperature during sedimentation and diagenesis. In deeply subsiding sediments we observe a series of catagenetic transformations. 2. In the absence of early cements compaction leads to a pronounced drop in porosity. The mixed-layers of the type montmorillonite-illite are transformed into wellcrystallized illites passing at moderate depths through ordered mixed-layer phases like allevardite and kalkbergite. The early iron-free carbonate cements became dissolved on a wide scale leading to a strong development of secondary porosity. The carbonate cement deposited later includes ferriferous dolomite, ankerite and siderite. The later structural inversion of the Triassic basins during the Upper Mesozoic led to a renewed outcropping of the edges of the basins and of other less subsided basins. The recent subterraneous waters which are little mineralized and
42
Chapter 2 • Characteristic Features o f the Region
often of neutral pH are able to dissolve the carbonate, sulfate and halite cements as well as feldspars and iron oxides. In the Oued el-Mya and Ghadames Basins the Triassic formation waters were intensely altered by the penetration of acid solutions originating in deeper strata. The acid nature of these waters is caused by carbonic acid and other organic acids resulting from the decarboxylation of organic matter and from its thermal transformation. These acid waters led to the destruction of feldspars and to the formation of a large quantity of kandite as well as to widespread dissolution of carbonate cement and of iron oxides in the early cement. 3. Composition and mineralogy of the carbonate cement are controlled by the composition of the pore waters and by the presence of certain elements. Furthermore, the carbonate formed from the decomposition of organic matter is characterized by an isotopically light carbon. The carbonate resulting from the bacterial reduction of the sulfates is generally depleted in iron which is fixed in pyrite as the same process also furnishes H2S. The carbonate forming in the absence of H~O during, e.g., bacterial or thermal decarboxylation, may be iron-bearing to iron-rich (Irwin et al. 1977; Coleman et al. 1979). The sequential variations in the mineralogy of the carbonate cement reflect the compositional evolution of the interstitial pore waters during burial of the sediments. These variations may be ascribed to changes in the degree of dissolution of minerals, to mineral and organic reactions and to the authigenesis of minerals. The iron content of the carbonate cements of the intercalated argillaceous layers, e.g., increases frequently with depth of burial (Boles and Frank 1979; Irwin 198o; Matsumoto and Fijma 1981). It could be tied to the transformation of the clay minerals. The iron-rich nature of the late carbonates is controlled by their mode of origin as the latter is derived from the bacterial fermentation or the thermal decarboxylation of the organic matter. The occurrence of early carbonate cement which has a low iron content is related to the presence of pyrite and the reduction of sulfate. 4. The formation of authigenic clay minerals in the sandstones studied reflects the conditions characteristic of the entire complex of diagenetic reactions: eogenetic, mesogenetic and telogenetic. However, establishing the exact framework defining diagenesis and catagenesis of the clay minerals is rendered difficult by the complexity of the relationships between the various characteristic features of the transformational processes. The transformation of kaolinite, chlorite and illite probably illustrates the attempt of the original detrital materials to achieve an equilibrium with the interstitial waters. The authigenesis of illite may be controlled by the composition of the waters, in particular their K+- and AP+-concentrations. Kaolinite and chlorite are formed by reactions between the primary material with the products of the bacterial degradation of the organic matter. Chlorite is encountered virtually everywhere and is formed frequently as a result of the enclosing lithofacies. Illite then forms from waters derived from alkaline strata at greater depths.
2.2.6 Triassic Sediments of the Oued eI-Mya Basin A schematic map of the typical paleoenvironments of the clastic Triassic sediments (Lower Series, T1 and T2) is presented in Fig. 2.12.
2.2 . Lithofacies and Sedimentary Environments o f the Oil- and Gas-Bearing Formations
43
Barrier island lateral accretion and progradation I
~
~-
.........~
Fs
I
-
SL
Fig. 2.12. Model of paleogeographic and deposition al paleoenvironments for Triassic clastic sediments in north 0ued el-ivlyaBasin
The Lower Series was deposited on a flooded erosion surface over a Paleozoic substrate, a marine surface on which were developed vast lagoons, volcanic flows, tidal flats and barriers. Truely marine conditions prevailed only over the northern part of the region. The sandstones of the Lower Series are not as uniformly continuous as those of the T, and T 2. They are generally lenticular or sometimes form wedges interrupted by the volcanic rocks. A great potential for the accumulation of hydrocarbons in lithostratigraphic traps may develop under these conditions like in the sandstones in the upper part of the Lower Series in TKT-1, BKZ-1 and BKZ-2. In certain parts the volcanic activity and the sedimentation of the Lower Series (especially in its upper portion) took place concurrently as, e.g., between Takhoukht and northwest Guellala. As a consequence the flows mentioned influence the sedimentation of the Lower Series and thus the thinning and thickening of the latter. The discordant surface on the Hercynian folds appears to possess a less pronounced relief at the onset of sedimentation of the Lower Series. The source of supply of the material of the Lower Series must have been smaller hills as throughout the series there is no individual sandstone member which is thicker or of greater extent. In certain areas there is only an argillaceous succession directly above the unconformity surface. The maximum thickness of the volcanic rocks was observed in the Guellala area (GLA-2). This volcanic "ridge" appears to have restricted the deposition of the sandstones of T-i in the northern regions, leading also to a thinning of this unit from E1Hojira (EH-I) in the north towards Guellala in the northeast and Takhoukht (TKT-2) farther to the south.
44
Chapter 2 • Characteristic Features of the Region
Unit C encompasses the basal sediments of a marine transgression. As the sea level continued rising the eroded material coming from the tidal flats, the intertidal zones, the lagoons and the volcanic expanses of the Lower Series was deposited on the coastal sea floors under a shallow water cover. Most of these materials were derived from the reworking of the Cambro-Ordovician sediments and transported along the coast by coastal currents. • The sandstones of Unit B were deposited in a fluvial-deltaic environment or, more precisely, in the coastal barrier islands and the delta mouths. • The sediments of U n i t A were deposited in the numerous anastomosing small rivers on the mud flats of Unit B. The transgression coming from the north actually raised the river level and transformed the channels into a meandering network reducing the coherence of the fluvial network. In the M~gadine area (MGD) the sediments were deposited under marine conditions. •
Tectonism exerted a strong influence on the sedimentation of the Triassic rocks and controlled the distribution of the sedimentary environments. Along the edges of the basin structural disruptions led to the development of elongate horsts and grabens. 2.2.7 Triassic Sediments of the Ghadames Basin and the Northern Flank of the Illizi Basin
The Triassic sediments of the Ghadames Basin are sandstones, conglomerates, siltstones and shales with intercalations of dolomitic marls and salts. In the upper part of the succession thick layers of anhydrite and salts are widely developed. The argillaceous-arenaceous lower beds or the so-called Argillaceous-arenaceous Lower Triassic (TAGI) encompass terrigenous volcano-sedimentary rocks, the latter in particular in the northwestern part of the Ghadames Depression. The section is characterized by intense facies variations, the detrital rocks rapidly becoming replaced laterally by argillaceous sediments. The continental origin of these sediments is indicated by the presence of organic debris (mostly coal), the enrichment of the sediments in quartz and the kaolin cement of the sandstones. The argillaceous-calcareous part is made up of carbonates and terrigenous rocks intercalated with volcanics. The shales are of lagoonal or lagoonal-lacustrine origin. We sometimes observe here intercalated thin beds of sandy-silty material. The upper argillaceous-arenaceous layers, the so-called Argillaceous-arenaceous Upper Triassic (TAGS), in turn is made up of terrigenous rocks and evaporites. In its lower portion we find essentially sandy-argillaceous sediments of fluvial and deltaic origin. The terrigenous rocks are particularly variable laterally and frequently become argillaceous. Towards the north and northeast they become replaced by haline and argillaceous facies. The thickness of these strata decreases towards the south and southeast until almost complete disappearance in areas HAD-l, EOS-1, etc. The upper part of this horizon (S4) consists of thick layers of halite intercalated with shale and anhydrite beds in the northern part of the depression. The thickness of $4 decreases towards the south until complete disappearance in GT-1, HAD-l, WT-1, etc. The lateral distribution of the thickness of the Triassic rocks is characterized by an increase from south to north from which we may delineate in the southern part
2.2 . Lithofacies and Sedimentary Environments of the Oil- and Gas-Bearing Formations
45
of the depression a series of potential paleostructures within which the Triassic deposits, in particular in the lower portion, are little representative of the general development. The peculiarities observed in the Triassic sediments outline several different facies resulting from the frequent variations of the conditions of sedimentation. The majority of the Triassic rocks were formed under continental to lagoonal conditions. The TAGI was deposited over most of the region, with the exception of the lower part, under continental conditions leading to the deposition of fluvial and deltaic sandstones as well of conglomerates with intercalations of shale beds. Towards the northwest shales were deposited with sandstone intercalations at the base in a calm fluvial environment. The argillaceous-calcareous sediments were deposited in a persistent continental environment, in particular in the northern and western parts of the region, whereas the carbonates were deposited in lagoons. Towards the south of the depression we note a general decrease of the thickness of the entire succession. Throughout the sedimentation of the sandy-argillaceous and of the argillaceous-calcareous strata intense volcanic activity was developed in the region. During the deposition of the lower sandy-argillaceous sediments of the Lower Triassic the zones of sedimentation remained restricted to the northerly areas. During this period a continental regime prevailed in the southwest of the Ghadames Depression where mainly sandstones and conglomerates with shaly intercalations were deposited in a delta environment. In the northerly regions a continental and lagoonal environment resulted in argillaceous rocks with intercalations of layers of dolomite, anhydrite and sometimes of halite. The terrigenous sedimentation is preserved only in the northern parts of the region, giving way progressively to lagoonal deposits which are developed throughout the depression. This reduction in thickness of the Triassic sediments and the facies variations described are evidence of a single cycle of sedimentation during a period of instability with frequent changes of environments. This instability is underlined by the eruption of submarine andesites which took place only during the deposition of the Triassic terrigenous rocks. A pronounced unconformity is observed within the TAGI, heralding the establishment of a new sedimentary regime with a closed submarine environment under an arid climate. The first evaporites make their appearance here, allowing us to place the upper limit of the Triassic terrigenous deposits along this unconformity. In addition to the Triassic evaporites the two carbonate horizons D and B mark an increase in the sedimentary cycle. Horizon B, furthermore, should be considered as part of the evaporitic cycle and not as part of the Liassic. The eruption of the andesites observed in the central portion played an important role in the tectonic development of the province. Under the great weight of the sediments deposited in the depression, widespread fracturing developed around the periphery of the basins in the less stable zones which in turn aided the andesitic eruptions. The latter, as well as the Austrian orogeny, contributed to the modification of the shape of the basin. Its present appearance is the result of a long evolution starting at the end of the Paleozoic and lasting into the Tertiary. The reduced thickness of the sediments and the facies variations are part of the same sedimentary cycle during a period of instability in which numerous "rocking" movements occurred. This instability is also indicated by the submarine andesitic eruptions in the unstable zone which took place repeatedly only during the detrital Triassic.
46
Chapter 2 . Characteristic Features of the Region
2.3
Geological Parameters for the Formation of Non-Structural Traps and Their Mechanisms Experience from geological studies shows that the depressions inherited from the platforms like the Oued el-Mya, Ghadames and Illizi offer the best insights for research on deposits within zones of stratigraphic unconformities and lithological patterns resulting from facies changes.As described already there are two main structural stages in the sedimentary succession of the Triassic Province, viz. the Paleozoic and the Mesozoic, separated from each other by the regional Hercynian unconformity (Fig. 2.13). On the crest of the Hassi Messaoud dome Cambrian sediments crop out on the Hercynian unconformity surface whereas on the Tilhremt rise even older sediments of the Lower Cambrian are to be found (Fig. 2.13). The slopes of these large domes are covered over the Hercynian unconformity surface by successively younger rocks of Paleozoic age. The roofs to the Paleozoic reservoir rocks are made up above the unconformity by impermeable Mesozoic rocks and because of this we observe below the unconformity traps of the stratigraphic type. In the upper structural level the presence of lithological traps at the base above the unconformity is more likely. This is observed in the northern part of the Oued el-Mya
Fig. 2.13. Hercynian unconformity
2,3 • Geological Parameters for the Formation of Non-Structural Traps
47
Depression where there is a south-facing wedge of productive sandy horizons in certain parts of the Triassic and Permo-Triassic (Fig. 2.4). Along this wedge zone of reservoir rocks we may observe lithological traps in the basal rocks of the Mesozoic complex. It is even possible to locate stratigraphic traps resulting from the closure of reservoir rocks by internal unconformities. The Triassic sediments on the whole, whether they wedge out against the unconformity or whether they are affected by tectonic movements, may include non-structural traps. The presence of the appropriate structures and of faults indicates that the conditions for accumulation of hydrocarbons are favourable. Hassi R'Mel, a field with a surface area of some 2 ooo km ~, represents a closed structure within the regional sedimentary wedge. The lithological traps and those in sedimentary wedges are generally connected to structural features. In faulted regions the faults play an important role for the traps. They may form veritable screens for the hydrocarbons and sometimes even act as passageways for their migration. The stratigraphic traps are more characteristic of the Paleozoic sedimentary complex in view of the presence of numerous wedge zones and angular unconformities within the respective basins, Lithological traps, on the other hand, are formed mostly in the argillaceous to sandy Triassic formations where they result from facies variations within these rocks themselves. The general features of the geological structures and of the lithofacies controlling the formation of non-structural traps have been described already in the preceding paragraphs. Below we shall illustrate the importance of non-structural traps for the formation of the productive zones in the Oued el-Mya, Ghadames and Illizi Basins.
2.3.1 Non-structural Traps and Methods of Investigation The stud), of non-structural traps is based on the detailed analysis of all geological and geophysical data. The detailed approach depends on the structure itself and on its geological evolution as well as on its petroliferous properties. The most favourable factors are: L presence of reservoir rocks over a considerable thickness of stratigraphic section, truncation by the Hercynian unconformity and presence of an impermeable cover; 2. facies variations within the productive complexes, and mostly in those of Triassic; 3- contemporaneity of the main phase of hydrocarbon formation and of the development of stratigraphic and lithological traps; 4. development of the major tectonic elements at the same time, favouring preservation of these traps. The investigation is carried out in several stages: choice of the period of paleoreconstruction, solution of the problems relating to the correlation of the productive layers, establishment of geological sections and preparation of maps outlining zones of wedging and tithological substitutions, delineation of local traps and estimate of the eventual potential. In the first stage of this investigation it is advisable to prepare a map of the angular unconformities and the facies changes. Usually all these data are compiled from
48
Chapter 2 - Characteristic Features of the Region
lithostructural profiles (paleo- and actual), from paleostructural maps and from uncovered paleogeographic maps of the time of formation of the angular unconformities. It is advisable to choose as the period to be reconstructed a time after which the stratigraphic traps of the Paleozoic and the reservoir-rocks of the Triassic had formed and after which they were no longer affected by later tectonic movements. In this case we are dealing with time of deposition of the uppermost terrigenous sediments of the Triassic when a stable tectonic situation prevailed throughout the study area, as indicated by the very fine-grained sediments. As a result of this, a shale layer of constant thickness and chemistry was deposited throughout the area, to be replaced in the upper portion by the evaporites of $4. These shales are clearly recognizable on the radioactivity logs. Their top was used as a reference horizon for the paleoenvironmental reconstructions. In other words, all reconstructions were carried out for the time of deposition of the uppermost terrigenous Triassic rocks preceding the accumulation of the $4 evaporites. 2.3.1.1
Preparation of Lithostratigraphic Sections and Maps as well as Unveiling of Wedging and Lithological Substitutions The preparation of paleoecological sections is carried out with the aim of determining the limits of stratigraphic wedges and lithological substitutions as well as the probable migration direction of the hydrocarbons from the onset of the main phase of oil and gas formation. The sections of the present geological situations give indications as to the preservation of non-structural traps. The choice of the vertical and horizontal scales represents a key question during the construction of profiles. In order to obtain a clearer impression, the profiles should be prepared with a horizontal scale of t : 500 ooo and a vertical one of 1 : 5 ooo. Because of the large degree of vertical exaggeration deformations of the slope of the beds (dip of strata) and thicknesses become inevitable. However, the boundaries of the sedimentary wedges remain unaffected. At the same time one can mark the composition of the rocks in the profiles, which is important for outlining lithological traps. The profiles are prepared in the following way: to start off, we draw paleosections utilizing lithological data and the results of seismic investigations of the present geological situation of the area studied. Then we draw a section of the present situation. For this purpose, after entering the data from drilling and seismic exploration, we draw into the profiles the structural surface of the Hercynian unconformity, the limits of the sedimentary wedges and the thickness of the various stratigraphic members. The resulting network of profiles forms the base for paleostructural and paleogeological maps, the paleostructural maps furnishing information on the formation and limits of possible development (extension) of wedge-type reservoir zones. As maps of the ancient relief they also serve as a topographical basis for the preparation of paleogeological maps, increasing the reliability of the latter and making them more detailed. Geological boundaries may be joined up between the profiles taking into account the slope of the beds and of the paleorelief. The paleogeological maps of the stripped-off situation offer the possibility of determining the limits of the wedge reservoirs and also facilitate the outlining of the extension of stratigraphic traps. In the sections one can determine the zones of replacement or of changes in facies and delineate their
2.3 • Geological Parameters for the Formation of Non-Structural Traps
49
extent. Between the individual profiles the boundaries of the facies replacements have to be extrapolated. The delineation of local traps was facilitated by combining the stripped-off paleogeological map with the map of the present structure. In this way we were able to exdude certain disturbed portions of wedge zones where the traps have been destroyed by later tectonic movements.
2.3.2 Lithostratigraphic Traps The detailed stratigraphic correlation of the various lithological units in the Triassic north of the Oued el-Mya Basin and the determination of probable extensions of sandstones T1 and T~ revealed the presence of lithostratigraphic traps like the one resulting from the wedging out of the Lower Series or from coverage by volcanic rocks as observed in drill holes BKZ-1, BKZ-2 and TKTq (Fig. 2.14). Variations in permeability resulting from facies changes lead to the establishment of permeability barriers favouring the retention of hydrocarbons even at the levels of the flanks of steep structures like, e.g., between HLJ-2 and KG-2 and between HEB-3 and HEB-1 (Fig. 2.14). The nature of the lithostratigraphic traps varies with the level of the reservoirs located above the volcanic rocks (T-1 and T-z) as well as below them (Lower Series). In the upper reservoirs (T-1 and T-z) penetration of saline waters from the Saline Series ($1-S2-S3...) higher up has led to blockage of these sandstones which in petrophysical aspects are better sorted than those of the Lower Series. The same saline solutions almost completely destroyed the permeability of the sandstone T-z which could have represented, among the Triassic reservoirs, a storage area of some importance for the accumulation of hydrocarbons. In the Lower Series low permeabilities are developed throughout because off L the enormous weight exerted by the overlying volcanic rocks; z. the concentrated solutions derived from the overlying volcanic rocks and flowing towards base within the Lower Series; 3. the absence of a pronounced relief which could have acted as a source area for the sedimentation of the Lower Series by shedding coarse clastic elements leading to the development of thicker and more extensive sandy formations; 4. the argillaceous cement of the sandstones; from the point of permeability the sandstones of the Lower Series are generally rather poor but the development of fractures has largely improved their reservoir qualities and compensated the low primary permeabiliD: The presence of strong facies variations with certain trends could support the formation of traps of this nature. Within the Lower Series there are good chances for non-structural traps as the sandstones are lenticular. In the northeastern Gueltala area (GLNE-5; Fig. 2.12) the presence of hydrocarbons is restricted to the upper part of the Lower Series which in general contains productive horizons, the equivalents of which, however, towards the southwest become argillaceous and thus non-productive. Around Hamlet el-Beida, hole HEB-1 produced oil from Tq whereas in HEB-3 farther to the south in a structually higher position than HEB-1 the same sandstones
Chapter 2 • Characteristic Features of the Region
50
~ " ~3-~ ~.-~ .uo
~
~..~
.~
~88~S
.~
.e: >.
5
-~ %",~N~ I~ - - - ~ l
.
N
c~ o ~ o
~
~-~
~
£
~.f~ •
::: .':-: £-
o .~ ~
i ,% . . . .
m
O'
~
m
:=o
0
/
""
%
r
o
,.,;.:..;.,,.
~.
~ ~-'P,'-,;
}
?-'-'i "~
0
2.3 • Geological Parameters for the Formation of Non-Structural Traps
5:
of T-: are argillaceous and compacted. A similar phenomenon is encountered at Kef el-Argoub (KG): the sandstones of T-: produced oil in KG-2 and KG-4 whereas the same sandstones become clearly argillaceous in HLJ-z in structurally more favourable position and the T-: sandstones are non-productive at Hassi Ladjouad. These phenomena represent encouraging signs for the search for non-structural and structural traps.
2.3.3 Lithologicai Traps In the upper argillaceous-arenaceous Triassic (TAGS) of the Triassic Province two zones probably containing non-structural traps are normally delineated in the border areas of this region. The first zone occurs in the southern part of the Ghadames Depression in the vicinity of the Illizi Basin (see Fig. 2.:5, Zemlet Mederba area (ZM)). It extends with a general E-W trend over about :1o km along the inner margin of the wedging-out zone of the reservoirs. This region is characterized by a rapid rise of the sediments towards the south and the entire sedimentary complex of the Lower Mesozoic is wedging out rather abruptly to become replaced by the thinner Zarza'itine Series. From a "petroleum" point-of-view this zone is of interest because of important indications of oil, gas and water encountered in a number of holes (SOH% EOS-:, ZM-1) in the south of the Ghadames Depression. It is obvious that the lithological traps related to the wedging out of the reservoir rocks upslope could form in different parts of this zone. The second, about 70 km long zone is restricted to the southeast of the study area along the regional fault, i.e. to the Sedoukhane (SED, SEDE) and Dimnet (DIMN) regions (Fig. 2.:5). In plan this zone coincides with the analogous zone within the Lower Triassic but overlies it. From a petroleum point-of-view this zone is highlighted by the same factors. The regional fault borders the western, poorly studied limb of the Tartrat fold and a series of fields, in which economic quantities of oil and gas have been produced from the upper argillaceous-arenaceous Triassic lies on the eastern limb of this fold. Drill hole data show a thickening of the productive sandstones towards the west in the axial direction of this fold. It is highly probable that favourable reservoirs could have formed also on the western limb of this fold. These observations favour the formation of non-structural traps as a result of wedging-out and erosion of the reservoir beds in the upslope direction. Another zone, located in the western parts of the Triassic Province, exhibits signs of lithological traps. Drill hole data had already revealed some time ago an approximate line of regional wedging of the reservoir rocks of the argillaceous-arenareous Triassic (TAG). This boundary passes throughout the northern part of the Oued elMya Depression, forming a wide arc with the convex side pointing to the south. It may be traced between drill holes HRS-: and CHA-: over the southern closure of the Tilhremt rise (Fig. 2.14). Then it runs southward to the interior of the depression in the area of hole DR-: to turn to the northeast, passing along to the western slope of the Hassi Messaoud dome between holes GEC-1 and GEB-:. The boundary is probably sinous. South of the limit of extension of the reservoir rocks of the TAG the succession is replaced completely by impermeable argillaceous formations with a thickness not exceeding :oo m. The entire wedging-ont of the argillaceous sediments of the corn-
52
Chapter 2 - Characteristic Features of the Region
o i~
~
~
~'~
%
! ~
c~ J::J
r~
%
zW ..5 ,< .,Q
,~
"-J E o
I
o
Ii II
<
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N ~=
.Q
/:.°]
+<:
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0
~
--
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J
+
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+
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2.3. Geological Parameters for the Formation of Non-Structural Traps
53 L~ O
L~
t
~4
o
54
Chapter 2 . Characteristic Features of the Region
plex in question takes place near the southern edge of the Oued el-Mya Depression, 5o-2oo km from the wedging-out of the reservoir rocks. We thus observe in the northern part of the Oued el-Mya Depression a little south of the line Hassi R'MeI-Hassi Messaoud the regional boundary of the wedging-out of the TAG reservoir rocks.With these wedging zones and those in which permeable fades become replaced laterally by impermeable ones traps of the lithological type may be associated. As indicated above, all deposits located so far are tied to structural or structurallithological traps. Along the line of wedging-out of the reservoir rocks there are also traps of the lithostructural type as in general all Iithological and lithostratigraphic traps form under participation of structural elements, the substitution of reservoir rocks by impermeable ones mostly taking place on the flanks of structural deformations. Although most of the lithological and lithostructural traps of the type considered here are rather small, they occur generally in bands (or chains) along the wedge zones and their total reserves may be considerable. Thus, the analysis of the general situation presented here shows that the majority of the terranes set out are found in two regions. The first region encompasses a large group of terranes associated with the Tilhremt rise and its southeastern slope. The second one covers the zones adjacent to the parts of the Oued el-Mya Depression and to the northernmost part of the wedging zone of the reservoir rocks. In the former region the deposits of the TAG occur at topographically higher positions than those of the latter region. The analysis of the distribution of the hydrocarbon phases in the study area suggests that in the former region area one will find mainly condensategas deposits and in the latter mainly oil. The main interest for exploration lies in local areas in which zones of the different complexes overlap in plan, increasing the stratigraphic reach of the probable productive levels. Other examples of lithotogical traps in the Illizi Basin are shown in Fig. 2.15.
2.3.4 Stratigraphic Traps The boundary of the wedging-out of the micro-conglomeratic shales on the eastern limb and in the centraI portion of the Oued eI-Mya Depression is considered as the limit of this paleodepression. Prior to the deposition of the salt-bearing layer S4, the depression was a structurally complex basin with a NE-SW running axis. On all levels the depression exhibits a complex structure on the paleostructural map of the Cambrian deposits in the axial portion of which we observe a hump with an amplitude of 50 m which is also evident on the recent structural map. Its flank is complicated by a fault with an throw of 5o-7o m. Farther east of this fault the hump is also developed but of smaller extent and with a low amplitude. The hump is separated from the flanks of the depression by folds. The Taconic phase of the Caledonian orogeny is only poorly developed in the Oued el-Mya Depression, the sedimentary succession of the Ordovician being virtually not eroded at all and there is no recognizable angular unconformity. All stratigraphic wedges had originated during the Hercynian erosion, but the majority of the stratigraphic traps formed outside the boundaries of the depression or in its peripheral zones. The lithologicaI traps predominate close to the boundary of the depression near
2.3 . Geological Parameters for the Formation of Non-Structural Traps EG-I
KA-Ibis ZES-I
HTB-I
AKF-1
RYB-1
55 HFR-1
S Depth
- 10(3
1000
-20c
2000
-30£
3000
-40c
~0OO
-50£
5000 5000
Fig. 2.16. Geologicalcrosssection(north-south)in the GhadamesBasin(horizontalscale= i : 2300ooo; vertical scale = 1: 275ooo)
the hump of el-Agreb-Hassi Messaoud (Fig. 2.14). A long geological period is characterized by facies changes over relatively short distances and the ages of the formations possibly containing traps of the non-structural type vary widely, covering the Ordovician (Azzel, Oued-Saret, Dalle de M'Kratta), the Upper Silurian and the Lower Triassic. The depths of the respective reservoirs range between 3 650 and 4 050 m. Variations in the reservoir characteristics took place where the main tectonic elements intersected each other. In the uplift zone ofMessdar-Rhourde el-Baguel the contours of an older uplift coincide with the border zone of the non-structural traps which are characterized by the micro-conglomeratic shales. The northern and western edge pass through the flank of the Bozema Depression and respectively the Dorbane fold. Prior to the deposition of the salts the study region possessed a complex tectonic structure. Along the Cambrian surface a large anticline is clearly developed, the eastern limb of which is complicated by a graben (Fig. 2.16). This graben becomes more evident towards the south and disappears northward. Its displacement reaches loo-15o m. On the down-thrown side, at present with a positive relief, there are local folds. The northern tip of the uplift is also characterized by fauks with low displacements. Because of continued sedimentation the structure of the Ordovician reservoirs very much resembles that of the Cambrian (Fig. 2.16). The appearance of the interior of the graben has changed and the older folds in the area of the present uplift of Nezla, Messdar and Rhourde el-Baguel (Figs. 2.16, 2.17) have become more clearly developed. The amount of throw along the faults bordering the graben in the area around drill hole Ld-1 is reduced. The reservoirs in the Hamra Quartzites and the E1-Atchane sandstones which had been eroded along the Hercynian unconformity are wedging out around this uplift. In the Triassic sediments around the northern edge of the uplift a lithological trap might be developed. The reservoirs lie at a depth of 3 ooo-4 200 m and the quality of the Triassic reservoirs improves towards the north. Other zones in which stratigraphic traps formed in the Illizi Basin are shown in Fig. z.15.
56
Chapter 2 • Characteristic Features of the Region
~ -,.-, o o
B- ~'~, .~ ~ "~ o ~ c ..~.~ q - ~
" - ~ ~ ~ ..b,~
'.~ :-~
.~
~,~
.~
~
~
~
~
Jn~!JaJu I xnasaJ9 o1!61V se!Jl
'
o
C
~ : ~ •
~ ? < %
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~.~
o
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2,3 - Geological Parameters for the Formation of Non-Structural Traps
57
c
NO 0
e-
oN
)
~J
E
i
~5 t
i
~
NO
:
{:)-
v
.--
8 o
o
58
Chapter 2 • Characteristic Features of the Region
2.3.5 Traps Related to Volcanic Deposits Traps may also form in sandy reservoir rocks interlayered with volcanic formations. During the Triassic, numerous volcanic eruptions took place in the Triassic Basin and between the various eruptions of the volcanic phase sandy sediments were delposited which as intervotcanic reservoir rocks may represent good stratigraphic traps for the accumulation of hydrocarbons. Genetically these traps may be considered as stratigraphic but their transformation into mixed structural-stratigraphic traps could have taken place during subsequent tectonic movements. The best example for such traps is found in the Guenafid area (GEF-1) as shown in Fig. 2.17. Oil accumulations were located in this reservoir in the form of a thin sand lens between 2o m of shales and 6o m of volcanic rocks. The abnormally high reservoir pressure of 6oo kg cm -2 at a depth of 3 475 m in this deposit may be ascribed to a well-developed cover sequence. It is, however, not impossible that also above this there may be large accumulations of hydrocarbons. Anyhow, the volcanic rocks are restricted to the northwestern part of the Erg and such traps may only be found in this part of the region.
2.3.6 Traps Resulting from Differential Compaction of Sediments A particular type of non-structural traps has been observed, resulting from differential compaction in the sandy-argillaceous Triassic sediments.
2.3.6.1 Argillaceous Intercalations These are not regional trends but variations on a structural scale. In the sediments of the arenaceous-argillaceous Triassic it is frequently observed that structural features are accentuated by differential compaction. The large sandy channels like those encountered on the Sef Fatima structure (SF-1) may be resistant to compaction whereas the peripheral shales are strongly compacted (Fig. 2.17). Such structures may lead to the trapping of oil mid-flank on the structures in the crevasse-splay sandstones with sloping aquifers whereas the top of the structure is completely empty. The drill holes targeted on the Triassic in the Eastern Erg only explored the smnmit areas of these structures and this type of prospect was not recognized. It has to be pointed out, however, that these crevasse-splay sandstones are considerably thinner than the channel sandstones. 2.3.6,2
Wedges on Limbs of Structures This situation is well developed on the Berkane structure (BRK-1) where the Triassic structure is derived from a Paleozoic structure present on the Hercynian unconformity as a high (Fig. 2.17). This wedging-out exhibits two aspects:
2.3 • Geological Parameters for the Formation of Non-Structural Traps
59
Table 2.2. Behaviour of various transition and waste zones in respect to reservoir features Characteristics
Transition zone
Waste zone
Type of fluid
Oil and water, water
Oit, oil and water, water, oil and mud
Productivity
Good
Poor
impregnations ofoit reservoir
No reservoir
Poor oil saturation
No impregnation
Strong oil saturation
Poor oil saturation
Calculated oil column
Poor
Important
the beds are deposited on the flanks, not reaching the top, an idea outlined above; the levels reaching the top of the structure lose their reservoir characteristics (especially their permeability) and despite important oil shows produce water in a so-called waste zone. In such a waste zone oil is readily produced down-dip of the summit holes.
2.3.6.3 The Waste Zone Concept If we look at a section through a stratigraphic wedge trap we encounter, according to Schowalter and Hess (1982), from bottom to top (Fig. 2.17): the water zone; the transition zone; the oil zone; at the top the"waste zone" corresponding in actual fact to the passage from a reservoir to a non-reservoir zone on top of a wedge marking the limit of oil accumulation. In this zone the petrophysical features porosity, permeability and oil saturation decrease gradually to the top of the reservoir zone. Because of this, in a curve showing the relative permeability it appears that the waste zone may exhibit on tests the same behaviour as a transition zone, i.e. producing oil and water. This waste zone concept is of great importance as it is d e a r l y evident that whether we encounter a waste zone or a transition zone has important implications for exploration. In the first case we have to look for an extension of the reservoir in the down-dip direction, but in the second case in the up-dip direction. It is also of prime importance to be able to locate one's position. Because of this, the authors mentioned above have tried to characterize the behaviour of the respective types of zones (see Table 2.2).
2.3.7 Morphological Traps Paleomorphological and Stratigraphic Conditions. In the centre of the Eastern Erg (Ghadames) and in particular in the south there are sufficiently good prospects
60
Chapter 2 • Characteristic Features of the Region
8
0
G o
2.3 • Geological Parameters for the Formation of Non-Structural Traps
61
for finding non-structural traps. In this region of the Triassic Province the Paleozoic succession (Ordovician to Carboniferous) has become successively eroded in a northwesterly direction before becoming covered by the terrigenous sediments of the Triassic (Figs. 2.16, 2.18). The detailed geological studies allow us to suspect the highly likely presence of stratigraphic traps in connection with the Hercynian unconformity. The reservoir rocks truncated by the Hercynian erosion, viz. mainly the Devonian and Carboniferous sandstones, may occupy, depending on their resistance to erosion, topographically high, intermediate or low positions. The topographically higher areas may exhibit morphological features of hills, cuestas or spits. The formation of such a relief is the natural consequence of the renewed subsidence of the basin after a period of epeirogenetic uplift. The reservoir rocks of intermediate relief, because of their lower resistance to erosion compared to those of the non-reservoir rocks, may take the form of submerged slopes or escarpment cliffs or eventually of sink holes in topographically low-lying areas. These pateomorphological traps, like numerous other stratigraphic traps, require above the unconformity the presence of impermeable isolating layers to form partly or completely a barrier against the vertical or lateral migration of hydrocarbons. In this part of the basin the sediments covering the Hercynian unconformity are essentially made up of sandy sediments of the arenaceous-argillaceous Lower Triassic. This situation reduces the possibility of the development of a truely impervious layer for such traps. There are, however, throughout this region areas in which the sandstones have been subjected to the same pronounced facies variations. Additionally, the volcanic rocks deposited directly on the Hercynian unconformity in this area may form a good cover sequence. This situation is observed in general on the geomorphologically lower-lying areas or, in other words, in the paleolows of the northwestern part of the Eastern Erg.
2.3.8 Diagenetic Traps In certain instances, the reservoirs themselves, having already lost their reservoir properties because of diagenetic and catagenetic processes, may react as a barrier. Such barriers are tied to the unconformity but may be also of other origins such as: • secondary cementation of reservoir rocks by minerals like halite, anhydrite, carbonates, silica, etc.; • transformation of petroleum into resins; • alteration of feldspars and tufts and neoformation of clay minerals; • compaction under the weight of the overlying sediments or pressure solution. The identification of such dia- or catagenetic barriers is of great importance in the search for traps as the accumulation of hydrocarbons may have taken place below a more pronounced barrier located above the unconformity. This type of dia- or catagenetic trap is developed in the north of the Oued el-Mya Basin in the regions of Haniet el-Beida (HEB), Kef el-Argoub (KG) and Hassi Lad)ouad (HLJ), as illustrated in Fig. 2.14.
62
Chapter 2 - Characteristic Features of the Region
2.4
Brief Petroleum Geology In the sedimentary cover of the Triassic Province we distinguish three oil- and gasbearing complexes, viz. the Cambro-Ordovician-Silurian, the Devonian-Carboniferous and the Triassic, which are separated from each other by impermeable layers of regional extent. The first of these complexes is distributed throughout the province with a maximum thickness of 17oo re.Within the Cambrian reservoirs we distinguish, from bottom to top, four lithozones, viz. R3, Rz, R~ and Ri, which possess the best reservoir characteristics. They consist of heterogranular sandstones which are highly compacted, silicified and fractured. Because of this fracturing their open porosity may frequently reach lo-17% whereas their permeability at 2 darcy is elevated. The Ordovician reservoirs are made up of three sandstone horizons, viz. the A1-Athan, the Hamra and the Ourgla, with petrophysical parameters similar to those of the Cambrian. The reservoirs of the Lower Silurian are formed essentially in the depressions of the province. The Devonian-Carboniferous complex is developed in particular in the Ghadames and Illizi Basins as well as on the slopes of the Tilhremt and Mlal rises and in the southwest of the Oued el-Mya Depression where its thickness is reduced. The reservoirs of the Permo-Triassic or of the Lower Series in the Oued el-Mya Depression have an effective thickness of 2o-25 m. The middle complex (T1) is mainly developed over the Tilhremt rise and in the northeastern part of the Oued et-Mya Depression, being made up of two reservoirs, viz. C and B. The upper complex (T~), on the Tilhremt rise and in the Oued el-Mya Depression, corresponds to reservoir A with an effective thickness of 25-35 m, a porosity of 15-25% and a permeability of i darcy. The main oil reservoirs of the Triassic Province are contained in the Cambrian reservoirs and the main gas reserves in the Triassic ones. As source formations we have recognized the widely developed Silurian and Devonian shales and, to a lesser extent, those of the Ordovician and of the Carboniferous. In the province considered we have the giant oil fields of Hassi Messaoud within the Cambrian and the gas and condensate fields of Hassi R'Mel in the Triassic, In the Paleozoic and Triassic sediments underlying the salt-bearing formations of the Triassic Province we observe, as a result of the temperature logs of the various drill holes, that there is a strong connection between thermal flux and tectonic factors. The large positive structures and the local uplifts complicating them are heated more strongly than the negative structural elements. The development of saline (evaporites) beds has led to a regional lowering of the geothermal base and leveled out the deep thermal flow by thermal exchange along the surface of these beds. The reservoirs of the province are characterized by an abnormally high reservoir pressure resulting from the pronounced isolation of its reservoirs. The oils from all these complexes are rather light (o.79-o.85 g cm-3), low in sulfur (o.oo5-o.3%), methane-bearing and possess a low content of resins and asphaltenes.
Chapter 3
Main Factors of Reservoir Compaction
The properties of accumulation and filtration in the sandy-argillaceous rocks of the Saharan Platform encountered at different depths vary considerably, even when the respective rocks were deposited under similar conditions. With increasing depth of burial of the rocks there is a general deterioration of the reservoir characteristics. However, in addition to the general lowering of porosity and permeability of the rocks with depth, there are intervals along the sections where the properties do not exhibit abrupt modifications but remain constant, and in some instances they even increase again. This takes place in sections between z ooo and 4 ooo m. The reservoir features are a function of numerous factors and, primarily, of the composition, the degree of maturity of the materials and of the facies conditions during their formation. For relatively narrow granulometric spreads the amount of argillaceous cement is among the main factors controlling porosity and, in particular, permeability. The correlation coefficient between permeability and clay content of the reservoir rocks at o.7-o.9 is comparatively high. As the grain size and the quantity of pelitic material are correlated with each other, the study of the respective relations permits us to roughly calculate the limits of variation of the permeability of sandstones and siltstones as a function of their granulometry. 3.1
Evolution of Density, Porosity and Permeability with Depth Density as well as absolute and effective porosity vary in an irregular manner with depth (Fig. 3.1). In the upper part of the section there is a relatively rapid increase of density and a more pronounced drop in porosity. At depths between 2.1 and 3.5 km the rates of variation of these parameters decrease notably and even exhibit a reversal in direction because of the appearance and development of secondary porosity leading to a certain decompaction. Below 3-5 km these variations become less important. This confirms the observation that the Paleozoic sandstones usually lying at depths of 3.5-4.7 km may represent good reservoir rocks containing commercial reserves of hydrocarbons. The fields of correlation between porosity, permeability and density show that over the entire section the mean values of effective porosity of the Devonian and Cambro-Ordovician reservoirs, generally occurring below 3.2 kin, exhibit little variation. At a depth of 4.o-4.5 km porosity varies between 6 and 15% with a mean of 8.5%, whereas permeability varies between o.1 and 50 mD with a mean of 15 mD. The greater spread of permeability values compared to those of the porosity may be explained by the fact that permeability reacts more sensibly to changes in pore structure and geometry than porosity (Fig. 3.z).
64
Chapter 3 • Main Factors of Reservoir Compaction
Density gradient (D.G.) 0
10 .
20
0.5 '~
30
40
Density echelon (D.E.) 50
0
10
I 1.5 20 Porosity ~ __~ , ~ gradient(PG.)
20
30
200
'/
~
~ \ ~ / \ \ ~ [ ~.~ I ~ ~ {
~ ~ ~ =
Sandstoneporosity Siltstoneporosity ~ Sandstonedensity = 5iItstonedensitiy
E
~2 e-
K
1
l
[ '\ o/
/
n/
40
50
400
600
'
Porosity
echelon (PE.)
Echelonsfor ~
Sandstone porosity
-o----e- Siltstoneporosity ~ ~ Sandstonedensity = = Siltstonedensitiy
[~'k/
p.E.:looh2-hl100
i l1\ ",
. . . . . . .P1-P h2-hl PG 100
=100~2~ ;, (%/lOOm) D.G. =
P2 - 01 (%l100m) 100 h2 _ h----~'
Fig. 3.1. Generalizedcurvesforporosityand densitygradients and echelonsof sandstonesand siltstones with depth in the Oued el-MyaBasin (based on statistical processingof porosityand density data)
It has also been established that extent and cycles of variations of the porosity gradients of the sandstones and siltstones are different in the various regions of the study area. The porosity gradient in the Triassic deposits exhibits modest variations in the order of o.5% per loo m, values that may nevertheless change from one region to the next. This situation is explained primarily by facies changes as well as by the type of secondary porosity developed in the interval considered. In the sandstones of the Upper Paleozoic the gradient of (primary) porosity increases relatively rapidly to 1.3% per loo m in certain parts of the regions at depths of z.9-3.6 km. However, this drop of primary porosity is largely compensated by the development of secondary porosity and because of this the mean gradient of porosity varies little between 3.1 and 3.8 km and sometimes even is reversed (Figs. 3.1, 3.2). We thus observe an irregular porosity gradient pattern because of facies changes and differences in solution porosity. The greatest stability of the porosity gradient is achieved in Lower Paleozoic sandstones where this parameter decreases considerably at depths below 3.8 km with a rate of o.15% per loo m. This situation is explained by the fact that sandstones become compacted under the weight of the overlying strata by grain displacement and deformation so that porosity remains relatively constant. The increase of the porosity gradient between z.6 and 3-3 km is caused, in addition to grain displacement, by the concurrent development of chemical compaction by precipitation of silica and secondary carbonate cement derived from underlying strata. The abrupt decrease of the porosity gradient below 3.8 km results from an increase of the contact surface between grains and ensuing drop of tension in the contact zone. In the Hassi Messaoud field the proportion of concave/convex grain contacts reaches 5o-6o% of all intergranular contacts. In the Ahnet Mouydir-Gourara basins the porosity gradient in the Upper Paleozoic sandstones grows to o.8% per loo m whereas in the Lower Paleozoic it decreases gradually to o.35% per loo m.
3.2 • Relationship Between Structure, Gravitational Compaction and Pressure
65
Solution
0.3 0.2 0.1 I
i
Porosity _
5 I
t
15 1
I
(%) 25 I
Permeability I,
35 I
.......
10 I
I
I
(roD) 102 103 I
I
I
I
i
l
i
i
Density (g cm -3) 2.0 2.4 2.8 I
I
I
I
I
I
1.5
2.0
"~ 2.5 r0J a
3.0
3,5
4.0 I ,
! fields
Fig. 3.2. Generalizedfield correlations of porosity,permeability and density with depth and distribution in the Oued et-Myaand GhadamesBasins (based on statisticalprocessingof reservoir petrophysicaldata) The analysis of the development of porosity reveals the same situation as outlined above. In the upper sandstone horizons down to 2.5 k m it drops gradually from 19o m per 1% to 85 m per 1%. Below this it increases to reach 5o0 m per I% at a depth of 4.2 km. The density gradients have been established in the same way and reveal the same trends for the compaction of the section as those of the porosity.
3.2 Relationship Between Structure, Gravitational Compaction and Pressure Solution of Granular Reservoir Rocks 3.2.1 Gravitational Compaction and Pressure Solution During the initial stages of diagenesis the grains of terrigenous material are mechanically brought together. During the more advanced stages, the burial of the rocks together with the growth of pressure and temperature lead to the mutual interpenetra-
Chapter 3 . Main Factors of Reservoir Compaction
66
tion of the grains which becomes of increasingly greater importance. The interpenetration is followed by the dissolution of the minerals. The sediments studied lie at great depths and their temperatures reach 15o °C, favouring a pronounced gravitational pressure and the interpenetration of grains. The intensity of this process has, however, not yet been evaluated. It is clearly evident that the more the grains of detrital material are interpenetrated, the more sutured intergranular contacts are developed. In order to quantitatively evaluate the gravitational pressure, we have determined the proportion of these sutured contacts from all types of intergranular contacts. The respective data for the Hassi Messaoud field are presented in Table 3.1. The table shows that the greatest gravitational pressure is developed in Zone Ra where the rocks exhibit the highest proportion of sutured intergranular contacts. Gravitational pressure is also high in other lithozones of the productive Cambrian, but is lower in zones R2 and R3 than in z o n e R a despite the fact that they occur at greater depths than in zone Ra. This is explained by the increasing role played by the argillaceous material, the high content of which keeps the grains from approaching each other. This also becomes evident when we compare the number of intergranular contacts with that of the proportion of the argillaceous cement in the different horizons (Table 3.1). In the reservoir rocks of different grain sizes with an argillaceous cement content below 15% the amount of secondary quartz varies widely but reaches a maximum. When the reservoir rock contains more than 15% cement the amount of secondary quartz usually is low with only o.l-o.5%. This may be related directly to the permeability of the respective rocks as the proportion of secondary quartz (silica) grows with the permeability. Part of this silica is precipitated from the formation waters during the displacement of the latter. The more water circulates through the pores, the more silica will be precipitated. It has to be pointed out also that in the presence of fractures the reduction of reservoir pressure as well as the drop in partial pressure of CO2 become more obvious and spread more rapidly in waters within beds of higher permeability. This contributes, furthermore, to a greater precipitation of silica. It has been observed that higher quantities of silica are found in fractured areas, this being also evidence for its precipitation from formation waters. It is possible that deep thermal
Table 3.1. Petrographic features of some compacted Ordovician reservoir rocks in Hassi Messaoud oil field
Cambrian sedimentary units
Portion of sutured intergranular contacts (%) Argillaceous No.of cement (%) (Mean observations without with total values in brackets)
regeneration regeneration Alternation zoDe
57
4
61
10
(23)
R~
70
1
71
11
(10)
7
Ra
63
13
76
8
(8)
50
R2 R3
66
5
71
16
(15)
12
50
0
50
24
(30)
5
]0
3.2 - Relationship Between Structure, Gravitational Compaction and Pressure Solution
67
waters play a certain role in this process. There is also another source of silica as this mineral becomes dissolved at contact points to become precipitated in the pore space. It could be shown that the depth of interpenetration of the grains depends on the granuIometry of the rocks. Microscopic determinations of this value were carried out for classes of granular reservoir rocks which differ in granulometry. The proportion of argillaceous cement in the Cambrian reservoir rocks at Hassi Messaoud does not exceed 15%. Histograms of the distribution of the percentage of penetration (5o-7o measurements per class) and the respective means are plotted in Fig. 3.3 and, for Hassi Messaoud, given in Table 3.2. They show that the interpenetration of the grains in the respective rocks grows when their dimensions become larger. This is explained by the fact that the total surface area of the intergranular contacts decreases with the increasing size of the grains. This follows from the relation that the effective pressure on the
80-
I
70" 60" 50" 40" 30" 20"
I
I
Width of regeneration borders Depth of grain interpenetration Median grain size from
M d = O. 1-0.2
Md
thin section
measurements
1o. 070 -
Md=0.2-0.3
60" 40 30-
~
~
////v///////
o.
/
60 • e-
Md = 0 . 3 - 0 . 4
l
50-
.24o
[
.~ 30 "~- 20
,\N?
I
10 0
/
! I Md = 0.4-0.5
20 10 0
Md = 0 . 5 - 0 . 6
20 10 0 30 20 10 0 30 20 I0 0 0.0t
I
t
Md = 0 . 6 - 0 . 7 ~-2~. ~. • I I Md = 0 . 7 - 0 . 8
\\\\
\\\\
I 0.03
0.05
0,07
0,09
0.I 1
Dimension (mm) Fig. 3.3. D i s t r i b u t i o n
of the percentage
of penetration
0.13
0.15
0.17
68
Chapter 3 - Main Factors of Reservoir Compaction
Table 3.2. Correlation between median grain size and grain size ofintergranular pores in compacted
reservoir rocks (all values in ram) Reservoir rocks classed according to granulometry
Median size after thin section measurements
Median size of intergranular pores evaluated graphically
Values GraphicallyMeangrain measured calculated dimension in thin values sections
Depthofinter- Widthof penetration overgrowth of grainsfrom rim mechanical compaction
Closestpacking Same,buttaking ofdetrital mate- into accountgrain rial but nointer- interpenetration penetrationor from mechanical regeneration compaction
Same,buttaking into account graininterpenetration following overgrowth
0.1 0.2 0.2-0.3 0.3-0.4 0.4-0.5 0.5-0.6 0,6-0.7 0.7-0.8
0.072 0,013 0.073 0,020 0,021 0,027 0.028
0.014 0.024 0,033 0.043 0.053 0.062 0.072
0 0 0 0 0,001 0.003 0.008
0.72-0.25 0.79 0.25-0.37 0.31 0.37-0,50 0.44 0.50-0.62 0.56 0.62-0,75 0.69 0.75-0.87 0,81 0.87M.00 0.93
0.022 0.025 0,028 0.032 0.035 0,032 0.031
0.006 0.012 0,020 0.024 0,030 0.032 0.040
(i.e. the difference between the geostatic pressure acting on the framework of the rock a n d the p r e s s u r e w i t h i n the d e p o s i t c o u n t e r a c t i n g c o m p a c t i o n ) surface of the intergranular contacts of coarser-grain rocks increases, leading to a more intense solution of quartz and thereby to a deeper interpenetration of the grains. The pressure at the contact points of grains in finer-grained rocks decreases also because of the overgrowth referred to above. Table 3.2 presents only mean approximate evaluations which may have different expressions in different rocks. We, nevertheless, may state that mechanical compaction acting on granular reservoir rocks around 3 ooo-3 5o0 m will not lead to the complete elimination of all pores unless we take the overgrowth of quartz into account. It would be interesting to analyse the behaviour of the quartzitic reservoir rocks lying at a depth of 4 o o o - 4 5oo m which in petrographic composition and granulometry are similar to those of zone R a. Calculations show that the effective pressure in Cambrian rocks of the same age lying at depths of 4 ooo-4 5oo m amounts to 5oo kg cm -2 at a densitiy of 2.3 g c m -3 and the excess pressure of the suspected oil is 45o-5oo kg cm -2. The loss of permeability of the reservoir rocks possessing a permeability of >5o mD is not large, amounting to 1.o-1.5% when the pressure increases from 3ao to 5oo kg cm-L We have thus to concede that because of petrographic analogy at depths of 4.o-4.5 k m the reservoir rocks do not lose much permeability when the rock temperatures increase only little. It is obvious that the mechanical compaction is also a function of the duration of the influence exerted by the effective pressure on the framework of the rocks. Table 3-3 presents the characteristics of rocks comparable in granulometry and percentage of argillaceous cement, illustrating the extent of quartz solution under pressure. To judge from the importance of sutured intergranular contacts, the younger Triassic and Jurassic rocks are less compacted than those of the Cambrian. Two other factors, viz. depth and petrographic composition, will also exert a notable influence, but it will be very difficult to evaluate them quantitatively without special investigations. It could
3.2 - Relationship Between Structure, Gravitational Compaction and Pressure Solution
69
Table 3°3. Correlation between age, day content and compaction level in various oil fields
Field
Age of rocks
Mean depth (m)
Proportion of sutured intergranular contacts (%) without with overovergrowth growth
total
Proportion ofargillaceous cement (%)
Mean size of grain (mm)
No. of observations
Hassi R'Mel
Triassic
21 O0
41
3
44
15
0.17
3t
Hassi
Jurassic Cambrian
1 900 3500
50 54
0 8
50 62
17 21
0.15 0.18
20 42
Messaoud
be assumed, e.g., that the greater compaction of the Jurassic rocks compared to those of the Triassic is caused by the higher content of ductile argillaceous debris in the former (Table 3.3, Fig. 3.4). Because of this, in the search for reservoir rocks possessing the best potential in sequences intensely transformed during categenesis one should take into account their mineralogical composition and in particular the content of quartz and clay. Experience from studies in polymict reservoir rocks containing 20-35% ductile debris shows that their accumulation and filtration Characteristic drop sharply at depths below 4 500-5 ooo m. It has to be pointed out that mechanical compaction of rocks differing in grain size will lead to variable losses of their reservoir characteristics with time. In fine-grained rocks the pores, because of their small dimensions, are rapidly reduced during interpenetration of the grains, explaining the accelerated rate of loss in reservoir properties. Because of this the search for better reservoir rocks in sequences undergoing more intense diagenesis should concentrate on coarse-grained horizons as these will maintain open pores over a longer time because of the larger dimensions of the latter. It is obvious that the formation of overgrowth rims will make the pores and the channels narrower. From the data discussed above (Table 3.2, Fig. 3-3) we can evaluate the influence exerted by overgrowth on the pore dimensions. The overgrowth rims tend to become enlarged in coarser-grained rocks and we have to assume that the process of grain penetration is followed by the formation of overgrowth rims. Considering the shrinkage of the pores resulting from compaction and overgrowth we have determined graphically how this process would affect the volume of the pores if overgrowth would encircle all grains bordering the respective pore. The results are shown in the last column of Table 3.2. They illustrate that when the above-mentioned conditions are met and the pressure solution of quartz operates at a medium rate the greater part of the pores will be entirely eliminated by their processes in fine- to mediumgrained sandstones. In contrast to this in rocks with a mean pore size above 0.5 mm the pores will not close up completely and their dimensions will actually grow in coarse-grained rocks. It is obvious that these are only rough approximations, but they are clear evidence of the fact that coarse-grained sandstones and gravel beds represent the best reservoir rocks in successions undergoing intense diagenesis. The calculations performed also show that fine-grained rocks are the first ones to be compacted by the process of quartz crystallization. This process is thus regular and any difference in timing will result from the differences in rock types encountered.
70
Chapter 3 . Main Factors of Reservoir Compaction
As a first approximation the variations in accumulation and filtration properties of the reservoir rocks down in a section that has undergone early diagenesis depend largely on the main characteristics established during sedimentation. The best reservoir rocks will form on the shallow continental platform and their good filtration properties result primarily from their low content in argillaceous cement. This question
Fig. 3.4. G r o u p s o f r e s e r v o i r
L
Age of rocks
I
r o c k s o f d i f f e r e n t m e d i a n sizes
II ,
(E)
D
T
J
I
;ToI
---
0.020 I
J
0.030 1
0.010~ 4 0.020
~t
I
'6
'
I I
~"'"
""
.-~'
-3
/
'-6j
0.030 4 I
0.010 ~
8 .q£.
.o
oo ot j
.6
0.030 I O.OLO
i
-I
-...j . CO ~~ ~,~._ _ / ,
5"
0,020 -
"~
"
,, I ....I I
o 0.010
l
-6 1
~~ . * "
o.030-
&
3
-
q
0ool
1
I I °°1°t ~ 0.020j 0.030 "t
• ~ ~ ~ ~ ~ "~
o.o~o-~..~ -ff
0.020 -
-4
0.0,04 I
.~"
13,
,.,.e
6
__Vj _~3:
~
" ~
Depth of grain
interpenetration
Ratio between depth of grain
penetration
t "3 I
]
-6
anOm t .........
I
3.2 • Relationship Between Structure, Gravitational Compaction and Pressure Solution 100
I.
II.
•
"~
75
gl,~.
r-'T~
=
" ' "
-'oM o • o o'~N~=*-~
•
"
~
~o
.--
• . -; ,-"-,t; o
Ill.
7z
Correlation coefficent for I = 0.94 II = 0.96 Ill = 0.93 IV = 0.89
IV.
'
:.O oOo -
o• oo o¢ A A N ~ a ~ 2 .0• • • , o AK**z~aaa
50
O u
6,~" IN
~
x
\
\ k
25 A I. Reservoirs of barric type from Palaeozoic, Oued Mya Basin A IL Reservoirs of deltaic type from Trias, Oued Mya Basin o IlL Reservoirs of fluvial type from Palaeozoic, Ahnet-Mouydir-Gourara Basin • IV. Reservoirs of marine type from Ordovician, Ghadames Basin 0
10
20
3'0
Porosity (%)
Fig. 3.5. Correlation diagram of porosity with mineralogical (chemical) maturity (increase of silica, especially secondary) and with structural maturity related to different facies-environment situations
has been intensely studied with the aim of establishing a quantitative link between the tithofacies conditions of the reservoir rocks and their accumulation and filtration properties under the same diagenetic regimes. The correlation plot established for reservoir rocks formed in different environments in different basins (Fig. 3.5) shows that porosity grows with increasing facies energy, all other conditions remaining unchanged. This increase in sandstone porosity takes place in the following sequence: tidal bars > delta > fluvial > marine. The succession is justified from the point of view of textural maturity of the reservoir rocks as evident from the high degree of sorting and rounding of the primary detrital material (mostly quartz) in the same order. In other words, under identical diagenetic conditions as expressed by the same amount of total and secondary quartz porosity, will be preserved best in tidal bar sandstones, followed by deltaic sandstones and finally in fluvial and marine sandstones. 3.2.2 Structural Arrangement of Sandstones
In order to obtain a more representative estimate of the compaction of reservoirs we have studied the contact index calculated as an average value of straight, concave/convex and sutured contacts. Sandstones containing much ductile cement suffered more intense compaction. Independent of their proportion of ductile (argillaceous) cement the Triassic reservoirs have suffered only weak compaction with depth (Fig. 3.6a, c). Their grains were displaced by shifting of one grain against the next at the contact points and by rotation (mechanical rearrangement).
72
Chapter 3 • M a i n Factors o f Reservoir C o m p a c t i o n
100%tangentialcontacts
100%embayedcontacts
100%long contacts 200 0 10 i
210
220
230
240
250
260
270
280
5
10
15
20
25
30
35
40
8 '
6 '
4
2
b
.OK20 (%)
2
2.5'
4
290Density (gcm "3) 45Montmorillo. nite in interlayered(M-I) minerals (%)
3
o
,
~ ~
A
1Reservoirprosity • 2 Adjacentshaledensity A 3 Montmo~'ilionite-lllite proportionin adjacentshale A 4 Polassiumcontentin
•e -, •
~Ax
3.0" A
A
t
3.5" .
iV"
-
: :"
4.0.
*•
4
° 8*xl#~° 6
10
14
I? 22 Porosity (%)
2'6
3'0
34
36
Fig. 3.6a,b. Mechanical compaction features of Saharan reservoirs, a Contact type plot for sandstones from s Cambro-Ordovician, Hassi Messaoud oil field (cement content <1o%), 2 Devonian, Stah oil field (cement content <12%),3 Triassic, Hassi R'Mel gas field (cement content <16%). b Correlation curves of 1 reservoirs porosity,2 adjacent shales density,3 montmorillonite/illite proportion in adjacent shales, 4 potassium content in shales also
3.2 • Relationship Between Structure, Gravitational Compaction and Pressure Solution
Fig. 3.6c, d. Mechanical compaction features of Saharan resservoirs, c Correlation curves of contact index with depth for Triassic, Devonian and CambroOrdovician reservoirs, Oued etMya Basin; 1 sandstones with 2o-3o% ductile cement, 2 sandstones with 2o-lo% ductile cement, 3 sandstones with less than 2o% ductile cement. d Comparative correlations of tight packing index for Devonian sandstones with ao-15% ductile c e m e n t from 1 Ghadames Basin, 2 0 u e d el-Mya Basin, 3 Ahnet-Mouydir-Gourara Basin. Average number of long, concave-convex, sutured and embayed contacts per grain (tight packing index)
2
73
Average number of contacts per grain (contact index) 3 4 5 6
2.0
2.5
~.
3.0
3,5
4.0 Average number of long, concavo-convex, sutured and embayed contacts per grain (tight packing index) 15 30 45 I
i
I
2.0-
":N?i.: ?:.% : \"
,,
,"
A
,~
A
A
2.5"
.C
~. 3.0-
3.5-
• Ghada • OuedMyaBasin " Ahnet-Mouydir-GouraraBasin
4.0
"."
..[:~: " i: I ' " •:ii-.
74
Chapter 3 - Main Factors of Reservoir Compaction
The Devonian and especially the Cambro-Ordovician reservoirs (Fig. 3.6a,c) are distinguished by the number of concave/convex contacts and sutures as well as by fracturing in particular in the Hassi Messaoud field. In this type of reservoir the rearrangement of the quartz grains in the sandstones take place also by flexures tied to the presence of mica, and by elastic and plastic deformation resulting from the clays typical of the Devonian and Carboniferous reservoirs in particular of the Illizi Basin. We have then undertaken a comparative correlation of the packing index with depth to reveal the influence of the thermal flow on the rearrangement (or compaction) of the Devonian sandstones, i.e. of a certain geological interval, with a limited spread of ductile argillaceous cement (lo-15%) and a nearly constant granulometric composition, i.e. a set petrographic parameter in basins with different thermal gradients: Ahnet-Mouydir-Gourara, Oued el-Mya and Ghadames (Fig. 3.6d). It becomes clear that the sandstones of Ahnet-Mouydir-Gourara at a higher geothermal regime are visibly more compacted than those of the Oued el-Mya and Ghadames Basins which are presently found at identical depths, i.e. there is a direct connection with the geothermal gradients. In contrast to this the sandstones of Oued el-Mya and Ghadames with similar thermal flux are at the same level of compaction judging from the contact index. In all cases the role of the temperature manifests itself probably in an acceleration of the processes of chemical compaction and in a decrease of the hardness of the grains. Thin section counts show that the number of grain contacts on a grain in the highly compacted Cambrian sandstones on average is about 5. Furthermore, nearly three quarters of them are of the sutured type which leads to a new local (interporous) distribution of silica. The mean interpenetration depths of the grains during gravitational compaction in medium-grained sandstones (Table 3.2) serves as a basis for the graphic determination of the amount of silica dissolved along the grain contacts. Morphology and dimensions of the pores are controlled by shape and dimensions of the detrital grains bordering them. Between grains of 0.2-0. 4 mm diameter the mean pore size would be o.l-o.3 mm, i.e. the mean pore diameter is about half that of the detrital grains. At these depths of 1.8-2.6 km the pores are mostly isometric. At depths of 2.6-3.6 km we observe a general shrinkage of the pores in the reservoir rocks and they acquire more or less sinous outlines, eventually becoming split up into smaller ones. The number of pores in cracks and interpore channels amount to an average of 70% of the pore space. When the size of the detrital grains drops to o.o5-o.1 mm the pore diameter will range between o.o2 and o.o7 mm. At depths of 3.6-3.9 km most of the pores are transformed into narrow channels essentially isolated from each other. The mean pore size is 25-33% of the mean grain size. The elongation of the pores, the ratio between largest and smallest dimension, is 2.5-5.o. In well-sorted sandstones and siltstones containing relatively little allogenic argillaceous cement all epigenetic transformations of the detrital matrix frequently lead to a volume decrease and to a complication of the configuration of the pore space and, because of this, to a deterioration of the reservoir properties. During burial of the sediments under the influence of the increasing weight of the overlying strata and of the warm formation waters the detrital grains in the rocks become dissolved primarily at the grain contacts and are deformed mechanically, leading to a denser arrangement. Quartz is dissolved at a higher rate than the feldspars and becomes transported in the alkaline environment. Due to this dissolution along the contacts a certain amount of silica is liberated which is partly deposited on the detrital grains as
3.3 • Reservoir Compaction by Silicification from Other Sources
75
new overgrowth. The proportion of secondary silica in the studied sandstones ranges from traces to above 3o%, achieving a maximum at 4.o-4.7 km depth. The overgrowth rims on quartz in the san&tones of the Saharan Platform are essentially highly irregular, usually covering the grains forming the pore partly and rarely on all sides. The thickness of the overgrowth zones is generally o.m-o.o3 mm, sometimes reaching o.15-o.3o mm and rarely even o.7 mm. The transformations in the rock matrix described above are the main cause of the deterioration of the reservoir properties during burial of the sediments. Thus, in rocks lying at depths of 1.6-2.6 km the mean porosity rises to 22% at a permeability of t8o mD, ranging between 12 and 30% and o.5 and 2 ooo mD respectively. For rocks at 2.6-3.6 km depth the mean values of porosity and permeability are 2o% and 5o mD respectively, for those at 3.6-3.9 km 12% and 20 mD respectively. They decrease at depths of 4.o-5.o km to 6% and 0.5 mD respectively. We shall present later other diagenetic transformations which contribute to the establishment of these petrophysical properties. 3.3
Reservoir Compaction by SUicification from Other Sources The compaction of reservoirs by silicification is the most important process characterizing the diagenetic transformations of the Saharan reservoirs. It started during the early phases of lithogenesis and continues until the appearance of oil and gas in the traps and, eventually, in the water-saturated horizons. Under these conditions we can distinguish several generations of secondary silica. Without exception secondary silica is encountered in all sandstones studied. It is recognized under the microscope in thin sections from some inclusions of dust around detrital quartz grains, the lack of atlogenic impurities and by the development of rhombohedral faces in the pore space in optical continuity with the detrital quartz grains. When there are no dust inclusions around the grains it is frequently difficult to define the boundary between the detrital grains and the zones of secondary overgrowth when the optical orientation is continuous. Because of this the proportion of secondary silica is difficult to determine. However, in every thin section there will be portions where the boundaries between detrital grains and overgrowth rims of secondary quartz are clearer than in others and where the error in calculation will be minimized. The proportion of secondary quartz in reservoir rocks differs from bed to bed. However, the total amount of secondary quartz in the beds and in particular the proportion of completely silicified layers varies considerably in the different sections. We may distinguish four classes of layers: (1) little silicified ones, mostly containing 2-5% secondary quartz, (2) moderately silicified ones with usually 7-1o%, (3) highly silicified ones with mostly 12-15% and (4) very highly silicified ones with mostly 17-2o% secondary quartz. The silicification of the sandstones as an expression of their mineralogical (chemical) maturity strongly influences the characteristics of the reservoirs (Figs. 3-7,3.8). In view of the irregularity of the silification any connections between the primary composition of the sandstones and their reservoir properties are rather complicated. There is some regularity in the relation between permeability and grain sizes as, all other parameters being equal, the permeability increases with the grain size.
76
Chapter 3 - Main Factors o f Reservoir Compaction
• L Oil-bearing reservoirs 25. ° II. Water-bearing reservoirs II
20 '~
I
25" 20
•
5-
15
o0l %
ooo*o e° • •
O°o o
o•
8 I0
o~
5~
°o
10
a
leo •
•
10"
,•
II
I
I
% %
\oo
",
oo
°°°
o
o
o
e
5-
%o
•
%oo °o o
o •
30
b
•
ee
o °o° * e o O*°oe • ° oo ° o ~ % o o o ov
20 Porosity (%)
•
• 'o:..
o~ o o
•
oee • •
° ••
~e
o o
•
• eeo • • •
•
• • •
•e
•
•°
• •e
eo
! •
10 100 Permeability (mD)
I 000
Fig. 3.7. Correlation curves of a porosity and b permeability with secondary silica for Palaeozoic reservoirs of Oued el-Mya Basin: I oiI-bearing reservoirs, I1 water-bearing reservoirs. Striving for representative results, the studied sandstones were selected so that they have a similar medium grain size (0.2-0.3 mm), and a close argillaceous cement content (lO-15%) with very little or no carbonate cement
30-
with y silica with iry silica with lary silica
20"
o
10"
0"
0
1
10 Permeability (roD)
1O0
t 000
Fig. 3.8. Impact of silicification (secondary) on reservoir quality, z Reservoir rocks with 1-5% secondary silica, 2 reservoir rocks wiht 5-1o% secondary silica,3 reservoirs with lo-2o% secondary silica
3.3 • Reservoir Compaction by
Silicificationfrom Other Sources
77
3.3.1 Silicification from Underground Waters The silicification of sediments little affected by diagenesis under surface conditions is a pronounced frequent phenomenon entailing the dissolution of silica by surface waters, as well as hydrolysis of silicates with the subsequent redeposition of the silica in the discharge zone of the surface waters (Taylor 195o; Thompson 1959; Weyl 1959; Trumit 1968). These cases of silicification are widespread under geochemical conditions favourable for the dissolution of silicates. Our observations have shown that secondary quartz is abundant in those beds in which the quantity of argillaceous cement is negligible. This situation is also well developed on the scale of thick beds in thin section studies. It is obvious that layers of more argillaceous rocks are characterized by a lower permeability and vice versa. Consequently, in less argillaceous (more permeable) strata, deep-seated mineralized waters and formation waters should be able to migrate provided the other conditions are favourable. To some extent these conditions probably existed in a number of provinces cut by a dense network of fractures. We may assume that a certain amount of secondary quartz could be deposited from thermal and formation waters. These processes could take place in permeable layers in which formation waters and deep waters can circulate as well as a result of the lowering of the partial pressure of CO2 under these conditions. In any case the presence of gypsum, anhydrite and baryte in numerous sections of the rocks studied is evidence of the penetration of mineralized waters derived especially from the salt-bearing sections of the Triassic. We have also tried to find a connection between the permeability of the sandstones and the distance of these horizons from tectonic disturbances. It has been noted that the mean permeability of reservoir rocks in holes situated less than 5oo m from fractures or faults differs very little from that in rocks from other holes (Fig. 3.9). We may suppose that certain formation and thermal waters could be displaced along fracture zones but that the circulation of these waters could take place essentially within layers possessing a certain permeability (slightly argillaceous beds). Under these conditions some of the silica could be deposited in the pores from these solutions as shown above. The presence of faults and fractures taken as a sign of silicification of reservoir rocks is one precondition but not the only one. This criterion becomes real in those cases where the beds are permeable and little argillaceous. In permeable beds we consequently have to analyse the relation between the content of secondary quartz and the fractured zones. In these cases old but reactivated fractures like, e.g., those of the Triassic and Cretaceous are of particular interest as they possess a higher flow potential and thus do not represent screens isolating displaced reservoir blocks. The silicification of reservoir rocks may take place under different geochemical conditions favouring the dissolution of silicates. These processes take place where rocks are buried to great depths under high pressures and temperatures. Thus a number of the Paleozoic sandstones in the basins studied possess a geothermal gradient of 25 °C km -~ whereas the temperature of the deposit itself sometimes reaches 15o °C. The equilibrium solubility of quartz in water of z5 °C is about 5 mg 1-~, increasing relatively slowly with rising temperature to zo mg 1-~ at 5o °C and to 6o mg 1-1at loo °C (Willey 1974). The solubility also increases with growing hydrostatic pressure, i.e. by
78
Chapter 3 - Main Factors of Reservoir Compaction
A
a
60
°°1
50
b
50. 40,
30
30,
20
20.
10
I0.
i
,
10
i
J
102
,
J
103
J
J
104
Kihi (mD H)
i
i
10s
10
702
103
104
105
Kihi (roD p)
Fig. 3.9. Histograms of permeability indice in Hassi Messaoud boreholes at various distances from tectonic disruptions, a Wells at distances less than 5o0 m from tectonic disruptions, b Wells at distances more than 50o m from tectonic disruptions
about 35% per kbar (Wiltey 1974). The following data have been determined by our laboratory for the silica content in sediments from the northern part of the Saharan Platform: 5-1o mg 1-1 in surface waters; 5 - m mg t -~ in thermal waters; 1-2 mg 1-1 in shallow coastal waters; 11-12 mg 1-1 in fluvial waters. It is well known that meteoric waters with lo-~5 mg dissolved silica per 1 circulate down to depths of several hundred meters, representing one of the sources of cementation of sandstones at shallow depths. However, it is obvious that the above-mentioned silica contents are too low to explain all the secondary silica encountered in the formation waters of the Saharan basins at depths of several thousand meters. The Cambro-Ordovician formation waters of Hassi Messaoud contain about 15o mg dissolved silica per 1 and the Albian formation waters of the same field 15 mg t 1. In the Berkaoui field the silica content of the Albian formation waters is 2 mg 1-1and that of the Triassic waters 14 mg 1-1,In the Hassi R'Mel field (gas and condensate) the Triassic formation waters contain 7 mg dissolved silica per 1. It goes without saying that the 15o mg dissolved silica per 1 in the formation waters of the Cambro-Ordovician reservoirs may represent a source for cementation and compaction of the sandstones whereas the 7 mg silica per 1 in the Triassic formation waters will not have any effect on the cementation of sandstones containing 6-1z% of secondary silica. We thus still have to locate an important source for the cementation of the sandstones.
3.3.2 Silicification Through Transformation of Feldspars, Detrital Illite and Other Micaceous Components Silicification following the transformation of detrital itlite and secondary kaolinite has been noted by a number of geologists (Vatan 1962; Millot 1964) and was confirmed by the following observations: the kaolinite in the pore volume of the sandstones exhib-
3.3 . Reservoir Compaction by Silicification from Other Sources
79
its indications as to its formation in situ. The transformation of illite into kaolinite should be accompanied naturally by the liberation of silica. It can be assumed that it will be enclosed in the affected pore itself or in the neighbouring pores. This process takes place during hydrolysis of silicates and leads to a lowering of the pH, to the migration of elements like Na, Ca and Mg and to the immobilization of others (A1,Fe, Ti, etc.). The quantitative assessment of this process, however, shows that it cannot by itself account for the 5-1o% of secondary silica frequently observed. The transformation of itlite into kaolinite can actually liberate 20% of silica from the amount of illite affected. The mean content of clay in the sandstone cement rises to above 15%, but it frequently is only lo-t5% or even lower. If we assume that all clay minerals would have been derived from detrital illite then, on the basis of the chemical formula of illite, viz. (K, Na, My) (A1, Fe, Mg)2 [(AlxSi4-x)Olo](OH)2with o.5 > x > o.75, the above-mentioned transformation should result in 2% secondary silica in the total rock. Such an amount of primary illite, however, could not have been contained in the sandstones of the study area and there is frequently even detrital kaolinite in addition to the newly formed one. Judging from the mineral composition of the shales in which such secondary transformations would be reduced to a minimum there could have been only 15-3o% detrital illite as a primary constituent in the sandstones of the Hassi Messaoud field whereas the amount of illite detected in the Cambrian siltstones is 3o-5o%. Consequently, only 1% of secondary quartz could have resulted from the transformation of illite to kaolinite. Thus, widely developed kaolinization of illite and feldspars is accompanied by the liberation of silica in the form of silicic acid which is mostly consumed in the formation of overgrowth on quartz grains. The extent of this silicification, however, is controlled by other factors: the chemical nature of the environment created and especially the permeability of the rocks. The lower the permeability, the more intense silicification wilt be. The same origin of silica is frequently assigned to the recrystaltization of detrital micas in an alkaline environment. The recrystaltization of allogenic micas (or hydromicas), muscovite and biotite results in the liberation of SiO2 whereas the alkaline environment created by this also favours some degree of leaching of the quartz grains. In rocks of low permeability even this dissolved silica starts becoming deposited on neighbouring grain as overgrowth rims because of its low mobility in the pore space. A characteristic feature of this redeposition of secondary silica resulting from the transformation of clayey matter and feldspars is the discontinuous nature of these rims. At places the quartz grains are transformed into monolithic aggregates reminiscent of a quartzitic texture (Plates 1, 2) and one can make out welt-developed crystal faces. However, in addition to the overgrowth rims there are small pyramidal quartz crystals oriented at right angles to the surface of the grains on which they develop. Under the SEM (Plate 2) the quartz grains exhibit an elongate shape and are frequently oriented parallel to the stratification of the beds. There are also overgrowth rims with straight outlines completely enveloping the detrital quartz grains. There is even a well discernible second generation of secondary silica especially in sandstones cemented by carbonates with a poikilitic texture. This generation is characterized by particularly regular crystal faces on the quartz grains (Plate 2). The formation of the different overgrowth rims is here linked to the recrystallization of the carbonate cement
80
Chapter 3 - Main Factors of Reservoir Compaction
Plate 1. Morphological features of quartz, a, b Monolithic quartz aggregates; c progressive crystal growth and formation of growth rims on quartz; d, e discontinuous growth rims on quartz; f welldeveloped crystal faces on quartz (/eft)
Plate 2. Morphological features of quartz, a Discontinuous growth rims on quartz; b, h quartz crys- i~ tals grown and oriented parallel to stratification; c, d quartz crystals grown and oriented at right angles to the surface on which they developed; e, f several generations of growth rims on quartz grains; g, h particularly regular crystal faces on quartz which frequently are encountered in carbonate-cemented sandstone
3.3 • Reservoir Compaction by Silicification from Other Sources
81
8z
Chapter 3 . Main Factors of Reservoir Compaction
(siderite, dolomite and calcite). The mass of the carbonate cemem is enriched in free silica furnished by the corrosion of the quartz grains and of other detrital compounds as well as from the transformation of the initial argillaceous cement. In other words, during recrystallization of the carbonate on compaction (or decompaction) the calcite crystals grow individually. A poikilitic texture is formed and the primary calcic mass clears itself of impurities and of the silica, thereby supporting the mineralization and reconstitution of the crystal faces of the quartzes. This process is accompanied by an opening of the pore space and consequently the overgrowth on the quartz grains is no longer limited in volume and is able to advance until crystal faces appear. 3.3.3 Silicification by lllitization of Smectite Illitization of smectic and of mixed-layer minerals takes place under liberation of some silica. The mixed-layer minerals must have existed in large quantities in the reservoir rocks studied, a fact that is confirmed by their occurrence in younger or diagenetically less altered strata of all basins. The process started during the early phases of diagenesis and is controlled initially by the thermal regime and then by pressure. Silica liberated here is deposited on detrital quartz grains in the pore spaces, representing one of the sources of secondary silica as a result of the presence of large quantities of mixedlayer compounds in the initial argillaceous cement. The transformation of smectite into illite accompanied by the liberation of silica takes place in argillaceous strata in direct connection with the diagenesis of organic matter. The organic acids and the carbonic acid set free play an important role in the extent of mobilization and distribution of the liberated silica. Quantitative estimates show that the silica liberated during the transformation of mixed-layer components of the illite/smectic type or of smectite alone into illite may a m o u n t to 2.2 g SiO 2 per loo g of clay (Towe 1962). Clayey components account for up to 6o% of the sediments (argillaceous siltstones). Even if the entire clayey matter was composed of mixedqayer minerals and if under this hypothesis all the silica liberated was transported towards the reservoir rocks, only 1.3% of quartz or secondary silica would be deposited (Towe 1962). The development of dioctahedral illites, characterized by elongate and lamellar shapes of the newly formed crystals, could perhaps be considered as indicator of the above-mentioned process. The older authigenic itlites may be distinguished by their more elongate crystals and in particular by their rightangle orientation against the bedding planes, i.e. in the b-axis (Plate 3). The illite formed from the mixed-layer components of the illite smectite (I/S) type mostly exhibits fibrous growth (Plates 3, 4) whereas the illites formed by transformation of kaolinite are characterized by a lamellar habitus (Plate 4). On the whole, the clay minerals with an elongate crystal habitus make up a continuous sequence of mixed-layer minerals of the smectite-illite type with a dioctahedral structure between the two extremes smectite and respectively iltite of the sericite type.
Plate 3. Development of authigenic illite, a Chlorite; b-d development of elongate and lamellar crys- 1~ tals at right angles to the stratification (b-axis); e - h authigenic illite with fibrous crystal habit developed by transition from I/S mixed-layer minerals
3.3 • Reservoir Compaction by Silicification from Other Sources
83
84
Chapter 3 Main Factorsof ReservoirCompaction
3.4, Main Controls of the Compaction of Reservoir Rocks
85
3.4
Main Controls of the Compaction of Reservoir Rocks Compaction due to increasing geostatic pressure, aside from possible tectonic compression, leads to a large reduction in porosity and one can note a priori an increase in density. The loss in porosity resulting from displacement and mutual approach of the grains, from the deformation of plastic grains and from pressure solution, to some extent is compensated by the formation of secondary diagenetic porosity. Not considering fracturing, the secondary porosity is formed by the partial or total dissolution of the matrix components between the grains and other debris, as well as of the cement of the respective sandstones. Where fracture porosity is developed like, e.g., in the Hassi Messaoud field, it may account for up to 4o% of the total porosity. The secondary porosity is formed virtually all along a section and its development does not abruptly stop at a certain depth. This secondary porosity accounts for the greater part of the total porosity over the depth interval from 2.6 to 3.8 km. The modal values of secondary porosity in the Cambro-Ordovician reservoir rocks of Oued el-Mya range between 25 and 5o% of the total porosity. Secondary porosity makes up an important fraction of the total porosity even in reservoirs now lying at lesser depths. This situation is well developed in the Paleozoic reservoirs of the Ahnet-MouydirGourara basins which have been uplifted to depths of o.6-z.7 km by the Hercynian orogeny. Naturally, secondary porosity would have formed when the region considered was at greater depths than now, but it was preserved after uplift. Nevertheless, the gradients of porosity and density vary within the same region. Such variations are particularly characteristic of the Iltizi Basin and the Ahnet-Mouydir-Gourara Basins. In addition to other factors like facies types, this observation is obviously an expression of variations in geothermal gradients.
3.4.1 Thickness of Sandy and Silty Reservoir Rocks The preservation of initially porous reservoir rocks especially at greater depths is above all a function of the thickness of the succession. The greater this is, the smaller will be the effect of compaction (Fig. 3.1o). This phenomenon has been noted in numerous reservoirs studied and is encountered in different geological contexts. Furthermore, within a certain layer the porosity will increase from the top and the bottom towards its centre (Fig. 3.11). The statistical treatment of the respective data (Figs. 3.1o,3.11) shows for different geological complexes in various basins of the Saharan Platform that: 1. the relative growth rates of porosity, i.e. the preservation of the accumulation potential of originally porous reservoir rocks, as a function of depth are more sensitive than the growth rates of permeability (Fig. 3.1o). Neglecting the other factors of compaction this phenomenon may be explained by the fact that permeability depends primarily on structure and geometry of the pore space;
•~ Plate 4. Developmentof authigenic illite, a-d Illitization of I/S mixed-layerminerals. Note fibrous habitus of illite; e-h developmentof authigenic lamellar iltite crystals from kaolinite
86
Chapter 3 -
I
II
,
III
Sandstonesfrom: ----I Ahnet-Mouydir GouraraBasin(PZ) II lllizi Basin(D+C) . . . . Ill Oued-My --x--~ tV GhadamE
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Fieldofchangef0rsandstones from(PZ),Ahnet-Mouydir Guaral ~eldofdlangef0~sand~nes ~om(D+Q,IlliziBasin
Fieldof[hangefor sandstones from(PZ),Oued-MyaBasin Fieldof changeforsandstones from(D),GhadarnesBasin
Fig. 3.10. Curves of relative change of porosity and permeability with thickness of beds (based on average statistical data). I Sandstones from Ahnet-Mouydir-Gourara Basin (PZ), // sandstones from Iltizi Basin (D + C), lII sandstones from Oued el-Mya Basin (PZ), IVsandstones from Ghadames Basin (D) Distance from center to bed roof (In)
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~ // ///i /
.. , "%
.//
./ '"* J 20 Distance from center to bed bottom (in)
Fig. 3.11, Curves of relative change of porosity and permeability within sandstone beds (based on average statistical data). I Sandstones from Ahnet-Mouydir-Gnurara Basin (PZ), II sandstones from Illizi Basin (D + C), HI sandstones from Oued el-Mya Basin (PZ), I V sandstones from Ghadames Basin (D)
3.4 • Main Controls of the Compaction of Reservoir Rocks
87
2. the relative growth rates of porosity and permeability as a function of the bed thickness in the Paleozoic reservoir rocks in the Ahnet-Mouydir and Gourara basins are lower than those in the Paleozoic reservoirs of the Oued el-Mya and Ghadames Basins (Fig. 3.11). This situation is in perfect agreement with all diagenetic features of compaction observed in the various basins. It goes without saying that the anomalous thermal flux in the former group of basins may lead to a reduction in the stability of the grains and the matrix of the sandstones and to an intensification within them of the processes of chemical compaction and pressure solution.
3.4.2 Argillaceous Diagenesis and Related Abnormal Formation Pressure Not considering other factors the preservation of porous reservoir rocks at greater depths in Paleozoic and Triassic reservoir rocks is explained by the appearance of an abnormally high formation pressure which acts as an absorber reducing the effective constraints suffered by these rocks. In the Oued el-Mya and Ghadames Basins and to a lesser extent in the Illizi the Silurian, Devonian and Triassic shales contain a large amount of mixed-layer days of the smectite-illite type, accounting for up to 2o-25% in many regions. It appears that the development of an abnormally high formation pressure in the deeply buried sediments results from a rapid compaction of the silty-argillaceous and sandy rocks in geosyndinal basins (Bradley 1975; Hower et al. 1976). This is caused by the very rapid accumulation of sedimentary material followed by the burial as characteristic of the Triassic Province. At temperatures around loo °C the mixed-layer minerals of the smectite-illite type react with the potassium of the feldspars associated with them and with the interfoliar water when the illitic component increases in the mixedlayer minerals (Boles and Frank 1979; Huang et al. 1993). The detailed correlation between porosity of the sandstones, density of the associated shales, proportion of smectite and illite in the mixed-layer compounds and potassium content of the shales with depth in the Paleozoic of the Oued el-Mya Basin exhibits a concordant behaviour of these parameters (Fig. 3.1z). As already mentioned, a certain decompaction takes place at the depth of z.8-3.8 km, corresponding to mesodiagenesis, within the reservoir rocks because of the development of secondary porosity. In this interval even the density indices vary only very little and sometimes decrease only weakly, suggestive of the appearance of a zone of abnormally high formation pressures. The increase of the illitic compounds in the mixed-layer clay minerals with depth is accompanied in the shales by an increase in the K20-content coming from the feldspars and the interfoliar water. The transformation of smectite into illite liberates much water from the clay structure in a number of successive stages which depend on the rate of temperature and pressure variations, i.e. on the rate of subsidence (Burst 1976). These variations may lead to a volume increase of the silty-argillaceous horizons and to the stabilization or reduction of the density of the rocks observed at depths somewhat below 3.o km in the shallowmarine formations, in particular in the marine Devonian and Carboniferous sequences of the Illizi Basin and in the deltaic rocks of the Lower Paleozoic in the Oued el-Mya Basin. An enormous amount of water is also liberated during the diagenesis of the Triassic and Jurassic evaporites. These waters are accumulated in the adjacent permeable sandstones which prove to be under abnormally elevated formation pressure. We must understand that the excess pressure resulting from the liberation of water during
88
Chapter 3 • Main Factors of Reservoir Compaction 2.0
2.1
2.2
2.3
2.4
2.5
2.6
2.7
2.8
2.9 Density (g cm-3)
0
5
10
15
20
25
30
35
40
10
8
6
4
0
0 1(20(%)
45 Montmorillonite in interlayered (M-I) minerals (%)
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o I Reservoirporosity • 2 Adjacent shale density • 3 Montmoril[onite/illite proportion in adjacent shale 4 Potassiumcontent in shale
2.5
~
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3.5
o
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4.0-
6
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14
18 22 Poros~y(%)
26
30
34
36
Fig. 3.12. Correlation curves of 1 reservoir porosity,2 adjacent shales density,3 montmorillonitelillite proportion in adjacent shales, 4 potassium content in shales, with depth for Triassic and Palaeozoic sediments, Oued el-Mya Basin
diagenesis of the shales and the decompaction of the sandstones because of the development of secondary diagenetic porosity may be superimposed onto each other in time and space. This is the key phenomenon in understanding the migration of the hydrocarbons and their accumulation in the traps of the Saharan basins. One of the essential factors for the appearance of abnormally high formation pressures in this province lies in the fact that during the Mesozoic thick layers of sediment were deposited and buried at rates exceeding the ability of the entrapped waters to adapt to the growing overlying weight. Another factor controlling the formation of an excess pressure would be tectonic activity. The development of an abnormally high formation pressure requires an impermeable cover sequence. It is clear that this cover has been affected in several regions, but in particular in the Illizi Basin and the basins in the southwestern Saharan Platform, by fractures and other features leading to the re-establishment of normal formation pressures. This suggests that most of the zones mentioned above as exhibiting excess formation pressures will be young in age, probably belonging to the Cenozoic.
3.4 - Main Controls of the Compaction of Reservoir Rocks
89
3.4.3 Early Development of Overgrowth Rims One of the main factors counteracting compaction in the quartz sandstones is the formation, prior to the important Mesozoic subsidence of the sediments, of a quartzose matrix up to incomplete occlusion in the form of overgrowth rims. These quartz rims represent a solid matrix opposing further consolidation of the rocks at great depth. This phenomenon is one of the main factors facilitating the conservation of good reservoir properties at greater depth in all basins studied and especially at Oued el-Mya and Ghadames. The Hassi Messaoud field is a well-known example in which, in addition to leaching initiated during the Hercynian orogeny, the good reservoir properties result from the early formation of quartz overgrowth rims up to incomplete occlusion prior to the Mesozoic subsidence. These rims strongly countered the subsequent compaction of the reservoirs. The other geological factors remaining unchanged, early diagenetic formation of a solid skeleton may be recommended as a criterion in the search for reservoir rocks little compacted at greater depths in the Saharan basins.
3.4.4 Presence of Mesozoic Evaporites The Saharan Platform is a geological structure with a double-level setup: sedimentary formations of the Paleozoic and Mesozoic separated from each other by a long interruption in sedimentation and a deep erosion in connection with the Hercynian orogeny, each level possessing its own rules for compaction and decompaction of the reservoirs. The Mesozoic evaporitic formations, and especially the salts like halite, play a double role in the geological history of the province. At first, these formations create a magnificent regional cover for the reservoirs of the Lower Triassic and sometimes of the Paleozoic. Then the weak thermal isolation lowers the heating of the sediments and the degree of diagenetic transformations and compaction, i.e. they contribute to the preservation of excellent properties in the underlying reservoirs. The sediments underlying the salt-bearing formations are continental arenaceous-argillaceous sediments, carbonates and sulfates. The reservoirs in the Triassic sandstones, in particular in the lower part of the succession, possess good reservoir characteristics which, in addition to other geological factors, may be attributed to the low degree of heating and to the moderate catagenesis of the sediments as a result of the weak isolation by the overlying thick beds of Triassic and Jurassic salts. Finally, the relatively moderate compaction of the Triassic and sometimes of the Paleozoic reservoirs is attributed, in addition to other factors, to the low density of the saks and to the presence of a low geostatic load. This circumstance may be considered as one of the reasons explaining the existence of Devonian and Cambro-Ordovician sandstones with a porosity in the range of 15-2o% at a depth of 3.4-4.z km in the Triassic Province. Such porosities are encountered outside the Triassic Province only at depths of 2.z-2.6 kin. In addition to the favourable role played by the salt-bearing formations, the lowering of the degree of compaction in the reservoirs may also be ascribed to abnormally high formation pressure, to the development of secondary porosity and to the formation of overgrowth rims on quartz prior to the Mesozoic subsidence.
90
Chapter 3 • Main Factors of Reservoir Compaction
3.4.5 Temperature and Pressure Studies of the evolution of temperature and pressure gradient with depth allow us to appreciate the role of these two factors. The geothermal gradient of the Saharan basins ranges from 0.7 °C / loo m in the Triassic sediments to 3.5-6.0 °C / lOO m in the Paleozoic rocks, in particular in the basins of the south and southwest. The gradient is generally moderate in the Oued el-Mya and Ghadames Depressions at 1.8-3.4 °C / lOO m, but rather variable in the Illizi Basin with 2.0-4.2 °C / lOO m. The temperature gradient, however, fluctuates considerably within the same region. The lowest gradient is encountered in the salt-bearing parts of the section. The value itself is a function of the thermal conductivity and of the rate of thermal flux. It is obvious that the thermobaric fluctuations depend on the variations of the other physical properties of the rocks with depth. The Cambro-Ordovician reservoir rocks of the Oued el-/vlya Basin with a geothermal gradient of 3.8-4.5 °C / loo m encompass medium-grained reservoir rocks with a porosity of 8-1o%. In the Ghadames Basin, however, the geothermal gradient in the Devonian sandstones is 2.8-3.1 °C / loo m and their porosity about 16%. A gradient of 1.9 °C / lOO m is associated with a porosity of 22% in the Upper Paleozoic sandstones of the same basin. It has to be kept in mind that the spontaneous modification of the pore space by thermal expansion or by compression of the rock matrix will not change the porosity index to a great extent. However, the indirect influence of the temperature on the porosity manifests itself by the velocity of the chemical reactions leading to the compaction of the rocks by cementation. The pressure exerted on the rocks will alter their physical properties, its influence varying with type and structure of the rock in question. There is always a limitation for the compressibility of a rock at the fracturing limit of its mineral components, but in the Saharan basins we very rarely observe crushed grains even at depths below 4 kin. This is explained by the fact that a certain portion of the purely physical transformation is replaced essentially by chemical processes especially during pressure solution.
3.4.6 Authigenic Transformation of Argillaceous Cement The clay minerals making up the cement of the reservoir rocks on the Saharan Platform are essentially made up of authigenic forms of kaolinite, illite and chlorite. Cementation is porous and of the contact-type and rarely film-like. Our SEM studies have shown that sandstones cemented essentially by epigenetic kaolinite are characterized by a better communication between the pores than sandstones cemented by fibrous illite. This peculiarity becomes even more evident when we compare them to sandstones in which the cement includes modest quantities of illite, detrital chlorite or mixed-layer minerals of the iltite-smectite type which reduce the communication between the intergranular pores even more. The filtration properties of reservoir rocks are thus directly related to the mineralogical composition of the argillaceous cement (Fig. 3.13). However, the authigenic nature of the main clay minerals in the reservoir rocks show that composition, structure and distribution of the sedimentary argillaceous cement have been subjected, especially in highly permeable reservoir rocks, to important postsedimentary transformations. These transformations are controlled by the initial char-
3.4 • Main Controls of the Compaction of Reservoir Rocks
91
acteristics of the reservoir and in turn they themselves exert a great influence on the latter. This is explained to a great extent by the physical and chemical properties of the reservoir rocks, by the dimensions and the arrangement of the clay minerals in the pore space and by the type of cementation experienced by the reservoir rocks. We have noted during our detailed studies of these factors in various sections of the basins on the Saharan Platform that in a number of cases, despite similar or identical values of depths, grain size and the amount of argillaceous cement (in the absence of carbonate cement), the reservoir characteristics can be quite different, with permeability ranging from a few dozen millidarcy to several hundred. Considering that the factors mentioned are here more or less constant the reasons for these differences can only be mineralogy, texture and structure as well as nature of the argillaceous cement. In the cements of the reservoir rocks studied here the altogenic and diagenetic clay minerals (mixedlayer, illite, chlorite and kaolinite) frequently are only a few fractions of microns in size, the particles themselves being isometric, sometimes lamellar elongate but also exhibiting poorly defined shapes with vague outlines. The latter, together with flat grains of other minerals like feldspars, biotite, etc., are arranged in the pore space of the sandy-silty materials mainly in an ordered fashion parallel to the stratification, thereby leading to relatively regular homogenous cementation. It results in an abrupt drop in permeability of the reservoir rocks especially at right angles to the stratification, the permeability in this direction being 5o-65% of that parallel to the bedding. The structure of the allogenic and partly authigenic minerals is generally irregular and the type of cementation may be porous, of the contact-type or mottled. At the same time diagenetic clay minerals appear in the coarse-grained reservoir rocks. They may be distinguished by their relatively perfect shapes and structures, the crystals lying frequently in the pore space in an unordered fashion. The cement may be porous, film-like or of the contact-type. In some cases it leads to an increase of the permeability of the reservoir rocks. The diagenetic kaolinite possesses a rather low sorption capacity (3-5 mg eq-1 per loo g of rock), is of large size (up to lO gm or more) and exhibits individual or aggregated particles with smooth or flat faces arranged in ordered fashion within the pore space of the reservoir rocks (Plate 5). To some extent the kaolinite particles play the same role as the fine-grained silty particles. The replacement of dispersed detrital clay minerals by kaolinite with a higher sorption capacity and an ordered arrangement in the pore space leads to an increase in rock permeability (Fig. 3.13c) even if its proportion is equal to or slightly above that of the allogenic argillaceous cement. An abrupt increase in the content of epigenetic kaolinite in sandstones containing little detrital argillaceous cement (Fig. 3.13c) can, however, lead to a deterioration of the reservoir qualities of the respective rocks. On the whole, the formation of kaolinite at the expense of allogenic argillaceous cement frees part of the pore space and thereby supports the formation of porosity within the cement. This porosity proves to be sufficiently effective in view of the weak absorption capacity of diagenetic perfectly ordered kaolinite. Cements made of anthigenic illite are characterized by a higher arrangement density of the individual particles which do not occur in such large particles as kaolinite and chlorite. It has to be remembered that, on the whole, elongate particles with a micaceous structure frequently containing a certain amount of swelling layers are characterized by a higher dispersion than chlorite and kaolinite. However, the rigidity of the micaceous mineral particles does not lead to a structure of the pore as complicated as in the case of chlorite. As the micaceous species sometimes contain important quan-
9~
Chapter 3 • Main Factors of Reservoir Compaction 30 ~ - ' ' I I, II Oil-(gas-)bearing reservoirs
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Fig. 3 . 1 3 b . Impact of diagenetic mineral cement transformation on reservoir quality, Saharan Basins. A Diagram illustrating the impact of the illitization on reservoir quality. B Porosity and C permeability plots against illite content in clay cement in reservoir rocks from Hassi Messaoud (*) and Hassi R'Mel (0) fields. Reservoir rocks involved in the correlations were selected a manner, that their other petrographic features are similar or very near, e.g. average grain size M.D. = o.2-o.4 ram, clay cement content = 10-15% (of total mass of the rock), with very little or no carbonate cement
3.4 - Main Controls o f the Compaction o f Reservoir Rocks ,
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Fig. 3.13c. Impact of diagenetic mineral cement transformation on reservoir quality, Saharan Basins. A Diagran~ illustrating the impact of the kaolinization on reservoir quality. B Porosity and C permeability plots against authigenic kaolinite content in clay cement in reservoir rocks from Oued el-Mya (~) and Ahnet-Mouydir-Gourara (o) basins. Reservoir rocks involved in the correlations were selected a manner, that their other petrographic features are similar or verynear, e.g. average grain size M.D. = 0.2-0.4 mm, clay cement content = m-15% (of total mass of the rock), with very little or no carbonate cement
30
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Fig. 3.13d. Impact of diagenetic mineral cement transformation on reservoir quality, Saharan Basins. A Diagram illustrating the impact of the chloritization on reservoir quality. B Porosity and C permeability plots against authigenic chlorite content in day cement in reservoir rocks from Oued el-Mya (e) and Illizi-Ghadames (0) basins. Reservoir rocks involved in the correlations were selected a manner, that their other petrographic features are similar or very near, e.g. average grain size M.D. = 0.2-0.4 mm, clay cement content = lO-15% (of total mass of the rock), with very little or no carbonate cement
94
Chapter 3 . Main Factors of Reservoir Compaction
3.4, Main Controls of the Compaction of Reservoir Rocks
95
titles of swelling layers they possess interfoliar water in addition to adsorbed water and crystal water. The elevated hydrophilic nature of the micaceous minerals will, as we know, exercise a large influence on porosity and permeability of the reservoir rocks. The widespread process of illitization characteristic of the late diagenesis of the reservoirs of the Saharan Platform reduces the adsorption potential and the ion exchange cap acity of the rocks compared to their primary capacity which in turn lowers the water retention potential of the rocks thereby increasing the effective pore volume. Furthermore, illitization of the detrital argillaceous material and in particular of the mixedlayer mineral and even of the micaceous debris results in secondary microporosity within the cement and thereby leads to an increase in the porosity of the reservoir rocks. However, this newly formed illite with its characteristic reticular structure notably disturbs the permeability of the reservoir rocks due to growth of the individual grains (Fig. 3a3b). This is caused by the high dispersion ability of "fibrous" illite frequently blocking the small pores and the channels between the pores, in particular during entrainment of hydrocarbons. As we know, the process of illiti-zation takes place in an alkaline environment leading to corrosion and partial dissolution of the quartz grains which concurrently results in a certain increase in the intergranular pore space. Chlorite in the reservoir rocks of the Saharan Platform is of rather diverse origin, composition and structure, influencing the reservoir characteristics in two ways. The authigenic well-crystallized chlorite with large particles (up to 5 pro) has a sorption capacity of lo-4o mg eq-' per loo g of rock. When this forms in the pore space of the rocks by replacement of smectite and mixed-layer compounds with a large sorptive and swelling capacity the volume of the space and communication between the pores will grow. On the other hand, when it assumes the form of films and incrusted fringes on quartz grains in the sandy-silty rocks the chlorite reduces the section area of the interpore channels and diminishes the permeability of the respective rocks (Fig. 3a3d). Lamellar allogenic chlorite of small grain size (o.2-o.5 pro) together with the other allogenic minerals of the primary cement reduce the permeability of the reservoir rocks. To sum up, the diagenesis of the reservoirs studied has entailed a pronounced transformation of the latter under participation of mineralized subterraneous waters. Of the primary allogenic argillaceous cement only negligible quantities are preserved. The transformation of the cement was particularly important in coarse- to medium-grained sandstones possessing originally a higher permeability facilitating the circulation of chemically active underground waters. Subsidence of the sediments favoured their compaction as well as silicification and authigenic transformation of the argillaceous cement. These processes caused a radical modification of the mineral composition as well as of the structure and texture of the rocks. As a consequence, in numerous zones of the succession a secondary microporosity developed which to some extent compensated for the negative influences of compaction and silicification on the accumulation properties. The influence exerted by the processes mentioned on the filtration properties is rather variable. Whereas kaolinization will improve the permeability of the reservoir rocks, the formation of fibrous illite leads to the opposite result. In each
Plate 5. Morphologicalfeaturesof kaolinite(dickite) and chlorite,a, b Kaolinite;c, d dickitewith welldevelopedcrystal facesand of large size (up to lO pm and above);¢-h authigenic chlorites developed in the pore space of sandstones frequentlyin the form of films or as encrusted fringes around quartz grains. Note that photo g exhibits two chlorite varieties of different origin
96
Chapter 3 • Main Factors of Reservoir Compaction
case considered, one should investigate each of these processes separately as well as the extent of all transformations on the whole. Such an assessment can only be based on detailed mineralogical, lithological and petrographic analyses.
3.4.7 Early Invasion of the Reservoirs by Hydrocarbons We have dearly established the great difference in extent of the diagenetic transformations between the productive (oil-saturated) and non-productive (formation water-saturated) horizons in all basins of the Saharan Platform studied. These differences are especially manifest in the extent of the mineral transformations and in the degree of structural perfection of the neMy formed minerals. Furthermore, in non-productive water-saturated horizons the amount of authigenic and lamellar-elongate or fibrous illite as the most stable mineral phase exceeds that of secondary authigenic kaolinite. However, in horizons filled by hydrocarbons this trend is not observed. It appears that in this case illitization of detrital argillaceous material and of authigenic kaolinite had started but did not continue to the end. SEM-investigations have revealed that authigenic elongate and lamellar illite crystals developed at the edges of detrital micas or secondary kaolinite whereas a large argillaceous mass remains intact. In contrast to this, nearly the entire detrital argillaceous mass of the cement has been transformed in the water-saturated strata into secondary forms mostly of illite and to a lesser degree of chlorite. The regularity mentioned for illite could not be observed for kaolinite. Authigenic kaolinite is encountered in horizons filled by hydrocarbons as well as in those saturated with formation waters. The explanation for this is that kaolinite is a less stable argillaceous phase than elongate or lamellar illite (Lanson 1996). Quartz overgrowth rims of the first two generations are developed virtually in all horizons whatever their type of saturation. The third generation with the overgrowth rims controlled essentially by secondary silica derived from the transformation of the argillaceous cement and from other sources dominates more clearly in water-saturated horizons than in the horizons filled with hydrocarbons. It is quite obvious that the differences in the extent of the diagenetic transformations are a function of their duration. The arrival of oil or gas in the reservoirs prominently diminished the velocity of all reactions in the aqueous environment due to the protective action of the hydrocarbons. This is due to a loss of the solution potential of the water and the stop or slow-down of the translation movements of the water molecules in the pore space occupied by hydrocarbon molecules. The protective action exerted by the organic molecules makes itself felt mainly by ion exchange reactions during which the hydrocarbon molecules occupy the exchangeable positions on the surface of the clay minerals where they are fixed solidly not only by polarization forces but also by the so-called Van-der-Waats forces which are more intense in the case of larger organic ions, This was also controlled by preferential sorption of organic ions from aqueous solutions on the clay minerals. The main factor for interaction between the clay minerals and organic ions is the sorption of the latter on the surface of the clay minerals and in particular on those parts of the surfaces where there is a deficit in positive charges because of structural defects. The structural defects result from the increase in dispersion of the minerals, from hydratization of their surfaces and from isomorphous substitution of the main cations in the crystal lattice by ions of a lower valence.
3.5 . Silica Solid Phase Transformation: A New Concept for Sandstone Diagenesis
97
During the neutralization of the active centres on the surface of the clay minerals by organic molecules, the concentration of the elements in the aqueous environment of the pores drops. Consequently, the difference between the intensity of the diagenetic secondary transformations within productive horizons filled by hydrocarbons and the intensity of those in non-productive water-saturated strata allows us to determine timing and duration of the formation of the deposits in the Saharan basins. Our quantitative comparisons of the impact of authigenic kaolinization, chloritization, illitization and silicification on the characteristics of reservoirs filled by hydrocarbons or saturated with water in the Sahara are presented in Fig. 3.13a-d. In preparing this illustration we also took into account other geological and petrological factors so that the correlations are made for reservoir rocks with similar or identical parameters. The arrival of the hydrocarbons in the reservoir rocks slowed down these reactions and the hydrocarbons themselves also modified the petrophysical characteristics of the environment, in particular in the contact zone with water. This phenomenon contributed to the decompaction of the reservoir rocks by dissolution in the transition zones, of cementing materials and of the matrix components by the products of incomplete oxidation and transformation of hydrocarbons like, e.g., the transformation of an oil deposit into a gas-condensate system. 3.5
Silica Solid Phase Transformation: A New Concept for Sandstone Diagenesis as Revealed in North African Basins This section is written together with Mrs I.I. Plyusnina.
In spite of considerable advances in sandstone diagenesis research in the last three decades, several questions remain unanswered. One of the most important problems is the silica budget in the course of sandstone diagenesis. The quality of oil and gas clastic reservoirs is much controlled by quartz cementation. New obvious petroleum seeps and less obvious subsurface structural traps and reservoirs are increasingly difficult to find. Subtle diagenetically controlled stratigraphic and lithologic traps and reservoirs are progressively becoming the prospects of the future. In this respect, oil and gas prospecting and development in North Africa are no exception. Many mechanisms have been suggested to explain quartz cementation processes in sandstones. Amongst these, pressure solution has been the focus of considerable attention as a principal process for supplying silica for quartz cementation (e.g. Taylor 195o; Heald 1956; Thompson 1959; Wey11959; Houseknecht 1984,1988). These observations ted to the elaboration of theoretical geochemical models of intergranutar pressure solution and a limited number of experimental simulations (Ernst and Blatt 1964; Heald and Renton 1966; Renton et al. 1969). There is still no consensus as to the importance of pressure solution. Some researchers insist that pressure solution considerably reduces the porosity of sandstone (e.g. Houseknecht 1984, 1988) whereas others conclude it is of limited importance (e.g. Sibley and Blatt 1976; Blatt 1979). The idea that clay mineral diagenesis and particularly the illitization of smectite might be an important source of authigenic silica have been previously discussed (Siever 196z; Towe 1962; Keller 1963; Fuchtbauer 1967; Dunoyer De Segonzac et al. 1968; Burst 1969; Perry and Hower 197o; Hower and Eslinger 1973; Schmidt 1973; Yeh and Savin 1973; McBride 1989). Hower et al. (1976) and Blatt (1979) also confirmed that dis-
98
Chapter 3 • Main Factors of Reservoir Compaction
solved silica could be a product of clay mineral diagenesis, but they consider it is possible that silica might remain undetected in shales and that this silica therefore can not be a major source of cementation in adjacent sandstones. Subsurface waters have been considered also as a significant source of authigenic silica for cementation (Davis 1964; Sibley and Blatt 1976; Blatt 1979) if enough water can be moved through the rock. According to Blatt (1979) cementation by precipitation of quartz from vertically circulating groundwaters could be efficient when sands are located close to the surface so that flow distances are minimized. Dissolution of feldspars and accessory silicate minerals during diagenesis could be another significant source of silica. Fothergill (1955) and Hawkins (1978) reported a statistical correlation between the occurrence of diagenetically kaolinized feldspars and quartz overgrowths. These phenomena, however, are generally believed not to be significant sources of silica (Blatt 1979). It is clear that none of these existing mechanisms fully and unequivocally accounts for the amount of diagenetic silica found in sandstones as quartz overgrowth. Each of the reported mechanisms might explain some quantity of the observed silica in a given sandstone and the relative merits of silica sources vary from one bed or region to another. Which particular mechanisms apply depends on specific geological (tectonics, facies, hydraulics) and geochemical conditions. Whatever the sources of silica, it is too difficult, even impossible, to account for the diagenetic silica budget observed in many sandstones. In Saharan basins many sandstones contain 7-2o% authigenic silica, and some very heavily silicified Cambrian and Ordovician some sandstones contain 20-35% of quartz overgrowth, without significant evidence of intergranular pressure solution, It would be difficult to justify the silica budget in such cases even if one could suggest that one or more of the recognized mechanisms producing authigenic silica have acted at their maximum. Something must, therefore, be missing in silica budget in sandstones. To date, there have been few experimental studies of the kinetic factors involved in quartz nucleation and we can only speculate about the role of authigenic silica produced by one or more of the existing mechanisms. For this reason we propose a new mechanism for quartz authigenesis and sandstone compaction via solid phase transformation processes even under moderate temperatures and pressure over relatively long periods of time in open geological systems. In this proposed model the aggregation of the grains occurs through diffusion, sliding, screw dislocation and face-to-face movements of big angle boundaries, these processes being irreversible. The impetus behind these transformations is the decrease in surface energy, and in open geological systems such processes of self-organization could justify in most cases the sandstone consolidation observed in the sedimentary cover. The aim behind development of this concept is to contribute to a better understanding of the diagenetic evolution of sandstone petroleum reservoirs.
3.5.1 Samples This study of silica diagenetic evolution in sandstones is based on several hundreds of core samples selected from various boreholes in some nine North African basins: the Oued el-Mya, Ghadames, Illizi, Triassic, Timimoune, Ahnet, Mouydir depressions and Amguid-Hassi Messaoud and Idjeran-M'Zab anticlinal systems (Fig. 3.14). Stud-
3.5 . Silica Solid Phase Transformation: A New Concept for Sandstone Diagenesis
99
ied samples represent all sandstones present, particularly those from oil and gas producing reservoirs from the Cambrian to the Triassic with the exception of the Perm which is absent in the region. Samples cover a range of subsurface depths from 15o m (Paleozoic sediments highly uplifted during the Hercynian Orogeny) to 5 7o0 m (Lower Paleozoic sediments deeply buried during the Mesozoic). The samples belong to reservoirs of various mineralogical and textural maturities with different facies origin. Cambrian sandstones are quartzitic and comprise thick (6oo-1 ooo m in the east, center and west of the Sahara), cross-bedded, non-fossiliferous, probably fluviatile sandstones with a porosity range of 2-2o%. Lower and Middle Ordovician shallow marine interbedded sandstones commonly range from loo to 7oo m in thickness and cover approximately one-half of the Sahara Plate area. Upper Ordovician (subangular to subrounded) quartzitic sandstones of glacial origin form are found in irregular (3o-45o m) patches across the province due to depositional as well as to Hercynian erosional patterns. Porosity ranges from o.5 to 12%. Silurian sandstones are localized, marine, generally fine grained, locally quartzitic and interbedded with siltstones and dark gray shales. Porosity ranges from o.5 to ~3%. Lower Devonian sandstones are shallow marine and continental, plant-bearing, interbedded with dark gray silts and shales, range from fine to coarse grained, are sometimes argillaceous or silt)" and are cemented with quartz cement. They are one of the most widely deposited sandstones in the study area. Porosity ranges from 2 to 24%, and permeability from lo to 5oo roD. Middle and Upper Devonian sandstones are
Boundary of highly ordered structures (synclinorium, anitclinorium...) Boundaryof first ordered structures (domes,depressions,,.)
Fig. 3.14.
Precambrian basement
Location map of principal Saharan basins and the fields studied as part of this project
loo
Chapter 3 • Main Factors o f Reservoir Compaction
mainly of marine and lesser lithorat facies and are fine to coarse grained, subrounded, sometimes friable, interbedded with siltstones, limestones and claystones, generally subangular, with quartz cement and locally varying from argillaceous to quartzitic. Porosity ranges from 2 to 22°~,permeability from 20 mD to 7 darcies. Lower and Middle Carboniferous sandstones were deposited in a shallow marine environment and are areally restricted to the east, center and west of the Sahara. These sandstones are thin bedded, fine to medium grained, varying from friable to quartzitic, with angular to subrounded grains and with siliceous to calcareous cement. Porosity is generally in the range of lO-18%. Late Carboniferous sandstones formed in a regressive shallow marine environment, and are distinguished by calcareous, clay and silica cements. They are very fine to medium grained, interbedded with dark shales and with thin limestones. Porosity ranges from 1o to 20%. Triassic sediments are deposited only in the eastern and northern parts of the Sahara, in the Triassic Province. Triassic sandstones are mainly a combination of deltaic to prodeltaic facies represented by upper deltaic fluvial system deposits, lower distributary channel rivers, mouth bars, beach sands and tidal fiat and prodeltaic shales. Typical marine Triassic shales are restricted in extent in the study area. Triassic Lower,Middle and Upper sandstones are stratigraphically separated by shale intervals, but occasionally also by carbonates or volcanic interbeds. They are often medium grained, and range from friable to cemented with clay mineral, carbonate, anhydrite and siliceous cement. Porosity ranges from 12 to 22%. With respect to mineralogical composition and maturity, Triassic sandstones are mature sublitharenite (Lower Triassic), sublitharenite to submature subarkose (Middle Triassic) and submature subarkose (Upper Triassic). 3.5.2 Methods
Thin section analysis, including point counting (3oo counts per thin section), provided information on mineralogy, structure and porosity. Thin sections were impregnated with blue resin to distinguish porosity and stained with sodium cobaltinitrate to differentiate high potassium feldspar, Alizarine Red S for calcite and potassium ferricyanide for ferrous carbonate. Air porosity and permeability measurements were made on most samples. These data were used for correlations, and other classic petro-physical measurements (including using mercury and inert gases) were systematically applied to all samples. Some of the porosity and permeability data are provided in diagrams. Cathodoluminescence (CL) petrography was widely applied to sandstone samples representing all types of reservoirs with respect to stratigraphy, facies, depth and geographic distribution. CL combined with standard optical microscopy has permitted quantitative determination of quartz overgrowths. Examples of sutured contacts, concavo-convex boundaries, long boundaries, all standard criteria of pressure solution, are revealed to result from optically continuous quartz overgrowths (Sippet 1968). Moreover, it was possible to distinguish up to four different generations of quartz overgrowths around the same detrital grain of quartz. Scanning electron microscopy (SEM) was the principal tool used to reveal quartz crystallographic fabric and features such as twinning, dislocations, interstices, junctions, discontinuities, convergences and overgrowths. Crystallographic determinations of facets were fullfilled by goniometry. X-ray diffraction was applied in order to iden-
3.5 • Silica Solid Phase Transformation: A New Concept for Sandstone Diagenesis
1Oi
tify and estimate authigenic mineral types, abundances and structural features. Infrared spectroscopy, as a sensitive method to determine the structural order of minerals, was used to reveal significant modifications of quartz doublets in the 8oo-78o cm -1 frequency range. Silica concentrations in pore waters particularly and also in other subsurface and surface waters were determined in dozens of samples as part of this study, using atomic absorption spectrophotometry as well as inductive coupled plasma. This was essential for silica budget estimations. Present-day and paleothermal profiles were obtained from measured temperatures in boreholes, vitrinite reflectance (Ro) in interlayeredwith-sandstones shales and from basins thermal modeling. 3.5.3 Results 3.5.3.1 General Thermal Data Geothermal gradients in the Saharan basins vary from 0.7 °C / loo m for the Triassic sediments to 3.5-6.0 °C / lOO m for the Paleozoic sediments. In Oued el-Mya a n d Ghadames Basins gradients are generally moderate at 1.8-3.4 °C / loo m, but more variable ( a . o - 4 . a °C / l o o m) in the Illizi, Ahnet, M o u y d i r a n d T i m i m o u n e Basins. However, within any particular basin, temperature gradients change fairly evenly. In the Oued el-Mya Basin paleotemperature distribution is between too and 125 °C at the base of Ordovician, and between 8o and 11o °C at the base of Silurian. For Ghadames and Illizi Basins, these paleotemperatures are from 237 to 187 °C at the base of Ordovician, 125 to 163 °C at the base of the Lower Devonian, lo7 to 148 °C at the base of the Middle and Upper Devonian and 89 to 128 °C at the base of the Carboniferous respectively. Lower Paleozoic sediments in the Ahnet, Mouydir, Reggane, Tindouf and North Timimoune Basins are characterized by a high level of diagenesis (Ro = 1.2-1.6%). In deeply b u r i e d areas in the T i n d o u f and Reggane Basins, Ro reaches 2-3% in the C a m b r o - O r d o v i c i a n sediments. On the o t h e r hand, the Sbaa s u b - b a s i n (South Timimoune Basin) is distinguished by relatively moderate thermal gradients and Ro is somewhat ~o.9-Lo% in the Silurian and Devonian sediments.
Plate 6. (Next pages) Quartz overgrowth features viewed under cathodoluminescence (a, b, c, d) and I~ in plane polarized light (a', b', c', d'). All scale bars are o.1 ram. All photomicrographs (a, b, c, d) show little or no evidence of pressure solution since apperent sutured contacts, concavo-convex boundaries and long boundaries are seen to result from optically"continuous quartz overgrowhts,a, a' Detrital quartz (DQ) grains are simply "floating" in a mass of authigenic silica. Early Ordovician sandstone, well BH-3m,depth 1563.4 m. b, b' Individualquartz grains (right part of the photomicrograph) showing authigenic overgrowths equal in volume to the original detritaI grains or more. The precipitation oflate authigenic kaolinite (K) (brightly blue luminescing area) impeded quartz overgrowth. CambroOrdovician sandstone, well GS-7,depth 3252 m. c, c', d, d' Quartz grains showing several (two, three or more) distinctive successive generations of quartz overgrowths. These different generations are not visible in plane polarized light (c', d'). The fractures in quartz detritat grains are filled with authigenic overgrowths. Individual detrital quartz grains here are also doubled (in volume) by quartz overgrowths (rhombohedral crystal in the right part oft, c'. c, c' EarlyOdoviciansandstone, wellSAF-1,depth 2 884.7 m; d, d' Early Triassic sandstone, well OEM-1, depth 38z6.4 m
loz
Chapter 3 • Main Factors of Reservoir Compaction
3.5 - Silica Solid Phase Transformation: A New Concept for Sandstone Diagenesis
lo3
io4
Chapter 3 • Main Factors of Reservoir Compaction
3.5.3.2
Petrographic and Petrophysic Features Authigenic quartz is present in all the sandstones studied in the form of quartz overgrowths. Quartz overgrowths are generally recognized in thin section by distinct "dust" rims, the scarcity of inclusions, and euhedral crystalline enlargements in the pore spaces. The common absence of "dust" rims makes the quantification of authigenic quartz difficult, but it was possible in most thin sections to find areas where the boundary between detrital grains and secondary enlargements was sufficiently distinct to minimize microscopic quantification error. However, cathodoluminescence (CL) microscopy was largely applied to provide more accurate quantitative estimation of quartz overgrowths. In addition, cathodoluminescence techniques were used to estimate the contribution of pressure solution-precipitation to sandstone cementation. Plates 6 and 7 show some examples of Sahara sandstones (with different amounts of quartz overgrowths) viewed in cathodoluminesence and in plane polarized light. Some distinguishing features can be determined: 1. In Plate 6a,a',b,b', there is little or no interpenetration of detrital quartz grains, they are simply "floating" and surrounded by quartz overgrowths. This can be seen in thin sections as well as in individual sandstone (quartzite) layers. A few small isolated poresqnartz overgrowths have been inhibited and pores remain filled by authigenic kaolinite (blue luminescence). This occurs in various stratigraphic units (at least in Paleozoic sediments) and at depths exceeding 1500 m. It is clear that here the contribution of pressure solution to sandstone quartz cementation is insignificant and none of the recognized mechanisms, nor even several of them, could account for all the authigenic silica present. 2. Some individual overgrown quartz grains are doubled in volume (Plate 6b,b' right part and Plate 6c,c' right part). Several (three or more) distinctive generations of authigenic quartz overgrowths can be distinguished (Plate 6c,d). This suggests differing geochemical and geological environments related to different stages of authigenic quartz formation. It could also be related to precipitation of different sources of silica, as well as to inclusions trapped under diverse conditions. It is not inconceivable that the various hues of cathodoluminescence showing overgrowth zonation are likely due to the degree of ordering of the lattice rather than to silica origin or to trace amounts of activators. 3. Plate 7a,a' shows a moderately lithified Lower Devonian sandstone. This stratum was uplifted during the Hercynian Orogeny and subjected to intensive erosion during the Upper Carboniferous and the Perm so almost all the early clay, feldspar and carbonate cements were leached. This epigenic dissolution of cement is extensively developed and was a major process in creating secondary porosity and the reservoir properties in the Saharan basins (Makhous et al. 1994). However, though Early Devonian sandstones still remained elevated in this area at very shallow depths (240-200 m) from the Triassic up to present, there is little evidence of considerable silica precipitation (in near-surface conditions) resulting from waters circulating in these porous and permeable sandstones over a long time, as suggested by Sibley and Blatt (1976) and Blatt (1979). The marked silicification
3.5 - Silica Solid Phase Transformation: A New Concept for Sandstone Diagenesis
lO5
is nonuniform and rather local and seems to be due to other mechanisms. The role of clay or carbonate cements is obvious in inhibiting quartz overgrowths (Plate 6b,b' and Plate 7b,b'). These figures show how the filling of pore space by late carbonate or by late clay cement has stopped further authigenic quartz overgrowths. 4. Sprunt et al. (1978) have studied the cathodoluminescence of two sequences of naturally deformed quartzites. They found that metamorphism appears to homogenize the luminescences and that the "color" is related to the metamorphic grade, with low temperature causing red luminescence and high temperature causing blue luminescence. They also stated that strain is a factor in the cathodoluminescence of quartz. Zinkernagel (1978) presented an extensive discussion of quartz cathodoluminescence features. He defines three types of quartz, based on cathodoluminescence. They are: violet-luminescing quartz from plutonic, volcanic, and contact metamorphic rocks, rocks which have undergone fast cooling; • brown-luminescing quartz from low-grade metamorphic rocks, and from highgrade metamorphic rocks which have been slowly cooled; non-luminescing authigenic quartz. •
•
The majority of the Saharan sandstones we have studied show a predominance of brown-luminescing quartz overgrowths. This fact, in the light of the above results from Sprunt et al. (1978) and Zinkernagel (1978), could be considered as a supplementary argument in favor of moderate quartz diagenesis and consequently little pressure solution in the studied Saharan sandstones. The quantity of quartz overgrowths depends to a certain extent on the lithology and varies from one stratum to another. However, in terms of total anthigenic quartz content, and especially the quantity of completely silicified horizons, the various oil fields are substantially different. They are mainly divisible into four groups: (1) poorly silicified, most horizons containing 2-5% authigenic quartz; (z) moderately silicified, most horizons containing 7-1o% anthigenic quartz; (3) heavily silicified, most horizons containing 12-15% authigenic quartz; (4) very heavily siticified, most horizons containing 17-2o% anthigenic quartz. The younger, and less deeply buried sandstones generally have less quartz cement than older Cambrian and Ordovician sandstones (Plate 8c,d). There is also a correlation between the character of intergrain contacts and the grain-size of the rock: the larger the grain size, the greater their "interpenetration" (Plate 8a,b). When clay cement is present in sandstones, quartz overgrowth cementation is restricted and distorted crystal forms result (Plate 8; compare e, f and g, h). Figure 3.15a,b,c shows the correlation between "percent authigenic quartz" and porosity (a), permeability (b) and clay content (c). The data are from reservoirs with mean grain size (Din = 0.2-0.5 mm) and similar argillaceous cement contents (5-15%), with little or no carbonate cement. In addition, Fig. 3.15d shows a correlation between the porosity and permeability of Sahara Plate sandstones and the mineralogical and textural maturity of reservoirs with different facies origins. Figure 3.15d dearly dem-
106
Chapter 3 - Main Factors of Reservoir Compaction
3.5 . Silica Solid Phase Transformation: A New Concept for Sandstone Diagenesis
m7
Plate 7. Role of early depositional and late authigenic cement in inhibiting quartz overgrowths, viewed under cathodoluminescence (a, b) and in plane polarized light (a', b'). All scale bars are o.l ram. c c Carbonate cement, a, a' A moderate, nonuniform authigenic siliceous cementation of a Lower Devonian sandstone uplifted during Hercynian Orogeny (Late Paleozoic) and still elevated at 2 4 5 . 4 m depth up to present. There is evidence of early cement (carbonate, days, feldspars) and epigenic leaching, probably in near surface conditions. The sandstone is undercompacted in consequence of the "buffer" effect of early cement preventing quartz overgrowths until this cement was leached in the Late Paleozoic, well MDqoL b, b' Late carbonate cement filling pore space and stopping further overgrowth developement. Early Ordovician sandstone, well SAF-z, depth 2 8 8 6 . 6 m
onstrates that porosity decreases with higher sandstone compositional maturity, i.e. with higher content of total and, in particular, authigenic quartz as a result of the diagenetic and catagenetic transformations (e.g. transition from volcanic fragmental sand to arkoses to protoquartzites). Porosity increases (for equal amounts of quartz) the higher the energy of sandstone depositional regime, i.e. in the sequence barrier --~ deltaic --~ fluviatile -~ marine. The same order is found when these correlations are compared to increasing textural maturity, as defined by the highest degree of sorting and roundness of the original detrital quartz. Thus, under the same barrier diagenetic conditions higher porosity is found in barric sandstones, followed by del-
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m8
Chapter 3 • Main Factors of Reservoir Compaction
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Fig. 3.15c,d. Correlation of authigenic quartz (% b.v.) (c) with clay coating (% b.v.) on quartz grain surfaces, d Correlation of porosity (% b.v.) with silica (% b.v.) (detrital quartz + authigenic quartz) in reservoirs of different paleoenvironments
taic, fluvial a n d m a r i n e s a n d s t o n e s respectively. The h i g h e s t q u a r t z c o n t e n t s are obs e r v e d in rocks w i t h t h e lowest clay c e m e n t c o n t e n t s (Plate 8, c o m p a r e e, f a n d g, h). This is t r u e over different scales, r a n g i n g f r o m t h i c k r o c k s e q u e n c e s to t h i n sections. Sequences w i t h h i g h e r clay c e m e n t c o n t e n t s are c h a r a c t e r i z e d b y lower p e r m e a b i l i t y a n d vice versa.
Plate 8, Correlation of quartz grain compaction and regeneration features. All scale bars are o.z5 ram. a, b The dependence of grain "interpenetration" on grain size; b greater level of grain interpenetration in coarse-grained sandstone (b) compared to fine-grained sandstone (a). a Ordovician sandstone, well GBC-1, depth 3355.20 m. b Ordovician sandstone, well GBC% depth 3329 .20 m. c, d Compaction and silicification features as a function o burial, c Upper Silurian sandstone (well MGD-1,depth 3224.60 m) is significantly less consolidated than d. d Ordovician sandstone (well TEG-1, depth 4098.0 m). e, f, g, h Role of clay cement preventing silicification: e clay cement content = 24%, authigenic quartz content = 3% (Late Silurian sandstone, well HAL4, depth 1208.70 m); f clay cement content = 22%, anthigenic quartz content = 4.5% (Early Devonian sandstone, well STAH-17, depth 2981.5 m); g clay cement content = 4%, authigenic quartz content 16% (Early Devonian sandstone, well DID-I, depth 1944.o m); h clay of an Ordovician sandstone from well TEG-1 at depth 411o.o m
3.5 - Silica Solid Phase Transformation: A New Concept for Sandstone Diagenesis
lo9
zlo
Chapter 3 • Main Factors of Reservoir Compaction
3.5.3.3
SEM Investigation Results High resolution SEM examination of the sandstones allowed observation of several relevant features and patterns of silica transformation in sandstones. Significant changes in the morphology of quartz grains in the sandstone strata are observed with depth. Despite their significant geological age, in shallower horizons we observed a low level of grain consolidation (Plate 9a,b and Table 3.4). Grains are well sorted with an average diameter of around 300 ram. They clearly show loose "cement" material over the rounded surfaces of non-contact zones on their surfaces. The area of grain contact associated with shearing increases with depth, the contact length increasing from 15 to 15o ram. Faceted quartz grains (Plate 9c) or crystals of euhedral habit (Plate loa,d) are commonly observed. In the development of this habit rhombohedral positive r{lon} and negative z{o111}faces are clearly prevalent, with little development or complete absence of prism faces m{lmO}. A pseudocubic habit is developed when grains have a rhombohedral (z) face (Plate loa). 3.5.3.3.1 Quartz Twinning During sandstone evolution quartz crystals packing approaches closest packing. In the early stages of transformation (shallow depth) crystals are confined by the rhombohedral face. This is confirmed by the corresponding habit of adjacent forms: flat trihedral pyramids (Plates 9g,h, loa,b). Measurement of the angles between the"c" axes of the quartz cement crystals gives values around 7o-85°. The contact faces, as already mentioned, are those of the rhombohedron (Plate loa, e,g), a twinning which is closest to the Esterel or the Reihenshtein-Grisental law. In more deeply buried horizons, sandstones are significantly more consolidated, with reduction of the angular characteristics between the"c" axes of contacting cement crystals (compare Plates 9e and lOj). Plate loi,j shows convergence of crystals by the laws of Brazilian-Dauphin~ twinning, illustrating more advanced stages of sandstone evolution. It is quite clear that the habit of the grains becomes more elongate with the development of prismatic facet forms. Plates 9e and 1ok show a uniform sandstone formed by aggregates of faceted grains coupled along both prismatic {112o} and rhombohedral {lo11} faces; this would appear to be Brazilian twinning. Combinations of the Brazilian-Dauphind twinning can also be observed. Numerous microphotographs confirm that the presence of clay coating hinders quartz faces formation and aggregation of the grains into a uniform mass (Plates 9,b and loc). Deeply buried or in ancient horizons, and in the absence of clay cement coating, sandstones have become uniform quartzite rocks, with almost indistinguishable individual grains except for fractures in the crystal (Plates 9d, e; lol). 3.5.3.4
Infrared Investigation In support we have analysed data obtained by infrared (IR) spectroscopy of quartz in sandstones. The IR method is sensitive to the structural order of minerals and has been used to demonstrate significant modifications of the quartz doublet in the 800-780 cm-~
3.5
• SilicaSolidPhaseTransformation:A NewConceptfor SandstoneDiagenesis
111
frequency range. It is known that in highly crystallized quartz the typical ratio of the doublet intensities is/8oo > ITso (Fig. 3.16). Analysis of quartz from sandstones of the Sahara Plate demonstrates that the ratios of intensity of the doublet bands change with increasing depth and degree of quartz m e t a m o r p h i s m : 18oo >> t78o and Isoo < 178o (Fig. 3.16). The latter reflects the previously unnoticed transformation of the quartz crystalline structure to a less crystalline structure. The nature of the doublet in the 8oo-78o cm -t spectrum of a quartz is due to qs Si-O-Si pulsations of the SiO 4 tetrahedron. Two components (8oo and 78o cm -1) are produced by both the cis- and transgroup of Si20 z. The cristobalite and tridymite spectra are known to include a band in the vicinity of 78o cm -1, which is accounted for by a trans-group. X-ray studies of the structure of quartz, cristobalite and tridymite reveal at least one more type, which is not composed of Si207 ring groups. The structures have the following corresponding bands: 695 (and 392) cm -1 (quartz, cis-isomer), 625 (and 396) cm -1 (cristobalite, transisomer), and 560 (and 380) cm -* (tridymite, cis-isomer). These are the most vulnerable to reduction of crystallinity, and consequently their band intensity varies widely. It is known that cis- and trans-isomers differ in their chemical and physical properties; transformation of trans-isomers into cis-isomers and back requires counteraction of re-bonding. The trans-isomers are usually more stable than cis-isomers, demonstrating a higher energy of x-bonding. Actually, the 790 cm -~ band of crystobalite and tridymite is more stable than the 625 and 56o cm -~ bands. Consequently, it was noticed that with
Fig. 3,16. Infrared spectra of quartz in Triassic, Devonian, Silurian and Cambro-Ordoviclan sandstones from Oued elMya Basin (MGD-1,TEGq, OS-x, GBC-z and OCT-1boreholes). Evolution of quartz doublet 8oo-78o cm-I for sandstones with increasing burial depth: z etalon of quartz; 2 depth 243.55 m; 3 depth 1153m; 4 depth 1558.6o m; 5 depth 1679.3o m; 6 depth 1691.3o m; 7 depth 2 043 m; 8 depth 2978 m; 9 depth 3 o04 m; lo depth 3008.20 m; n depth 3157.8o m; z2 depth 3zn.8o ra; I3 depth 37oo.7o m; ~4 depth 392o.2o In; i5 depth 3996.30 m
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112
Chapter 3 • Main Factors of Reservoir Compaction
3.5 . Silica Solid Phase Transformation: A New Concept for Sandstone Diagenesis
113
~1 Plate 9. Sculpture and morphology of quartz grains at different stages of agglomeration. SEM images of sandstones of the Sahara Plate. a, b (b is a magnified part of a) Individualized quartz grains, well ALR-16, depth z 624.3 m, Middle Devonian (D~). c Faceted quartz grains, well OTLA-1, depth 19zo.o m, Early Devonian (D~). d Closely agglomerated quartz grains, well BH-3oa, depth r398.m m, Middle Ordovician (O0. e Agglomeration of quartz grains, quartzite-like sandstone, well MGD-~,depth 3 955.o m, Cambrian (e). f High density of outlet dislocations on quartz grain cleavage, well GS-1,depth 3r76.zo m, Early Ordovician (Q). g, h Crystal growth steps on quartz grain, well STAH-I7, depth z 9o4.o m, Early Devonian (D1). i Overgrown interstices of quartz grains with Burgers (spiral) dislocation mechanism. The spiral growth of grains is localized around triangular lunules (arrow), well HR-ro, depth 2142.4o m, Late Triassic (T3). j Overgrown junction of quartz grains with Burgers (spiral) dislocation mechanism, well RDC-4, depth 3 799.4 m, Early Ordovician (02) Table 3.4a. Sedimentological and mineralogical characteristics of some studied samples by SEM,shown in Plate 9 (remark: the quantitative mineralogical composition was determined by X-ray diffraction of sample powder and oriented preparation, using mixed multimineral standards; accuracy ±1-2%) SEM Well image
Depth Age (m)
Sedimentological characteristics
Mineralogical composition (%)
Basin
a,b
ALR-16 2624.3 Middle Medium- to coarse-grained, Devon. poorly sorted, rounded with siliceous and clay cement, shallow marine sandstone
Quartz 74;Siderite5; Pyrite 2; Clays 19
Illizi
c
OTLA-1 1920.0 Early Medium-grained, sorted, Devon. rounded, quartzose sandstone
Quartz 79; Siderite 3; Feldspars3; Clays 15
Azzene high
d
BH-301
Quartz 64; Clays 23; Siderite 13
Abner
e
MGD-1 3955.0 Cambr. Medium-to coarse-grained, very Quartz 72;Clays 14; poorly sorted, poorly rounded, Dolomite 8; Feldshallow marine sandstone spars 6
Mouydir
f
GS-S
3 ] 76.2 Early Medium-grained, poorly Ordov. sorted, rounded, very quartzose bar sandstone
Quartz 78; Feldspars 2; Clays 20
Oued eI-Mya
g, h
STAH-17 2 904.0 Early Coarse-to medium-grained, Devon. poorly sorted, rounded, shallow marine sandstone
Quartz 82; Feldspars 3; Clays 15
Ghadames
i
HR-10
2142.4 I_ate Mediumgrained, sorted, argilTriassic laceous, alluvial sandstone
Quartz 73 Feldspars 8; Oued Dolomite 6; Clays 13 eI-Mya
j
RCD4
3 799.4 Early Medium-grained, sorted and Quartz 78; FeldOrdov. well rounded, alluvial sandstone spars8; Clays 14
1 398.1 Middle Fine-to medium-grained, very Ordov, poorly sorted, angular, marine sandstone
Oued eI-Mya
114
Chapter 3 . Main Factors of Reservoir Compaction
3.5 • Silica Solid Phase Transformation: A New Concept for Sandstone Diagenesis
115
Table 3.4b. Sedimentological and mineralogical characteristics of some studied samples by SEM, shown in Plate lO (remark: the quantitative mineralogical composition was determined by X-ray diffraction of sample powder and oriented preparation, using mixed multimineral standards; accuracy +1-2%)
SEM Well image a,b
Depth (m)
Age
Sedimentological characteristics
Mineralogical Basin composition (%)
DL-405
830.6 Middle Devon.
Coarse-to medium-grained, Quartz sorted, rounded, quartzose, Siderite shallow marine sandstone Clays
79 6 15
Illizi
BH-301
1 398.1 Middle Ordov.
Coarse-to medium-grained, Quartz sorted, rounded, quartzose, Clays marine sandstone
84 16
Ahnet
d
HR-10
2149.0
Late Trias.
Medium-grained, sorted, rounded, slighttyshaly, alluvial sandstone
e,f,g
MGD-1 2979.5
Late Trias.
Coarse-to medium-grained, Quartz sorted, rounded, veryquart- Clays zose, deltaic sandstone
92 8
Oued ei-Mya
h
O5-1
2786.9
Early Ordov.
Coarse-to medium-grained, Quartz sorted, rounded, quartzose, Clays bar sandstone
90 10
OuedeIMya
GS-5
3 392.23 Early Ordov.
Medium- to fine-grained, poorly sorted and rounded, shaly, alluvial sandstone
Quartz 86 Feldspars 3 Clays 11
AmguidMessaoud anticlinal system
GS-S
3176.2
Early Ordov.
Medium-grained, poorly sorted and rounded, very quartzose, bar sandstone
Quartz 78 Feldspars 2 Clays 20
AmguidMessaoud antictinal system
k
ME1
2914.0
Cambr.
Medium- to coarse-grained, quartzose, very compacted, alluvial sandstone
Quartz Clays
86 14
tdjeraneM'Zab antictinat system
1
GS-S
3392.0
Early Ordov.
Medium-grained, poorly sorted, rounded, quartzose, bar sandstone
Quartz Siderite Clays
87 3 10
AmguidMessaoud anticlinai system
Quartz 76 Feldspars 7 Dolomite 5 Clays 12
Oued el-Mya
• Plate 10. Evolutional character and aggregation grade for quartz grains with depth and age. SEM images of sandstones of the Sahara Plate. a, b (b is a magnified part of a) Cement has not yet completely filled spaces between grains, they are clearly individual, Area of contact is up to loo microns and uneven, revealing the surface of original grain. In a, grains are bordered by a"coating"-oriented growth of a grain by rhombohedral faces {lo11} (arrow); it shows also large-angle boundaries of grains. Grain size about 3oo-6oo microns.Well DL-4os, depth 830.6 m, Middle Devonian (D2).c Individnalizedgrains, coated by chlorite cement which "prevents" faceting; grain size 15o-3oo microns. Well BH-3m, depth 1398.1 m, Middle Ordovician (02). d Isometricity of faceted grain wirth orientated coating of rhombohedral faces; some grains have a polycrystal coating forming individual portions of future faces (arrow). Size up to 600 microns. Well HRqo, depth 2249.0 m, Late Triassic (T3). e, f, g, h Quartz crystal individuals are clearly identified, with chiefly rhombohedral faces. Spiral depressions are formed by three rhombohedral faces of quartz grains (arrows) arranged after the right-hand or left-hand, 3~ and 3~, screw axis, thus making a fragmem of quasi-closest packing. Photomicrograph g shows a twinning (double arrow) which is closest to the Esterel or the Reihenstein-Grisentallaw. Mean size 15o-3oo microns, e, f, g Wetl MGD-1, depth 2979.5o m, Late Triassic (T3). h Well OSq, depth z786.9 m; Early Ordovician (07). j Quartz monocrystalline individuals with fiat boundaries of convergence, pyramid faces are clearly seen {11zo}. Brazilian-Dauphin4 twinning. Well GS-5, depth 3176.2 m, Early Ordovician (O1). k Complete mergence of grains into a monolithic quartzite-like rock; well-faced quartz crystals (rhombohedral faces) in pores. Well ME-z, depth 2914.o m, Cambrian (C). I Convergence of grains with linear boundaries with spiral growth of grains around triangular lunules (ar~vw).Well GS-5, depth 3392.0 m, Early Ordovician (O~)
116
Chapter 3 • Main Factors of Reservoir Compaction
quartz metamorphism the intensity of 8oo and 390 cm -1 bands of Si20 z cis-groups decreases as the crystalline structure of quartz becomes more compact. Thus, the observed modification of the intensities of 8oo and 78o cm -~ bands of metamorphogenic quartz of Saharan sandstones is due to the increased"contribution" of Si207 trans-groups to vibrations, and can be interpreted as a typomorphic feature of quartz metamorphism. 3.5.4 Discussion 3.5.4.1
Possible Sources of Silica According to existing hypotheses, authigenic quartz can originate from several different sources: precipitation of silica in pores in near-surface conditions (Davis 1964; Blatt 1979), diagenetic transformation of clays and feldspars and other aluminosilicates (Siever 1962; Towe 1962; Keller 1963; Dunoyer de Segonzac et at. 1968; Hower and Eslinger 1973; Hawkins 1978; Blatt 1979), flow of silica-bearing waters from depth (Taylor 195o; Heald 1956), pressure solution (Trurnit 1968; Sibley and Blatt 1976; Robin 1978; Houseknecht 1984) and illitization of montmorillonite (Keller 1963; Dunoyer de Segonzac et al. 1968; Burst 1969; Hower and Eslinger 1973). Silicification of poorly consolidated sediments in near-surface conditions by precipitation from subsurface waters occurs following dissolution of silica and hydrolysis of silicates by surface waters and subsequent reprecipitation in the sedimentation zones by ground water discharge (Thompson 1959; Trurnit 1968). Silicification phenomena can occur under a wide range of geochemical conditions which facilitate dissolution of silicates. The chemical stability of sandstone is a function of temperature and pressure changes during burial; in the Paleozoic sandstones of several basins in the Sahara Plate geothermal gradients are 25 °C km -1. Many of the studied reservoirs are buried to great depth with formation temperatures up to 12o °C and higher. 3.5.4.1.1
Precipitation of Silica from Subsurface Waters We present some dissolved silica data here for waters from various environments in northern Algeria: subsurface waters 5-1o ppm, thermal waters 5-20 ppm, shallow nearshore water 1-2 ppm, fluvial surface water 11-12 ppm. Meteoric waters with lo-15 ppm dissolved silica are known to circulate to depths of many hundred of meters and therefore could present a cement source (Btatt 1979). Although surface water contains some silica in solution which might promote quartz cementation at shallow depth, it is clear that this silica content is too little to account for the amount of authigenic quartz present in Saharan oil-field sandstones at depths of several thousand meters. In the Hassi Messaoud oil field formation water from the Cambro-Ordovician reservoirs contains about 15o ppm dissolved silica, formation water from the Albian reservoirs in the same field contains 15 ppm, while in the Berkaoui oil field Albian formation water contains lO ppm, Triassic water 3.6 ppm and Hassi R'Mel (gascondensate) Triassic formation water contains 5-7 ppm silica. It is obvious therefore that the 15o ppm of dissolved silica found today in Cambro-Ordovician reservoirs could
3.5 • Silica Solid Phase Transformation: A New Concept for Sandstone Diagenesis
127
present a source for sandstone cementation, but the 5-7 ppm of silica in Triassic formation water is not great enough to account for the cementation observed in these sandstones (6-12% of authigenic quartz). Using the same mode of calculation applied by Sibley and Blatt (1976) for Tuscarora orthoquartzites and Blatt (1979) for estimation of the volume of water needed to account for the amount of quartz overgrowths observed shows that some lo 5 to 5 x lo 7 pore volumes of water would have to circulate through these sandstones. This leaves the question unanswered as many Saharan sandstones contain lo-3o% and sometimes up to 35% of quartz cement associated with porosity ranging from 0.5 to 16%, and often with little or no evidence of pressure solution. These sandstones range in age from Cambrian to Triassic, most of them are buried to depths ranging from 1500 to 4 ooo m, and many of them were uplifted after deposition to nearsurface position, a condition necessary to promote quartz precipitation from waters as suggested by Blatt (1979). However, many sandstones uplifted in the Late Paleozoic as a result of Hercynian Orogeny in central and western Sahara which remained elevated in nearsurface conditions up to the present do not show a significant content of authigenic quartz, as was previously demonstrated (Plate 7a). Now, it is obvious that the bulk of quartz overgrowths in Saharan sandstones is not attributable to precipitation from subsurface waters at least because of hydraulic and geochemical considerations. Accepting that dissolved silica in pore waters may in some cases present a source for sandstone cementation by quartz overgrowths, the major source of authigenic quartz is still to be identified. 3.5.4.1.2
Diagenetic Transformation of Clay Minerals: Detritai Iliite Tronsforrnation into Kaolinite The transformation of detrital illite into authigenic kaolinite is observed in all the sandstones studied. This process of illite kaolinization, with silica produced as a byproduct, is widely recognized by geologists (Millot 1964). This process occurs as a result of silica hydrolysis involving a pH reduction, some element migration (Na, Ca, Mg) and immobility of other elements (A1, Fe, Ti...). The quantity of silica liberated reaches 2o% of the dissolved silicate and 15% of potassic alkali is also released. This process may lead to a porosity increase following alkali silicate dissolution. The silica released could be reprecipitated in the absence of a transport medium, thus giving a quartz overgrowth. However, on adjacent grains we can observe quartz grain corrosion resulting from the parallel action of released alkali. These two processes could partially explain the characteristic reservoir heterogeneity in most Saharan basins, but they cannot explain the high authigenic quartz content in the reservoirs studied. In ma W of the sandstones studied quartz cement is up to 17-2o% and the clay content in these rocks only rarely exceeds 15%, but is more usually between 5 and 8%. So if all the detrital clay is illite we derive only 2-2.5% of silica from the total mass transformed. What is more, the iltite transformation into kaolinite is never complete, hence one always observes in the cement a non-transformed detritaI illite, and sometimes detrital kaolinite. A quantitative estimation on the Hassi Messaoud sandstones which contained initially about 3o-5o% of detrital illite (that is, the illite content in the non-transformed siltstones in the same site) shows that only 1% of authigenic SiOa could be derived from the above transformation.
1~8
Chapter 3 - Main
Factors of Reservoir Compaction
3.5.4.1.3
Diagenetic Conversion of Montmorillonite or Interlayered Montmorollonite-Illite to Pure fllite This process may be an important supplier of silica to pore solutions. The free silica released by this mechanism is estimated as z.z g of quartz or chert per lOO g of clay transformed, producing 1.3% of quartz or chert in an ideal case (Towe 1962). 3.5.4.1.4
Diagenetic Alteration of Feldspars and Volcanic Rocks Feldspar dissolution is observed in all the reservoirs studied. Some correlation exists between the occurrence of diagenetically kaotinized feldspars and the amount of quartz overgrowths (Fothergilt 1955; Hawkins 1978), but this phenomenon is generally believed to be quantitatively unimportant as a source of silica (Blatt 1979). Volcanic rock fragments were found to be altered diagenetically in Triassic and some Paleozoic sediments of various basins thus presenting a possible silica source. However, volcanic fragments generally are not produced in the cratonic and shelf environments where most orthoquartzites occur so their diagenetic influence is only local. 3.5.4.1.5
Flow of Silica-Bearing Deep Water by Faulting Tectonics Sandstones with little clay cement are a favourable environment for the circulation of silica-bearing deep waters, especially in provinces with a network of faults (Taylor 195o; Heald 1956). In these conditions some silica could precipitate in the available pore space. In the stud?, area, it was particularly important to examine closely ancient faulting, i.e. Triassic and Cretaceous communicating faults which do not seal between different reservoir blocks. However, studies of the relationship between sandstone permeability and the distance ofboreholes from fractures and faults have demonstrated that the mean permeability of sandstones from boreholes located within 500 m from faults is practically the same as the mean permeability of sandstones from other boreholes. 3.5.4.1.6
Intergranular Pressure Solution A comprehensive petrographic investigation of Sahara sandstones and quartzites has revealed a limited occurrence of intergranular pressure solution even in ancient or deeply subsided sediments. Widespread examination using cathodoluminescence petrography of Cambrian, Ordovician, Devonian, Carboniferous and Triassic sandstones and quartzites from depths ranging from a few hundreds of meters to 5 km has shown that the standard criterion of pressure solution (sutured contacts between apparently detrital parts of adjacent quartz grains) occurs relatively rarely. The great bulk of interpenetrations between adjacent grains occur in fact as compromise boundaries between overgrowths (Plates 6, 7). These overgrowth-overgrowth contacts could not be a source of silica to pore solution (Btatt 1979). The major factors which imply a limited role of pressure solution as source for sandstone cementation in the studied area are: (1) early quartz overgrowths probably at shallow depths, (2) formational overpressure, (3) relatively moderate thermal regime and moderate gravitational stress due to the presence of thick salt formation in the upper part of the stratigraphic column, and (4) cathodoluminescence color of quartz overgrowths.
3.5 • Silica Solid PhaseTransformation: A New Concept for Sandstone Diagenesis
119
In the Sahara sandstones petrographic evidence suggests significant quartz cementation prior to much compaction. Petrographic evidence of early quartz cementation of sands has also been reported by numerous other researchers (Dapples 1959; Siever 1959; Millot et al. 197o; Bucke and Mankin 1971; Blanche and Whitaker 1978). In the Sahara Triassic Province (Oued el-Mya, Triassic, Ghadames and North Illizi Basins), the relationship between quartz overgrowths and evaporites which have penetrated from the overlying Triassic formation into fractured Paleozoic sandstones (salts are deposited in fractures intersecting quartz overgrowths) demonstrates that much quartz cementation in these sandstones (origin other than pressure solution) occurred prior to significant Mesozoic subsidence. Such early significant quartz cement equalizes stress through the rock volume and thus reduces stress concentration at grain boundaries (Sibley and Blatt 1976) and impedes the possibility of pressure solution. So, early quartz overgrowths formed a solid skeleton which resisted subsequent rock consolidation during further subsidence. This phenomenon is one of the factors governing conservation of relatively good reservoir properties at depth in the study area. Saharan basins enclose thick Silurian, Devonian, locally Carboniferous and lesser Triassic shales. These shales were composed initially of a large amount of smectite and interlayered smectite-illite day minerals. Today lo-2o% of smectite-illite in these shales at a formation temperature of 8o-12o °C is quite common. Extensive shale diagenesis, particulary the smectite to illite transformation, produces a water volume equal to approximately one-half the volume of the original smectite according to Powers (1967), as well as the normal expansion of water with increased temperature that accompanies increased burial depth (Magara 1975). This could create overpressured zones in muds and undercompaction in associated sandstones in geoclinal sediments (Bradley 1975; Burst 1976; Hower et al. 1976). In northern and eastern Saharan basins, as well as thick Paleozoic shales, thick Triassic and Jurassic evaporites (up to 2 ooo m) occur universally. The water released during evaporite compaction could create additional stress, leading to formation overpressure and undercompaction which damps, in particular, pressure dissolution. On the other hand, Triassic and Iurassic evaporites provide a regional impermeable seal due to their abundances, uniformity and widespread distribution, conditions required for the conservation of abnormal "excess" fluid pressures. Impermeable seals are also provided by Silurian, Devonian, Carboniferous and Triassic shales as well as by a Triassic volcanic seal overlying Early Triassic or Paleozoic sediments at the level of the Hercynian unconformity. Furthermore, our basin modeling, in particular burial histories, shows that in northern and eastern Saharan basins in the Mesozoic, sediments were rapidly deposited and buried probably more rapidly than the depositional fluids could adjust to the added load. In such situations, the rapidly increasing weight must be partly supported by the fluid that is trapped as a result of the decrease in porosity and permeability during compaction (Blatt 1979). Less favorable conditions for pressure solution occurrences in Saharan sandstones are also indicated by the presence of thick Triassic and Jurassic evaporites in the sedimentary column of the Sahara Plate as a result of their high thermal conductivity and low density compared to clastic sediments. This results in a relatively lower thermal regime and lower geostatic stress. Supporting evidence of a low thermal regime in this province is also derived from quartz cathodoluminescence colors, as was previousely cited.
1~o
Chapter 3 • Main Factors o f Reservoir Compaction
The circumstances above explain why there is little evidence of pressure solution even in Cambrian, Ordovician and other sandstones buried to 5 ooo m and more. The large amount of authigenic quartz viewed in cathodotuminescence microscopy, despite little evidence of pressure solution criteria, suggests that a more appropriate mechanism for silica authigenesis in sandstones needs to be identified.
3.5.4.2 Silica Dissolution-Precipitation A common approach in geology is to analyse processes as a function of temperature and pressure. However, many tong-lived geological processes, in particular the diagenetic processes in sedimentary rocks, develop under moderate temperatures and pressures. As is known, the concept of silica pressure solution in sandstones is shared by many authors (Heald and Renton 1966; Renton et al. 1969; Houseknecht 1984,1988; McBride 1989; Sibley and Blatt 1976; Williams et al. 1985a). Fundamental to supporters of this concept are the laboratory experiments by Fairbairn (1954) in which quartz sand compressively loaded and heated transforms to quartzite. There are several contradictions in the concept of pressure solution. For instance, it is widely believed that stylolites are the result of this mechanism. From a solid state point of view they are most likely associated with the formation of the broken-sinusoidal type of grain boundaries, probably due to their migration. Grain boundaries with a crystallographic disorientation, due to a grain's tendency to regular orientation, can locally bend (Poirier 1983). Silicic acid can get into underground and surface waters as a result of direct dissolution of siliceous rocks and minerals, the weathering of silicates, volcanic activity and microorganisms. The main mechanism suggested for silica extraction from marine waters is the biological activity of diatoms, radiolaria, silicaflagetlates and sponges, and correspondingly the main silica supply occurs in the biogenic zone. The concentration of silicic acid in modern river water is usually within lO-2O ppm, in marine water o.5-3.o ppm and in subsurface water lo-4o to aoo-3oo ppm, depending on the conditions of water saturation (Bogomolov et al. 1967; Schenak and Migovich 1969). The highest content was found in formation waters at depths of 2-3 km, and in hydrothermal waters: from 200-4oo to 5oo-7oo ppm. The content of silicic acid in some sodium carbonate-bicarbonate brines can be as high as 2 70o ppm at pH lo (Jones et al. 1969)- However, natural waters usually show a concentration of silicic acid considerably lower than the solubility limit of amorphous silica under the same conditions (Fournier and Rowe 1962; Table 3-5). Consequently, modern subsurface waters are undersaturated with respect to amorphous silica, while marine water is unsaturated also with respect to quartz (Fournier and Rowe 1977). Some conditions, however, are known which allow mineral crystallization from solution. The main agents of crystallization are the crystallization power and the crystallization rate; crystallization is possible in systems out of equilibrium. The measure of a system's deviation from equilibrium is called the dri~fng force of crystallization, and its actual expression is supersaturation and supercooling. The most important parameter atlowing the growth of crystals from solution is solubility. The concentration of saturated solution quantitatively determines the solubility of the substance under particular conditions. Crystals do not grow from unsaturated solutions, crystals mainly dissolve in them.
3.5 • Silica Solid Phase Transformation: A New Concept for Sandstone Diagenesis
lal
Table 3.5.
Thermodynamic data for mineral forms of silica (all values in k] mo1-1)(O'Connor 1958; Van Lier et al. 196o;Fournier and Rowe1962;Siever1962;Moreyet al. 1967;Walther and Helgensont977) Mineral silica form
Formation Formation Dissolution Solution Energyofacti- Activation enheat freeenergy heat freeenergy v a t i o n o f d i s s o - e r g y o f p h a s e lution in water transformation
Amorphous silica
-898.353
-849.469
cz-Cristobalite -907.510
-853.667
19.2
19.0
c~-Quartz
-858.812
21,7 22.3
21.4 22.9
917.257
] 4.0
15.5
74.5 75 67
25.1
Thus, the crystallization of cristobalite and quartz directly from natural waters seems unlikely. However, the partial precipitation of silica in amorphous form is possible in areas where fluvial and marine waters mix. Precipitation of silica as gels followed by the formation of opal-like silica is possible in practice only in areas of intensive volcanic activity.The heat and standard free energy of solution were calculated from the solubility of quartz in water found analytically to be within the temperature range 25-473 °C (Van Lier et al. 196o). Analogous calculations were made for cristobalite and other silica forms (Fournier and Rowe 1962; Table 3-5)The activation energy of the phase transformations of silica was estimated on the basis of the equations describing the kinetics of phase transformations (Emanuel and Knorre 1984). Similar results were obtained in the laboratory under hydrothermal conditions (Mizutani 1966). The hydroxylation-dehydroxylation of silica occurs during phase transitions.The rate of this process and, correspondingly, the activation energy depends on the activating agents, the most important of them being hydroxyl-ions catalyzing dissolution, polymerization and phase transformations of silica. In anhydrous conditions at temperatures of 45o-6oo °C and a pressure of 4 kilobars the activation energy is 197 k] mole -1. In the presence of lO% H20 it drops to as little as 21 kI mole -1 (Naka et al. 1976). The activation energy necessary for the t r a n s f o r m a t i o n of opal-A --+ opal-CT and opal-CT --~ quartz in anhydrous conditions is 89 k] mole-', while in water vapor the activation energy for the transformation of opal-CT ~ quartz is equal to 13.2 kJ mole-l; OH- and water catalyze the transition. According to numerous supporters of the pressure solution mechanism, the following should occur: SiO, gel --~ opal-A ~ silica solution --~ opal-CT --~ silica solution ~ quartz; as well as: quartz (sand) --~ silica solution --~ quartz (sandstone) -~ silica solution --~ quartzite. This cycle of reactions is unrealistic in the subsurface in practical terms due to the low solubility of the mineral forms of silica in water and, consequently, the low solution saturation, i.e. the failure of mineral forms of silica to achieve the saturation necessary for crystallization. The following well-known fact also contradicts the pressure solution mechanism: at great depths, where permeability- is very low, authigenic quartz overgrowths continue to grow. No breaks of Si-O bonds should occur in the solid phase transition of opal-CT (acristobalite) --~ a-quartz, as there is a martensite type of phase transition occurring through cooperative movement of atoms without any breaking of Si-O bonds (Plyusnina ~986).
i22
Chapter 3 . Main Factors
of ReservoirCompaction
3.5.5 Solid Phase Process
Modern physico-chemistry considers two main solid phase processes to be of great importance to geology: phase transitions and agglomeration. Solid phase transitions take place by interaction between adjacent solid matter particles; they can occur with the participation of gas or liquid phases or with the simultaneous participation of both phases. Solid phase transition can be considered in those cases where the primary phase is amorphous and displays polymorphism, while the crystalline phase is finegrained, disordered and initially deformed. The driving force behind crystallization in amorphous gel-like bodies, considered to be super-cooled liquids, is the redundant free energy. The driving force behind the growth of coarser crystals in a fine-grained phase is the redundant free energy of crystals boundaries. Boundary energy may be considered to be a kind of interphase strain. When grains become coarser strain decreases and the driving force of crystallization weakens. At higher temperatures the processes are sharply accelerated. 3.5.5.1
Solid Phase Transformations
Solid phase transformations can be considered as a set of"chain reactions". During these reactions, an intermediate phase (or a state with an increased reaction activity) is created which activates the adjacent reagent lattice. Polymorphic transitions, particularly in low temperature conditions, are accompanied by the disordering of the crystal structure and the fine dispersion of matter and, more importantly, the formation of an initial metastable amorphous phase. Ostwald (1935) formulated the following rules for physico-chemical systems: (1) in any process, the initially originated state is not a state of highest stability with the least free energy, but rather a state of lowest stability with its free energy close to the initial energy; (2) if there exist, between the initial and the final state, a number of relatively stable intermediate states, these states will follow each other in the decreasing order of free energy (rule of stepwise transitions). Therefore, the driving force of crystallization will change through the various stages of the process (Ostwald's steps) due to changes in the volumetric proportions of phases and shrinkage stresses. Solid phase transformations proceed mainly by diffusion. Various mechanisms have been suggested to explain the process of diffusion, in particular point and line defects and ion exchange. Diffusion most often takes place due to the migration, supersaturation and mergence of vacancies and the formation of pores (Friede11964). As already demonstrated, for relatively recent sedimentary rocks and weathering crusts (20-20o Ma), the characteristic sequence of silica transformation is: silica gel -~ opal-A --~ opal-CT --+ chalcedony -~ quartz. This is a well studied, polymorphous transformation of the solid phase transition type (~z-cristobalite -+ o¢-quartz) (Plyusnina 1983, ~986, 199o). One aspect of the solid phase transformations is the gradual transition of one silica phase to another and the gradual transformation of morphological features down the stratigraphic column with increasing burial depth and age (similarly in weathering crusts). The observed sequential cristobatite-quartz ratio would be unlikely to occur in rocks with repeated dissolution and crystallization from solution (Plyusnina 1983).
3.5 - Silica Solid Phase Transformation: A New Concept for Sandstone Diagenesis
123
3.5.5.2
Agglomeration Sandstone studies have disclosed that in older sedimentary rocks (2oo-25o Ma) another type of solid phase transformation, agglomeration, takes place. This phenomenon can be simulated in the laboratory only at rather high temperatures. However, despite the relatively moderate temperature and pressures characteristic of buried sediments, in our opinion the duration and the specific features of the process of ageing justify the interpretation of these transformations in sandstones as the process of agglomeration. Agglomeration is an irreversible process of transition of an unstable system into a more stable state by the spontaneous consolidation of a dispersed porous body (Tretyakov 1978). This process can be accelerated as well as decelerated by changing burial diagenetic parameters. The final outcome of agglomeration is the formation, from a conglomeration of sand grains weakly bound by forces of adhesion and friction, of a solid monolithic quartzite rock. Essentially, agglomeration is the elimination of the pores in a porous body through their filling with material. Then the"final" density of the rock is reached; resistivity, hardness, chemical stability and thermoconductivity increase, while gas and water permeability decrease. The driving force of the process of agglomeration is surface energy. At the start, the system consists of sand with a large internal interphase surface far from thermodynamic equilibrium and possessing elevated reserves of free surface energy. Each system tends towards reduction of interphase surfaces and this is equivalent to a reduction of the surface energy and, consequently, the total energy of the system. The sandstones of the Sahara Plate illustrate the different stages of agglomeration. In the initial stage, consolidation of a body of grains takes place by the gliding of particles along grain boundaries in response to the great surfaces and excessive free energies that create pressures which tend to compress a body of grains to reduce its free surface. This compressive pressure is given as p ~ ( a / p ) P (1 - Po), where p is surface tension at the boundary of a solid phase, r is grain radius, P is porosity under a given temperature or in a given age interval and Po is original porosity. Therefore the first phase of sand consolidation is related to movement of whole grains. This movement stops when a compact packing of grains is achieved. After this, compaction is possible only through diffusion, not through the movement of grains. In the initial stage of agglomeration, grains are agglomerated to each other, which results in an increase in their contact area and convergence of their centers. During this initial stage grains still exist as separate entities (Plate 9a,b and Fig. 3.17a). In the intermediate and final stage, consolidation takes place and pores which were once in communication shrink and become isolated (Plate 9d,e and Fig. 3.17b,c). 3.5.5.3
Diffusion The main physical process during agglomeration is diffusion, i.e. a mass transfer leading to an equalization of ions and molecule numbers in a system. The strength of diffusion in a given direction, say the X axis, is described quantitatively by Fick's law: f = - D d~N , dx
124
Chapter 3 • Main Factors of Reservoir Compaction b
c
2 d
e
Pore
Pore
Direction of movement - - - - Vacancies - Material particles
Fig. 3.17. Schematic sketch of agglomeration phases: a initial, b intermediate, ¢ final, d Scheme of pore-filling between grains under solid phase agglomeration: comact line of grains shown by double dotted line, grain boundaries before agglomeration shown by dotted line, grain boundaries after agglomeration shown by continuous line; diffusion direction of vacancies and particles are shown by dotted and continuous arrows, respectively, e Scheme of dosed pore filling in grains under solid phase agglomeration. ~, 2, 3 Pore variously sized, subjected to filling in an analogous manner as the intergranular pore space, f Scheme of agglomeration with participation of the liquid phase, involving a capillary pressure and grain consolidation
wherefis the n u m b e r of partides moving across the unit area in one second and d N ! d x its gradient in the X axis direction. The coefficient of diffusion D is expressed in cm 2 s-I. Its temperature dependence is exponential: E
D = Doe RT
,
where D o is the diffusion coefficient at T = ~o and E is the process activation energy. The activation energy in the surface layer, along the grain boundaries a n d in the volume is correlated: Esurf"< Egr. bond.< Evol. a n d Dsurf"> Dgr. bond.> DvoLFrenkel a n d Pines (1945) suggested that the solid state agglomeration is related to viscous flow or creeping into pores, especially d u r i n g heating, due to surface tension in grains possessing some surface bending. During this process the surface free energy decreases. They also showed that this p h e n o m e n o n is due to substance redistrib u t i o n d u r i n g directed v o l u m e a n d surface self-diffusion. Using the w e l l - k n o w n T h o m p s o n formula it was shown that the pressure above the curved surface of a phase
3.5 - Silica Solid Phase Transformation: A New Concept for Sandstone Diagenesis
125
and the vacancies concentration ~ in the crystal body near its surface depend on the curvature of this surface: ~=¢0(1+
2r•
)
or
A=~r-~o=
2°~° ; r-~ 0 ,
where 4o is the equilibrium vacancy concentration near a flat crystal surface, ~r is the vacancy concentration near the curved surface having a curvature radius r, a is the surface tension coefficient, Vo is the volume of vacancies in the crystal, K is the Boltzmann's Constant, T is the absolute temperature and A is the v a c a n c y concentration near a curved surface. The vacancy concentration increases near surfaces with small r, i.e. high curvature, and the vacancy concentration near a convex crystal surface (r is positive) is less than near a concave one (r is negative). The higher the curvature of a surface in the crystalline phase, the higher the concentration of vacancies near this surface, i.e. a gradient of concentration of vacancies exists (Fig. 3-17a)- With time, and in particular with heating of sandstone masses during burial, this gradient tends towards uniformity. Therefore a directed diffusionat flow of vacancies develops from higher to lower concentrations which is equivalent to counter-flow of particles (atoms, ions) or diffusion in the opposite direction. Figure 3.17a,b shows the overgrowth mechanism in the pore space. Grain boundary contacts have a concave surface and small r of curvature. The vacancies move to the grain surfaces and dissipate there (the place of their dissipation can be also the block borders in crystals, dislocations, microfissures and other defects). Particles are diffused towards the crosspiece, enlarging it and filling the pore. The coefficient of surface diffusion being greater than that of the volume diffusion, the atomic flow is mainly directed to the surface of agglomerating grains. Consequently, the preferential transformation of grains at the level of their surfaces is seen. As part of the flow directed to the surface of the crosspiece is taken away from the intergrain contact area, the grains become yet more closer. The process of small pores filling through their integration is more advantageous in terms of reducing the internal interphase surface and, consequently, the surface energy. With time the vacancy oversaturation decreases during agglomeration and the system moves closer to equilibrium. If grains are moistened, a liquid capillary develops and a concave meniscus is formed with a small negative radius of curvature. Capillary pressure raises the liquid in a capillary expelling it into pores and bringing grains together as the sandstone consolidates or "shrinks" (Fig. 3.18c). The time required for this type of agglomeration is directly proportional to the surface tension on the liquid-solid phase interface, and is inversely proportional to the viscosity of the liquid phase and the size of the solid phase particles. The study of these phenomena in sandstones of different origin seems rather promising. Reaction rate constants increase with temperature such that a lo °C temperature increase increases reaction rates (k) by 2-4 times: E
k = Ae
RT
,
where A is the theoretically calculated constant, E is the activation energy, R is the universal gas constant and T is the temperature.
lz6
Chapter 3 • Main Factors of Reservoir Compaction
Sandstone grains with a distorted crystal structure (clearly demonstrated by outcrops of dislocations on grain shears and assessed at 4 x lo 7 cm -~ (Plates 9f, lob) show a much higher coefficient of diffusion and accelerated mass transfer compared to similar processes in perfect crystals.
3.5.5.4 TwinnmgandMechanismofTrans~rmation Aggregates of weakly bonded sand quartz grains constitute an unstable system. The initially disordered mutual orientation of grains later becomes crystallographically ordered with time and burial. Different kinds of deformation contribute to sandstone transformation during diagenesis. Mechanical deformation appears to produce a greater effect initially: it causes change of the mutual position of numerous particles (grains) of the sand, leads to changes in their form and size and contributes to change of interaction forces between them. Twinning is one of the manifestations of orientation regularities in grains. A preliminary condition for this is that the oriented faceting of quartz grains is analogous to the closest packing. However, the most active deformation is plastic deformation as realized by translational sliding. Sliding propagates along certain directions and preferably along closely packed layers of the crystal structure of quartz (plane {lO11}).Consequently, one kind of plastic deformation is twinning, which is particulary characteristic for closely packed crystal structures. The closest-packed planes are planes with minimum surface energy, and for this reason they are most commonly either faceting planes or slide planes. It has been observed that closest-packed layers are typical of intergrown crystals, or twins (Belov 1976). As mentioned above, one dislocation type is the closest-packing defect: twinningplane ABCABCABC
1
ABCABABABC . packingdefect
The packing defect can be considered as a combination of twinning planes. Therefore, deformation-displacement dislocations are due to sliding along crystallographic planes. On completion of deformation the deformed part of the crystal (or a neighbouring grain) may become the mirror reflection of the non-deformed part (or another grain), and thus twinning is achieved. Extraneous particles (distinctly observable in thin sections), overgrown on the "seeding" grains and regenerated fragments are particulary favourable to the formation of strained slide twins. The crystal straining that occurs under near-surface conditions leads to a local twinning of the crystalline quartz structure and to reorieutation of the twins formed with respect to the host crystal (Gordienko et al. 1966; Karyakina 1974). Quartz twins, by analogy with sticking growth twins, develop when grain closing and consequent additional mutual orientation occur under the action of their own electrostatic field. The energy required for the formation of twinning juncture-twinning across the twinning plane is extremely low (Lawson and Nielsen 1959).
3.5 • Silica Solid Phase Transformation: A New Concept for Sandstone Diagenesis
127
3.5.5.5
Role of the Dislocations in Recrystallization The plastic deformation of quartz generates stress and facilitates the accumulation of a significant amount of elastic energy, which is a major drMng force for recrystallization concomitant with agglomeration. Both types of plastic deformation - sliding and twinning - are due to the movement of dislocations, which renders the crystal highly plastic. Moreover, it is precisely the high-density dislocations, visible as triangular depressions in the grain intergrowth planes, that are responsible for twinning (Plates 9f, g,h, loa,b). The mechanism of dislocations formation by plastic deformation of crystals was suggested by Frank and Read (1952). Dislocations are known to be linear defects of various type capable of generating a proper stress field and, consequently, prone to migration under the action of external forces. The resultant effect is the "slipping" of atomic planes - a plastic deformation, i.e. the breaking and remaking of the bonds between the dislocation line and the adjacent atoms. The creeping of boundary dislocations in the direction normal to the slide plane is effected via attachment of the vacancies to or their breaking away from the plane boundary. This process is directly connected with the diffusional mass transfer and plastic deformation and is ultimately associated with the ageing and the instability of sandstone as a system. A specific role should be attributed to screw dislocations, sinbe they facilitate recrystallization (Verma 1953). In the process of compaction, deformational dislocations in sandstone grains move elastically and interact to form stable articulations or nodules (Fig. 3.18). Nodules of screw dislocations are usually extremely strong (Poirier 1983). Dislocations are characterized by excess energy which is made up by the energy of broken or distorted bonds in the dislocation nucleus and the energy of weak stresses around dislocations (elastic stresses). Due to the excess energy in the area of the dislocation nucleus, the substance porsesses an increased chemical activity and thus growth, dissolution and oxidation always originate in defect areas and proceed more intensively along them. An active screw dislocation, once sprung into existence, becomes involved in the spiral crystal growth. As is known, the growth dislocations are virtually indistinguishable from the dislocations produced by strain, and the latter are typical of sandstone grains. As a screw dislocation has gained access to the surface of a crystal, a step with height equal to the perpendicular component of Burgers vector is produced (Plate 9g,h,i). The preferential crystallization in the direction along this step initiates a spiral growth (Mntaftschiev 198o). Consequently, the spiral growth may be envisaged as centered around a certain defect produced by a slight mutual displacement of the atomic planes within a crystal (or crystals), which in fact is screw dislocation. According to theory, the growth of imperfect crystals under weak oversaturation (presumably, one may refer to the diffusional overgrowth of quartz grains) is affected only by screw dislocations (Burton et al. 1951). As mentioned above, spiral pyramids with a screw step (whose height is a multiple of the crystal celt size) are formed. With a step parallel to the closest-packing direction along which the growtl~ rate is minimal, polygonal spirals are formed in such a manner that the face symmetry is retained. It is assumed that the triangular forms (Plates 9i, me,f,h) are most likely associated with partial dislocations which are produced when the atoms become displaced from
128
Chapter 3 , Main Factors o f Reservoir Compaction
d
30
0
1 ,
I0 ,
I00 ~
I000 30
I
I l Oil(gas)-bearing zone I1 III Water-bearing zone
o ~ ' ~
•
o
.o & - . : . .
20
20 •
•
o
oe
mm
~0
• 0
• Oe
•
0° ~ o
•o
o~
,/4Zo
~o
o o~ %.-,,% ,~ ..o o.o % o O , , 0 o-&..: ~ o • . . "~'%Oo o
.~
o o--~ o. oooo
.
I
• "
CI.
10~
e oooo
o
~
~
O
0
0~ O •
•eo° o o o
-
o
~
~
0 O I
1
I0
2o-ey~..t,-.o
0 0 ~0
O~oUO
0 O0~,JO O O OU
0
•e oZ o.;z
, ~
~ , ~ 2 o _ ! )0
0
,°o°•°
el•l@
ii
• A
~ 0
<
~
O0 O
u O
o/
oc
-
/ i/
I
I0 Permeability (roD)
I00
0 I000
Fig. 3.18, a-¢ Development of dislocations forms. Meeting of three dislocations with formation of a node; this is possible if the sum of three Burgers vectors is zero: B12+ Bz3+ B13= o. d Correlation between porosity, permeability and zonality of oil- and water-bearing sandstones of the Sahara Plate; A sandstones with l-5% of authigenic quartz, B sandstones with 5-to% of authigenic quartz, C sandstones with ~o-2o% of authigenic quartz
their regular sites of normal crystal structure to those of a twinned structure, or when the alternating order of closest-packed sheets becomes disturbed. Spiral depressions, or void dislocations, have also been observed to occur and, accordingly, spiral growth of grains localized around triangular lunules (Plates 9i, lol). At the early agglomeration stages, the spiral depressions are formed by three rhombohedrat faces of quartz grains, arranged after the right-hand or left-hand, 31 and 32, screw axis, thus making a fragment of quasi-closest-packing (Plate loe,f,g,h). They are also observed to form at a later stage - the stage by which the recrystallization takes place and the intergranular gaps become filled (Plates 9i, 1ol). The dextra- and sinistrarotatory growth spirals are associated with the right-hand and left-hand helical arrangement of SiO 4 tetrahedra for the positive and negative rhombohedron. The arrange-
3.6 • C o n c l u s i o n s
~29
ment of rhombohedral faces of one type generates the filling of the intergranular space by a screw dislocation mechanism with a given helicity sign. One can identify similar forms in a number of images where the juncture of grain intergrowth is clearly observed (Plate 9i,j). The vicinal forms are scupltural figures pyramids, typically occurring in slow growth. They were first discovered on the faces of the principal rhombohedron of quartz crystals (Verma 1953). The formation of vicinal forms is mainly due to the phenomenon of crystal intergrowth. The spiral nature of vicinal form structure has been clearly established (Lemmlein 1973), and the concentration of vicinal forms commonly observed on the splits of quartz grains is considered as a criterion of an active surface enabling the intergrowth of a quartz grain according to the spiral mechanism. During the ageing of sandstones, the grain boundary migration is observed, stimulated by forces of different nature, among them internal stresses. This process may be regarded as recrystallization in itself. The distinction between recrystallization, grain growth and ageing is not by any means always clear (Van Bueren 196o). During intergrowth, free energy usually diminishes with the decrease of the total area of boundaries due to the growth of grains free from dislocations, which is accompanied by the formation of an equilibrium polygonal or even monocrystal rock, quartzite. To illustrate the data discussed above, we have produced correlation diagrams of porosity and permeability with authigenic quartz in oil- and water-bearing reservoir rocks of the Sahara Plate (Fig. 3.18a). It is obvious that these characteristics are significantly modified as a result of agglomeration of the sandstone in the early phases of diagenesis. With continuation of agglomeration and the consequent reduction of porosity and permeability the water- and oil-bearing capacity of reservoirs should decrease. However, in some advanced stage of diagenesis the processes of agglomeration were probably decelerated but not completely stopped due to filling of the traps by hydrocarbon, which resulted in neutralization of the active centers on the surface of quartz grains by organic molecules. 3.6
Conclusions The main distinction of a closed system, which is a state of inner equilibrium, from a system open to flows of matter and energy (natural systems are such systems) is its behavior in time. Time is a"forgotten" parameter which plays in many geological processes a directing and determining part in the tendency for geological systems to achieve a more stable state. Sandstone transformations can be related to a range of irreversible processes of self-organization that occur under conditions far from equilibrium and are spread along time. Such processes play a constructive role in physics, chemistry and biology (Prigozhin 1985). The appearance and development of coherent structures and the processes of self-organization are quite likely possibilities in open geological systems. In such systems time becomes of primary importance. The main factors controlling a sandstone compaction are: shales diagenesis and subsequent overpressures, the presence of evaporite formations which permit heat loss due to their high thermal conductivity, and early quartz overgrowths which resulted in a solid sandstone skeleton before Mesozoic subsidence. These overgrowths have impeded further reservoir compaction. Mapping such early quartz overgrowth devel-
13o
Chapter 3 • Main Factors of Reservoir Compaction
opment, other geological parameters being equal, is to be recommended as a reliable criterion for petroleum prospecting in Saharan basins. Silica sources are very variable. All sources considered in the literature may have contributed to sandstone silicification, i.e. pressure solution, precipitation from subsurface waters, diagenetic transformation of clay minerals, conversion of interlayered clays to illite, alteration of feldspars and volcanic fragments, precipitation from silicabearing deep waters. However, none of these mechanisms, or probably even several of them, could account for all the authigenic silica observed in sandstones (quartzites), in particular in Saharan basins. In this respect the described theory of solid phase transformation of silica presents an alternative and comprehensive mechanism for sandstone consolidation. During surface transport sand grains obtain more or less isometric shapes and active dislocation surfaces. Grain aggregate transformation occurs over a wide range of time and appears to be a solid-state phenomenon. At depths ranging from 15o to 5 ooo m we observed solid-state grain aggregate transformations: increased grain packing, increase in grain contact areas and authigenic quartz contents, development of orientated low-energy facets on grains, twinning under Esterel and (later) Dauphin~-Brazil rules, gradual decrease of porosity and permeability, and finally junction of grains into a dense quartzite rock. This aggregation-agglomeration occurs through diffusion, sliding, screw dislocation and face-to-face movement of big angle borders, these processes being irreversible. The impetus behind this transformation is the decrease of surface energy, and in open geological systems such processes of self-organization provide a comprehensive mechanism for silica diagenetic evolution in sandstones.
Chapter 4
Reservoir Decompaction and Formation of Accumulation Capacity (in Secondary Porosity) of Reservoir Rocks
4.1 Decompaction Due to Solution of Binding Compounds of Sandstones (Carbonates, Silicates, Aluminosilicates) 4.1.1 Relation in Time and Space (as a Function of Temperature) of Organic and Mineral Diagenesis After Formation of Reservoir Properties The carbon dioxide produced during decarboxylation of the organic matter is not the only factor responsible for the formation of secondary porosity, although on a regional scale in the typical sedimentary basins it may explain a certain portion of the secondary porosity volume. The representative petrographic data, e.g. in the Cambro-Ordovician sandstones of the Hassi Messaoud field, show that the dissolved feldspars here account at least for 5% in volume. The amount of acid necessary to dissolve the feldspars depends on their mineralogical composition and on the type of dissolved aluminium. For this purpose each mole of mineral dissolved requires at least 1 mole of protons. In the sandstones, carbonate cementation took place probably because of the low solubility of CO2 in the formation waters migrating upwards. Taking into account the solubility of CO2 and the distribution of the organic acids we may conclude that the importance of these acids for the formation of secondary porosity has been high when they formed only locally. Other possible sources of the acids required for the dissolving reactions could be the inverse reactions of leaching (MacKenzie and Garrels t966) in the clays and pyrolysis of organic carbon by oxygen contained in the waters as well as the synthesis of organic acids from kerogen (Hoering 1982).
4.1.1.1
Origin and Role of Organic and Carbonic Acids in Diagenesis The analysis of numerous brines of the Saharan fields during exploitation studies revealed evidence of notable quantities of carboxylic acids (Table 4.1), the most frequent ones being monofunctional acetic and propionic acids. The respective brines are characterized by a notable alkalinity controlled by the presence of short-chained organic acids. Experiments on the solubility of aluminosilicates like feldspars, zeolites and clay minerals show that the concentrations of alumina in solutions may be raised considerably by bifunctional organic acids (Carothers and Kharaka 1978). The extent of this effect depends on the pH, with the greatest increase of solubility being encountered at pH = 5. The increase of porosity during progressive burial of the sediments is limited by the ability of aluminium to be transported by the interstitial fluids. In the ab-
132
Chapter 4 . Reservoir Decompaction and Formation of Accumulation Capacity
Table 4.1. Concentrations of short-chain aliphatic acid anions in formation water from Saharan basins Basin
Well
Oued el-Mya OCT-1
Age
mformation (°C)
Acetate (mg 1-I) Propionate (mg 1-1)
Ordovic.
105.6
72.4
HAL-1 GLNE-5 RN-5 MGD-1 EKN-t BEL--1 KB-4 Ni-5 HRS-1 QNR-1 GT-1 HR-81 HR-104
Ordovic. Early Devon. Early Devon. Early Devon. Silur. Silur. Carbonif. Carbonif. Triassic Triassic Triassic Triassic Triassic
98.9 107.8 100.0 104.4 77.0 78.0 79.0 101.7 82.2 71.1 71.0 96.0 99.0
116.2 214.5 261.0 123.1 254.7 474.5 304.2 87.6 1t .4 7.3 4.4 6.1 4.4
0.2 26.7 30,4 16.2 21.4 38.7 6.0 5.4 0.3 0.2 0.3 0,1 0.2
Ghadames
MDR-tbis NZ-1 Rbq STAH-1 STAH-3 MRK-t MRK-16 ELB-1
Cambrian Ordovic. Cambrian Early Devon. Middle Devon. Early Devon. Middle Devon. Triassic
119.0 100.0 90.0 122.0 110.0 115.0 110.0 76.0
2.2 4.6 12.4 27.5 t 29.5 84,7 229.4 47.0
2.7 6.0 27.8 16.5 27.0 12.4
Illizi
TFT-203 AMA-5 ONT-102 RCL-12 TG-37 ZR-188 EAL-16 TG-16 DL-177 DL-241
Ordovic. Ordovic. Early Devon. Early Devon. Early Devon. Middle Devon, Middle Devon. Middle Devon. Middle Devon. Carbonif.
82.0 80.0 93.0 82.0 85.0 84.5 71.6 82.0 62.3 41.7
294.8 t 76,8 127.6 272.2 184.6 375.1 84.7 211.2 46.5
29.2 14.6 25.4 18.4 21,5 41,9 8.7 14.8 4.9
4.9
1.15
-
sence of complexing agents like the organic acids the akuninium concentration in the pH range of 3-8 is only about ] mg 1-~ (Siebert et al. 1984). Because of this, any component dissolved from a mineral will be precipitated in the immediate vicinity of the respective grain, rather leading to a reduction in porosity than to an increase. The increase of the mobility of aluminium caused by the bifunctional acids contributes to the transport of atuminosilicates as well as of carbonates. In an inorganic-organic system other organic solvents may also be active during diagenesis. In the brines of the fields one frequently observes a discrepancy between the dissolved carbon in general, the amount of carbon in the form of bicarbonates and the carboxylic acids. It should be pointed out that in these brines also phenols were detected. Evidence of the origin of the organic acids from kerogen comes from the analysis of the kerogen itself (Mason 1983). Experiments show that the aqueous thermal degradation of kerogen at moderate temperatures of loo °C will produce different quantities of bifunctional acids depending on the type of kerogen present (Vitorovic 198o). The third type of kerogen with a tow oxygen content generates the largest quantity of organic acids. The maximum formation of organic acids takes place in the tempera-
4.1 • Decompaction Due to Solution of Binding Compounds of Sandstones
133
ture range of 8o-120 °C when they form a complexing agent for aluminium as well as a blockage for the alkalinity of the formation waters. 4.1.1.1.1
Minerol Oxidants Mineral oxidants react with the organic matter of the deposits. During, e.g., the transformation of smectite to illite, the reduction of iron plays a certain role (Fig. 4.1). Ferric iron is considered a strong oxidant especially during early diagenetic processes (Berner 198o). Sedimentary rocks contain also numerous other mineral phases with an oxidizing potential. Among these we observe clay minerals, oxide phases on the surface of mineral grains as well as polysulfides. A number of factors show that during diagenesis there will be interactions between the clay minerals and the organic matter.
Diagenesis of Clayey Matter. We know that at lower temperatures the stability of the organic acids is controlled by bacterial activity. As the temperature rises the bacteria will perish, either as a result of the temperature increase itself or because of the toxicity of phenols derived from the organic matter. The high-temperature limit of the stability of the organic acids is controlled by thermal decarboxylation. On the whole,
I00
0
4- -A- 11t-~-
80.
4-
Onordered phase
4-
Little-ordered phase
4k-
Well-ordered phase
-A- 41:-li-l-41_ 4i-4k~s~-~
20
.--
~k-
4-
44-4t4- 4-~4-A- ~5- 4 - 4 4 - -~- 4 - 4 -
o 60.
; 40
~6 40.
60
!
4~p~
~i-
8
g ca >
-&-& ~ _ ~ -A4- 4-
4-4d,-
eO.
Ca Ca
_~.
.6-A_4-
N
"0
(~ 20"
: 80 -&
4Y,-ac~-N o
4X-
-6--6-4-
494% u~
o
~
, o
, ~
, o
. . . . . . . m
o
m
o
, tn
o
,
,, , ,
m
o0~r--~
..... m
, ~.
m
, ¢-1
Fe2+/Fe ~+ (%)
Fig. 4.1. Transformation of mixed-layer clays of illite-smectite types (unordered) into ordered phase of vermiculite-chlorite type vs. transition of Fe 2+ into Fe 3+. Paleozoic sediments of the Illizi Basin
134
Chapter 4 . Reservoir D e c o m p a c t i o n and F o r m a t i o n o f A c c u m u l a t i o n C a p a c i t y
short-chained acids become unstable above 2oo °C. The maximum concentration of organic acids is encountered at the same temperature as the ordering of the mixedlayer minerals of the smectic-illite series (Hower 1981). The transformation of disordered into ordered phases of the allevardite type is accompanied by a liberation of iron from the octahedral position in the lattice structure (Hower et al. 1976). Whether the iron is set free and then reduced or whether it is reduced within the structure of a clay mineral to become split off then as ferrous iron (Almon 1974), an electron transfer must have taken place. Oxidation of the organic matter dispersed in the shales and the intense later drop in valence of the iron appear to be the main processes. Chemical analyses of shales from the Saharan basins show moderate to elevated iron contents. It is thus easy to understand why paragenetic minerals enriched with ferrous oxide, like chlorite, siderite and chamosite, formed during diagenesis became so ubiquitous under these conditions.
Fig. 4.2. Schematic generalized distribution of clay minerals on local scale: chlorite local abundance increases towards the sand/shale contact; immediately above is a zone of enhanced porosity; a zone enriched in kaolinite overlies the zone of enhanced porosity
Clay mineral species (%) 0 50 100
~ T 3000 -
u 'ttio n
3100
-
j~~/
Zone of local enhanced porosity
3200 -
O °o ° ° J ~ Zone of anomalous diagenetic chlorite and Fe-rich carbonate
3300 -
=:
:
: :
: • :
lllite Mixed-layer clays Kaolinite Chlorite
4.1 . Decompaction Due to Solution of Binding Compounds of Sandstones
135
Figure 4.z illustrates the typical mineral zoning as a function of depth observed in numerous petroleum-bearing formations of the Saharan Platform. The correlation between porosity and newly formed minerals shows that porosity increases towards the shalesandstone contact. Within the section the content of ferrous authigenic minerals like chlorite, chamosite and siderite is part of the general mineral zoning which also affects minerals such as kaolinite, illite, chlorite and the mixed-layer clays. Such a trend is developed especially where shales act as a source for the iron as well as for the organic solvents. The kaolinitic zone overlying the section represents the limit for the transport of aluminium resulting from the dissolution of the matrix particles of the rock. 4.1.1.1.2
Source of the ]ieids With the increase of temperature during burial the organic matter is subjected to decarboxylation, a process that also takes place in the same way under experimental conditions (Robin and Rouxhet 1978; Rouxhet et al. 198o; Johns 1982). The quantitiy of CO2 produced by the decarboxylation of organic matter in a shale of defined composition depends on the amount and type of the organic matter present. The type kerogen III is the organic matter most prone to the production of COv This type of kerogen which is thermally least mature may contain up to 25% oxygen (Tissot and Welte 1985). According to infrared spectroscopy data (Robin and Rouxhet 1978) the carboxyl groups may contain about 25% oxygen. It is conceivable that an average of 6% of the oxygen in the kerogen may be transformed by decarboxylation into COz. Comparable values were obtained experimentally (Harwood 1977). The amount of secondary porosity formed under the influence of the liberated carbonic acid depends not only on type and concentration of the organic matter but also on the shale/sandstone ratio in the succession considered (Table 4.2). When the dissolution takes place by the above-mentioned directions the amount of CO2 produced by decarboxylation from the organic matter in the shales is large enough to account for a moderate amount of secondary porosity by dissoMng carbonates and feldspars according to the following formulae: CaCO 3 + CO2 + H20
> Ca 2+ + 2HCO;
zKAISi3Oa + zCO2 + 11H20---9 AI2SizOs(OH)4 + 2K++ 4H4SiO4 + 2HCO~ . A balance calculation for different shale/sandstone ratios and types of organic matter shows that in many basins the CO2 formed by decarboxylation of organic matTable
4.2. Secondaryporosity (%) from decarboxylationof kerogen (after Bjorlykke1984)
Type of organic matter
Oxygen
Oxygen in -COOHgroups
Volume of dissolved calcite
Volume of dissolved feldspars 4.4
Ill
25
25
1.6
El
75
13
0.5
1.4
I
I0
7
0.2
0.5
136
Chapter 4 • Reservoir Decompaction and Formation of Accumulation Capacity
ter is not sufficient to explain the entire secondary porosity. In these cases it has to be assumed that the dissolution of the carbonates and feldspars took place under the influence of other organic acids of the acetic and oxalic type.
Other Functional Groups in Kerogens. The oxygen in the organic matter is contained not only in the carboxylic groups but also in other functional groups like ethers, ester complexes, phenols and carbonyl (Rouxhet et al. 198o). From these groups the oxygen is usually split off during higher levels of thermal evolution of the kerogen, probably in the form of COz and H20, although the types of these functional groups during diagenesis are not much known. Inverse Reaction of Leaching. Leaching may take place along the following path with the complementary acid possibly forming during burial (MacKenzie and Garrels 1965): silica + clay + cations
, ~ alumosilicates + water + H +
During the reactions leading to chlorite or an authigenic aluminosilicate, protons in particular will be present in excess because abundant hydroxyl groups form part of the brucite sheets of the chlorite. Numerous successions from the Saharan basins offer clear evidence of a decrease in the chlorite content of the reservoir rocks at the expense of the kaolinite. From this we may suspect that a reaction splitting off the acid may have taken place under the effect of a mineral transformation (Boles and Frank 1979). It should be noted that contrary to commonly held notions the leaching to which the shales are subjected represents rather a trap than a source for these cations.
Mineral Oxidants and Aqueous Pyrolysis Reactions. In the organic matter, concentration and distribution of oxygen in the functional groups clearly limit the amount of C02 produced by simple elimination. Thus, in a certain type of organic matter where the carbon may react with oxygen from an extraneous source, the amount of C02 produced may be much larger than if only the primary oxygen contained in the kerogen were available. In the reducing conditions of deeper horizons it is difficult to establish what had been the source although the oxygen of the H~O could be a potential source. So far it has not become completely clear whether such an aqueous pyrolysis has taken place, although in his experimental maturation of organic matter Hoering (1982) has shown that the exchange of hydrogen between organic matter and H20 is a rather frequent process. Mineral oxidants like ferric iron in the shales have been considered as potential oxidizers of the organic carbon to form CO2 (Eslinger et al. 1979). Again, an extraneous source of oxygen may be available: CH20 + 4Fe3+ + H20
'~
C02 + 4Fe2+ + 4H+
It should be noted that according to this formula 1 mole of protons is produced for every mole of iron reduced. A reaction of this type undoubtedly is of great importance for bringing the materials into solution. The uncertainty, nevertheless, is great as we do not know at a given time how much ferric iron was in the detrital authigenic clays and, consequently, what quantity was available for the reaction with the organic matter.
4.1 . Decompaction Due to Solution of Binding Compounds of Sandstones
~37
Another potential mechanism for the oxidation of kerogen would be the reduction of sulfates or sulfur: SO~- + 2CHaO + H +
> H S - + 2HzO + 2CO a
.
The notable contents of soluble sulfates in the formation waters of the Saharan reservoirs underline the necessity to evaluate the role of this mechanism during the oxidation of kerogens and thus in the increase of the secondary porosity. Role of Meteoric Water. In certain basins the supply of meteoric waters may lead to
the dissolution of cement in the sandstones and to an increase in the secondary porosity. Such a situation appears to be developed in particular in the Illizi Basin where in the outcrop zones the supply of surface waters and their eventual discharge in buried horizons are important phenomena. 4.1.1.1.3
Origin of Organic Species in Reservoir Waters Aliphatic acids, the most important constituents of reservoir waters, are found in the formation waters, organisms and sedimentary rocks in such high quantities that they may probably represent a source of petroleum (Eisma and Jurg 1967; Hunt 1979). The fact that the concentrations of dissolved organic acids in shallow subterraneous waters and formation waters with temperatures below about 80 °C are relatively low shows that the organic species, whatever their origin, and in particular that of the acetates, whether formed by the organisms or from the organic matter in the rocks, have been degraded by bacteria. This degradation of the organic species and their transformation to CO 2 and hydrocarbon gas probably stops at temperatures of about 8o °C (Davis 1967; Carothers and Kharaka 1978). The maximum concentration of acid (organic) anions occurs in the reservoir waters at temperatures above 8o °C (Fig. 4.3, Table 4.1). The most probable source of these acid anions is the thermocatalytic degradation of kerogens containing an excess of aliphatic acids (Eisma and lurg 1967). The quantity of oxygen especially in type III kerogen is high, the atomic O/C ratio being in the order of 3o%, and appears to be sufficient to explain higher concentrations of the acid anions encountered in the study area. High concentrations of acid anions are also encountered in the gas deposits of the Sahara and in those where liquid hydrocarbons, another potential source of organic acids, probably no longer exist. Kharaka et al. (1983) have shown that a defined portion of the natural gas of the sedimentary basins formed by decarboxylation of these acid anions. A certain portion of these acid anions may have acted as a source for liquid hydrocarbons. Oxygenated as well as nitrous and sulfuric compounds have been detected in small quantities throughout the oils and in highest concentrations in the lubrifactants and sedimentary facies (Hunt 1979). Oxygen is present mainly in the resins, asphaltenes and waxes of high molecular mass where its content represents more than 5% of the total mass of the resins. The proportion of heteroatomic heavy compounds usually is below lo% of the oil, but it nevertheless increases with decreasing oil density, reaching values above 5o% in the bituminous sandstones (Tissot and Welte 1978). The anions of the organic acids may be formed through thermocatalytic or bacterial degradation of resins and asphaltenes as in the case of the kerogens.
138
Chapter 4 . Reservoir D e c o m p a c t i o n and F o r m a t i o n o f A c c u m u l a t i o n C a p a c i t y 500 -
500 Zone I
[
e o
o e
Zone 2
Zone3 300
e
300 -
I .o
oN;
!:.
~.% • :V:. -.V " 0%\.
• !'."
o
.1"
"i ..
-a 100
.
it
i
.!o-.: [~'o
t
-. % "." ol
-
"
100 -
i
"
"
"~ 50
50-
~/.o!
Oo/oil
o
.
o
uo
¢p:
ol
FI
20
/o 0 70
50 a
" 20-
I o Triassicreservoirs I • Paleozoicreservoirshighlyuplifted by Hercynianorogeny Paleozoicreservoirsmoderatelyupliftedor non-uplifted I , , I 0 90 110 130 150 170 190 Formation temperature (°C)
Triassic reservoirs
]~ Aliphaticacidanions 3000
.
H2CO-
Paleozoic reservoirs and Triassic reservoirs with immediate contact with Paleozoic (S, D) source shales
r 3 000 I I
3-
~ 2000
~
I- 2000
1000
I000
u
0 80
b
[H, 0
100 120 140 160 80 100 120 140 160 Formation temperature (°C)
Fig. 4.3. Distribution of aliphatic acids and total bicarbonate in oil-field waters, Saharan basins. a Concentrations of aliphatic acid anions (C2-C5) in formation waters from different Saharian basins, in various geological situations. Note that: 1. the highest concentrations are between 80 and lao °C and decrease with increasing temperatures, 2. the lowest concentrations are related to Triassic reservoirs (organic matter-poor shales and low level of kerogen maturation), 3- intermediate concentration are characteristic for paleozoic reservoirs highly uplifted by Hercynian orogeny (involving relatively moderate temperatures and conditions less favourable for organic acides generation), b Average concentrations of C~-C 5 aliphatic acid anions and total bicarbonate in oil field waters for Triassic and Paleozoic reservoirs. All samples with temperatures lower 80 °C have been averaged together; for samples with temperatures higher than 8o °C, averaging is over 20 °C intervals
4.] - Decompaction Due to Solution of Binding Compounds of Sandstones
139
4.1.1.2
Summary of Organic and Mineral Diagenesis The diagrams presented in Fig. 4.4 are a schematic integration of the organic-inorganic diagenesis which may be used for establishing the observed succession of transformations by which porosity increases. The first stage corresponds to eogenesis and immature to semi-mature mesogenesis (2 and 3) prior to the migration of the hydrocarbons. Towards this phase in the upper zone already certain overgrowth rims of quartz are encountered and the sandstones undergo some silicification. The grains of silica with carbonate inclusions for the time being remain unaffected. The appreciable overgrowth rims appear at the end of this stage prior to the deposition of secondary clayey (and mixed-layer) cement. The cement then occupies a certain volume but by no means all of the primary space of the remaining pores. Schematically, a complex kerogen molecule is made up of an aromatic core surrounded by an aliphatic layer, the mono- and bifunctional carboxylic acids as well as the phenols being tied to the peripheral portions of the molecule. As temperature increases with growing burial the bonds with these peripheral functional groups are broken. As the reactions towards ordering of the crystalline structure of the mixedlayers of the smectite-itlite type in the shales (100-110°C) coincide in time, temperature and space with the formation of the highest concentration of organic acids at 8o-12o °C in the brines of the petroleum-bearing formations we may propose a possible mechanism for the origin of the bifunctional organic solvents. It is clear that the mineral oxidants undergo reduction and that the kerogens concurrently are oxidized, which together with the thermal degradation of the kerogens represents an efficient process generating groups of peripheral organic acids. Dissolved in water these acids may be transferred from the source rocks into adjacent reservoir rocks by waters furnished by compaction and transformation of shales. In the second stage (4), corresponding to mature mesogenesis, the previously generated organic solvents have already migrated into the sandstones. In the lower part of this interval the organic bifunctional acids form metallo-organic complexes with the aluminium contained in the clayey cement and the feldspars transporting the products of this reaction upwards above the sandstones. This removal of aluminium from the argillaceous cement (and the mixed-layer minerals) contributes to the formation of secondary porosity. At this stage a defined dissolution of feldspars might have probably taken place in the presence of organic solvents. In the fluid-rock system kaolinite starts to become deposited in the open pore space as the chemical equilibrium changes and the organic aluminium complexes become unstable. As one moves upward in the sandstone succession the abundance of kaolinite increases stratigraphically, this mineral becoming generally more frequent and occurring in greater quantities in the upper part of the section. Thus, the transfer of material resulting from the complexification of aluminium in the lower zone into the upper zone appears to be efficient enough. At the onset of the second stage the kerogen molecule is made up of an aromatic core surrounded by an aliphatic belt. At this stage the peripheral organic acid groups have already been split off and transported away. As temperature increases the bonds between the aliphatic hydrocarbons and the aromatic core are broken and the resulting hydrocarbons migrate into adjacent reservoirs.
14o
Chapter 4 • Reservoir Decompaction and Formation o f Accumulation Capacity Mineral transformation and textural development
Quartz intergranular pressure solution and quartz overgrowths 1
2
Open-packed clean to muddy quartz sand, fresh feldspar and undeformed muscovite, with interstitial illite. (Minor amounts of glauconite, biotite, chlorite, zircon and rutile not shown)
3
iQ
solution 1
Sequences of stages in the development of quartz pressure solution
Q Ca,
A CQ.
Rhombs of calcite, dolomite or siderite attached to and replacing quartz, kaolinite and any remaining feldspar.Authigenic pyrite, mainly cubic
B Rest of pore space filled by large crys~cals of poikilotopic calcite. Further growth of siderite and pyrite, some of which changes to marcasite
D ~
Quartz euhedra grow into pore spaces. Feldsparsmostly destroyed. Kaolinite occupies much pore space.Muscovite frays, forming sericite and kaolinite. (lllite in~reasesin crystallinity
Schematic illustration defining detrital quartz (D.Q.) quartz cement (Q.C.) and overlap quartz (O.Q.)
Different generations of quartz overgrowths
D
e
t
r
e silica
i
t
~
Silica dissolved by pressure solution to give sutured contacts with reduction in pore space.Feldspars partly dissolved. Muscovites bent
~
Schematic diagram of major features of quartz overgrowths
Key Ca =Calcite Ch =Chert D =Dolomite F = Feldspar I = Illite
K =Kaofinite Q =Quartz M = Muscovite Se = Sericite Ma =Marcasite Si = Siderite Po = Pore space Py = Pyrite
Fig. 4.4a. General diagenetic effects within Saharan sandstones: quartz intergranular pressure solution, quartz overgrowths, and mineral transformation and textural development
4.1 - Decompaction Due to Solution of Binding Compounds of Sandstones Mineral authigenesis and pore system evolution
~4~
Kerogenevolution
Non compacted sandrock entirely composedby detrital minerals B
i
t
Mixed-layerclaysgradual conversion.Kaolinitedeve~ lopment. Detrital quartz coating mainly by authigenic chlorite.Slight quartz overgrowths.Some non-ferroancarbonate cement precipitation or replacement.
3
='~ : " q ~ /~7;
)
~ ~ / ,
4 ~
~
/
\
5 ~ , ~ i ~
'
Predominanceof authigenic illite and chlorite as most stable clay minerals. Intensivecarbonate (including ferroan)dissolution. Quartz pressure solution/precipitation. Somepossiblegain of secondaryporosity
Legend Nonferr0an [ ~ - - ~ Mixed-layerclay!~ ~ " Icarbonate Mica ~ Ankerite ~~ Kaolinitedikite I-~-'~--I CNorite [---'--] lllite ~ Feldspar Chert ~ Quartz
Silica Ferroan carbonate Pores
~
!
Furtherconversion ofmixed-layerclaystowardsilliteauthigenesis.Authigenickaoliniteprecipitation,Authigenicchloritedevelopment,Mixed4ayer'sordering corrensite,allevardite.Feldspardissolution.Significant quartzovergrc~c~thsdevelopment,chertdissolution.Carbonatecementdissolution/precipitation.Somesecondaryporosity development. Significant reductionof mixedlayerclaysandtheirfurtherordering~ kalkberg.lmproving x ~ crystallinityofilliteandchlorite. Emergenceofdikite.Feldspar dissolution.Intensivecarbonate cementleashing(especiallydolomite)andpredpitation,aswell asferroancarbonates.Quartz pressuresolutionandothersilicatedissolution.Further(intensive)porosityenhancement.
u
,acidss A--Z
! Organic acids+ liquid hydrcarbons
0
!
Fig. 4.4b. General diagenetic effects within Saharan sandstones: mineral autigenesis and pore system evolution, and kerogen evolution
14a
Chapter 4, Reservoir Decompaction and Formation of Accumulation Capacity
In the third and final stage (5), corresponding to super-mature mesogenesis, after migration of the hydrocarbons only the aromatic core of the complex kerogen molecule will remain. A particularly notable modification of this stage is the deposition of defined quantities of illite and chlorite as well as the intense dissolution of quartz during pressure solution. With continuing burial of the reservoir rocks temperature will increase further and the organic fluids trapped in the reservoirs evoive in situ. The thermal decarboxylation of hydrocarbons results in methane (CH4) as well as in CO2. The thermal degradation of hydrocarbons also forms methane to leave the respective residues in the former contact zone between water and oil as "dead oil". During the same phase the silicate grains together with other components of the cement undergo strong dissolution. The CO~ generated by thermal decarboxylation of organic matter probably increased in many cases the concentration of C02 to a level where its exceeds the concentration in the carbonates. The weak acids present in a system dominated by carbonic acid (CQ) attack the carbonates, creating a higher porosity. In zones without an abundant source of organic fluids there is only a minimal late dissolution of cement. It is obvious that the above-mentioned mechanism is not the only one possible, but its integrated nature makes it attractive. 4.1.1.3
Formation of SecondaryPorosity During Diagenesis Secondary porosity results from chemical, physical, physico-chemical, biochemical and biophysical processes which lead to leaching and compaction of reservoir rocks as well as to the formation of fissures and open cavities. It may take place in the sedimentary shell: (1) under conditions of accumulation of sediments until they undergo greater burial (eogenic porosity), (a) at any depth of burial above the zone of metamorphism (mesogenetic), and (3) during uncovering of the sediments in outcrop after the final phase of burial (telegenetic). Secondary porosity may form in sandstones of any type of mineral composition, textural peculiarity and age. It is encountered especially in sandstones which have undergone relatively long-tasting deep burial and thereby lost their primary porosity. In old sandstones the largest part of the secondary porosity is derived from mesogenetic leaching of carbonates like calcite, siderite and dolomite (Plate 11) as welt as of feldspars (Plate 12) and clays. This decarbonatization eliminates the sedimentary carbonate compounds and the diagenetic carbonates present as cement or replacements of other minerals. A large part of this mesogenetic decarbonatization may be ascribed to the decarboxylation of organic matter during its maturation in shales adjacent to the sandstones, a process that leads to the formation of carbonic minerals and feldspars. Decarbonatization culminates during mature mesogenesis, i.e. at a stage where it notably surpasses carbonization. Because of this, more secondary porosity will form when the primary porosity has been lost. Fissures and non-reducible tamellar porosity evidently assure a sufficient access for the fluids resulting from decarboxylation so that the process of leaching may be initiated even in sandstones of low permeability. An enormous amount of carbonates will be displaced upwards by the solutions away from the diageneticatly mature sandstones to become deposited, at least in parts, in imma-
4.1 . Decompaction Due to Solution of Binding Compounds of Sandstones
143
Plate 11, Carbonate cement in sandstones which suffered intense solution by acids. Note that the carbonate cement in photo c undergoingdissolution is of secondaryauthigenic origin
ture to moderately mature sandstones. In terrigenous sediments undergoing continued burial a large portion of their carbonates will be transformed in a cyclic manner to become displaced further upwards leading to an environment of the sandstones in carbonates at shallower depths. The primary migration of hydrocarbons normally starts when the secondary porosity forms, as during maturation of the organic matter the main phase of hydrocarbon formation takes place just after the culmination of the decarboxylation. This close association of hydrocarbon sources and reservoirs in time and space favours the accumulation of hydrocarbons in the secondary porosity.
144
Chapter 4 • Reservoir Decompaction and Formation of Accumulation Capacity
Plate 12. Siliceouscement (feldspar) undergoing intense dissolution by acids (mesogenetic)
4.1.2 Thermodynamic and Stoichiometric Regime of Formation of Carbonic and Organic Acids and Their Role in the Establishment of Secondary Porosity in Reservoir Rocks In Table 4.3 the calculated pH values for the organic acids for which dissociation constants are available in the literature are compiled. It shows that the concentration of H+-ions which 1 mole of an organic acid can produce ranges from lo -144 mole 1-1 for chloroacetic acid to lo -2.52mole 1-1for trimethylacetic acid. Carbonic acid may furnish
4.1 . Decompaction Due to Solution of Binding Compounds of Sandstones
145
Table 4.3. Dissociation constants and approximate pH values for some organic acids
Acid
Dissociation constant K1
Concentration of H +
Approx. pH
K2
Formic
10 -3.75
I 0 -]88
1,88
Acetic
10-4Js
10 -1.38
2.38
Propionic
10 -487
10 -2,49
2.44
n-Butyric
10 .4.82
10 -2'41
2,41
iso-Butyric
10 .4.82
l 0 -2"41
2.37
Trimethylacetic
10-5.05
10 -2.52
2.52
2.37
Diethylacetic
10
4.75
] 0 ~'37
Chloroacetic
10 -2.87
] 0 -~44
1.44
Glycolic
10 -3.83
] 0-t 92
1.92
Oxalic
10 1.3
] 0 -4.82
l 0 -0"65 - 10 -2"14
0.65 - 2 , 1 9
Carbonic
10 63s
i04o2s
10 4"~6 -- 10-s~
3.2
Table 4.4. Calculated pH values for some organic acids and their Acid dissolution efficiency Carbonic
-5.1
Approx. pH
[H+I
_>3.2
6.3 x 10-4
6x
Efficiency
Propionic
2.44
3.6 x ] 0 4
Acetic
2,38
4.2 x 10-3
7x
Formic
1.88
1.3 x 10 -2
20 x
Oxalic
0.65
2.2 x 10 t
350 x
10-3"16--10 "-Sa moles of H+-ions. Thus, on a stoichiometric basis organic acids may furnish up to three orders of magnitude more H+-ions than carbonic acids. The dissociation constants at z5 °C and the calculated pH values for a number of organic acids either found experimentally by Surdam et al. (1982) or recognized in underground waters by Carothers and Kharaka (1978) are represented in Tables 4.3 and 4.4 which illustrate that the capacity of organic acids to furnish H+-ions is up to 35o times higher than that of carbonic acid. Anions are another factor controlling the stoichiometric efficiency of organic acids. In the case where carbonic acid acts as a solvent for carbonate cement the carbonate ions in solution result from the solvent (acid) as well as from the mineral dissolved. In contrast to this, for acetic acid as the solvent, at least half of the carbonate ions produced during dissolution will be substituted by acetate anions; calcium acetate and not merely calcium carbonate alone is also a product of the dissolution. Calcium acetate is three times more soluble in water than calcium carbonate (Table 4.5). This actually favours the stay of the calcium ions in solution, hindering their precipitation in the form of calcium carbonate. Thus, the higher solubility of the salts of a
~46
Chapter 4 • Reservoir Decompaction and Formation of Accumulation Capacity
Table 4.5. Solubilityof calcium T(°C) Calcium acetate salts corresponding to acids (acetic acid} (in g 1-~) at different temperatures. (after Seidel11965) 0 374.0 10 360.0 20 347.0 25
342.0
30
338.5
Calcium formate (formic acid)
Calcium carbonate
161.5
0.081
-
0.070 0.065
166.0
0.056 0.042
40
332.0
170.0
0.044
60
327.0
t75.0
0.038
80 84
335.0 338.0
179.5
-
85
329.0 311.0
-
-
184.0
-
90 100
297.0
Table 4.6. Comparison of Conversion Gibbs' free energy (AGr°) for carbonic, formic and acetic acids (A1{° represents reduction in the volume of solid phase K-spar > illite products when compared to solid phase reactants) K-spar ~
kaolinite
Anorthite--3 kaolinite
AV~
Converting agent
AG~ (kcal tool-1)
-15.3
H2CO3 HCOOH CH3COOH
+102.6 +95.6 +23.73
-12.5
H2CO3 HCOOH CHsCOOH
+22.98 + 15.8
--1.26 H2CO3 HCOOH CH3COOH
-4.28 -15.0 -165 -36.9
certain acid represents an important factor for intensifyingthe dissociation of the same acid on its own, whereas this is not the case for carbonic acid on its own as suspected by other authors (Lundegard et at. 1984). In Table 4.6, which was compiled on the basis of data on molar volumes and Gibb's free energy values published by Robie et at. (1979), implications for the conversion of certain silicates, for their potential for generating secondary porosity and for a certain thermodynamic regime are compiled. The right-hand column in Table 4.6 presents the reaction energy (AG~') for each conversion series at 25 °C. Comparison of these values shows that AG~) for acetic acid is below that of the other acids, indicating a more advantageous thermodynamic regime. There are two ways that dissolution is possible (Fig. 4.5). The first one is based on the generation of CO2 by decarboxylation of organic matter in the source rocks, the second one on the formation of H + by dissociation of organic acids. In Fig. 4.5 the
4.1 - Decompaction Due to Solution of Binding Compounds of Sandstones
t47
Decarboxylation (carbonic acid hypothesis) ".'.'." zone of :...,.. '-" i:emen1:ati0n,.:
613C = f r o m - 2 t o + 6 % 0
8 ~3c = f r o m - 2 0 t o - 3 0 % o
l
?>i-i i i i.i i i i ,.i.i-i 2Ca2+ + CMO2- + COO2- __>2CaCO3
I )
~i13C= from-2 to -12%0
,Sands?~ane i,i-
iiil;iiiiiil;i;i;iil;ii,iiiill
i i.i.i.;.i,>i-.I.I.;C iiiil;;i;i;iiiiil;iiiiil;il;ill I-/.C i->i-i i.>; i i. . . . . . . . . . . . . . .
2 ) HCMO~ --->H + CMO3-
/
/
HCoO~ -~ H+ + COO2-
CaCMO3+ H2C°O3 --->Ca2++ HCMO~ + HC°O~
(~I3C = from -2 to -8%o
.- .z0ne0f .v... i i.ldissoluiion.}i. -Shale-
I
Dissociation (Organic acid hypothesis)
...'Zon~ o f . . . . dernentati0n : i i i sail;ds}0ne { .i.
O ~-tO
C°O2 + H20 --->H2C°O3 Decarboxylation of organic acids form C002 (50-150 °C)
6~3C : f r o m _ 2
to +6% °
2Ca2+ + CMO2- + 2RCOO- -e CaCMO3+ Ca(RCOO)2 insoluble ~/ soluble
/
H+ + CMO~-
CaCMC03 + H÷ ~ Ca2+ + HCMO~ ' ' .'.
' -.-...' .' •
}i.i-.-z0nei0f}i-i }
. d~sso~ut~on- v -, "-S-6ale- -
RCOOH --~ RCOO- + H÷ (/opt.= 80-120 °C) organic acids
Fig. 4.5. Generation of secondaryporosity due to carbonic or organic acids may be distinguished by the 6~3Cvalues of resulting CaCO3cements, as shown by decarboxylation versus dissociation mechanism. CM carbon of mineral (sedimentary) origin; C0 carbon of organic origin
mechanism for the formation of the acid ions (CO~, H +) in the source rocks, their dissolving influence on the carbonate cement of the adjacent sandstones and the precipitation of the dissolved carbonate in the cementation zone are presented schematically. If decarboxTlation were the dominant mechanism then the (reprecipitated) carbonate cement in the cementation zone should be characterized by a lighter 613C isotopic composition, inherited from the C02 derived from kerogen, compared to that from a carbonate cement of sedimentary mineral origin. If dissociation were the dominant process then it should be controlled by the generation of H + and the carbonate cement would be characterized by an 613C isotope composition similar to that of marine carbonates. Furthermore, ions derived from the organic acids and in particular acetates should be encountered in the formation waters.
148
Chapter 4 • Reservoir Decompaction and Formation of Accumulation Capacity
4.1.2.1
Distribution of Carbon Isotope Composition in Carbonate Cement of the Saharan Reservoirs in Time and Space Studies of calcite concretions collected in shallow sandstones (eogenetic zone) have shown a considerable enrichment in 13C,the said concretions being characterized by fairly low temperatures of precipitation. These values are indicative of sedimentation at a slightly deeper water-sediment interface as a by-product of sulfate reduction (Fig. 4.6).
Fig. 4.6. Stable isotope range (613C) for carbonate cements
-I0
-7.5
-5.0
-2,5
0
+5
I
I
I
I
I
I
+I0
o N
I,.I-I
It" .
I
E K
>_'~
r~
f o Triassic carbonate cement • Paleozoiccarbonate cement • Cabonate associated with sulfate
4.1 . DecompactionDue to Solution of Binding Compoundsof Sandstones
149
In the cement of the Saharan sandstones the residual ferruginous sandstones exhibit a wide spread of fi13Cfrom -lo%o to +6%0, values which agree with those of Irwin et al. (1977). The authigenic ferruginous carbonate cement with heavy fi13C (o to 6%0) apparently was precipitated in the zone of bacterial fermentation (immature mesogenesis) whereas the authigenic ferruginous carbonate cement with lighter ~13C (-5 to -t o%0) rather crystallized in the zone of thermal decarboxytation of organic matter (mature mesogenesis). Cements with fi13Cintermediate between these extremes represent a mixture of the two types. Hmvever, the mineralogical and diagenetic situation is much more complicated than that as the fi13C is not only controlled by diagenetic zoning but also by the type of acid present, viz. carbonic or organic, as outlined above. In the zones of semi-mature and mature mesogenesis we can actually distinguish three fields. The first one is characterized by light fi~3Cvalues (-3 to -12%o) characteristic of well-crystallized carbonate cements in particular in the Triassic reservoirs resulting from decarboxylation and dissolution of the cement by carbonic acid. The second field covers heavy 8~3Cvalues (o to +6%0) in the carbonate cement which is also well-crystallized and developed especially in Paleozoic and Triassic reservoirs lying in immediate or close contact with argillaceous Paleozoic source rocks [Silurian (S), Devonian (D), Carboniferous (C)]. The latter situation obviously corresponds to the generation of (essentially) organic acids and their eventual activities during dissociation. This may be explained in particular by the intercalation of the above-mentioned reservoirs with Paleozoic argillaceous beds rich in rather mature matter. Nevertheless, carbonic acid is also present here and its influence on the resulting fil~C has to be taken into account. A third zone may be delineated between the two above ones, being characterized by intermediate 8~3Cvalues of -2.5 to +2%0. The isotopic composition of the carbonate cement is actually the intereference product of a multitude of factors including diagenetic zoning and type of reagent acid. The respective models explaining the assembly of the geological and geochemical aspects of the hydrocarbon-bearing Saharan provinces will be discussed in detail below. 4.1.3 Model of Differential Dissolution and Redistribution of Carbonate Cement with Compaction/Decompaction of Reservoirs in Space as Based on Carbon Isotope Data 4.1.3.1 Triassic Reservoirs of the Northern Triassic Province
Investigations carried out with the aim of establishing characteristics of the carbon isotope composition of the carbonate cement have shown that the Triassic reservoirs overlying the Lower Paleozoic silty-san@ complex (Cambrian-Ordovician) exhibit the following features: heavy carbon isotopes (fi13C= o to +6%0) are associated with initial carbonate cement; in sandstones with a high solution porosity, the pores of which preserve traces of the carbonate cement, the isotope composition of the latter is also hea W (fi13C = o to +6%0);
15o
Chapter 4 - Reservoir Decompaction and Formation of Accumulation Capacity
low porosity is observed in sandstones with recrystaltized basal cement which is dominant in the upper parts of sand bodies and contains isotopicalty light carbon (813C = -2 to -lo%o). The basal cement results from the crystallization of carbonates transported in solution. A rather interesting distribution of the 613C (and of the carbonate cement type) becomes recognizable when we consider the shale beds separating the Triassic sandy reservoirs (in particular A, B and C) from each other (Fig. 4-7)-The sandstones directly overlying the shale horizons are dominated by residual carbonate cement with clear indications of dissolution, high porosity and heavy 613C. In contrast to this, in sandstones (with low porosity) directly underlying the shale horizons (see Fig. 4.9) recrystallized massive cement with a light isotope composition is dominant. This illustrates that the shale horizons separating the Triassic reservoirs play a double role. Firstly, they are the source of the carbonic acid (from decarboxylation of organic matter), the dissolving action of which becomes clearly evident in the carbonate cement of the sandstones directly (or closely) overlying these beds. Secondly, these horizons represent a barrier holding back the dissolved carbonate which is rising from the lower zones under the influence of acid solutions. Under these conditions the reservoirs become blocked directly below the shale layers. There is a peculiar observation: on a regional basis and throughout the Triassic reservoirs there is an isotypicatly heavy carbon in the carbonate cement of the reservoirs in the lower part of the sandy-argillaceous Triassic succession (Lower Series, ReservoirC and Lower Argillaceous-Arenaceous Triassic), i.e. in the lower part of the sandstone beds closest to the Hercynian unconformit?: Along the latter, probably acid solutions enriched in carbonic acid derived from the decarboxytation of organic matter in source rocks of the Silurian and Devonian located to the northwest, southeast and south of the Triassic Basin and the northern part of the Oued el-Mya Basin have been migrating. It has to be underlined that during the later stages of diagenesis of organic matter the migration of the hydrocarbons followed the same routes, i.e. the Hercynian unconformity, from the argillaceous source rocks of the Silurian and Devonian in the above-mentioned areas towards the Triassic (and also Paleozoic) reservoirs in the Triassic Province including the Hassi Messaoud, Ait Kheir fields, etc. Because of their relative instability only few of the organic acids which could have formed in the same Silurian and Devonian source rocks managed to reach the above-mentioned reservoirs. We have to stress the rote played by the diagenesis of clays and organic matter in the Triassic itself in transformation and redistribution of carbonate cement in its own reservoir rocks. Judging from its degree of maturation the organic matter of the Triassic shales has not yet reached the phase of hydrocarbon generation. Their thermal regime lies thus generally outside the main phase of formation and stability of the organic acids (8o-12o °C). The organic matter of the Triassic shales, however, is just at the stage of decarboxylation where it could have furnished a great amount of carbonic acid which would have participated prominently in the dissolution and reprecipitation of carbonate cement in the sandy reservoir rocks with shale intercalations. Judging from the similarity in the variations of isotope composition and porosity resulting from the dissolution and reprecipitation of the carbonate cement the carbonic acid was produced from the organic matter in the Triassic shales in exactly the same way as in the Silurian and Devonian shales. The difference in the role of the Triassic and the Paleozoic
4.1 • Decompaction Due to Solution of Binding Compounds of Sandstones
151
r~
©
o c~
c~
c~
c~
%
8
o
o
t
~o kl.
152
Chapter 4 • Reservoir Decompaction and Formation of Accumulation Capacity
O
o
N
r~ o E
~r M.
4.1 • Decompaction Due to Solution of Binding Compounds of Sandstones
253
:!1 + +.,'i +' ++"5".++ :+' ~,','~l~:+++!~. :P+++++++~+ .~.~
,% ,Nap atle(]
+~
.++
I ++!I
• ' "+
L L +
+ +
+
+°+J
+ + +
+
+~++:::+:::~
-+:~!
++'+-,'~"~,~,~,,]'~t
:.~]i+H,:
~+ ,,,/+ ++ !, : ++',+,~+ ++' +'t:+'+~s+',+!~++:+::++~+l+',;'~:,-,-+:++.,:+,-?+-++,;~+~+~+k,p
1
154
Chapter 4 • Reservoir Decompaction and Formation of Accumulation Capacity
(S, D) shales lies in the fact that the access of the carbonic acid formed from the organic matter of the Triassic shales to the intercalated reservoir rocks was easier than the penetration of the carbonic acid generated from the organic matter in the Devonian and Silurian successions. These are encountered in relatively elongate reservoir regions in the northern part of the Oued el-Mya Basin and in particular at Hassi R'Mel, Bordj Nili, Ait Kheir, Oued Noumer, DET, DEA, GA, etc. Nevertheless, the quantity of carbonic acid generated in the Silurian and Devonian shales is considerably larger than that formed in the Triassic shales as the former strata are much richer in organic matter. The processes diagenetically transforming the reservoirs by dissolution and redistribution of the carbonate cement have been studied in detail, with the giant gas and condensate field of Hassi R'Mel (HR) as an example, in the light of all geological, and in particular structural, sedimentary and tectonic factors (Figs. 4.8, 4.9). The results are described below. In the upper part of the succession in the HR field, and particularly in reservoir A, there is a notable but inhomogeneous cementation of the sandstones by secondary dolomite and anhydrite. Such a cemen)ation is developed to a lesser degree (in the form of spots) in the upper part of the other two reservoirs (B and C) below A. The two types of cement, viz. dolomite and anhydrite, are different in nature. Isotope studies show that the dolomite cement results from dissolution by acid solutions (H2CO3) and its transfer towards the upper part of each of the three reservoirs directly below the impermeable shale beds separating reservoirs from each other. The transfer of the dolomite cement is more or less progressive in nature. This is confirmed by carbon isotope studies as 613C gradually changes from heavy values in the residual cement of the lower portion of each reservoir towards a lighter value in the upper part, passing through intermediate values in the middle sections (Fig. 4.9). The above process is also observed when we take the three reservoirs as a group. Nevertheless, the described trend for the secondary carbonate cement is more evident in reservoir C, less in B and still less in A. This may be explained by the influence of solutions containing carbonic acid derived from Devonian and Silurian source rocks (shales) to the west, northwest, south and southeast of HR which followed the same routes of migration as the hydrocarbons, i.e. along the Hercynian unconformity towards the HR reservoirs. During their upward penetration in the sandstones the influence of the acid solutions naturally weakened to some extent. Under these conditions one should not forget the local internal sources of HzCO3 represented by the Triassic shale horizons separating the reservoirs from each other or the less important shale zones within each reservoir. In the northern part of the HR field there is a pronounced similarity in the sedimentary characteristics of the three reservoirs, expressed by: • a relatively constant thickness; • a rapid variation in grain size accompanied by a series of conglomerate beds; • rapid pronounced variations in the amount of cement and in particular of clay and carbonate cement strongly affecting the petrophysical properties of the reservoirs. Under these conditions the reservoirs exhibit a great petrophysical heterogeneity. Our research has shown that the factor controlling the petrophysical properties of the reservoirs, like in the two other zones and especially in the central zone, is a superimposed diagenetic dolomitization, in particular in the upper portions of the reservoirs.
4.1 • Decompaction Due to Solution of Binding Compounds o f Sandstones
255
.I i "6
i
i o
!
11111
" "= o-a
rX]
o.I
[ 2 ; 1 '
T
~;-=
i° ~.
o
-io
ii
156
Chapter 4 • Reservoir
Decompactionand Formation of Accumulation Capacity
As we have already pointed out, the secondary carbonatization of the sandstone cement owes its origin to the dissolution of carbonates by carbonic acid in the lower parts of the reservoirs and to their reprecipitation from brines entering the upper parts. The source of the respective H~CO 3 is decarboxylation of dispersed organic matter in the shale layers intercalated within the reservoirs and separating them from each other as well as in the Silurian and Devonian source rocks of neighbouring areas. We also have to evaluate the influence of structural heterogeneity on the diagenetic transformation of the reservoirs in the HR field by dissolution of the carbonate cement. This heterogeneity may be ascribed to faults and their extent as well as to the development of open fissures which at places might become important. One can observe that various faults found in seismic profiles are set out in an ordered network throughout a field. Although the throw of the faults is variable the respective variations stay in a fairly narrow range. The distribution of the faults and of their throw shows that with the exception of the northern tectonically more complicated zone the reservoirs A and C can be isolated from each other only locally. In contrast to this, reservoir B which is normally thinner and exhibits a greater sedimentary heterogeneity may be isolated more frequently, in particular in portions where the throw of the faults is greater. Because of the notable development of faults in the northern part of the field there is a structurally favourabte basis for the penetration of acid solutions, with H2CO3 derived from the shales within the Triassic reservoirs themselves as well as in the Devonian and Silurian shales in neighbouring reservoirs. These very solutions lead to the dissolution and subsequent upward redistribution of carbonates in particular in the reservoirs A and C in which there is a higher secondary porosity at the expense of carbonate cement and silicates (feldspars) becoming dissolved. This is exactly the same process which leads to the vertical diagenetic heterogeneity superimposed over the petrography of the respective sandstones especiallyin zones with a denser development of faults, in particular in the northern part of the field (Figs. 4.7, 4.9)R is this network of fractures developed in particular in the central zone where, furthermore, open fractures dominate over closed ones which assures a good communication between reservoirs and contributes to the preferential exchange of the fluids between them. The establishment of such a process here is, moreover, easier as in this zone the thickness of the shale beds separating the reservoirs from each other is negligible. 4.1.3.2
Triassic Reservoirs of the Ghadames Basin and Other Regions of the Triassic Province In the Ghadames Basin and certain parts of the Oued el-Mya Basin where the Triassic reservoirs rest directly (and unconformably) on Carboniferous, Devonian and Silurian shales the mechanism of dissolving and precipitating carbonate cement in the reservoir rocks (as indicated by the distribution of 613C) appears to be rather complicated (Fig. 4.1o). This situation may be ascribed to two different processes of dissolving carbonate cement. In addition to carbonic acid derived from Silurian, Devonian, Carboniferous and Triassic shales in the areas mentioned, the activity of organic acids was noted which formed in the Paleozoic shales. These acids were not subjected to the tong migration along the Hercynian unconformity. As pointed out in the preceding para-
4.1 - Decompaction Due to Solution of Binding Compounds of Sandstones
~57
~'u
~arc Bou Chaffra [emlet El Bazima
El Hamamit(EHTM)
.,%.%
~,%~e, ~ ~T~i~teLi~li~l~
,j~/ "N "~ \ "k
;econd.,middqe
'~
~ '~ ~elict,littie
~a~
Bir Retmaia (BRT-1)
aem
~%h
dem
181.4 TAG)
Age
Porosity (,%1
~ ~5~
........
=~
c~ O O"
~J~ ~ ~ ~ ~
Argillaceouscarbonate Anhydrite Quartzite Shale Limestone
~ Anhydrite cement ~ Eruptiverock TAGS- Upper ~ shale sandstone T A G I - Lower ]triassic complex
~
;291.~~ludan 298~
~-z ~ +2 ~=o~d~ll ry~ali~d Le~L~li~lo ~
¢
Sandstone SiltstoneMari
6 ~3C(~)
-~
TAC~
~s~.~ T ~ :263~ T~I I ;~ s ~
~
~
qTstaized
~
~
Ip ~
dem ~
,
illD~h
El Borma(ELB-I) A~'
~' lfl i~ 2~
t L'R~"'~'d~
I
0"'$I CementGeness
1
ta zed 2406.5 TAG1
24434 TAGI
i~P/sta&]ized
24623 2514.0 D3
iid~m
2S27.60s
ii~Jem
Fig. 4.10. Isotopic criteria diagenetic dissolution and precipitaion of carbonate cement (in Triassic and Paleozoic reservoirs) shown on block diagram of Erg Oriental (boreholes stratigraphy and lithology are based on Sonatrach's drilling data)
graphs formation and activity of carbonic acid differ in principle from that of the organic acids. In the first case decarboxylation of organic matter leads to the formation of CO2 (with lighter organic carbon) which is transformed into carbonic acid to react with the primary carbonate cement and depositing recrystallized secondary carbonate. The isotopic composition of the carbon is intermediate, i.e. between that of the organic and of the mineral-derived carbon, with values of-2 to -12%0 depending on which of the carbon sources dominated (Fig. 4.1o). In the second case the organic acids were derived from thermal carboxylation of organic matter. Their ultimate dissociation furnishing H+-ions reacting with the primary carbonate cement leads to the reprecipitation of carbonate inheriting an isotopically heavy carbon with 813C = 0 to +6%o.
'
158
Chapter 4 - Reservoir Decompaction and Formation of Accumulation Capacity
Under the above-mentioned geological conditions the influence of the carbonic acid as well as of the organic acids on the carbonate cement of the reservoirs is different and, depending on the origin of the carbon, there are two distinct types of recrystallized secondary carbonate: one with isotopicatly light carbon derived from organic matter when H2CO3 is formed and the second one with normal to heavy carbon (of mineral origin) in the case of organic acids. Under the microscope these two types of carbonate are indistinguishable from each other. The two above-mentioned processes are superimposed onto each other and overlay in time and space. For the first process the temperatures range from 50 to 15o °C and for the second one from 80 to 12o °C. Because of this the secondary recrystallized carbonate cement ultimately forming represents a mixture of the two types derived from different mechanisms. Consequently, the isotopic composition of the carbon in the cement with 6~3C= -2 to -8%o represents a mixture of the two types, i.e. the cement is isotopically lighter than if there were only the process of dissolution under the influence of organic acids, but it is heavier than in the case of the sole action of carbonic acid. This type of formation of the isotopic composition of the secondary carbonate cement was actually observed in the Triassic reservoirs of the Ghadames Basin and certain parts of the Oued el-Mya Basin where the 613Cof the secondary cement of the reservoir rocks ranges mostly between -1 and -8%0. Under these conditions in each case one has to establish the degree of crystallization and the genesis of the carbonate may this be of residual origin (with signs of dissolution) or secondary well-crystallized in nature, taking into account that the residual carbonate may eventually also have recrystallized during a secondary process (Fig. 4.1o). Such a mineralogical diagnosis is dictated by the need to distinguishing the eventual participation in the dissolution of the primary carbonate cement of organic acids and of H2CO3in order to obtain a correct interpretation of the isotope data. The various processes in the reservoir rocks implying dissolution and redistribution of carbonate under participitation of H2CO3 and organic acids may be summed up as follows: • In the lowest parts of the strata (or of the sand bodies) the carbon isotope composition of the residual cement (of mineral, sedimentary origin) is always heavy (fiI3C = o to +6%o) irrespective of its mechanism of dissolution. This carbonate is very often associated with good porosity of the solution type. However, the amount of residual carbon is much lower where organic acids were developed (because of their stronger dissolving ability). Consequently, in this case we observe the highest values for porosity of the solution type. The amount of remaining (residual) cement is on the whole higher where carbonic acid had been developed and the associated porosity, although still elevated, is much lower than in the case of the organic acids. • In the central parts of the sand bodies residual carbonate cement is characterized on the whole by a lighter isotope composition of 6~3C= -4%0. Under these conditions the residual carbonate cement, even when it shows signs of dissolution, nevertheless possesses a matrix characterized by secondary recrystallization and is distinguished by an overall good porosity, although on average it is lower than at the base of the succession. Furthermore, when secondary or residual carbonate cements recrystallize from primary (poorly crystallized) carbonate under the in-
4.2 • Decompaction-Cornpaction by Intergranular Pressure Solution
159
fluence of organic acids, the isotope composition of their carbon will be shifted towards heavier values with 613C = o to +4%0. In the case where residual carbonate cement of secondary origin forms under the influence of carbonic acid at the expense of primary (poorly crystallized) carbonate the isotopic composition of its carbon will be shifted towards lighter values of 6~3C= o to -4%o. This serves to confirm that the multiphase recrystallization of carbonate cement in sandstones is possible under the combined influence of carbonic acid and organic acids with all its ultimate effects on porosity and isotopic composition of the eventual carbonate cement. In the upper part of the sand bodies (generally below a shale cover) the development of carbonate cement is quite intense and its degree of crystallization high, naturally leading to low porosity values. The isotopic composition of the carbon is shifted here towards lighter values, but they are on the whole slightly heavier than the carbon in the carbonate cement of the sandstones from the northern part of the Oued et-Mya Basin. The 613C values here are controlled not only by the combined influences of organic acids and carbonic acid but also by the degree of recrystallization of the carbonate cement. An important role is then assigned to the thermal regime within the succession and to its evolution, which in turn controls extent and limits of the generation of these types of acids and of their combinations, the temperature ranges being 5o-15o °C for the former and 8o-12o °C for the latter.
4.1.3.3 Paleozoic Reservoirs Transformation and dissolution of carbonate cement in the Paleozoic, i.e. Silurian, Devonian and Carboniferous, reservoirs in all basins are similar in nature to those in the Triassic reservoirs of the Ghadames Basin and certain parts of the Oued el-Mya Basin as these Paleozoic reservoirs alternate with numerous thick rather mature shales rich in organic matter. These beds are able to generate large volumes of H2CO3 as well as of organic acids which have immediate access to adjacent reservoirs. 4.2
Decompaction-Compaction by Intergranular Pressure Solution (of Quartz Grains) and Removal of SiO2 by Alkaline Solutions Sand and sandstones lose their primary porosity under the influence of three main processes: mechanical compaction, pressure solution and cementation. In nature the porosity of graded sands varies on average between 35 and 45%. During the initial stages of burial mechanical compaction of quartz sands may reduce the primary porosity by more than one third and in the case of sands rich in lithic fragments by even more than that. Such a primary reduction of porosity is accompanied by a reorientation and rearrangement of its components and by a plastic deformation of lammellar grains. With continuing burial the sandstones undergo pressure solution, a process usually affecting the detrital quartz grains especially in quartz sandstones. The silica dissolved at the grain contacts may be transferred from these sites of dissolution by circulating interstitial fluids or it might become deposited in the form of quartz cement
260
Chapter 4
• Reservoir Decompaction and Formation of Accumulation Capacity
in areas not submitted to the pressure of the same grains which have not suffered dissolution. The fate of the silka depends on the chemical and hydrological characteristics of the interstitial fluids. In numerous cases of quartz sandstones, pressure solution and the cementation by quartz associated with it may notably reduce the primary porosity. Cementation takes place especially during diagenesis and at relatively large depths of burial. This process depends on the appropriate geochemical and hydrological conditions under which silica may again be transported in the form of cement era,eloping usually quartz, feldspars, carbonates, various clay minerals, etc. The formation of one or more minerals may considerably reduce the primary porosity. When cementation took place prior to mechanical compaction or before the end of pressure solution these processes may be arrested by blockage of the pore space by cementation alone. Among three main processes leading to a reduction in porosity, pressure solution has become the object of extensive studies (Renton et al. 1969;Manus and Coogan 1974; Sibley and Blatt 1976; Robin 1978; Bjrrkum 1996). Until quite recently pressure solution was the "best" known mechanism for explaining the increase in the amount of silica in the interstitial solutions and during cementation (Weyt 1959). However, Marius and Coogan (1974) and Sibley and Blatt (1976) as well as a number of other authors suggested that pressure solution is less important and there are suggestions in the literature according to which it appears that for the reduction of primary porosity in quartz sandstones pressure solution is not really as important as so far assumed. A study of pressure solution undertaken to show which mechanism really controls the porosity has three aims: (1) to evaluate quantitatively the spatial distribution of pressure solution and quartz cementation on a regional scale in the sandstones, (2) to evaluate the influence of certain geological variables like median grain size, composition of clay binders, temperature, etc. on pressure solution and (3) to evaluate its role in quartz cementation during evolution of the sandstone porosity. We know that the sources of silica deposited in the form of quartz cement may be of an internal nature, like pressure solution, stylolitization or solution of quartz grains of the framework of the rock, or of an external nature, like products of leaching caused by meteoric waters, transformation of clay minerals and/or dissolution of siliceous components in adjacent silty-argillaceous sediments. This serves to show that it witl be difficult to evaluate these sources objectively. For the purpose of this presentation we shall accept by convention that pressure solution is the dominant and possibly only source of silica. Although this hypothesis may" not be perfectly justified, it nevertheless enables us to calculate the minimum quantities of silica formed during diagenesis of the rock itself. Thus, the quartz cement (overgrowth rims) was considered as the only "adsorption zone" for silica within the rock. This hypothesis is admissible for the sandstones studied as the only other minerals with authigenic silica deposited during diagenesis are clay minerals, feldspars accounting only for minute quantities. Anyhow, the silica forming part of the authigenic minerals is in equilibrium with the silica supplied by the formation waters and with that activated in the transformation of the detrital clays and of the feldspars. The ultimate error referred to is negligible in all cases and will not influence the overall nature and distribution of the silica in the various reservoirs.
4.2 . Decompaction-Compaction by intergranular Pressure Solution
161
4.2.1 Pressure Solution and Q u a r t z C e m e n t
Pressure solution and quartz cement to a great extent depend on a number of geological variables and in particular on grain size of the quartz and temperature. The influence of other parameters like grain sorting, clay cement content, composition of framework grains, presence of early cement and burial history of the reservoirs also have to be considered. The volume of quartz dissolved by pressure solution denotated as OQ or overlay quartz and the volume of quartz cement (QC) are expressed as a percentage of the overall volume of the detrital quartz (DQ) (Fig. 4.11). In this context, the utilization of the same approach by Houseknecht (1984), i.e. of the ratios OQ/DQ and QC/DQ, excludes in the sandstones considered any influence of the variability of the volume of detrital quartz on the volume of dissolved quartz and quartz cement. A quantitative evaluation of quartz cement and pressure solution was carried out on the basis of a study of thin sections under the optical microscope and with the aid of cathodoluminescence.
a
1
2
utShiiO2en~c
Differentgenerations of quartzovergrowths
b
Latesilica B
E core C
~.x
kamellargap
'"'" Styloliteand fracturing Fig. 4.11. Quartz intergranular pressure solution and quartz overgrowths, a Sequences of stages in the development of quartz pressure solution (modified after Siever 1962). b Schematic illustration defining detrital quartz (D.Q.), quartz cement (Q.C.) and overlap quartz (O.Q.).e Schematicdiagram of major features of quartz overgrowths
162
Chapter 4 • Reservoir Decompaction and
Formation of Accumulation Capacity
4.2.2 Factors Controlling Pressure Solution Influence of Quartz Grain Size. Correlations observed reveal that there is a clear statistical correlation between mean grain size and volume of quartz dissolved by pressure solution, i.e. the dissolution of finer grains is more intense than that of coarse grains (Weyl 1959; Renton et al. ~969; Sprunt and Nur 1977). This results from the decrease of the distance required for diffusion along the grain contours caused by the decrease of their size. However, sometimes there is a certain dispersion of the regression lines which results from the influence of the other factors to be described below.
Influence of Grain Sorting. There is no clear-cut correlation between pressure solution and sorting. Although there is even a slight trend for a negative correlation, this results less from the intergranular space in the poorly sorted sandstones and not from the direct influence of sorting on the quartz cementation. Thus, sorting does not exhibit a detectable influence on pressure solution within the sandstones. Influence of Clay Mineral Content. A number of authors like Thompson (1959), Heald and Renton (1966) and Sibley and Blatt (1976) have shown quantitatively that thin isolated sand intercalations or continuous sandstones possessing abundant clay "bridges" on their quartz grains have suffered more pressure solution and eventual removal of SiO2 than sandstones with a smaller amount or total absence of clay"bridges" in the form of a film. This was also observed in the Saharan sandstones (Plate 13). It is, however, difficult to establish a quantitative statistical correlation between the amount of clay components and pressure solution. There is at the same time in the Saharan sandstones a clear inverse statistical correlation between the amount of clay minerals in the form of a film and the quartz cement (Fig. 4.12~Plates 13,14). This is in agreement with observations showing that the clay cement impedes in the sandstones the formation of overgrowth rims of quartz (Millot 1964; Heald and Larese 1974). It appears that the degradation of illite contributes to the formation of strong alkaline solutions catalyzing pressure solution. Weyl (1959) suspected that the clays create an efficient environment for diffusion and thereby reinforce pressure solution by the removal of the silica from its place of dissolution. It is also possible that the clay "film" impedes the overgrowth rims of quartz and thereby supports the preservation of relatively small contact zones between the individual grains. Tension in such small contact zones wilt be much more concentrated than along wider contacts which appear during the formation of quartz overgrowth. Influence of Sandstone Matrix Composition. Some of the sandstones investigated contain a notable quantity of feldspar. It has been confirmed that there is a clear correlation between the composition of the sandstone matrix and pressure solution. The
Plate 13. Intergranular pressure solution of quartz. The contact zones between the quartz grains are I~ clearly discernible. The silica dissolved in the contact zones appears to be "exported". Despite the obvious dissolution of quartz the amount of quartz cement is negligible. Photo h clearly shows dissolution features on quartz. The"films" envelopingthe quartz grains are made up of authigenic chlorite
4.2 • Decompaction-Compaction by Intergranular Pressure Solution
163
t64
Chapter 4 • Reservoir Decompaction and Formation of Accumulation Capacity
Fig. 4.12. Correlation of secondary silica with authigenic day coatings on quartz grains in oil- and water-bearing reservoirs
30 o
I •
Oil-bearing reservoirs
II o
Water-bearing
r e s e r v o i r s ( O u e d M y a Basin)
o
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o oo
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•
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i
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10 2; Authigenic clay coating on quartz (%)
30
grain size controls the contents of quartz and feldspar. We have observed the following correlation: with decreasing mean grain size the percentage of quartz drops and that of feldspar increases, the respective correlation coefficient being o.7. This entails that the main controlling factor for pressure solution is rather the grain size than their mineralogical composition. Where there are contacts between quartz and feldspar the quartz will be preferentially dissolved in most cases and it is thus more susceptible to pressure solution than feldspar. A number of the sandstones investigated contain a notable quantity of metamorphite fragments (including schists and micas), but we have not been able to observe any correlation between the latter and pressure solution. However, in samples containing abundant ductile lithic fragments we have noted a decrease of quartz cement as well as of pressure solution. Actually the schist fragments are deformed plastically during compaction and efficiently reduce the contact zones between quartz grains and thereby also pressure solution. The same schistose fragments cover the surfaces of the quartz grains and in this way preclude the formation of overgrowth rims as in the case of clay cement.
Influence of"Early" Cement. The presence of calcite, dolomite and anhydrite in the form of early cement strongly hinders pressure solution and thereby contributes to the preservation of intergranular space. This observation may also be made for early quartzose cement. The latter contributes to the formation of a solid framework in the sandstones which obstructs their later compaction by equalizing tension along the boundaries of the quartz grains of the framework.
Plate 14. Relationship between quartz and clayS'bridges'.Photos a under polarized light (PL),b under I~
SEM,f PL, g PL and h PL exhibit intense developmentof quartz cement under near-complete absence of clay "bridges". Photos c SEM,d SEM, i PL and j PL show in contrast to this the lack of quartz cement where clay"bridges" are abundant
4.2 • Decompaction-Compaction by Intergranular Pressure Solution
165
i66
Chapter 4 • ReservoirDecompactionand Formationof AccumulationCapacity
Influence of Sandstone Texture. The rate of compaction of sandstones with small irregular grains is twice that of a sandstone with small rounded grains. Pressure solution and eventual overgrowth of quartz lead to the situation that the shape of angular grains differs little from that assumed by rounded grains after a comparable degree of pressure solution. In the sandstones pressure solution and simultaneous quartz overgrowth form much more compact aggregates than those resulting from cementation alone. Sandstones with grains of siliceous schists succumb much more rapidly to pressure solution than monocrystalline quartz and will not recrystallize completely. However, the outlines of these grains are obliterated and the aggregate formed represents a solid continuous mass with the appearance of a siliceous schist. As a result of pressure solution, differences in texture and composition of the sandstones may result in pronounced variations of the eventual porosity. 4.2,3 Silica Budget The application of various modern methods like cathodoluminescence and SEM studies enabled us to quantitatively evaluate the silica budget in the sandstones. Figure 4.i3 presents correlation diagrams of the parameters pressure solution (OQ) and detrital quartz (DQ) with quartz cement (QC) and detritat quartz (DC) for four groups of reservoir rocks of the Saharan Platform. In these diagrams sandstone samples plotting below the i : i line may be called'~sitica exporters", i.e. in these sandstones the amount of quartz dissolved by intergranular pressure solution is larger than the amount of quartz cement. The samples above the i : 1 line may be called ~;silica importers" as they contain more quartz cement than could have been furnished by pressure solution. A negative balance indicates that silica has migrated outside the sample whereas a positive balance indicates that a certain amount of silica has migrated into the sample, i.e. that it has been supplied by outside sources. Within the limits of each group of samples in the diagrams the relative position of the individual sample points is controlled by the grain dimensions. Sandstones with smaller grains are located in the lower portion of the "exporter" field whereas coarsergrained sandstones are found in the upper part of the ~importer" field. The relative positions of these groups of data are also controlled by thermal maturity. Samples from zones of low thermal maturity fall farther into the direction of the upper part of the '~importers" compared to samples from zones of higher thermal maturity. In Fig. 4.i3 and Table 4.7 the silica budget for the Cambrian sandstone complex of the Oued el-Mya Basin studied here is presented. Of the i63 samples studied 34 are'timporters", i26 "exporters" and three are in equilibrium. Using average values for OQ/DQ, QC/DQ and detrital quartz in general (its percentage of the overall sandstone mass) we have calculated the silica flux in percentage of the total sandstone mass, with negative values for exported silica and positive values for imported silica, as shown in Table 4-7. Assuming that the samples studied are representative of the areas considered, e.g. of the outcrop area of the Cambrian sandstones of the Oued el-Mya Basin, we may conclude that 7.26% of the sandstone mass was exported as SiO2 as a result of pressure solution. When carrying out the same evaluation on the results from other successions we may conclude that for the Devonian of the Ghadames Basin, 4.92% silica has been
~67
4.2 • Decompaction-Compaction by Intergranular Pressure Solution
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34 Importers 3 Balanced 126 Exporters
29 Importers 4 Balanced 122 Exporters
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Fig. 4.13. I n t e r g r a n u l a r p r e s s u r e solution, a C a m b r i a n Oued el-Mya, b Devonian Ghadames, c Carboniferous Illizi, tt Ordovician Ahnet-Mouydlr
Table 4.7. Silica budgets of quartzose sandstones in Saharan basins (Si flux expressed as a percentage of current rock volume; negative values indicates volume of silica exported and positive values indicates volume of silica imported)
Complex
Mean
Si flux
Net
Mean DQ
OQ/DQ-QC/DQ
(%)
0.062
-0.095
76.4
-7.26
0.081
-0.070
70.5
-4.94
0.158
0.052
-0.106
78.9
-8.36
0.153
0.214
0,061
72,8
4.44
OQ/DQ
QC/DQ
Cambr. Oued eI-Mya
0.157
Devon. Ghadames
0.151
Ca rbonif, lllizi
Ordov. AhnebMouydir
:68
Chapter 4 • Reservoir Decompaction and Formation of Accumulation Capacity
exported and 8.3% for the Carboniferous complex of the Illizi Basin. In contrast to this, in the Ordovician of Ahnet Mouydir 4.44% of silica dissolved by pressure solution has been imported. In the latter case the spread of the values for OQ/DQ and QC/DQ is caused mostly by the influence of pressure solution on the median grain sizes as well as by the presence of film-like "clay bridges" around the quartz grains. As the distribution of samples of a certain succession is more or less even over the "export" and "import" fields we may assume that silica has migrated locally from the "exporters", i.e. fine-grained sandstones, towards "importers", i.e. coarser-grained sandstones, and in this way on the local scale an approximate mass balance was maintained. Nevertheless, the overall data show that an important transfer of silica has taken place on a much wider scale and that this is closely related to temperature as the controlling factor for intergranular pressure solution. In zones of higher thermal maturity pressure solution probably was so efficient that it could act simultaneously as the source and driving force for the mass transfer of silica. In zones of low thermal maturity pressure solution was not particularly effective for the mass transfer of silica and the reduction of the primary porosity. We can advance here two hypotheses for explaining the routes of migration which could have contributed to this mass transfel. If we assume that within the limits of a formation there is a communication between the various sandstone bodies and the shale beds of variable thickness enclosing them, it follows logically that the silica liberated during diagenesis of the shales will migrate into the sandstones causing them to become "importers" in areas of low thermal maturity. However, this hypothesis also entails that in zones of higher thermal maturity the silica liberated by pressure solution should migrate into the enclosing shales or across them, a situation of rather low probability. If the sandstones should react as a more or less closed system then the silica dissolved by pressure solution in regions of higher thermal maturity should migrate within the limits of the sandstone bodies against the thermal gradient to become deposited as quartz cement in zones of lower thermal maturity. This latter hypothesis is rather attractive as it corresponds perfectly to the approximate mass balance observed actually in the Ordovician sandstones of the Ahnet Mouydir Basin. Of these sandstones, 66 may be considered as "importers", 59 as"exporters", whereas six are in equilibrium (Fig. 4.:3). Assuming that the samples analyzed from the various formations are volumetrically representative of the granulometry of the sandstones and of their thermal maturity we may draw conclusions as to the internal mass balance on a regional scale which are plausible on the basis of the geological data from this basin. Within each stratigraphic entity studied samples of fine-grained sandstones and those containing well-developed illitic"clay films" plot in the lower right of the respective diagram whereas those of coarser grain size and with few illitic "films" or none at all plot in the upper left. It has to be pointed out that the above reasoning requires the presence of a genetic link between the silica liberated by pressure solution and the silica deposited as quartz cement within the same formation. In most models of pressure solution this assumption appears to apply (Robin 1978). In the Cambrian reservoir rocks of the Hassi Messaoud field there is much more secondary silica in the upper parts of the sand bodies than in their lower parts whereas the distribution of the porosity behaves naturally in the opposite sense (Fig. 4.14). This is proof of an upward transport of the silica liberated by pressure solution under the
4.2 • Decompaction-Compaction by Intergranular Pressure Solution
De ~th
i Lithology
Scale
0246810m
1 Classification Morphoscopy i
1
Main constituents
Granulometry
Porosity
Observations
5 7.5j012S15
169
Ag, Formation Member Permeabilitymeasured horiz.-- v e r t . - lOmSOlOO~O~5~a
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Porosity Illite Kaolinite Clay Micas Heavy minerals
N ~ Anhydrite Secondary silicates ~ 3 Detrital quartz Feldspars [2~ Iron oxide Carbonates
[ ~ Medium grained Sandstones Coarse grained Sandstones ~ Pebble Figilites Detrital clays
Fig. 4A4. Petrographic log of borehole MD-Iso
alkaline conditions produced by the decomposition of the feldspars and micaceous schists. The various aspects of pressure solution outlined above combine in the geological and geochemical processes of formation of the respective reservoirs.
170
Chapter 4 • Reservoir Decompaction and Formation of Accumulation Capacity
4.2.4 Mechanisms of Silica Transport Numerous authors have studied the geochemical and hydrological problems of the transport of such a vast amount of silica within the sandstones during cementation and formation of secondary porosity (Land and Dutton 1978; Bjorlykke et al. 1979; Bjorlykke 1980; Land 1984 ). In their majority they agree that quartz cementation in the case of absence of pressure solution or in presence of an internal silica source requires at least lo4-1o 5 cm 3 of water per cm 3 of cemented sandstone. The authors are also unanimous in that such volumes of water cannot be derived entirely from the dehydration of silty-argillaceous rocks. Because of this a number of alternative mechanisms for the hydrological transport of SiQ in the sandstones have been advanced, ranging from "isochemical systems" of Bjorlykke et al. (1979) through water recycled by convection (Wood and Hewett 1982) to subterranean conditions under which silica is transported by intruding meteoric waters (Bjorlykke 1984). The calculation of the hydrological parameters necessary for establishing the abovementioned silica balance appears to be a rather complicated problem as intergranular pressure solution leads to a dynamic reduction in rock mass and volume and frequently also to a lowering of the porosity. These parameters have been simulated for different petrophysical and geochemical conditions. The first results show that for the large amount of secondary silica observed to become exported, 0.5 x lO4-O.5 x lOs cm 3 of water is required for every cm3 of the Cambrian sandstones of the Oued el-Mya Basin. If we assume, on the other hand, that the quartz cement in the Ordovician sandstones from Ahnet Mouydir resulted within the rocks themselves from pressure solution then the water flux necessary was small or virtually nil and this formation approached an isochemical system. A more extensive investigation of the problem shows that a certain amount of silica from a variety of sources was deposited in the form of"early" cement among which we also have silica produced indeed by pressure solution but in deeper buried sediments. In this way, silica may be furnished during early diagenesis by the flux of meteoric water (in a downward direction) which might be saturated from the leaching of rocks or by the flux of formation waters (in an upward direction) which might be saturated during pressure solution or during the diagenesis of day minerals or again by dissolution of various silicates in deeper diagenetic zones. The silica resulting from pressure solution during deeper burial should be displaced upwards within the boundaries of the bed or still higher up along the stratigraphic section probably by passing along faults or other tectonic disturbances. These various mechanisms cannot be evaluated fully and with certainity without a vast amount of petrographical data on a regional scale and/or isotopic data which could help in the definition of the geochemical aspects of quartz diagenesis. Whatever it might be, there remains one highly important question: can the silica mobilized by pressure solution in a deeply buried sandstone really migrate upward in the stratigraphic section to become deposited as quartz cement in shallow-buried sandstones? Such an approach could explain the regional variability which may take place during pressure solution or during quartz cementation within the boundaries of the individual sandstone complexes of numerous Saharan basins studied by us and in particular in the Ahnet Mouydir and Illizi Basins. In general, sandstones with little detrital clay actually possess an elevated permeability and exhibit a notable secondary silicification as if they were part of a subarkosic
4.2 • Decompaction-Compaction by Intergranular Pressure Solution
171
diagenetic suite of sandstones. Such a sandstone normally lies below and not above the impermeable silty-argillaceous rocks. This association of low and high permeabilities signifies that during diagenesis the supply of interstitial solutions was high.
Sources of Fluids Transporting Silica. Boles and Frank (1979) expressed the reaction of illitization by the following formula: 3.93K+ + LsNa + - montmorillonite
) illite + 1.57Na+ + 4.z8Mg 2+ + 4.78Fe2+ + 26.66Si 4+ + 570 z- + 11.4OH- + 15.7H20.
At a temperature of 75 °C this reaction is not yet complete. In order to dissolve the 26 moles of silica liberated d u r i n g this reaction we need 38 x lO3 moles of water (Walther and Helgeson 1977). It is thus obvious that even in the case of the silica supply being accompanied by dehydration of the enclosing argillaceous strata there has to be a supplementary source of water. In the case of the Devonian sandstones of the Tiguentourine field in the Illizi Basin the mean porosity is 13% and the mean content of secondary silica is also 13%. Under the temperature conditions prevailing in this region (7O-lOO °C) the solubility of quartz drops from o.o56 to o.o316 g 1-1.When the diagenetic fluid cools from lOO to 8o °C SiO2 will be deposited at a rate of o.oolo g / 1 °C (Holland and Malinin 1979). Then per cubic centimeter of sandstone a volume of 13o x lo 3 cm 3 of pore fluids will deposit o.29o g of silica for each temperature drop of 1 °C and we need actually 1.4 x lo 6 cm 3 of fluid at a drop of I °C to deposit the 13% of authigenic quartz observed in these sandstones. Even if the sandstones of the basins investigated are nearly always intercalated with silty-argillaceous rocks of a thickness well above that of the sandstones themselves, the water supplied during their dehydration will not suffice to transport such a vast amount of silica mobilized by pressure solution or derived from another source. Under this situation the polycyclical circulation of fluids by convection could be a highly probable mechanism explaining the transport of silica (Fig. 4.15), in partictflar by the
Fig. 4.15. Possible mechanism for mass transfer of silica (quartz) in moving pore fluids during diagenesis in reservoir sand units, a In fluids supplied by mechanical sediment dewatering and com-paction; b in fluids by recharge from seawater; ¢ in fluids convecting within a sandstone sequence realizing SiO2of various origin including pressure solution (from Wood and Hewett 1982)
S
• ' |_ |[ a Flu~dsmoveup-dlp out of basra JD
Approx,2km
oc 7 -r
l"e"~--J~*' ~at " /~ c ? paC*ing
Fluids recharged by seawater
convectionwithin C sand
T~
-\
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Chapter 4 • Reservoir Decompaction and Formation of Accumulation Capacity Legend Hercynian discordance []~]]~
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alkaline solutions resulting from the decomposition of micaceous schists and feldspars. This polycyclic circulation of fluids is a plausible mechanism for two reasons: (1) it is in agreement with solutions containing resins, (2) efficient heating shows that the status of the pore fluid will change dozens of times, (3) it is characterized by the great rapidity of its taking place which could favour the diagenesis of quartz as a distinct process over a reasonable scale of geological time. Wood and Hewett (1982) have shown theoretically that a layer of porous and permeable sandstone gradually buried under a geothermal gradient in the order of 15 °C km -1 will inevitably undergo slow convection in a unicellular system. Such fluids can transport dissolved silica on a kilometer scale much more rapidly than by diffusion alone. Under higher geothermal gradients, in particular in the Paleozoic deposits of the western and central Saharan basins, and in the presence of two highly different complexes the use of such a theoretical model shows that convection will take place not only as an unicellular system in a sloping saturated sandstone but also in an energetic multicellular and more complicated polyhedral system composed of small
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polyty~ic cells of near-hexagonal outlines (Fig. 4A5). Such small convective cells may also be envisaged for the flux of acid fluids within the pore volume as described above. 4.2.5 The Role of Pressure Solution
in t h e E v o l u t i o n
of Porosity
We shall now propose an approach for the evaluation, in the Saharan reservoirs, of the relative importance of the process of compaction and cementation in the reduction of porosity by presenting correlation plots of intergranular volume (VIG) vs. cement (Fig. 4.16). It is generally held that under the conditions of sediment accumulation on surface the VIG of well-graded sandstones is about 4o %. This VIG or porosity can only be reduced by compaction, a mechanical process reducing VIG to 3o%. Any further reduction by chemical compaction or pressure solution is a specially important process. The intergranular porosity of a sandstone is a function of the volume preserved after compaction and of its (the VIG's) portion filled by cement (Fig. 4.16). The inter-
174
Chapter 4 - Reservoir Decompaction and Formation of Accumulation Capacity
granular porosity thus is equal to VIG reduced by the amount of cement. The dotted line in Fig. 4.16 separates the sandstones according to the contribution of a certain process for the reduction of the porosity, i.e. compaction in the lower left field and cementation in the upper right. In order to obtain better conditions for the establishment of such a model we have retained only samples with average sorting indices, thereby eliminating the influence of other factors on the correlation. Figure 4a6g,f represent the diagrams for VIG vs. cement for the Ordovician sandstones of the Hamra (Hr-1) and respectively the Begura (BRq) regions. The samples of the two complexes scatter mostly, i.e. to 6o-7o%, in the lower left of the diagram, indicating that the horizons considered the process of compaction more important for the establishment of porosity than cementation. The mean values for all data are around 16-18% for VIG and 12-15% for cement (of all types), implying that 5o-6o% of the primary porosity has been reduced by the process of compaction, i.e. by mechanical compaction and pressure solution, whereas 2o-a5% has been eliminated by cementation. These data are in good agreement with the situation of these complexes on the correlation plots of OQ/DC vs. QC/DQ, with the latter being"exporters" of silica rather than "importers". The diagram for the Devonian sandstones of the Reggane (REG-3) and Oued Talah (ODTHq) regions in Fig. 4.16b,e shows that 65-8o% of the sandstones investigated plot in the upper right of the diagram indicating that here the reduction of the porosity was dominated largely by cementation and not by compaction. In these regions the mean values for VIG range from 20 to 24% and for cement from al to 23%, so 35-4o% of the primary porosity was eliminated by compaction processes and 5o-55% by cementation. In the OQ/DQ vs. QCIDQ plots these complexes are more "importers" of silica. In the Teguentour region (TEGq) and ECFq the correlations were investigated for two sandstone complexes of different ages, i.e. at Teguentour from the Carboniferous and Cambro-Ordovician and at ECF from the Carboniferous and Ordovician. tt is well known that these complexes of different age find themselves in different situations regarding the dominant influences of compaction or cementation. It is also known that the Carboniferous sandstones of the two regions have been affected to the same degree, more by cementation than by compaction. On the whole, the model of porosity reduction by compaction and cementation results from a combination of geological and geochemical processes in the formation of reservoirs. 4.2.6 Conclusion
Quantitative petrophysical and petrographic analyses with the aid of cathodoluminescence and SEM studies have shown that in the Saharan sandstones pressure solution has been influenced by a number of geological variables. In the complexes studied there is a linear correlation between median grain size and pressure solution by which finergrained san&tones suffer more pressure solution. The presence of stronger clay"films" on the quartz grains also favours pressure solution. Sorting and variations in the compaction of the matrix of the sandstones exert little influence on the process of pressure solution. The presence of "early" cement dominates pressure solution and c o n -
4.3 . GeneralizedModelsof the Transformation of Oil-Bearing and Reservoir Formations
175
tributes itself to the preservation of a relatively large amount of intergranular porosity potential. We can envisage that the sandstones notably cemented during shallow burial represent an exception and later deep burial notably increased the extent of pressure solution in them. The calculated silica balance shows that pressure solution is an important agent for mass transport but by no means the dominant one. Nevertheless, petrographic observations show that the greatest portion of the quartz cement could have been deposited prior to pressure solution. These observations suggest that certain sandstones could have behaved like silica "importers" during early and as "exporters" during later diagenesis. The intergranular porosity of the sandstones results partly from the primary intergranular volume eliminated by (chemical and mechanical) compaction and partly from the intergranular volume remaining after filling by cement. Within the boundaries of certain study areas like the Ordovician sandstones of Hamra and Begura, processes of compaction have been shown to be more decisive for the establishment of the final porosky than cementation. In the Devonian sandstones of the Reggane and Oued Talah the opposite trend is developed, i.e. it is the cementation that played the decisive role in the reduction of the primary porosity of the reservoirs. In the Devonian sandstones of the MD-lm region compaction and cementation are virtually equally important. In the Saharan sandstones temperature exercised a notable influence on pressure solution. On a regional scale the extent of pressure solution increases with greater thermal maturity: Fine-grained sandstones and sandstones of elevated thermal maturity exhibit more grain contacts, a more complicated (longer) geometry of the contacts characterized by mutual interpenetration of the grains and lower intergranular porosity than coarser-grained sandstones and those of low thermal maturity. The extent of pressure solution as well as the largest possible amount of quartz cement may be predicted on the basis of grain size and thermal maturity. The absolute volumes of cement, however, cannot be predicted. The polycyclical circulation of fluids by convection under these circumstances appears to represent a highly probable mechanism for the transport of silica, in particular in the case of waters the alkalinity of which is controlled by the decomposition of mica schist fragments and feldspars.
4.3 Generalized Models of the Transformation of Oil-Bearing and Reservoir Formations 4.3.1 Model of Diagenesis in Space and Time A general model of the diagenesis of the sandy reservoirs of the Sahara is presented in Fig. 4.17.The sequence of the diagenetic stages is controlled by temperature and time. Vitrinite reflectance in shales adjacent to or intercalated with the quartz sandstones indicates their exposition in time and space to temperature and may be used also as an indicator of the thermal maturity of the quartz sandstones. In this way the vitrinite reflectances correspond to the stage of mesodiagenesis. Mechanical compaction dur-
176
Chapter 4 • Reservoir Decompaction and Formation of Accumulation Capacity
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4.3 • Generalized Models of the Transformation of Oil-Bearing and Reservoir Formations
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ing the immature stages affects essentially the primary porosity, reducing it by about one quarter. Mesogenetic carbonatization, decarbonatization and decarboxytation of the organic matter in shales adjacent or intercalated with the sandstones taking place during the immature stage are at an exceptionally low level. Silica diagenesis, corresponding to the chemical compaction of quartz sandstones, is more efficient during the stage of medium maturity when it reduces mainly the primary porosity. Carbonatization reaches its maximum during the last two steps of medium maturity and fills mostly the primary porosity but also a notable part of the secondary porosity. Decarbonatization and decarboxylation of organic matter start to become active during the stage of medium maturity. Loss of porosity dominates over its growth despite the formation of a considerable volume of secondary porosity. Towards the end of medium-grade maturity nearly all effective primary porosity has been eliminated but the secondary porosity nevertheless remains important. Stage I of maturity is characterized by maximum decarbonatization and decarboxylation of organic matter. The generation of hydrocarbon fluids is overall moderate, increasing to its maximum towards the end of this stage. Carbonatization, quartz diagenesis and mechanical compaction proceed rather moderately. The establishment of secondary porosity is at its maximum and results essentially from decarbonatization and dissolution of feldspars. This increase notably compensates for the loss of porosity. The initial phase of stage II maturity corresponds to moderate to low levels of carbonatization, decarbonatization, decarboxTlation, mechanical compaction and quartz diagenesis, while the generation of hydrocarbon liquids takes place over a short period only. The loss of secondary porosity is minimal and its increase is equally negligible. The diagenesis of quartz functions is an important diagenetic process during the intermediate and final stages of stage I maturity. Under the physico-chemical conditions in certain cases this diagenesis leads to the loss of any reducible secondary porosity, whereas in other cases it results in an increase in this secondary porosity. During the intermediate phases of stage II maturity organic diagenesis leads to the formation of a great volume of methane. The zone of maximum active decarbonatization is situated below the zone of maximum active carbonatization, or, in other words, the carbonates dissolved are transferred upwards by aqueous solutions in the form of a "carbonate flux". The diagenesis of clay minerals in intercalated argillaceous beds may obviously change the chemical nature of the interstitial waters by modifying ionic ratios, alkalinity and pH so that the carbonates start to become precipitated. One could alternatively suggest that the formation waters migrating upwards from zones of higher pressure into zones of hydrostatic pressure can precipitate carbonates in intermediate pressure zones. In most of the petroliferous basins studied the distribution of the zones of decarbonatization of reservoir rocks reflects closely the rhythms of formation of C02 and organic acids in the related argillaceous horizons (or in the oils) resulting from the thermal decarboxylation of matured organic matter. Decarbonatization is more pronounced in quartz sandstones closely associated with argillaceous source rocks rich in organic matter. This underlines the fact that in the lower horizons the formation of C02 and of organic acids represents the main cause of mesogenetic decarbonatization. Other obvious indications of mesogenetic decarbonatization are: (1) the physico-
178
Chapter 4 • Reservoir Decompaction and Formation of Accumulation Capacity
chemical conditions in the periphery of zones of excess pressure, (2) massive dehydration of clay minerals (Savkevic 1971) and (3) the formation of hydrogen sulfide. In many cases mesogenetic carbonatization has taken place even in compact quartz sandstones. Fissures and lamellar porosity assure sufficient penetration for leaching solutions to start the process of decarbonatization. As secondary porosity is formed it favours the continued access of acid waters to the sandstones. The secondary porosity that may form during intense decarbonatization is limited by the content of carbonate cement and by the mechanical stability of the non-carbonate portion of the matrix. It is possible that quartz sandstones with carbonate cement will not undergo decarbonatization for one of the following reasons: (1) lack of chemical conditions conducive to leaching, i.e. absence or insufficient content of organic matter as a source of CQ, (2) very low leaching potential of the formation waters during the earlier stages of diagenesis and (3) lack of access for the formation waters which would have caused leaching in the sandstone horizons. In the history of sedimentary basins intense decarbonatization probably follows rather soon after the older sandstones attained the stage of maturity. These sandstones contain mostly sedimentary and eogenetic carbonate cement. The reprecipitation of such carbonate dissolved in underlying zones of medium maturity may initially produce solutions with a low carbonate content. Sedimentation followed by burial probably brings the carbonate compounds into the zone of maturity where they may undergo strong decarbonatization which in turn, in the sandstones "newly arrived" in the zone of medium maturity, could lead to the establishment of a more massive "carbonate flux". In this way the two zones of maximum decarbonatization and carbonatization could migrate continuously into the direction of the younger deposits. During this process carbonate cement may undergo repeated transformations in a large number of cycles. A large quantity of carbonate may accumulate in a mobile zone if the amount of organic matter present in the section is sufficient to maintain decarbonatization. A large part of the liquid hydrocarbons and eventually also of the natural gas is formed in diagenetically mature zones above a "peak" in the formation of secondary porosity. Such a situation is very favourable as quartz sandstones with secondary porosity should be susceptible to "trap" these hydrocarbons. 4.3.2 Generalized Model for the Transformation of Reservoir Rocks, Mass Transfer and the Formation of Reservoir Properties All diagenetic processes are combined into a model of the transformation of the cement of reservoir rocks distinguishing primary and authigenic phases, of the evolution of the field waters (in composition and volume), and especially of the organic and inorganic interactions leading to the formation of carbonic acid and organic acids which in turn initiate an important transfer of masses of carbonate and silica (Fig. 4.18). The above-mentioned processes are considered in their entirety as presenting evidence for the formation of reservoir properties as defined by the two phases of leaching of carbonate and silica cement under the influence of carbonic acid and especially of organic acids as well as by a phase of pressure solution of quartz with the eventual transport of silica in an alkaline environment. These three phases overlap in time and space thereby actually controlling the alternative nature of the upward transport of
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masses of carbonate and silica and leading to the cyclic transformation of the reservoir rocks under the conditions of permanent burial of the overlying sandstone horizons which are blocked or of low mineralogical maturity. The establishment of good reservoir properties in the Paleozoic strata by leaching after the Hercynian orogeny under the influence of regional epirogenesis may be considered as a special case of reservoir decompaction. Conditions favourable for the formation of good reservoirs are developed in the domes on ancient elevated zones where the top of Paleozoic sandstones and compact quartzites (especially of the CambroOrdovician) has been directly subjected to erosion, mainly of that following the Hercynian phase. The erosion which proceded under epigenetic conditions notably improved the reservoir properties regarding capacity of accumulation and filtration. It is exactly under such conditions that the group of reservoir fields of the EI-AgrebHassi Messaoud chain formed. The optimum depth of erosion of the various fields varies as a function of their primary characteristics and of the thickness of the productive sandstones. The large most productive portion of the Hassi Messaoud rise is tied to an erosive zone of 4o-12o m depth. In the E1-Agreb structure the best production is observed in an erosion zone of 30-70 m depth (Figs. 4.19, 4.20). The analogous conditions favouring the formation
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182
Chapter 4 . Reservoir Decompaction and Formation of Accumulation Capacity
of good Cambrian reservoirs by erosion also exist around the Dahar dome which makes this zone rather promising. In regions with sandstones with good reservoir properties at the base of the Triassic, these horizons form good traps whereas the underlying Cambrian sandstones have become non-productive. This situation is encountered in the Telrhemt dome where the Cambrian deposits around the edge of the structure indicated have suffered an erosion of the same magnitude as those of the Hassi Messaoud. However, despite this erosion the Cambrian sandstones do not harbour any economic hydrocarbon accumulations as these materials were trapped by the Triassic sandstones of the topographically higher Hassi R'Mel structure.
Chapter 5
General Geochemical Features of Generation, Migration and Accumulation of Hydrocarbons
5.1 Geochemical Characterization of Potential Source Rocks, Hydrocarbons and Burial Histories 5.1.1 Silurian Graptolitic Source Clays High geothermal gradients are known to be confined to the zones of an uplifted basement, especially in the Hoggar massif and Ougarta shield. The high gradients tend to increase along the E1-Biod ridge, whereas in the Ghadames Basin, with its thick Mesozoic deposits, the geothermal gradients are comparatively low. The regions around the Hassi Messaoud oil field (HM) exhibit relatively low gradients of 2.2-2.7 °C / loo m, and those around the Hassi R'Met gas-condensate field (HR) somewhat higher gradients, 2.7-3.2 °C / 100 m. Ultimately, the formation temperatures of these two oil pools are roughly the same, since the higher HR temperature gradients are counterbalanced by the deeper location of the HM reservoir. It is clear that a number of other factors are responsible for the type of hydrocarbons trapped in these two oil pools. Therefore, in order to gain a deeper insight into the origin of HR and HM hydrocarbons, one must necessarily account for a possible hydrocarbon sourcing from both the Silurian and Devonian shales. The Silurian organic-rich source rocks, located in the area of the Berkaoui oil pool, and in the northwest and north-east of HR, may be regarded as a potential hydrocarbon source for HR. Carboniferous oil-source rocks entered the oil-window in the Nezla (NL) region north-west of HR. These source rocks may be the source for an oil fringe around this gas-condensate pool. The HR hydrocarbons might be generated in the Silurian source rocks in the west (Berkaoui region) or in the south, and could migrate through the Triassic clastic rocks prior to filling the Cambrian HM reservoir. If so, the HM reservoir structure would be filled with hydrocarbons both by lateral and vertical migration from the adjacent Triassic rocks. However, it is highly likely that the HM hydrocarbons were generated in the Devonian source rocks in the western Ghadames Basin. This petroleum could initially migrate via the Lower Devonian sandstones and then via clastic Triassic rocks. On the whole, although the Silurian source rocks are found at present in the "gas window", they have never been strongly warmed up. This apparent contradiction may plausibly be explained by the uplift and erosion of the Paleozoic sediments during the Hercynian orogeny. In many instances, the present-day subsidence depth of source rocks is smaller than it was before the Hercynian uplift. In the Illizi Basin, in particular in its southern and western regions, the deepest subsidence took place prior to the Hercynian uplift. A slight uplift could not cause a substantial temperature decrease;
184
Chapter 5 • Geochemical Features of Hydrocarbons
consequently, the maturation of organic matter continued, although at a somewhat slower rate. Ultimately, the maturity level rose higher than might have been expected for the present-day temperatures. In the Ghadames Basin, however, an appreciable uplift took place, acting to inhibit somewhat the maturation of organic matter (OM). When, past the Hercynian orogeny, the subsidence started going down to the pre-Hercynian depth, the maturation continued at a greater rate. In reality, this maturation rate growth took place in the beginning of the Cretaceous after a thermal "pause" longer than loo million years. The regional estimate of the average total organic carbon (TOC) takes into account both the level of organic matter maturity (since the areas with overmatured organic matter show low TOC values) and the original depocenters. At present, only some areas with the average TOC in excess of 2% have been observed. These areas are, as a rule, associated with the original depocenters which are not at the present time overmature. Apparently, the organic matter-bearing sediments in the adjacent depocenters have suffered a rapid subsidence, which provided for the preservation of kerogen. In overmature depocenters, the organic matter contents were reduced at the expense of hydrocarbon formation; therefore, their source rocks are not necessarily expected to exhibit high TOC values. The available geochemical data provide evidence that the overmature areas with low TOC were initially highest in TOC. The maturation thus proceeded at the expense of depletion in TOC. Viewed in this aspect, two regions, the northern Sbaa and Ghadames, are of special interest. In the latter basin the original TOC contents were higher than the present-day residual TOC owing to the mature kerogen which was beyond its peak of hydrocarbon generation. This implies that the oil generation potential in this region is higher than might have been expected on the basis of the average TOC values only. This issue will be dealt with in greater detail in the following section concerned with modeling studies. In the north of the Sbaa Basin, the original TOC contents reach 9 % which is markedly higher than the present-day average (about 3%). This area may be promising as a prolific hydrocarbon generator. The other area of interest is the Triassic province. In the Berkaoui region, near the Gellala oil pool (Takhouht area included), a zone with both high initial and presentday TOC values is distinguished; it has been assigned to high-potential Silurian source rocks (Table 5.5) which might be feeders for the known Hassi Messaoud and Hassi R'Mel oil and gas pools. The effective-to-general shale thickness ratio is a factor which determines the ability (or inability) of hydrocarbons to migrate from source rocks. In this context, of definite interest for evaluating the amount of generated hydrocarbons were also the less thick shale beds in the Ghadames, Illizi, Triassic and Reggane Basins. 5.1.2 Devonian Source Shales
The Devonian source shales, when compared with the Silurian, appear to be markedly less mature than the Silurian; they have developed to the stage of a gas-and-condensate window only in certain locations of the area of their occurrence. This situation is most clearly observed in the Triassic Basin, where the Devonian thickness at depths greater than 3 km exceeds 1.o kin. Here the organic matter maturity in the Silurian and Devonian shales is markedly different. Indeed, in certain regions of the Devonian roof,
5.2 • Generation and Directions of Migration
185
the organic matter is immature (Ro = o.4-o.5), whereas in the Silurian roof the organic matter is overmature (R > 1.8). It should be kept in mind, however, that mature OM in the Devonian shales occurs within this range - a point which not should be disregarded in evaluating the gas-and-oil potential and the hydrocarbon migration routes. In turn, the initial and present-day contents of the TOC in the Devonian shales as welt as its occurrence in the basins are quite different from those of the Silurian. Thus, the content of TOC in the Late Devonian shales in the Mouydir, Ghadames and Illizi Basins is very high. Here the TOC values are frequently much greater than those of the respective Silurian rocks, especially in the north of the Ghadames and Mouydir Basins. On the other hand, the TOC contents in the Late Devonian shales tend to decrease westwards across the platform. All these features presumably bear relevance to the alteration in both the direction of spread and the quality of clastic material vis-avis the Silurian. On the whole, the Devonian shales, reaching 2 km in thickness, are of greater occurrence than the Silurian shales. Of special interest are the Ghadames and Illizi Basin areas where the thickness of Late Devonian shales is in excess of 5o0 km. To briefly summarize, the Late (and Middle) Devonian source shales are potentially of commercial interest in the Ghadames, South Timimoune and Reggane, whereas the Silurian source rocks are potentially of commercial interset in the basins of the Triassic Province, North Sbaa and North Timimoune. In these regions, these two source horizons display a similar potential for hydrocarbon generation. 5.2 G e n e r a t i o n and Directions of M i g r a t i o n 5.2.1 Generation in the Silurian Source Rocks Silurian Source Shales in Pre-Hercynian Time. There are two regions in which the Silurian is of interest as regards the migration routes in the pre-Hercynian time. One of these comprises the East Timimoune, Ghadames and Triassic Basins. In those areas, the migration proceeded northwards, to a region north of the Hassi R'Mel and Hassi Messaoud oil and gas pools. The hydrocarbons, generated within this region, could presumably be all degraded owing to the uplift and erosion during the Hercynian orogeny. Another region of interest is the region encompassing the Ghadames and Illizi Basins. The hydrocarbons generated within this region appear to have migrated to the numerous present-day oil pools (the F-6 horizon) of the Early Devonian. The amount of hydrocarbons generated by that time was not great; the essential point was that the hydrocarbons migrated to reach, for the most part, those oil pools. Moreover, the general routes of migration towards those oil pools have not changed much since the preHercynian time. Silurian Source Rocks in the Late Triassic. In the Late Triassic, the Silurian zones o f hydrocarbon generation extended, in all likelihood, over a larger part of the platform and local areas of the Ghadames and Illizi Basins. The migration along the E1-Biod ridge was directed southwards far from the buried zones to more uplifted and exposed zones, along the Ghadames-Illizi Basin boundary, invariably tending towards the resevoirs of the present-day oil pools at the F-6 horizon (Lower Devonian).
186
Chapter 5 • Geochemical Features of Hydrocarbons
Silurian Source Rocks at the Present Day, The migration has not stopped continuing towards the F-6 horizon oil pools such as Tin Foy4 and Tabankort in the Illizi Ba~ sin. The migration of generated petroleum proceeds also in the south of the Tindouf and Reggane Basins. Likewise, in the North Sbaa, E1-Biod and South Timimoune Basins hydrocarbons are generated and migrate toward exposed zones. 5,2.2 Generation in the Devonian Source Rocks Devonian Source Rocks in the Pre-Hercynian Time. In the pre-Hercynian time, the Late Devonian was not in possession of large sources capable of generating hydrocarbons. The hydrocarbons generated in the Timimoune Basin might have migrated in any direction. The hydrocarbons that migrated north and west should have been degraded during the Hercynian uplift, whereas the hydrocarbons that migrated in any other direction might have become entrapped. Devonian Source Rocks in the Late Triassic. By the end of the Triassic, the Late Devonian shales in the Illizi and Ghadames Basins started generating hydrocarbons which are at present confined to the Late Devonian reservoirs. In the west of the platform, the hydrocarbons migrated to both subsided and exposed zones. The former, in all probabilitT, were degraded, while the latter could be entrapped beneath the Triassic sak-bearing horizons. Devonian Source Rocks at the Present Day, The present-day Devonian source rocks do not differ much from those of the Silurian. In the Illizi and Ghadames Basins, the hydrocarbons appear to migrate towards the present-day Devonian oil pools in the south and towards the Triassic pools in the north. In the southern part of the Tindouf and Reggane Basins, petroleum might have been degraded on its exit to exposures. In the North Sbaa Basin, the hydrocarbons migrate in various directions in the HBZ area. This formation had not reached the "oil window" phase by the Late Devonian; nonetheless, it may hold promise owing to the migration from deeper horizons. 5.3 Geochemistry of the Triassic Province 5.3.1 Source Rocks in the East of the Province (Ghadames and Illizi Basins) On the whole, the Carboniferous system in the east of the Triassic Province (Ghadames and Illizi Basins) has low to medium contents of terrestrial gas-trend kerogen (TOC = o.1-1%). Because of the poor enrichment in organic matter and limited aerial spread, this system is believed to be a minor source of gaseous hydrocarbons in the Triassic Basin, despite the occurrence of fairly massive Carboniferous shales in the Ghadames and Illizi Basins. In the Late Carboniferous period, kerogen had not reached the main phase of hydrocarbon generation. The structural and stratigraphic traps confined to these sediments hold promise for hydrocarbons.
5.3 - Geochemistry o f the Triassic Province 2000'
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In the Ghadarnes Basin, there are thick Devonian shales (Fig. 5.1) containing large amounts of thermally mature amorphous organic matter with a high percentage of sapropelic compounds. These shales are believed to be an excellent source of liquid hydrocarbons. Extractive compounds of hydrocarbons and the distribution of liquid hydrocarbons in petroleum suggest that the Devonian facies were prolific producers of petroleum, high in both quality and density. The underlying Devonian and overlying Triassic reservoirs, communicating with the maj or Devonian migration routes, offer a very promising prospect. In the Illizi Basin, the Devonian source rocks are essentially congeneric with those of the Ghadames Basin, but are less massive (Fig. 5.1). In the Triassic Basin, the Devonian shales are depleted of organic matter and hold little promise as potential source rocks.
188
Chapter 5 - Geochemical Features of Hydrocarbons
Fig. 5.2. Isopach map of Silurian shales of the Triassic Province
The Silurian sediments in the western Triassic Basin (Fig. 5.2) contain large amounts of very mature organic matter and are therefore potential sources of hydrocarbons. They feature a high percentage of low-molecular normal paraffins. The Silurian shales in the southern Ghadames Basin and the northern Illizi Basin feature a moderate to good enrichment in organic matter (TOC = 0.5-1.5%). This amorphous organic matter is in the late stage of oil generation. The petroleum of Devonian and Triassic reservoirs immediately adjacent to these shales appear to be of Silurian origin. 5.3.2 Source Rocks in the North of the Province (Oued eI-Mya and Triassic Basins) The present-day overall distribution of the Paleozoic shales (Ordovician, Silurian, Early Devonian) was determined by the sedimentation conditions and by the rate of Hercynian erosion of various units in the region. Over the geological history, the thickest Paleozoic shales were localized in the south, south-west and west of the region. Here the primary thickness was at a maximum, whereas the Hercynian erosion rate was at its minimum. At present, the overall thickness of shale beds is 6oo-7oo m in the south, 280-660 m in the west, and 22o-46o m at the center (see Figs. 5.1, 5.2). The major source rocks in the Oued el-Mya Basin are Silurian and Devonian shales and, to a certain extent, Ordovician shales (Figs. 5.1, 5.2, 5.3). The organic matter of the
5.3. Geochemistry o f the Triassic Province .....-30
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19o
Chapter 5. Geochemical Features of Hydrocarbons
Ordovician shales is sapropelic. The Ordovician source rocks in the southern part of the region were the only to enter the oil window at the end of the Paleozoic, when hydrocarbons were mainly generated in these deposits, At present, the gas-and-condensate generation is possible here. In the rest of the basin, these rocks entered the main phase of oil generation in the Late Cretaceous only. The Silurian and Early Devonian organic matter is typically sapropelic, mixed, and humic (Fig. 5-3). It was intensely generating hydrocarbons in the south of the region in the Pateozoic, and in the areas of the other basins in the Mesozoic. At present, the organic matter may presumably persist either at the end of the oil window, or in the gas window and low-molecular hydrocarbons. The radioactive shales of the Early Silurian are good source rocks in the central and north-eastern areas of the region. The Late Devonian and Carboniferous periods have provided satisfactory source rocks in the south-west and north-east of the region. In the Triassic, ]urassic and Cretaceous columns, of common occurrence are rocks depleted of organic matter and exhibiting a low maturation level (Ro = o.3-o.5%), which resulted in the low yield of the source rocks unrealizable under such conditions. 5.3,3 Characterization of Petroleum in the Eastern Area of the Province
On the whole, all the oils appear to be chemically alike, which is suggestive that they are generated from the same or congeneric source rocks. The distributions of saturated hydrocarbons (Clo) and the pristane/phytane ratio do not feature conspicuous distinctions. However, an analysis of the gasoline C4-C z fraction has revealed certain distinctions among the oils. The following five petroleum types could be distinguished from the compositional analysis of normal, branched and cyclic C7 species: 1. Triassic: Hassi R'Mel, Oued Noumer, A~t Kheir, Ben-Kahla, Berkaoui, Guellala,
Takhoukht (TKT), Dra el-Tamra, Makouda, Hassi Keskassa. 2. Cambrian: Hassi Messaoud, Gassi Touile, Hassi Chorghi, Rhourde Nouss. Devonian: Tamadanet. 3. Devonian: Mereksen, Stah (F4, F6). 4. Devonian: La Recul6e, Tiguentourine, Gazel (F v F4, F6). 5. Triassic: E1-Bourma.
These distinctions bear relevance to the thermal maturity and trapping time; however, their relation to the source rocks is evident and cannot be questioned. More mature oils are confined to the Triassic, Nezla, Gassi Touile, Hassi Chorghi and Rhourde Nouss reservoirs; less mature oils are confined to the Triassic reservoirs El-Bourma and to the Devonian (F3) Stah oil pool. The paleotectonic analysis has revealed two structural groups for the trapped hydrocarbons. On the one hand, there are relatively ancient structures with a good cap rock which were existent in the ~l¥iassic (Rhourde el-Baguel, Messdar, E1-Bourma, Khamadat el-Bourma, Hassi Keskassa); on the other hand, there are recent structures that have formed in the late Lower Cretaceous (North Nezla, Gassi Touite, Hassi Chorghi, Rhourde Nouss). Based on the hydrocarbon ratios, in particular the alkane-to-cyclic ratio in the low-molecular fractions, two genetic
5.3 • Geochemistry o f the Triassic Province
19~
petroleum groups can be distinguished. The first group encompasses the oils of Triassic reservoirs North Nezla, Gassi Touile, Hassi Chorghi and Rhourde Nouss. The second group includes the oils of the Cambrian (Rhourde el-Baguel, Messdar) and Triassic (E1-Bourma, Khamadat el-Bourma, Hassi Keskassa) reservoirs. The oils of the Stah, Mereksen and North Tamadanet deposits are compositionally close to the second group of oils. In this classification, the hydrocarbon ratio reflects the decisive role of some or other of the components of the primary organic matter (steranes, tristerpanes, fatty acids). Thus, the oils of the first genetic group are associated with recent formations, whereas the oils of the second group should be assigned to ancient formations (irrespective of the reservoir age). This fact provides an explanation of the occurrence of various petroleum types in the Triassic formations, that is, the age of the structure, rather than the age of a reservoir, is mainly responsible for the distribution of different genetic petroleum types. This conclusion sheds a new light on the time of the regional hydrocarbon migration from the source rocks. The migration of oils of the first genetic group fell into a period between the Late Triassic and the Early Cretaceous, whereas the oils of the second genetic group became entrapped over the course of a period after the Early Cretaceous.
A study of genetic features of the Devonian petroleum in the north of the Illizi Basin (which merges into the southern flank of the Ghadames Basin) will be of prime importance for the gas-and-oil prospecting of the Ghadames Depression, at present poorly explored. Since the Ghadames Depression and the Illizi Basin constitute a single sedimentary basin in the Paleozoic, the presence or, conversely, the absence of genetic links between the oils of these two formations may prove to be crucial for the discovery of hydrocarbons entrapped in the Devonian formations of the Ghadames Depression. 5.3.4 Petroleum Types and Their Variations in the North of the Province An analysis of distinctive features of the occurrence of oils in the northern area of the platform enables their classification into two types:
1. Paraffinic type (saturated n-alkanes) occurring in the Laghouat, Gassi Touile, Hassi R'Mel and North Nezla oil pools. All these oils are confined to Triassic reservoirs, except for the North Nezla which belongs to an Ordovician reservoir. 2. Naphthene-paraffinic type (saturated cycloalkanes, isoalkanes, more rich in naphthenic and aromatic components than the former type), occurring in the Cambro-Ordovician Hassi Messaoud, E1-Gassi, Rhourde el-Baguel, Silurian Oulouga, Triassic Berkaoul and E1-Bourma oil pools. Such a distribution of oils of these two types reveals no significant criteria - either in geography or in age - for the reservoirs. Besides, geochemical variations have been established within each of the specified types, potentially associated with the depth of source-rock subsidence and the migration pathways. Compositional variations in paraffin have also been observed, presumably associated with the physical segregation during the migration. Certain oil fields (E1-Gassi, Berkaoui) were presumably fed from different source rocks. Variations in C1,C2 and aromatics are recorded even within the same oil pool (for example, Hassi Messaoud). These variations in chemical corn-
192
Chapter 5 . Geochemical Features of Hydrocarbons
position in the north of the platform are indicative of (1) eventual occurrence of different source rocks; (2) eventual catagenesis of different level; (3) eventual dissimilarities in the environmental and provenance conditions (aromatic species are presumed to be associated with marine conditions). In short, the geochemical history of the Saharan basins is quite complicated, and the processes occurring in one region cannot be safely extended to another one. Chemical Evolution of Oil. Two hypotheses of petroleum evolution may be suggested: 1. The primary and secondary migrations were coeval with the petroleum generation in the source rocks. The hydrocarbons were entrapped at the beginning of the oil pool buildup and continued to feed it until this deposit became closed or filled up to a maximum. 2. The petroleum generation started after a good cap had formed; during the course of subsidence, petroleum evolved to cracking stage followed by the gas formation as the submergence proceeded. 5.3.5 Conditions for Hydrocarbon Generation in the North of the Province The theoretical subsidence in various areas over time can be reconstituted with reference to the Hercynian unconformity taken for a zero level by adding successive layers of Mesozoic sedimentary layers. Since the Silurian source rocks (and, occasionally, the Devonian) closely underlie the Hercynian unconformity in the northern Sahara, the subsidence depth relative to the Hercynian surface (the bottom of a Triassic reservoir) is in fact the depth of source rock subsidence for a given area. Two major stages of hydrocarbon generation and accumulation can be defined in accordance with the two generation stages, Paleozoic and Mesozoic. Paleozoic Stage. During the Carboniferous period the hydrocarbons generated in the Ordovician and Silurian in the north of the basin (where the subsidence and high paleotemperature were the most favorable for the organic matter maturation) migrated north and north-east. This process was accelerating over the course of Hercynian orogeny as the north-eastern area of the region suffered uplift and the associated erosion. Consequently; the early hydrocarbon accumulations entrapped in the Paleozoic were destroyed in part (if not completely). Part of these accumulations were destroyed owing to the Hercynian erosion, while the other part, altered to a certain extent, was entrapped in the newly produced structural Ordovician and Silurian traps at the end of the Hercynian cyde. Mesozoic Stage. The Mesozoic stage of hydrocarbon generation and accumulation had its onset during the intense subsidence in the north-eastern area of the region, that is, during the sedimentation of terrigenic Triassic rocks and massive TriassicJurassic salt layers. In the course of subsidence, definite structures formed and evolved, concomitant with a partial or complete oil pool collapse. The hydrocarbons migrated from the destroyed traps to form new oil pools with a Triassic or Jurassic cap. Reservoirs may occur in the Triassic (Berkaoui, Guellala) or Pateozoic (Hassi Messaoud, El-
5.3 . G e o c h e m i s t r y
of the TriassicProvince
~93
Agreb) formations. Such oil pools could have started forming in the beginning of the lurassic, after the salt cap had developed sufficient density and impermeability. The subsequent subsidence of Paleozoic source rocks within the Jurassic and Cretaceous periods stimulated the generation, migration and accumulation of newly-formed hydrocarbons in the north of the region. The Triassic and Paleozoic reservoirs were filled up with the hydrocarbons generated by Paleozoic source shales. Most hydrocarbons of Paleozoic origin were located in the Triassic reservoirs owing to the lateral and, to a lesser extent, vertical migration.
Sumary of the Generation History. Two stages - Paleozoic and Mesozoic - are clearly distinguished over the course of tectonic evolution, with the predominance of subsidence processes. They are separated by the Hercynian orogeny. The area of deepest subsidence, characterized by an increase thickness of Paleozoic sediments, is localized in the south of the Oued el-Mya depression. During the course of the Hercynian orogeny, the north-eastern area of the region, much more elevated, was subjected to severe erosion. Over the Mesozoic period, the north-eastern area of the basin exhibited the highest rate of subsidence owing to the tectonic inversion, which resulted in the formation of a salt-bearing basin of TriassicJurassic age. It is seen therefore that the distinct conditions for the accumulation and maturation of organic matter were associated with the specific geological evolution in the north-east and south-west of the region. By the end of the Late Carboniferous period, when the sedimentation cycle was entering its final stage, the highest paleotemperatures were recorded in the south and south-west of the region, reaching 8o °C at the top and 125 °C at the bottom of the Gedinnian sediments. The lowest temperatures were observed in the north of the region: 80 °C at the Ordovician bottom and 50 °C at the Silurian bottom. By the end of the Cenomanian age, the paleotemperatures rose from loo to 125 °C at the Silurian bottom and from 8o to 11o °C at the Devonian. At the present time, the temperature remains at the same level, except for the north-western and south-western areas where it reaches 13o-15o °C. An analysis of the profiles of present-day temperatures and paleotemperatures in the Paleozoic sediments has indicated favorable conditions for the hydrocarbon generation mostly in the south and south-west of the region. As to the hydrocarbon generation in the central and northern areas, it might have taken place mainly in the Mesozoic. 5.3.6 Conditions for Hydrocarbon Generation in the Eastern Area of the Paleozoic Province
The complex evolution of the Ghadames Depression was, on the whole, a decisive factor for the accumulation conditions, for the organic matter type of the Paleozoic sediments, and the degree of their subsequent conversion, as well as for the eventual migration and accumulation of hydrocarbons. Geological processes of crucial importance are recognized to be the Hercynian orogeny and the erosion of Paleozoic deposits. Another important geological process was the regional subsidence in the Mesozoic, responsible for the formation of hydrocarbon pools. The subsidence occurred in par-
194
Chapter 5 • Geochemical Features o f Hydrocarbons
f
30° r--
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Fig. 5.4, Studied fields
allel with the sedimentation of massive evaporate layers which served as a cap for more than 5o% of the total depression area. The region, located south of the line drawn from Ain Sefra to Zaouia el-Kahla (north-east to south-west) (Fig. 5.4) is a zone where hydrocarbons were generated during the Paleozoic, where the depth of the source rock subsidence was greater than 2.o km. In cases where the subsidence was 3.5 kin, gaseous hydrocarbons formed. Tissot et al. (1975) believed that the Cretaceous period was the earliest time for generation and migration of petroleum in the Hassi Messaoud and Rhourde et-Baguel oil pools. They have put forward the following arguments: (1) Rhourde el-Baguel was an Early Cretaceous horst, with no clear structuration prior to the faulting; (z) the subsidence curves show the deepest burial to have taken place in the Mesozoic (see Fig. 5-5)Tissot and coauthors clearly underestimated the Paleozoic burial, since it seems to be highly improbable that the Silurian shales were buried a mere 1 km beneath the Devonian and Carboniferous sediments which at present are completely eroded by the Hercynian orogeny. Presumably, there were Devonian and Carboniferous sediments at least 2 km thick in the central, deepest subsided part of the depression in the Haiad area (HAD). The uplift that had occurred by the end of the Middle Devonian seemingly eroded the Late Silurian shales to the Middle Devonian on the flanks of the Ghadames Basin, but left the central part of the basin (that is, the eventual site of oil generation) less affected. The paleotemperature regime of the region is suggestive of the fact that the Silurian shales might have reached the oil generation temperature in the Early Carboniferous. Moreover, the uplift at the center of the basin was minimal, so that the necessary condi-
5.3 • Geochemistry of the Triassic Province 0
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tions for hydrocarbon generation were not interrupted by the Hercynian movements. Given the present-day overmature state of organic matter at the basin's center, the residual potential after the Mesozoic subsidence is not thought to be significant. However, the Silurian shales in the areas adjacent to Hassi Messaoud persist at a relatively low level of maturity ("oil window" level), since they were submerged to shallower depths.
Geological History of the Region in the Paleozoic. The sediment reduction of Paleozoic formations in the south-east of the region at the beginning of the Paleozoic and the absence of Cambrian sediments in the south of the basin (SED-1 area) may be explained by the fact that the comparatively elevated basement in this region suffered subsidence at a lower rate. The same rate of regional subsidence was typical of the central and north-western parts of the region. As to the overall thickness of the Paleozoic formations of the Messdar area (MDR) and Fort Lalleland (FLD), it was some 3.o km and somewhat greater at the centre of the depression. In the Rhourde Nouss (RN) and Keskassa (KA) areas, the Paleozoic formations also could be significantly thick. Thus, by the end of the Carboniferous the average depth of subsidence was about 3.o km for the Ordovician, 2.8 km for the argillaceous Silurian, 2.5 km for the Early and 2.2-1.5 km for the Middle and Late Devonian (see Fig. 5.5). The Hercynian orogeny entailed a very significant uplift of the western and northern parts of the depression, which resulted in the formation, in the central part of the area, of a depression which was progressively subsiding in the south-east direction to form a saddle in the SEDq area. The Paleozoic sediments were highly eroded in the northern and western areas where the Ordovician, Silurian and Early Devonian deposits were exposed directly beneath the Hercynian unconformity.
I96
Chapter 5 . Geochemical Features of Hydrocarbons
Organic Matter of Paleozoic Rocks. The organic matter suffered a profound alteration as early as in the Paleozoic by the action of high temperature and pressure caused by the regional subsidence. Having assumed the temperature rise to be 1 °C per 34 m, one can see that by the end of the Carboniferous the Ordovician rocks were "heated" up to 113 °C, the Silurian clays to lo7 °C, the Early Devonian to 98 °C, and the Middle and Late Devonian, to 9o and 69 °C. It is natural to assume that by the end of the Carboniferous the organic matter (of sapropelic and humic types) of Ordovician, Silurian and, possibly, Early Devonian was capable of generating liquid and gaseous hydrocarbons which, when entrapped, could form oil pools of industrial interest. Presumably, the eroded Silurian and Devonian sediments were also rich in organic matter, that is, they exhibited a high hydrocarbon potential. The regional Hercynian uplift and the subsequent erosion "inhibited" the hydrocarbon generation and caused a partial destruction of accumulated hydrocarbons. The Paleozoic cycle of hydrocarbon generation and accumulation terminated precisely during these processes.
Mesozoic (Triassic) Province. In the region north of the Mn Sefra-Zou~a el-Kahla (Fig. 5.4), the hydrocarbon was generated mainly during the Mesozoic. In the case of Silurian source rocks, represented mainly by radioactive shales, thickness lines were drawn with reference to the Ordovician roof. For the Devonian source rocks, the thickness lines were drawn with reference to the Early Devonian roof (Siegenian). The Lower Devonian roof appears to be the only reliable datum mark for intra-Devonian correlations. Geological History of the Province During the Mesozoic. During the Triassic and Jurassic, the entire Ghadames Depression was subjected to inverse tectonic movements which caused an intense subsidence in the northern, north-western and western areas, uplifted at the end of the Paleozoic.The salt-bearing basin that had formed in this region extended far beyond the depression boundaries. This fact was reflected in the occurrence of massive evaporites in the Messdar, E1-Kret and Fort Lalleland areas, where the evaporites exceeded 1.o km in thickness, much the same as in the Keskassa area (9oo m) also. Shallowest subsidence was observed in the Rhourde Nouss and Ektaia zones as well as in the south-eastern area of the region, that is, in the zones of deepest Paleozoic subsidence where no evaporites were deposited. The thickness of the Triassic and Jurassic sediments was greatest in the deepest submerged zones of the depression, reaching 1.5-1.9 km. Within the overall subsidence, the E1-Kter, Oued el-Tekh, Ha'iad and Ber Rebaia areas experienced the most intense subsidence. In the Cretaceous, the subsidence in the south-east, west and north-west of the region proceeded at a slower rate, with the resultant formation of the Ghadames Depression in its present-day configuration. Thus, the Ghadames Depression is a superposed tectonic Mesozoic structure whose closure took place during the Cretaceous.
Organic Matter and Hydrocarbons. A t the end of the Mesozoic subsidence, the organic matter, deposited in the Pateozoic, once again became subjected to high temperature and pressure and suffered alteration. In the intensely subsided zone, the Ordovician and Silurian sediments reached a depth of 3.8-5.5 kin, Lower Devonian 3.4-4.7 kin, Middle and Late Devonian 2.8-4.2 km and Carboniferous 2.2-3.5 km. Once again, making use of the present-day geothermal gradient of 34 m / 1 °C, we can determine that by the end of the Mesozoic the Ordovician and Silurian sediments
5.3. Geochemistry of the Triassic Province
197
were heated to 137-187 °C, Early Devonian to 125-148 °C and Carboniferous to 89-128 °C. One will infer therefore that during the Paleozoic cycle, only the Ordovician and, in part, Silurian sediments could have realized their petroleum-generating potential, whereas during the Mesozoic, all Paleozoic sediments, including the deepest subsided Carboniferous sediments, were involved in the hydrocarbon generation process (see Fig. 5-5)-A point to be noted is that the generation of liquid hydrocarbons in the Ordovician, Silurian, Devonian and, in part, Carboniferous sediments was at its completion by the end of the Cretaceous. In the course of regional subsidence and tectonic restructuration of the depression, the newly formed hydrocarbons kept accumulating in Paleozoic traps. Here were also trapped the hydrocarbons that migrated from destructed pools. The severe Hercynian erosion which caused the denudation of Paleozoic reservoirs was subsequently a decisive factor in the migration of Paleozoic fluids to the Triassic reservoirs. The massive evaporite layers and appropriate structural conditions favorably contributed to the buildup of large hydrocarbon accumulations in the Triassic reservoirs. The Mesozoic cycle of hydrocarbon generation was completed by the end of tile Cretaceous. The subsequent geological processes produced no substantial alteration in either tectonics of the region or distribution of the oil pools. For the Silurian source rocks, any Mesozoic trap that was formed in the pre-Aptian period stands a good chance of becoming filled up with oil from these rocks. The traps formed during the Austrian orogeny and closed in a post-Aptian period (during the Late Cretaceous or, possibly, the Tertiary) are probably gas-bearing. Promising traps or privileged structures are those located close to the subsided zones in which the Devonian source rocks have escaped erosion. The eastern province is one of such privileged zones because it remained active during the aleozoic and Mesozoic. By the end of the Cretaceous, the Ghadames Depression started generating gas. 5.3.7 Petroleum to Source Rock Correlations
Numerous analyses of oils from the Ghadames Basin (STAH, MRK, WT) possess isotopically light carbon corresponding to kerogen from marine algae and exhibit 5~3C values between -3o.1 and 3o.6%o. It has been noted that oils from pre-Carboniferous source rocks exhibit slightly heavier carbon isotopes than the corresponding kerogens. It can be suspected that the Silurian shales constitute the source rocks in the Hassi Messaoud area as well as in the central and southern parts of the Ghadames Depression (Wadi el-Teh region) although it cannot be excluded that the oil at Wadi el-Teh (WT-1) was locally derived from the Devonian shales. These shales contain amorphous organic matter in association with pyrites indicative of a reducing sedimentary environment capable of preserving kerogen which eventually will result in petroleum. The oils at WT-1 are too mature to be correlated to biological markers. Tissot et al. (1973) tried to use aromatic steranes as correlation criteria. This method, however, finds little support in the data which, on the other hand, indicate an increase of maturity with age. A classification of the oils into "Cambro-Ordovician", "Upper Silurian-Lower Devonian" and "Upper Devonian-Lower Carboniferous" proposed by some authors actually reflects only the effect of maturation and not any differences in
198
Chapter 5 • Geochemical Features o f Hydrocarbons
source rocks. Although the hypothesis possesses a low probability, it cannot be excluded that the oil at WT-t is a displaced accumulation which is forming today from supermature Silurian kerogen and is preserved in a trap on a structural grid. In order to throw light on the problematic correlation between oils and their source rocks we have investigated oils and bituminoids by mass spectrometry together with gas and liquid-phase chromatography. In the Hassi Messaoud and Rhourde el-Baguel fields the oils in the Cambrian and Ordovician sandstones are isotopically light with fi13C = -29.18 to -29.75%o, typical of marine organic matter largely transformed by microorganisms. The oils are characterized by iCI9/ iCzo = 1.o-1. 4, a generally low concentration of isoprenoids in the saturated hydrocarbon series (Pr/nC~z= o.2-o.3) and a clear preponderance of hydrocarbons with a low molecular mass in the n-alkane series (Table 5.1). The Devonian oils in reservoirs D1-T3 of the Illizi Basin are characterised by a fi13C with a rather narrow spread of-28.1o to -29.85%0 and their pristane/phytane ratio of
Table 5.1. Saharan oil isotope and geochemical features No.
Field
No.ofwell
Age
-~13C(%)Pr/P f
Pr//lC17 K=nC15_lz/nC25_27
I, Triassic Province basins I
Hassi Messaoud
152
E-O
29.75
1,2
0,3
2
Hassi Messaoud
152
E-O
29,51
1.0
0.3
5.5
3
El Gassi
7
E-O
29.90
1.2
0.2
4.8
4
Rhourde ei-Bague[
18
E-O
29.t8
1.6
0,2
5.7
5
Rhourde el-Baguel
25
E-O
29.26
1.4
0.2
5.5
6
Guellala
5
S
29.46
1.5
0.2
4.4
5
7
Guellala
8
Hassi R'Mel
9
4.1
T3
30.46
1.4
0.2
2.8
38
T3
30.47
1.8
0.3
3.0
Rhourde Chouff (condensate)
1
T3
28.52
10
Kef el-Argoub
4
T3
30.80
t.6
0,2
2.8
11
Takhoukht
I
T3
29.34
IA
0.3
3.6
2. llliziBasin 12
Tiguentourine
37
DI (F-6) 28.10
1.9
0.2
5.6
13
Tiguentourine
115
D3 (F-2) 29.30
1.7
0.3
3.8 2.0
14
Edjeleh
29
D2 (F-4) 29.13
1.3
0.9
15
Zarzaitine
65
D2 (F-4) 29.85
1.5
0.3
3.6
16
Mereksen
16
DI (F-6) 29.85
1.8
0.3
3.1
17
Tin Fouye
2
D3 (F-2)
28.91
1.5
0.3
4.5
18
Alrar
5
T3
28.44
-
-
T3
29.95
0.3
3.7
3.
Timimoun Basin
19
Hassi Lato
1
1.7
5.3 • Geochemistry of the Triassic Province
199
2.5 is the highest in the succession studied (Tables 5.2, 5.3, 5.4). This appears to be related to an increase in the proportion of kerogen of terrestrial origin in the total amount of organic matter in the source rock. The isoprenoid content here is also low, as it is in the oils of Cambro-Ordovician reservoirs. The oils from the Triassic reservoirs of fields like Hassi R'Mel (HR), Kef el-Argoub (KG), Takhoukht (TKT) or Hassi Lato (LT) are isotopically similar to the oils of the Lower Paleozoic. An exception are the isotopically relatively hea W condensates of Rhourde Chouff, Tiguentourine and Akrar (Table 5.1) as well as the oils of the Itlizi Basin in general (Table 5.4). The oils of the Paleozoic and the Triassic are exceptionally similar in composition to each other. This chemical homogeneity of the oils is characteristic not only of the vertical succession of the productive complex but also for the various sub-basins of the Triassic Province. The pristane/phytane ratio of 1.3-1.75 is typical and there is a clear dominance of low-molecular hydrocarbons in the n-alkane series with nCl~-vl nC2>~7ranging from 2.0 to 5.7 (Tables 5.2, 5.3, 5.4)- The ratio of pristane to nC v does not exceed o.2-o.3, indicative of sufficiently a high level of oil catagenesis and of the absence of influences of epigenesis on the composition of the oils.
Table $.2a. Pristanelphytane data for oils and bituminoids from Saharan basins (Illizi, Atlas and
Oued el-Mya) Field
Well
Depth (m)
Age
Pristane/ phytane ratio
Oil (0) or bituminoid (B)
Carbonif.
1.63
O
Devon. Devon.
1.44 1.50
O O O
Illizi Zarzaitine
ZR-t 15
523
Tiguentourine
TG-1 t 5 TG-128
1 055 1 022
- 1 053 - 1 039
lama
TAM-1
2082
- 2049
Ordov.
1.53
Amasak
AMA-3
1 998
-2008
Ordov.
1.47
O
Tin Fouye
TF
Devon.
1.65
O
Devon.
1.45
O
Devon.
1.65
O
Devon.
1.39
O
Devon.
1.48
O
El Abed Larache
EAL-12
Edjeleh
DL
1 246
Ohanet
OTN~I 19
Askarene
ASK
Gara
Gara-2
1 937
- 1 939
Devon.
1.48
O
Guelta
GLT
2 686
- 2 697
Devon.
1.47
O
Zarzaitine
ZR
Devon.
1.58
O
Rom
ROM-1
Devon.
1,63
O
Cretac. Eocene
2.53 1.63
O O
2 365
- 1 250 - 2 369
Atlas Oued Gueterine
GKN-1 GKN-1
3 250
- 3 254
Oued et-Mya Hassi Messaoud
MD
Cambr.
1.31
O
Rhourde eI-Baguel
RB
Cacnbr.
1£5
O
Haoud Berkaoui
OKJ
Gassi Touil
GT
3 468
- 3 520
Trias.
1.33
O
Trias,
1.62
O
200
Chapter 5 • Geochemical Features of Hydrocarbons
The isotopic composition of the bulk carbon in the oils is sufficiently uniform and varies only within a narrow range of -28 to -3o%o (Table 5.6). On the whole, the oils are isotopically light which is typical of petroleum genetically related to marine source rocks. For the purpose of a correlation the isotopic composition of bituminoids from Silurian and Devonian shales of a very high hydrocarbon potential has been investigated ('Fables 5.5, 5.6). For the correlation of oils and source rocks we have used the carbon isotope composition of oils and bituminoids within probable source rocks as a diagenetic feature
Table 5.2b. Pristane/phytane data for oils and bituminoids from Saharan basins (Ghadames and
Sbaa) Field
Well
Depth (m)
Age
Pristane/ phytane ratio
Oil (O) or bituminoid (B)
HFR
HFR-1 HFR-1 HFR-1 HFR-I RE-1 ZAR-I ANR-I
2687.30 3 276.05 3321.15 3979.95 3454.00 3991.55 3914.90
Carbonif. Devon. Devon. Devon. Devon. Silur.
1,70 1,70 2.50 2.25 1.60 1.65 1.71
8 B B B B B B
Keskessa
KA-1 bis KA-1 bis KA--1bis ZES-1 REA-1
2 706.10 3 072.70 3 213.40 2 818.80 2773.30
Trias. 5itur. Ordovic. Silur. Silur.
1.20 1.60 1.67 1.69
B B B 8 B
El 8orma
EL8-2 ELB-2 HT8-2 HTB-2
2 394,40 2 404,35 2 434,10 2 69t .90
Trias. Trias. Trias. Silur,
1,09 1.43 0.67 1,43
B B B B
Ghadames
Keskessa
KA-2
Trias.
1.62
0
Wadi Teh
WT-2 HTB-1
2420
-2447
Trias. Trias.
159 1,75
0 0
ELB ELB-9
2 419 -
- 2430
Trias. Trias.
1.50 1.50
0 O
Mereksen
MRK-16
2 77t
- 2 796
Devon,
1.35
0
Stah
STAX
2 719
- 2 726
Devon,
t .36
0
570.0 588.0 674.0 6525
Devon, Devon, Devon. Devon.
2.33 2.36 2.14 1.83
O O O O
995 1 013.25 - 1 025 593.1 614.0 -1 013 - 1 025 592 625 764 793
Devon. Devon, Devon. Devon. Devon, Devon. Devon, Devon.
1,94 1.92 2.18 1.83 2.12 1.80 1.92 1.69
O O O O O O O O
Devon,
1.93
O
El Borma
Sbaa Toat
TOT TOT TOT
ODZ-lbis Hassi Lato
LT-1 LT-1 LT- 1 DECH-I DECH-IW LT-1 DECH-I ODZ-1
Toat
TOT-1
555.0577,75671,0633.2 -
555
57.0
5.3 •
Geochemistry of the Triassic Province
zoi
for the source of the hydrocarbons. Under the influence of biological (fermentational) fractionation and diagenetic elimination of isotopically heavy groups like -COH or OCH 3 or polarization, an isotopic composition is established in the kerogen of the source rock which then is inherited by the oil. Given the direct link between the polarity of the compounds and their isotopic fl-factor, five fractions may be distinguished for petroleum and bituminoids which are in increasing polarity: hexane, hexane-benzene, benzene, benzene-methanol and asphaltene. As a result we have obtained an"isotopic portrait" of the different oils from the Triassic Province, i.e. of the Cambro-Ordovician, Devonian and Triassic (Fig. 5.6, Table 5.6). There are two different shapes of the isotopic curves of isotopically light asphaltenes, which is, as has been established, characteristic of the lithofacies of marine source rocks (Peters et at. ~986). The oil families mentioned above may be separated isotopically into two main groups distinguished clearly by their isotopic curves. In the probable source rocks of
Table
5.3. Average of pristanetphytane ratio for oils from Oued el-Mya fields
Field
Welt
Age
No. of samples analyzed
Hassi Messaoud
MD
Cambr.
11
1,07 - 1.79
1,4
EI-Gassi
GS
Cambr.
3
1.33 - 1.40
1.4 1.7
Pristane/phytane ratio
Zotti
ZT
Cambr,
4
1.36 - 1.82
EJ-Agreb
AR
Cambr.
3
1.l 6 - 2.0
1,5
Average: (21)
1.07 - 2.0
1.5
Nord-East Guellala
GLNE
Silur.
1
1.47
Ouloga
OLG
Silur.
1
1.18
Average: (2)
1.18 - 1.47
1.3
1.54 - 2.92 1.15 - 1.62
2.0 1.5
Haoud Berkaoui
OKP
Trias.
4
Ben Kahla
OKJ
Trias.
3
Org
Org
Trias,
1
GueLlala
GLA
Trias.
3
1.43 - ],75
1.5
Nord-East Guellala
GLNE
Trias.
2
1,27 - 2,27
1,8
N, Goussa
NGS
Trias.
2
1.47 - 1.50
1.485
Takhoukht
TKT
Trias.
I
1.4
Draa-Temra
DRT
Trias,
]
2.6 2.2
1.8
Garet-Ech-Chouf
GEC
Tfias.
1
Kef el-Argoub
KG
Trias.
2
1.33 - 1,94
1.6
Average: (20)
].I 5 - 2,92
1.6
2.1
Hassi R'Mel
KG
Djorf Oued-Noumer Air Kheir
ONR
Trias.
8
2.0 - 2.7
Trias.
2
1.67 - 1.68
1,675
-lrias.
5
1.83 - Z]8
2.0
Trias.
3
1.73 - 2,20
1.95
Average: (18)
1.7
].95
- 2.70
zo2
Chapter 5 - Geochemical Features of Hydrocarbons
the oils examined we have studied the isotopic composition of the various fractions of different polarity in the Devonian and Silurian bituminoids. The first group of oils exhibits exactly the same sharp-peaked curves characteristic of the isotopic composition of the bitmninoid fractions of the Saharan shales which also contain isotopically light asphaltenes. However, the isotopic "portraits" of the bituminoids from the Devonian shales differ sufficiently from those of the Silurian and follow the isotopic image of the oils of the second group which is characterized by a spread of fi13C values similar to that of the five fractions investigated (Fig. 5.6). This group covers oils from, e.g., the Stah
Table 5.4. Carbon isotope composition (-fi13C (%0)) of oils and their fractions in Saharan basins Basin
Field
Well
Age
Fraction
General compos.
Asphaltene
Satur. Atom. Timimoune
Guvette de Sbaa
TOT-1 DECH-1 ODZ-1 bis LT-1 ZR-115
D O D D C
29.5 29.8 29.6 29.4 29.2
29,1 29.1 28.9 28.1 28.6
29.4 29.4 29.2 29.1 28.7
Ghadames
Mereksen
MRK-9 MRK-16 Stab-43
D D D
29.4 29.4 29,5
28.1 28.t 27.9
29.1 29.0 29.2
29.1 29.0
282 28.2 28,4 27.6 28.2 28.1 28,1 27,9
28.6 28,9 28.8 28.0 28.8 28.3 28.7 28.4
Stah Illizi
Askarene
ASK-107 ASK-108
Guelta Amasak EI-Adeb Larach Tiguentourine
GLT-t01 AMA-3 EAL-I 2 TG-115 TO-128
D O D D C
Gara
Tamadanet
Gara-2 TAM-1
O
28.8
28.1
28.7
Ohanet-Nord Ohanet-Sud
OTN-119 OTS-133
D D
28.9 29.0
28.1 28,0
28.6 28.6
Ghadames
Wildgat Rom Gassi Touil Rhourde Baguel
8BK-1 Rom-1 GT-I RB-1
D Pz T
C
29.9 30.2 28.6 28.9
29.3 29.5 28.1 28.6
29# 29.8 28,7 29,5
lllizi
Edjeleh DL Tin FouyeTabankort TFT TFT-I Zarzaitine ZT-I
D D D D
29.1 29.1 29.1 29,0
28,5 28,7 28.1 27.8
29,1 29.4 29.3 29.0
29.1 28.5 29.0 29.0 29.0 28.9
Oued eI-Mya HassiMessaoud Guellali~ EI-Agreb EPGassi
MD GLA AR GS
E S C G
29,9 2937 30.16 30.08
29,4 28,53 29.03 29,18
29,7 28,89 29,21 29.31
29.21 29.31
Timimoune
DECH
D
29.68
29.04
29,07
25,57
5.3 • Geochemistry of the Triassic Province -30
-29
-28
203 -27
BenzeneMethanol
-26
~i13C (%0)
~.
,e ze e
~eXanzne-
Asphaltene
~
b
BenzeneMethanol 3enzene
" D " en v. o. n, J, a~
~
-lexane- ~ ]enzene 613C(%o)
-texane_,30
-29
-28
-27 ,
-26,
-25
Fig. 5.6. Correlation of oils to source-rocks system: Carbon isotope composition for five oil (a) and bituminoid (b) fractions in Saharan fields Table 5.5. Carbon isotope composition (-613C (%0)) for bituminoid fractions from Oued el-Mya oil source rocks Well
Depth (m)
Age
Saturated fraction
Aromatic fraction
Resins
ZTE
3155 3 223
Ordovic. Ordovic.
27.43 27.38
27.15 26.32
27.81 26.35
HBA-1
3 309
Silur.
26,77
27.06
26.67
HBA-1 Nord
3 101 3 240
Silur. Ordovic.
28.89 26.22
27.81 26.24
27,98 2723
SAF-1
2 667 2685 2 815 2450.3 2 539.5 2458.68 2 822
Ordovic. Ordovic. Ordovic. Ordovic. Ordovic. Ordovic. Silur,
30.68 31.11 26.72 30.03 28.45 29.89 30,23
29.27 29,80 27,24 27,28 28,23 29.38 28.15
29.56 29.96 27.60 29,28 27,36 28,81 27.76
EA-1
FDN-1
Chapter 5 - Geochemical Features of Hydrocarbons
204
Table 5.6. Carbon isotope composition (-813C (%0)) for five oil and bituminoid fractions from the Traissic Province (Sahara) ( Z o n e F3 corresponds to Middle-Upper Devonian. T A G I - Trias argilogr4seux infGrieur. T A G S - Trias argilo-gr4seux superieur. S I - s~rie inferieure
Field
Well
Depth/ performation
Stratigraphy
Fraction
System/ epoch
Age
Asphal- Benzene- Benzene Hexane- Hexane tene methanol benzene
Oils Hassi R'mel
HR-38
Triassic
Reserv.C
28.52
27.09
26.50
28.92
29,40
Oued Namous
ONE-8
Triassic
Reserv,B
28.35
26.62
26.06
28.02
28.82
Guellala
GLA-16 3394-3534
Triassic
[~ + Serie 29.09 infer,
27,14
26.68
27.76
28.73
28.68
27.35
27.05
28.56
29.30
29.31
27,48
26.72
28.35
28.90
Hassi KA-lbis Keskessa
2710 -2733
Triassic
Draa Tamra
DRT-3
3 597 - 3 605
Triassic
HassiMessaoud
MD-64
3 350- 3472
Cambrian
29.30
27.03
27.21
28.31
30.04
Rhourde 8ague[
RB-t2
2739 -3010
Cambrian
28.82
27.76
27.46
28.14
28.50
Rhourde Nouss
RN-48
2 726 - 2 730
Triassic
TAGS
28.25
28,02
27,92
28.t8
28,60
Nezla
NZN-7
2527 -2572
Triassic
TAG
28,45
28.21
28,12
28.46
29.12
Hassi Chergui
HC-5
2 461 - 2 478.5
Triassic
TAGI
28.74
28.32
28.20
28.26
28.60
Mereksen
MRK-17
2 771 - 2 776
Devonian ZoneF3
28.87
28,70
28.80
28.86
29.20
Hassi
flTG-13
1533- t 572
Triassic
28.95
28.87
28.88
29.13
29.27
TI
TAGS
Tonareg Stah
STAH-24 2722-2726
Devonian ZoneF~
29.t8
28,94
28,96
28.75
29.00
WadiTeh
WT-t
Triassic
TAG
29.40
29,t2
29.10
29.18
29,85
3724-3731
8ituminoids Rhourde Nouss
RNSE-1 2877.4
Siludan
Gothland 27.83
26,21
25.08
27.90
29.73
Takhoukht
TKT-1
3 842.5
Silurian
Gothland 27.30
26.88
25.27
27.84
28,85
Jejessa
KA-lbi
3072,7
Silurian
Gothiand 28.20
26,67
25.78
28.32
29.16
Wadi Teh
WT-1
3 760.1
Upper Devonian
Strun.-
37.28
27,12
26.90
27.05
28.50
Mereksen
MRK-3
2515.85
Upper devonian
27.54
27,33
26,80
26.77
27.12
Tiguentourine
TG-201
1044.27
Upper De- Strun,vonian Tournis.
28.00
27,16
27.26
28.06
28,06
Tournis. Famen.
5.3 • Geochemistry of the Triassic Province
205
(STAH),Wadi el-Teh (WT), Hassi Tuareg (HTG), Mereksen (MRK), Hassi Chorghi (HC) and Nerzla North (NZN) fields. It is interesting to note that all these fields are enclosed in sandy reservoir rocks intercalated with Devonian shales (D 2 + D3) rich in rather mature organic matter. This is convincing proof of the statement that the above-mentioned deposits were derived indeed from Devonian source rocks. The isotope curve of oils from the gas and condensate field Hassi R'Mel (HR) is rather sharp-peaked (Fig. 5.6) which allows us to connect it to Silurian rocks as the source of its hydrocarbons. In the Hassi Messaoud deposits (MD) and at numerous other sites along the western edge of the Ghadames Depression like, e.g., Rhourde elBaguel (RB) and Rhourde Nouss (RN), the isotope ,portraits" assume an intermediate position in Fig. 5.6 suggesting that they were filled by hydrocarbons derived from Silurian as well as from Devonian source rocks. In addition to the obvious supply from Silurian shales throughout the Triassic Province and the region mentioned, we have every reason to accept the highly likely contribution of Devonian shales of the western part of the Ghadamas Basin to the generation of the hydrocarbons filling, in particular, the numerous deposits in the east and south of the Triassic Basin. The geochemical data on the organic matter-rich shales of the Silurian and Devonian of the Triassic, Ghadames and Illizi Basins exhibit good to even perfect correlations with the oils of these basins. The distribution of Clo-Saturated hydrocarbons, their C4-C7 gasoline composition and the geological distribution of the respective source rocks show clearly that the oils of the northern and central Triassic Basin are genetically connected essentially to the black shales of the Silurian. The oils from the south of the Triassic Basin as well as from the south (and west) of the Ghadames and Illizi Basins on the whole are derived from a Devonian source, The oils from the deposits Hassi R'Mel (HR), Makouda (MK), Mt Kheir (AT), Oued Noumer (ONR), Djorf (D]), Dras Temra (DRT), Guellala (GLA), Takhoukht (TKT), Berkaoui (OKP) and Ben Kahla (OK]) were formed mainly in the Silurian organic matter-rich black shales of the central part of the ~l¥iassic Basin. Their migration probably took place along the Hercynian unconformity, filling on the way reservoir rocks of the Triassic as well as of the Cambrian. The lower structural zone separating the deposits of Rhourde el-Baguel (RB) and Messdar (MDR) from those farther west suggests that otis of these two deposits probably originated in Silurian and Devonian shales of the western Ghadames Basin. Furthermore, isotopic and chemical similarities show that the Silurian shales of the Ghadames Basin could also have been the main source of the hydrocarbons of Hassi Keskessa (KA), E1-Bourma (ELB) and of reservoir F-3 of the Stah deposit. The Devonian shales of Ghadames and Illizi which are rich in organic matter have furnished a food part of the hydrocarbons of these deposits. The Devonian shales of the Ghadames Basin generated much petroleum which could migrate westward to fill the Triassic reservoir rocks at Nerzla (NZN), Gassi Touil (GT), Hassi Chorghi (HC), Rhourde Nouss (RN) and Arzel. The oils of the deposits at Tiguentourine (TG), La Recul~e (RCL), Gazel, Mereksen (MRK), Tamadanet (TAM) and Stah (STAH; with the exception of reservoir F-3) could have been derived mainly from the Devonian rocks of the Illizi and the southern Ghadames Basins. Silurian source rocks underlying this region could also have contributed part of the oils in these fields.
Chapter 6
Burial History and Kinetic Modeling for Hydrocarbon Generation
6.1
The Models Numerous computer programs, such as MATOIL, GENEX, TEMISPACK, PDI, and others, have been widely used to reconstruct the burial and thermal histories of source rocks in order to estimate hydrocarbon yields (phase and amounts), identify abnormal pressure zones, predict the reservoir rock properties, and identify oil migration pathways (Welte and Yukler 1981; Galushkin et al. 1985; Berthold and Galushkin 1986; Doligez et al. 1986; Nakayama and Lerche ~987; Tissot et al. I987; Espitali4 et al. 1988; Welte and Yalcin 1988; Ungerer 199o; Ungerer et al. 199o; Lopatin et al. 1992; Galushkin and Kutas 1995; Makhous et al. 1995). The underlying algorithms on which the software packages are based have been described in numerous papers. Because the diversity and complexity of processes involved in basin evolution are extremely specific, software packages created at different times by various researchers are quite different. Furthermore, the results of the numerical analysis are dependent on the basic principles of the model. The purpose of this chapter is to outline the basic principles of the integrated galo thermal modeling program. The main goal of this part of the study has been to present the basic design of the integrated galo software package. In our thermal analysis, the sedimentary blanket, the lithosphere, and the upper part of the asthenosphere are considered together. This approach allows us to calculate the amplitude of tectonic subsidence by considering changes in the density distribution vs. depth in the lithosphere. Local isostatic response of the lithosphere on load is assumed, and then the comparison of relative variations in the amplitude of tectonic subsidence is calculated by traditional methods (removal of the water and sediment load on the basement surface), with the variations obtained by nontraditional methods (consideration of changes in the density profile in the basement) providing an additional opportunity to control the program's sequence of the tectonic and thermal events in the lithosphere. Incorporating the burial stage of the source rock sample in the fitting procedure of kinetic parameters of reactions controlling the organic matter maturation is another key feature of our program package for basin modeling. Consequentl); we can determine the energy spectrum of kinetic reactions for samples with relatively moderate levels of organic matter maturation (vitrinite reflectance Ro = 0.5-0.8%).
6.1.1 The Program Our modeling program consists of three main blocks: input data for basin structure and evolution, initial parameters for basin modeling, and numerical simulation
zo8
Chapter 6 • Burial History and Kinetic Modeling for Hydrocarbon Generation
Input data for basin structure and evolution ] Present-day sedimentary section
Measured present-day valuesof porosity
Present-day thermal profileversus depth
I
I ]
Measured values ofvitrinite reflectance
4,
__A
->
I I I 1 I I I I
I Initial parameters for basin modelling [ Lithology,sedimentation,erosion,and hiatus, tectonicand thermaleventsin the basement, initial heatflow,boundaryconditions
INumericallsimulationl t
I I
Calculation of rock porosity
L__~I--I
Densityof sedi- ] mentaryand ' basem;t r°cks ' I l Tectonic i L --J subsidenceof | the basement
Restorationof burial historyof the basin
-J;
conductivity, capacity,and generation
,
t
, I t
~t~ Solution of heat transfer equation Temperatureof sedimentaryand basementrocks l
Temperature] profiles versusdepth
[Calculationof] | vitrinite | ~'~. reflectance j--
Time-temperature and burial histories of source rocks Open and dosed pyrolysisdata of sourcesample
V
Kineticparameters estimationfor --I reactions of kerogenmaturation
V --[
Estimateof HCyield history and expulsionthresho,d determinatination
Fig. 6.1. General scheme of our basin modeling program. Arrows with solid lines demonstrate relations between various program units for given variant of basin modeling; arrows with dashed lines show the relation that are involved in correction and calibration of this variant (Fig. 6.1). The first data block contains geological, geophysical, and geochemical data describing basin structure and evolution, including information about present-day sedimentary section, measured values of porosity, temperatures, and vitrinite reflectance. The second block deals with preparing initial parameters for numerical simulation of thermal history of the basin: calculating the volumes of uncompacted sediments on the basin surface, estimating time and amplitude of tectonic and thermal
6.1 - T h e M o d e l s
zo9
events in the basement (thermal activation, stretching of the basement, etc.), calculating the initial temperature profile, and determining temperatures at the base of the computed domain. The third data block uses prepared parameters to carry out a onedimensional numerical simulation of burial, thermal, and geochemical evolution of the basin. The comparison of rock porosities, temperatures, and vitrinite reflectance computed in this block with corresponding present-day values from the first block, as well as the calculated curves of tectonic subsidence, are used to correct the initial parameters for our basin modeling (dashed arrows in Fig. 6.x show the corresponding relationships between these blocks). The third block includes the chemical-kinetic modeling package. Data of open- and closed-pyrolysis experiments are used here for restoring the kinetic spectrum of maturation reactions in source rocks. This spectrum is applied to achieve a numerical estimate of hydrocarbon yield and the expulsion threshold. 6.1.2 Burial and Thermal History Modeling 6.1.2.1
Input Parameters The input parameters for the model include the present-day sedimentary cross section, estimates of the amplitude and rate of erosion, the lithological composition and petrophysical characteristics of rocks, the structure of the lithosphere (basement) and its rock parameters, paleotemperature markers (vitrinite reflectance), paleoclimate, sea paleodepths, present-day surface heat flow, depth-temperature profiles, and information on the paleotectonics and the present-day tectonic setting of the basin. The evolution of the Oued el-Mya Basin is used to demonstrate the model. Table 6.1 presents the basin's main stages of evolution, which include sedimentation, hiatus, and erosion. The input data on basin evolution (see Table 6.1) assume that about z.z km of Silurian-Devonian sediments were eroded during the Permian Hercynian orogeny, nearly the upper limit of erosion amplitude. The presence of thick layers of these sediments in neighboring sedimentary sections supports this assumption. Some details of the problem of erosion amplitude assessment are discussed in the following paragraphs. 6.1.2.2
Burial History When sediments are progressively buried, they are compacted, and pore fluid is expelled. We consider compaction in this program with the following assumptions (Perrier and Quiblier 1974): (1) the volume of solid matrix is preserved throughout compaction and (2) porosity depends only on burial depth and can be expressed as
[1- P((z2)ldg 1 = [1- P(z2) ]dg 2 ,
(6.1)
where P(z) is the porosity at the depth Z, and d Z 1 and dZ~ are the thicknesses of the layer during burial at depths Z~ and Z . respectively. The backstripping procedure for
21o
Chapter 6 •
Burial History and Kinetic Modeling for Hydrocarbon Generation
every discrete sediment layer dZ is based on Eq. 6.1 and the exponential porosity-depth relationship (Sclater and Christie 198o; Deming and Chapman 1989; among others):
P(z)
= Po- z l B
,
(6.2)
where Po is the mean value of porosity on the upper loo-15o m of the sedimentary section, and B is the depth-scale factor. Equations 6.1 and 6.2 are used to reconstruct sedi-
T a b l e 6 . 1 . Main stages of Oued el-Mya Basin evolution ~
No. Stageof evolution
Geologic time Depth (Ma) (m)
1
Sed.
0
-
65
Hiat.
65
-
91
3
Sed.
91
-
93
125-
322
Im, dl, ml
12-18
0-
30
4
Sed.
93
-
97.5
322-
870
hi, an
12-13
30-
80
5
Sed. Sed.
97.5 - 113 113
- 119
125
Surface temp. Sea level (°C) (m)
2
6
0-
Rocktype
125
sn, lm
15
-
t5-
0 18
0
870 - 1042
cl, an
13 - 15
1 042 - 1 489
ct, sn
15
170
80 - 170
7
Sed.
119
-144
1489-2033
cl, sl, dl, ml
15-18
1 7 0 - 130
8
Sed.
144
-213
2033-2886
cl, dl, hl, an, mt
18
130-
9
Sed.
213
- 231
2886-3485
cl, hl, an
18
0 0
l0
Sed.
231
-243
3485-3540
cl, sn, hl
18
11
Sed.
243
-248
3540-3711
vl
18
0
t2
Eros.
248
- 286
2200-2200
-
15 - 18
0 0
0
13
Sed.
286
-360
3711 - 3 7 1 1
cl, sn
8 - 15
14
Sed.
360
-408
3711 - 3 7 1 1
cl, sn
7-
8
0-240
15
Sed.
408
- 428
3711 - 3 8 5 4
cl, sn
5 -
7
240 - 350
16
Sed.
428
-438
3854-3924
cl, sn
5
350
17
Sed.
438
- 590
3 924 - 4100
cl, sn
5 - 15
350 -
0
a Depth column shows present-day depths of the bottom (first number) and roof (second number) of the sedimentary layers; that is, the erosion amplitude. Sed. = sedimentation, Hiat. = hiatus, Eros.= erosion, an = anhydrite, cl = clay and shale, dl = dolomite, hi = halite, lm= limestone, ml = marl, sl = siltstone, sn = sandstone, vl = volcanics, No. = number o f the basin's evolution stage.
Fig. 6.2. Burial and thermal histories of the sedimentary section, Takhoukht region, Oued el-Mya Basin. a Pateodimate history based on literary paleogeographic data of the regeon, b Burial, thermal, and maturation histories resulting from basin modeling. Note that the considerable rise of isotherms in the post-erosion period is related to the Permian-Triassic thermal activation in the lithosphere. More moderate activation occurred here in the Cretaceous-Cenozoic. The temperatures of Silurian rocks did not exceed 85 °C during the pre-erosion period despite significant amplitude of erosion, c Tectonic subsidence of the basement surface calculated in local isostasy approach by removing of sediment and water load (solid line) and by consideration of variations in densities of basement rocks (dashed line). The coincidence of solid and dashed curves provides an additional opportunity to control the sequence of the tectonic and thermal eventes in the lithosphere. STRi and STR2 = streching periods of the basement; TACt and TAC2 = periods of thermal activation in the basin lithosphere, d Variations in sedimentation (>o) and erosion (
6.1 - The Models
211
Time (Ma)
a
400
600 1
10
200
1
I
i
i
I,
1
I
I
f
~f_.---J
Paleoclimate
temperatures
0
m
-.
i
i
/
i
K1 Bm-Ber
•
1
\
/ / /
SO°
~-,.
I
I
r
\
\\
i !
I
r I
L
J
"\i
I 1
1
Sedimentary layers boundaries Basement surface
~ ~ ,~
I
l I
[
Isolinesofvitrinite reflectance
/
Isotherms i
\
. ~
~
_
~
~t
/
I
\
I
i
t II
\ i
600
120o \
400
i
i ~,
J
200
0
=
]
I
[
......
. B a ~;p p. e .d . . tectonic
"%. %-. ~
i]
.............
~Sr=i%too,c
1; fi"
\
~ - .~ . ~ .
'
subsidence
[
Tectonic subsidence
-,,,~
",~,.~
STRI
STR2
I....
I
,_T_~_% , ,
CTT_..~A_~. ~........
I Io Islolc IPIT 'EK d 150 Sedimentation and erosion rates
IO0
~ so E ¥ o
.....................J LL
e¢,
-50 -100
,
600
I
,
I
400
,
,
I
200
,'
,
,
0
Chapter 6 . Burial History and Kinetic M o d e l i n g for Hydrocarbon Generation
z12
mentation rates during basin evolution (Fig. 6.2b,d). The values for Po and B given in Table 6.2 were calculated on the basis of world-average data for the m a i n lithological units (Sclater and Christie 198o; Gretener 1981; Beaumont et al. 1982; Goff 1983; Hutchinson 1985; Stockmal et al. 1986; D e m i n g and C h a p m a n 1989; Burrus and A n d e b e r t 199o; Nielsen and Balling 199o; Forbes et al. 1991) and for a combination of lithologies (see later additional features). C o m p a c t i o n is assumed to be irreversible during periods of erosion. We did not consider variations in porosity due to dissolution, cement deposition, or recrystaltization.
6.1.2.3 Heat Transfer Equation a n d Thermophysical Parameters The temperature distribution in the lithosphere (Figs. 6.2, 6.3) is obtained by solution o f the heat t r a n s f e r e q u a t i o n d e r i v e d in the f r a m e o f the e n e r g y balance s c h e m e (Carlslaw and Jaeger 1959; Paskonov et al. 1988):
Table 6.2. Petrophysical parameters of sedimentary rocks, Takhoukht region, Oued el-Mya Basin a No.
1
Po
0.429
B
Km
AI
Cv
On
A
(km)
(W m -I °C)
(°C-I )
(MJ m -3 K-I )
ig"cm -3)
(mkW m -~)
2.77
4.00
0.0027
2,872
2
2.66
0.816
-
-
2.696
2,73
0.578
-
3
0.572
1,91
3.49
0.0011
4
0.244
5
0.577
0.86
5.61
0.0050
1.943
2,30
0.050
1.39
3.71
0.0030
2.332
2.52
6
0.888
0.600
2,06
2.96
0.0017
2.575
2,71
1.465
7
0.635
1,88
2.82
0.0011
2.487
2,70
1.394
8
0.296
1.20
5.17
0.0043
1.993
2.32
0.209
9
0.354
1.24
4.72
0.0040
1.955
2,30
0.431
10
0,620
1.94
2.81
0.0015
2.462
2,66
1.549
11
0.500
3.27
2.01
0.0001
2.500
2.70
1.005
-
-
12 13 14
-
0.610
2.03
2.88
0.0016
2.549
2.68
1.516
15 16
0.684
0.684
1.84 1,84
2.24 2.24
0.0007 0.0007
2,324 2.324
2.69 2.69
1.968 6.699
17
0,684
1.84
2.24
0.0007
2.324
2.69
1.968
a No. = number of the basin's evolution stage (corresponding to No. in Table 6.1); Po= average rock porosity within the near-surface layer at the depth of 0-200 m; B = scale for porosity change vs. depth in the law P(Z) = Po exp (-z/B); Km =heat conductivity of the matrix rocks at the temperature T= 8 °C; At=temperature coefficient of matrix heat conductivity K(t)=Kin 1(1.0 + A I x T(°C)); Cv= vol-ume heat conductivity of matrix rocks; p m = density of matrix rocks; A = heat generation per unit volume. The values in this table were computed according to relative content of facies and data in Table 6.1.
6.1 - TheModels
213
c3(Cr7")1 & + c3(CvwVT) / OZ = c?[K(c3TI c~Z) I c?Z+ A(Z,t)]
(6.3)
,
where Cv = pCp is the volumetric heat capacity, Cp is the heat capacity per mass unit, p is the density, T is the temperature, t is the time, V is the velocity of water flow expelled from the compacted sediments with respect to the "steady" basement, Z is the depth, K is the heat conductivity and A is the heat generation per volume unit. Equation 6.3 has been solved in a coordinate system tied to the basement. At every incremental time step, dr, the computed column, including sediments and basement, is increased from above by the increment dZ, which is the sediment layer deposited during time interval dt. The thicknesses of each elementary sedimentary layer, dZk, under upper layer dZ decrease due to compaction according to the porosity-depth relationship shown in Eq. 6.1. The calculated column height is decreased from above by dZ if erosion took place during time dt. The grid points within the computation column then move to new positions, which are determined by the deposition of layer dZ, as well as by subsequent compaction of the column. Temperatures obtained in the previous time step are assigned to these moved grid points. The first grid point (n = 1) has the temperature determined by paleoctimate conditions at the basin surface during time interval dt (Table 6.1). The temperature distribution obtained in this way is the one used to solve Eq. 6.3 in the next time step. The convective component in Eq. 6.3 describes groundwater movement in sediments and is not considered in our analysis. Arguments for such a solution are discussed later in Section 6.1.4. The t h e r m o p h y s i c a l p a r a m e t e r s of s e d i m e n t a r y rocks in Eq. 6.3 are functions of porosity and temperature. Density (Ps), v o l u m e heat capacity (Cvs), and thermal conductivity (Ks) are specified by the values for the mineral matrix (Pro, Cvm, Kin) and
Fig,6.3. Numerical modeling of the thermal regime in the lithosphere of the Oued el-Mya Basin, Takhoukht region. Solid line ist the base of the lithosphere as determined by intersection of the current geotherm with solidus curve for peridotite (shown in Fig. 6.4). Long dashed lines are isotherms Moho = base of the crust; comp. transition = location of pyroxene peridotitegarnet peridotite compositional transition in the mantle. The considerable rise of isotherms is related to periods of thermal activation in the lithosphere in the Permian-Triassic and Cretaceous. Reduction of the crust thickness occurred during basement stretching (in the Ordovician-Devonian and Cretaceous) and during Permian erosion
Time (Ma) 600
400
250°
/
500°
.~ ~ 40
80
-.
~
"~.
_.--
I I\ \
60
\
/ \
20
~-
0
200
750o
.____Co rap.transition.......... ~ ~
\
)
.....
- -
214
Chapter 6 • Burial History and Kinetic Modeling for Hydrocarbon Generation
for water (Pro, Cvw, Kw) (Beck 1976; Sclater and Christie 198o; Oxburgh and ,4ndrewsSpeed 1981; Doligez et al. 1986; Deming and Chapman 1989; Ungerer et al. 199o), where ps(Z) = p~[1 - a Z ) ] + p J ' ( Z )
,
Cvs(Z) = Cvm[1 - P(Z) + CvwP(Z)]
, (6.4)
Ks(Z) = Km[1-P(Z)]Kw vtz)
For water, Cpw = 4186.8 J / kg °C, Pw = 1030 kg m -3, and Kw, in mW/ m °C, is defined by Kw = 0.565 + o.oo188T- o.ooooo723T~ (o < T< 137 °C) Kw = 0.602 + o.oo1313T- o.ooooo514T2 (137 < T< 300 °C) The thermal conductivity of the mineral matrix, which is inversely proportional to temperature (Doligez et al. 1986), is expressed by Km=Km(T=o°C)/(I+
AI×T)
.
(6.5)
The world-average values of Kin, Cvm, Pro, heat generation, .4, and A1 (temperature coefficient of matrix thermal conductivity), as well as values of the porosity parameters Po and B in Eq. 6.2 for the main lithological units, are taken from Sclater and Christie (198o), Gretener (1981), Beaumont et al. (1982), Goff (1983), Hutchinson (1985), Stockmal et al. (1986), Deming and Chapman (1989), Burrus and ,4ndebert (199o), Nielsen and Bailing (199o), and Forbes et al. (1991). These values, together with the algorithm described in section 6.1.4, are used to calculate the thermophysical parameters reported in Table 6.2 for the mixtures of lithological units given in Table 6.1. From Eqs. 6.4 and 6.5, the thermal conductivity of clays or sandstones can increase by a factor of 2-2. 5 during their burial history. In our examples of basin modeling, we adopted the characteristics of "typical" continental lithosphere, consisting of 19 km of granitic rocks, 20 km of basaltic rocks, and more than lOO km of mantle rocks, with a contribution to surface heat flow of about al m W m -2 resulting from radioactive decay in the crust and mantle (Smirnov 198o; Baer 1981). In our thermal analysis, we consider the lithosphere and part of the upper asthenosphere as containing partly melted peridotite. The depth to the top of the asthenosphere changes during basin evolution. The thermal contribution of the latent heat released due to melting of rocks is considered in the usual enthalpy approach. 6.1.2.4
Boundary and Initial Conditions The upper boundary condition for solution of Eq. 6.3 corresponds to the temperature at the basin surface (Z = o) at time t. This value, specified as the mean-annual temperature, is derived from paleoclimate data (Fig. 6.2a). The depth to the lower boundary of the domain for solution of Eq 6.3 (Zlow)is specified at the initial modeling stage by the intersection of the geotherm [Train(Z)] with
zl5
6.1 • The Models
"'~ Temperature (°C)
Fig. 6.4. Calculated initial and present-day temperature profiles in the lithosphere of the Oued el-ivlyaBasin, Takhoukht region. The solidus temperature of peridotite rocks with small content of H20 (Wyllie a979) were used to determine the base of the lithosphere in the basin modeling
0
~
400
,
i
20
I
800
r
,
M
I
,
,
...
1200
-@ e-
40
K
",\ 60
. . . .
I
I ----~ Initialtemperatures Solidusof peridotitewith
~22
/
J
\\&/
~l' ~
80
the solidus curve for peridotite rocks [Ts(Z)] reported by Wyltie (1979); (Fig. 6.4). The function Train(Z)is computed as the assumed geotherm of the basement for the coldest state of the lithosphere in the history of the basin. If intersection of Train(Z) and Ts(Z) does not occur, Zion, is assumed to be zoo km. The temperature at the lower boundary is specified as Tlow= Tmin(Zlow).For many basins, the coldest geotherm Train(Z) corresponds to the present state of the basin lithosphere; however, this is not always the case. In the Oued el-Mya Basin, the lowest heat flow occurred about 3oo Ma (Fig. 6.5). In the model, the depth to the lower boundary Zlo,,,increases from the initial value by the thickness of the sedimentary cover. The steady-state temperature, T = 7low, is maintained at this boundary. Following the procedures given for the choice of the parameters Zlowand Tlow,a temperature of 1 o5o °C was assigned to the louver boundary having an initial depth of 9o km in the Oued el-Mya Basin. For physical reasons, th e steadystate temperature at the louver boundary of the modeled region is preferred over a fixed steady heat flow (see Section 6.1.4). The initial temperature profile for Eq.6.3 is determined by solving Eq. 6.6, steadystate variant of the equation for the value of initial surface heat flow, Qo:
r ( z , t = 0) = To +
Oo
i a(z")×az"]×dE,'
,
(6.6)
I
where Z is the depth, To is the surface temperature at t = o, the time of the initiation of the basin subsidence (see Table 6n), A(Z) is the heat generation per volume unit in the basement rocks and K(Z) is the thermal conductivity of the basement rocks, we believe that in the depth interval from the base of the lithosphere (Zlit) to the low
z16
Chapter 6 • Burial History and Kinetic Modeling for Hydrocarbon Generation
Fig. 6.5. Computed variations in heat flow during the Oued el-Mya Basin modeling, Takhoukht region. The difference between the heat flows through the surfaces of the basement and sediments is due mainly to radioactive heat generation in sediments
I00
Heatflowon sedimentsurface Heatflowon basementsur~ce
80
E
0
I
60
600
400
200
0
Time (Ma)
boundary of computation (Zlow),where the mantle rocks are rheologically weak, temperatures increase linearly from the temperature at the base of lithosphere (Tlit) to the temperature of the lower boundary (Tlow) (Fig. 6.4). Initial heat flow, Oo, is estimated from the present-day values of surface heat flow in the regions with tectonic settings similar to that expected for the time of the basin initiation. For example, Qo ~ lOO-11o mW m -2 is typical of continental rift areas (Baikal Lake, African rift system); however, on rift shoulders, Oo can be nearly 65-85 mW m -2 (Smirnov ~98o). Estimations of the initial heat flow can be adjusted by analysis of variations in tectonic subsidence of the basement (see following sections). In the case of the Oued el-Mya Basin, Qo is approximately 52 m W m -2 (Fig. 6.5). Equation 6.3, together with its boundary and initial conditions, is solved by the implicit finite-difference scheme discussed later. 6.1.2.5
Thermal History and TectonicSubsidence The burial and thermal histories of a basin, based on the data in Tables 6.1 and 6.z, are presented in Fig. 6.z. The complex sequence of thermal and stretching events in the lithosphere for this example is shown in Figs. 6.2b,c and 6.3. Two methods are used to calculate the relative change in tectonic subsidence that controls this sequence of lithospheric events (Fig. 6.zc). These calculations allow us to link the variations in sedimentation processes with stretching and heating in the basin lithosphere (McKenzie 1981; Galushkin and Kutas 1995; Makhous et al. a995). In an isostatic framework, when the contribution from the collision of plate margins to tectonic subsidence is negligible, the upper solid line in Fig. 6.zc must coincide with the upper dashed line. The solid line represents the relative change in tectonic subsidence of the base-
6.1 . T h e M o d e l s
217
ment surface caused by variations in the load on the basement surface, such as the load of water or sediments. The dashed line in Fig. 6.zc is the tectonic subsidence caused by change in the depth distribution of the rock density within the basement column. This change can be due to thermal expansion or contraction during heating or cooling of the lithosphere, to stretching of the basement, or to displacement of the phase boundaries (Figs. 6.2c, 6.3; see also section 6.1.4). The sequence of tectonic and thermal events in the lithosphere, which is in accordance with the tectonic subsidence curves in Fig. 6.2c, is illustrated by the example of the Oued el-Mya Basin shown in Figs. 6.z and 6.3. Slight variations in amplitude of tectonic subsidence in the Cambrian-Ordovician (from 60o to 48o Ma; Fig. 6.zc) indicate only moderate variations in heat flow during the first period of lithosphere cooling from an initial heat flow of approximately 52 m W m -z (Fig. 6.5). In our example, basement subsidence in the Silurian and Devonian (44o-36o Ma),with deposition of about z 50o m of clays and sands, has been accompanied by basement stretching, with total amplitude of about 1.2 during 95 Ma (Fig. 6.2). Stretching rates were slow (V _
z18
Chapter 6 - Burial History and Kinetic Modeling for Hydrocarbon Generation
0
Vitrinite reflectanceRo(%) 0.5 0.7
0.3 i
l
t
I
Present-dayvitrinite reflectanceprofile
o
Temperature(°C) 40 80 t
I
~
NN~
I
i
r
i
l
I
120 1
I
I
Present-dayrofile
2
3-
4-
Calculated values Measured values
*
~N~
Calculated
values Measured values
\
] I
~ \
~,
Fig. 6.6. Vitrinite reflectance and temperature in the present sedimentarycolumn of the Takhoukht
region
ing stages, the crustal thickness was reduced from 35 to 29 km, with the value of 29 km including the thickness of sedimentary cover (Fig. 6.3). 6.1.2.6
Estimation of Organic Matter Maturation Temperatures within the sedimentary section calculated as a function of time are used to estimate the maturation level of the organic matter. The kinetic model of vitrinite maturation (Sweeney and Burnham 199o) is the primary method of maturity estimation. Comparison of measured vitriuite reflectance (Ro) with the calculated values and present-day temperature profile (Fig. 6.6) is used to control modeling parameters. The time-temperature index (TTI) (Lopatin z971; Waples 198o) is also computed as a simple method to describe the pateothermal regime. 6.1.2.7
Sensitivity to Input Variations 6.1.2.7.1
Erosion Amplitude The assumed erosion amplitude of z.z km cannot be considered precise if the measurement was based only on the available geologic data; however, this value can be
6.1 • The Models
219
partly justified by analysis of relative variations in the amplitude of tectonic subsidence. For this purpose, we made calculations for a model with a colder lithosphere in which the asthenosphere roof (with r = 1 ooo °C) in the pre-erosion period is located at depths below loo km. In this case, even a relatively weak thermal event (surface heat flow in the order of 63-67 m W m -a) should result in substantial uplift of the basement surface, which leads, in turn, to an erosion amplitude of over 3 ooo m. However, low heat-flow values are not consistent with the presence of volcanic flows in the section, which corresponds to a situation close to continental rifting. In another version, calculations show that when the 1 ooo °C isotherm is located at depths of from 55 to 70 km in the pre-erosion period, an erosion amplitude of 2-2. 5 krn requires an unreasonably high heat flow (250 mWm-2). In our model, a temperature of about 1050 °C was maintained at a depth of about 90 km (Figs. 6.3, 6.4). In this case, thermal activation of the lithosphere in the Permian (surf~e heat flow of approximately 90 m W m -2) results in erosion of 2.2 km. This erosion amplitude is geologically reasonable (Figs. 6.2, 6.5). The thermal and burial histories of the basin for a variant that is free from erosion in the Permian were simulated for comparison with the main model (Fig. 6.7). This variant was controlled by present-day temperatures and vitrinite reflectance (similar to Fig. 6.6) and, by coincidence of the tectonic curve, calculated by removing the sediment and water load, resulting from variation in the thermal state of the basement (as
600 0
Time (Ma)
400 ~
I
I
p •
I
I
"
....
-
"-
200
I
~
I
~
~
. I
0 I
,
I,
~
~
'
z
](2
]
Cen
' - K7 ;t-;,b-
K1 Bin-Bet
,..¢: OJ ¢3
---
4
i
~ [
Sedimentary layers boundaries Basement surface
~ ~,v
Isolines of vitrinite
.~
Isotherms i
- ~ I *,
reflectance
5 6OO
/ "~"
~ n'
ao.
I
4OO
~ d
t
720 ° %
\ ,.
/ I~
200
I
ii i
0
Fig. 6.7. Burial, thermal, and maturation histories of the sedimentary section in the Tahoukht region of the Oued el-MyaBasin resulting from basin modeling in the variant without erosion. The first thermal activation in the variant without erosion is tess intense than the thermal activation in the main variant in Fig. 6.z
220
Chapter 6 - Burial History and Kinetic M o d e l i n g for Hydrocarbon Generation
in Fig. 6.zc). The variant cited in Fig. 6.7 repeats the same sequences of stretching and thermal activation events as are shown in Figs. 6.2 and 6.3. The exception is the Permian event for which the surface heat flow in the variant free from erosion is 60 mWm-L Comparison of the two variants (Figs. 6.2, 6.7) shows that the influence of erosion on the thermal state of the sediments is limited. In both variants, temperatures at the base of the Silurian rocks did not exceed 85 °C during the Devonian and Carboniferous. Consequently, only minor changes occur in the maturity level of the organic matter in the source shales during the pre-erosion period. The limited effect of the Permian erosion on the thermal history modeling and the maturity of source shales in the northern Oued el-lVlya Basin is a direct consequence of low sedimentation rates during the Cambrian-Silurian. Total sediment thickness did not exceed 6oo m (Figs. 6.2b,d, 6.8); however, in the southern and southwestern parts of the Oued el-Mya Basin, where ] 3oo-2 ooo m of Cambrian-Devonian sediments were deposited by the beginning of the Carboniferous, the situation is the opposite. A considerable amount of hydrocarbons was generated prior to the Hercynian uplift, and subsequent erosion influenced total hydrocarbon output.
a
i
0
Time (Ma)
400
600
,
I,,
I
200 I
0
I
[
~
I
80.
I-...
~.4o. E / b
t
.
.
.
.
erosion Variant without erosion
0.8
,J
o.6 ¢-
/
/
/
/
/
~ 0.4,
" 0.2, >
....
Variant with erosion Variant without erosion
0.0 6O0
400
200
0
Fig. 6.8. a Temperature and b maturation histories of the Silurian source shales in the Takhoukht region computed for the main variant with erosion in Fig. 6.z (solid lines) and for the variant without erosion in Fig. 6.7 (dashed lines). The temperature of Silurian rocks did not exceed 85 °C during the preerosion period and, therefore, the present-day maturation level does not differ in either variant
6.1
-
The Models
221
6.1.2.7.2 Heat from Radioactive Decay The heat from radioactive decay in sediments influences their thermal state. The content of U and Th in the lower 70 m thick layer of Silurian shales of the Oued el-Mya Basin is 25 and lO ppm, respectively, and heat generation per unit volume is about 6.7 mkW m -3 (Table 6.2). Some researchers believe that the heat released by the radioactive decay in petroleum source rocks may accelerate the maturation of organic matter (Byakov et al. 1987). Our calculations for two variants of basin modeling confirm that the contribution of radioactive heat in the lower layer of the Silurian shales is local. We calculated data for a case where A = o.84 mkW m -3,which is typical for argillaceous rocks, rather than A = 6.7 mkW m-3,which was assumed in the previous example. The decrease in heat generation in a variant without anomalous radioactive decay caused the isotherms to move no more than 3o m, whereas temperatures within the Silurian layer were lowered by about 0.4 °C, and rrIvalues were lowered by approximately 1%. Thus, anomalous radioactive heat emission in comparatively thin layers of petroleum source rocks (H = lOO-15o m) cannot change the level of organic maturation in these rocks. The long duration of contact with the enclosing rocks, the thermal conductivity, and the heat capacity of rocks are responsible for spreading the thermal anomaly over a larger depth range and a subsequent decrease in its amplitude.
6.1.3 Modeling Maturation History and Hydrocarbon Generation 6.1.3.1 Reaction Kinetics
The modeling of kerogen maturation and hydrocarbon generation in source rocks is achieved using the algorithms described in Tissot and Espitali~ (1975), Tissot et al. (1987), Espitali~ et al. (1988), Welte and Yalcin (i988), Issler and Snowdon (199o), Ungerer (199o), and Ungerer eta]. (199o). We assumed that the maturation of organic matter in source rocks can be described by N Arrhenius first-order independent reactions:
Ki( t) = A i exp(-Ei/RT(t)). ,
(6.7)
where K i is the rate of ith reaction, A i is its Arrhenius or frequency factor (1 s-l), Ei is its activation energy (Kcal mol-l), R is the perfect gas constant, T is temperature (in Kelvin) and t is time. If Ai, El, and Xio (initial hydrocarbon potential for the ith reaction) are known for all N reactions, the amount of hydrocarbons formed and the rates of generation can be evaluated by these expressions (Tissot and Espitali~ 1975): N
dQ
N
S2(t) = -~- = 2 Xi0 x Ki(t)x exp[- I t Ki(t')x dr'] I to "
(6.9)
Chapter 6 • Burial History and Kinetic Modeling for Hydrocarbon Generation
zzz
6-
a
~ 1
2
Calculated values Measured * * * values
400
Total initial potential
O" 7 c~
H I 0 : 630 mg HC g-1TOC
fi 300
co
2oo ~
F ~2-
loo ~e~
0300
J
400 500 Heating temperature (°C)
40
J
I
'
I
0
'
50 60 70 Apparent energy (Kcal mo1-1)
80
Fig. 6.9. Rates of hydrocarbon yield during experimental pyrolysis and kinetic energy spectrum restored from these rates for the Silurian source shales of the Takhoukht region, a Rates of hydrocarbon generation during pyrolysis programming heating with the rates of (1) 5 °C min-*, (2) 15°C rain-~, and (3) 3o °C min-1. b Distribution of apparent activation energies vs. initial petroleum potential for the chemicalkinetic reactions. The time-temperature history of the Silurian source rock (see Fig. 6.8a) was considered when developing the energy spectrum (see text)
These equations are applied to determine hydrocarbon yield as a function of time, and the rates of hydrocarbon output ($2 curve) during both natural and experimental cracking of kerogen. The results of open-system pyrolysis (Rock-Eval II) have been used to specify the kinetic parameters controlling maturation. Hydrocarbon yield rates as determined by these experiments are shown in Fig. 6.9a. Both nonlinear optimization technique (Levenberg-Marquardt method: Press et al. 1986; Issler and Snowdon 199o) and linear methods are used to determine the values of the reaction parameters Ai, Ei, and Xio. This technique minimizes an error function by comparing the hydrocaibon release rates, Sv calculated by Eq. 6.9 and those rates measured in open-system pyrolysis. An example of the spectrum of activation energies obtained from this analysis is shown in Fig. 6.9b. Our algorithm for fitting the kinetic parameters El, Ai, and Xi uses the total timetemperature history of the source rock sample inferred from the level of organic maturation (solid line in Fig. 6.8b) and from open-system pyrolysis in which temperatures change linearly at rates of 5,15, and 3o °C min -1, ranging from 3o0 to 600 °C. For this purpose, the time integrals in Eqs. 6.8 and 6.9 are divided into two parts:
f tot
x dt'+S;' Ki(t' ) x dt' , d,'=,t0ft, K.[tq " '
(6.1o)
where to < t' < q is the time interval of heating of the source rock sample during its burial history, and q < t' < t is the time interval of the Rock-Eval programmed heating. This approach helps us to overcome, in part, the problems (discussed in Tissot et al. 1987; Espitalid et al. 1988; Issler and Snowdon 199o; and others) that are related
6.1 - T h e M o d e l s
223
to the necessity of using immature organic matter to determine chemical-kinetic parameters from pyrolysis experiments. The energy spectrum shown in Fig. 6.9b corresponds to a rather high initial hydrocarbon potential, HIo = 630 mg HC / g TOG which is typical for the marine type II kerogen (Espitali6 et al. 1988; Ungerer et al. 199o). 6.1.3.2
Effects of Uncertainties in Kinetic Parameters Determining the activation energy spectrum fi:om Rock-Eval pyrolysis data is an inverse task of mathematical statistics having multiple solutions. Some problems are related to the restoration of chemical-kinetic parameters of effective reactions for organic matter maturation in source rocks. For example, reactions with activation energies of less than 50 Kcal mo1-1 do not contribute to the Rock-Eval pyrolysis $2 curve because these reactions can occur during the burial stage and would not contribute to $2. Another problem relates to the use of variable values o f A i, There is no physical reason why all reactions in energy spectrum must have the same frequency factor as is usually assumed in chemical-kinetics modeling (Tissot et al. 1987; Espitali6 et al. 1988; Welte and Yalcin 1988; Ungerer et al. 199o). Spectra with variable values of Ai have a more irregular shape than spectra with a single-frequency factor, but the first spectra allow us to achieve a closer coincidence of the observed and calculated curves for S~ in the fitting of kinetics parameters. An additional problem is created with the inclusion of the geological stage of maturation in the fitting procedure of the kinetic parameters and multivalued frequency factors. This is a consequence of the need to establish an upper limit of the initial potential for hydrocarbon generation (HIo). For example, the spectrum in Fig. 6.9b is based upon the assumption that HIo is 630 mg HC/g TOC (typical for marine type II kerogen) (Espitali4 et al. 1988; Ungerer et al. 199o). It is possible to solve this problem if a maturation scale is available for kinetic analysis, provided that multiple analyses can be carried out on source rocks containing the same type of kerogen, but having various levels of its maturation. 6.1.3.3
Hydrocarbon Generation and Primary Migration The kinetic parameters Ei, , Ai, and Xio are used with the modeled thermal history of the source rocks to compute hydrocarbon output, using Eqs. 6.8 and 6.9, and to estimate the quantity of hydrocarbons expelled from the source rocks. The relative contributions of oil, gas, and coke to the total output of hydrocarbons are computed by using the standard three-fraction (oil, gas, and coke) and five-fraction (C1,C2-C3,C6-C15,CI5+,and coke) kinetic models for kerogen types I, II, and III (Tissot et al. 1987; Espitali4 et al. 1988). The onset of primary migration or the oil expulsion threshold is computed taking into account changes in porosity and rock density with depth and time. The critical determination is the threshold porosity Pth. Expulsion of liquid hydrocarbons from the source is assumed to commence when the pores become 20% saturated; that is, when Pth = 2o%. The threshold values for pore saturation are poorly defined and can range from lO to 6o%, depending on the organic matter type (Tissot et al. 1987;Espitali6 et al. 1988; Ungerer 199o).
zz4
Chapter 6 . Burial History and Kinetic Modeling for Hydrocarbon Generation
6.1.4 Additional Features of Our Burial and Thermal Modeling 6.1.4.1 Thermal Aspects of Groundwater Flow Within subsiding continental basins, groundwater can be induced by sediment compaction (pressure water) and by driving forces from topographic relief or from variations in fluid density. The intensive water flow can disturb the thermal profile and change the time-temperature history of buried source rocks. Maximal rate of pressurewater expulsion (Rpwe) can be estimated relatively easily using Eqs. 6.1 and 6.z for the case of one-dimensional consolidation of homogeneous sediments on immobile basement: , l-P(zb) Rpwe = S × L n ~ , where S is the sedimentation rate at the basin surface, Ln is the natural logarithm, P is the porosity of sedimentary rocks calculated as a function of depth (Z) according to Eqs. 6.1 and 6.z, Z is the current depth in sediments (Z = o is the sediment surface) and Zb is the depth of basement surface. The estimation for Rpwe assumes that there are no bars in the region between Z and Zb for vertical movement of water expelled from pores due to compaction. In the Oued el-Mya Basin (Takhoukht area), which was used to demonstrate the model, maximal sedimentation rates (5) are observed in the post-erosion period and did not exceed 149 m Ma -1 (Fig. 6.zd); therefore, Rpwe is 3 mm / year. For reasonable values of clay-sand thermal conductivity, such a value of Rpwe corresponds to a Peclet number Pe < o.o5, where the Peclet number is a nondimensional ratio between the heat-flow density components due to convection and to conduction (Bredehoeft and Papadopulos 1965). The calculations for a one-dimensional model of compaction of homogeneous sediments (Bredehoeft and Papadopulos 1965; Clauser and Villinger 199o) and a two-dimensional numerical analysis of the problem (Bethke 1985,1989; Deming et al. 199o; Person and Garven 1992) show that the vertical flow of groundwater having a Pe of o.1 has a negligible effect on the thermal profile in sediments. The cited authors concluded that expulsion of pore fluids arising as a consequence of typical sedimentation rates would not perceptibly perturb the thermal regime of sediments. In contrast, freshwater infiltration into basin strata can cause a significant distribution of the thermal regime of the basin; however, analysis of these fluid movements contains many uncertainties in formulating and applying hydrologic and paleohydrologic techniques to basin studies (Bethke 1989). The permeabilities of basin strata on regional scales are affected by heterogeneities such as the distribution of facies, interbeds, faults, and fractures. In paleohydrologic modeling, records of critical variables, such as past topographic relief, may have been destroyed by erosion. In addition, only empirical methods are available to assess changes in the hydrologic properties of sediments under conditions of increasing and decreasing effective stress during basin evolution (Bethke ~989). In the Oued el-Mya Basin, in particular in the Takhoukht region, the presence of several thick evaporite strata prevents the develop-
6.1 • T h e
225
Models
ment of wide-scale convection of groundwater in the sedimentary cover. Water movement along the quasihorizontal strata on a several hundred kilometers scale has a weak influence on vertical temperature profiles provided that the studied section is far enough from surface water infiltration sites. A similar situation is observed in Oued el-Mya and in the great part of the Ghadames Basins; however, the situation is different in the Illizi Basin, where the whole stratigraphic section (Paleozoic, Mesozoic, and Cenozoic sediments) crops out on the surface. Certainly, the fluid flow largely influences the thermal regime of this basin. 6.1.4.2
Calculation of Thermophysical Parameters far Mixed Lithologies The porosities oflithologic mixtures are calculated by the expression (Doligez et al. 1986) - v
1 - P(Z)
c;
z"il - PiZ
where Ci is the fraction of the ith Iithology unit in the rock, n is the number of fractions, and Pi(Z) is the porosity of this lithology unit at depth Z, computed from the world-average data (Sclater and Christie 198o; Deming and Chapman 1989; Burrus and Andebert 199o). The thermophysical parameters for a mixture of lithological units are calculated by the expressions (Deming and Chapman 1989; Ungerer et al. 199o) P~ =
Pm~G+
pm2C2 + ... + Pm~C~
,
p~,Cpm = p ~ C p 1 Q + pm~Cp2G + ... + PmnCpnCn , K m = KmIC~ + Krn~C2 + ... + KmnC n
The coefficient AI (temperature parameter of matrix heat conductivity) in Eq. 6.5 is determined using the values K m and K m i at a temperature of lOO °C: AI = o.ol{exp[Z tnCiLn(1 + looAli) ] - 1}
where C/is a fraction of the ith component in the considered rock, Pmi, PmiCp i, Kmi, and A1 i are corresponding characteristics of this fraction, 6.1.4.3
Latent Heat Effect The thermal effect of latent heat derived as a consequence of melting or solidification of peridotite rocks is considered as follows. First, the fraction of melted rock, f, is assumed to increase linearly with temperature, T, within the interval Ts < T < TI (Carlslaw and Jaeger 1959):
f-
T-~ TI-Ts
Chapter 6 • Burial History and Kinetic
z26
Modeling for Hydrocarbon Generation
Then the heat capacity of the rock Cp within this temperature range is replaced by Cp':
@'= @ + - - L
,
T I - Ts
where L is the latent heat of fusion (loo cat g-1 for peridotite) (Forsyth and Press 1971), the solidus temperature, Ts, is a function of pressure P (see Wyllie (t979) for peridotite with o.2% H20), and the liquidus temperature for peridotite is TI = Ts + 6oo °C and for basalt TI = Ts + 75 °C (Turcotte and Schubert 1982). 6.1.4.4
Lower Boundary Conditions Maintenance of heat flow of about lOO-12o mW m -2 at the lower boundary, Zlow,which is typical for continental rifting regions (Smirnov 198o) results in temperatures of about 2 0o0-3 ooo °C over a depth range of lOO-15o kin. These values exceed by 8oo-15oo °C the temperatures estimated by geophysical methods for such depths (Anderson 198o). The magnitude of the temperature difference affects the modeling results significantly; however, variations in the temperature at the lower boundary, Zlow,during basin evolution are approximately lOO-2OO °C and thus cannot significantly affect the thermal history of the basin. Because the lower boundary is located within a rheologicallyweak layer of the upper mantle, restoration of mechanical and thermal equilibrium takes place rather quickly (over 1-1o Ma with a horizontal heterogeneity of 5O-lOO km in width). Therefore, it is unlikely to expect large temperature variations at this boundary. For this reason, we prefer to maintain a constant temperature at the lower boundary rather than assign a constant heat flow. 6.1.4.5
Finite-Difference Scheme for Solution of Equation 6.3 The heat transfer in Eq. 6.3 is solved by implicit finite differences. This solution closely resembles that of Peaceman and Rachford 0955), but is modified to indude the variable values of the thermophysical parameters and variable steps dZ: (pCpT)~+I
dt
n+I n+l + (pW Cpw V T)k +1 - (pw Cpw V T)k_ 1 dZ k + dZk+ 1
r.-n+1 t,vn+l
*'k,k+l ',*k+l
T~+I ) 2 ~-~+~ (r~+~ _ r~_~~) *~k,k-1
dZk+l (dZk + d Z k + t )
dZ k (dZk + dZk+t)
A~+1 ,
(6.11)
where K~-I and K~{_~Iare the thermal conductivities averaged for the neighboring pair of elementary depth layers dZk, dZk+~ and dZk, dZk_l at the (n + 1)th time step. In Eq. 6.11 dZ k + dZk+ 1
Kk'k+l = dZk + dZk+l
Kk
Kk+l
dZ k + dZk_ 1
and
Kk'k-I - dZk ÷ dZk_l
Kk
Kk-1
The tridiagonal equation system for unknown values of T~+~,T~+I and Tff+1,T~_+~is obtained from the difference scheme (Eq. 6.11) and supplemented with the boundary
6.1 - T h e M o d e l s
227
conditions on the surface and at the base of computed domain. The tridiagonal equation has been solved by the usual driving method (Samarskiy and Gulin 1989), and the temperature distribution on the (n + l)th time step is obtained as a result. For constant thermophysical parameters and absence of convective members in Eq. 6.3, the implicit finite-difference schema (Eq. 6.11) is unconditionally stable; however, the parameters depend on porosity, temperature, depth, and time. In a real situation, one cannot analyze equation stability. The equations can be solved only empirically by comparison of solutions obtained for different time and space steps. The validity of the finite-difference scheme also was checked against the analytical solution for temperature distribution T(Z,t) and heat flow Q(t) from Carislaw and Jaeger (1959). We considered the following variants: (1) a steady-state deposition of homogeneous sediments on the basement having similar thermal characteristics, (2) erosion of a homogeneous half-space, (3) cooling of the half-space with homogeneous initial temperature T = Ts, and (4) a temperature distribution within homogeneous half-space with periodically variable surface temperature. In addition, we compared our numerical results with temperature and heat flow data computed by semianalytical methods (Golmstock 1979,1981; Galushkin and Smirnov 1987) for homogeneous sediments deposited on homogeneous basement with thermophysicat characteristics different from those of the sedimentary cover rocks. In all the considered cases, an appropriate choice of dZ and dt steps has allowed us to achieve a coincidence of computed and analytical temperature distributions and heat flows with an accuracy of up to o.1%. The analysis has shown that one must be careful in estimating erosion to avoid too large a step in dZ and dt. For example, the analytical surface heat flow exceeds by 30% the flow computed by the difference scheme, when the homogeneous half-space erodes by 9 m m / year and is numerically simulated with the step dZ = 50 m and dt = loo ooo years. The same difference does not exceed 1% for a step dZ = lO-2O m (dr = 20 000-40 ooo years). The correct numerical simulation of natural sedimentation and erosion processes requires sufficiently small values of time and space steps when applying a finite-difference scheme. In our calculations, the total number of time steps averages 1 ooo. Space steps, dZk, change linearly within each lithological unit and as a continual piece-linear function of depth for the whole column, including the sedimentary' cover and the basement. Space steps, dZ, can change from values of 1-4 m at the surface of computing domain up to 1-3 km at the bottom of the basement, at depth Z = loo--2oo km. 6.1.4.6
Tectonic Subsidence
The tectonic subsidence of a basin is the displacement of the basement surface compared to a reference level generally adopted as an initial surface position. In our model, the local isostatic response of the basement on load is assumed. Then the tectonic subsidence consists of two parts (Sclater and Christie 198o; McKenzie 1981): ZTs, originating from load on the basement surface (sediments, water column), and ZTb, originating from processes that change the density distribution vs. depth in the basement column (due to stretching, cooling, and heating of the basement): ZT = ZTs + ZTb .
(6n2)
Chapter 6 • Burial History and Kinetic Modeling for Hydrocarbon Generation
228
A t=O
B
b
t
i
'
Zw(O) %~._ V~zATER
I
A1 t=O
B1 t
v
| Zw(t) 1
i zt(t)
" - - 3 ~ 7 . - - 4 - ....
c (0)
a4s,
I
A~ S T 1~ E N 0 S A
I
H E R,E
Compensation depth
| | ! I
B
A S T H I
A|
,I
~
E N O S P H E R I,
Compensation depth
B|
Fig. 6,10. Principals of tectonic subsidencecalulation,a Situationwith load on the basement surface when the columnsAA (time t = o - onset of basin evolution) and BB (time t > o) include water, sediments, and basement surface; columns AIA1 (time t = o - onset of basin evolution) and BIB1 (time t > o) include only basement rocks. Equality of the weight of AA column to the weight of A1AI and of the weight of BB to the weight of B1B1leads to Eq. 6.13for calculatingthe first part ZTs of tectonic subsidence.In addition, the equality of the weight of the A1Az column to the weight of the BiB1 column leads to Eq. 6.15for calculation the second part ZTb of tectonic subsidence
The equality of the weight of column AA to the weight of column AIA1 in Fig. 6.1o leads to expression for ZTs determined as the position of the basement surface after removing the water and sediment load:
ZTs(t)-ZTs(O)= pa - P s s ( t ) + P a - P w [ Z w ( t ) - Z w ( O ) ] Pa
.
(6.13)
Pa
In Eq. 6.13,ZTs(t) is the tectonic subsidence amplitude, t is time and t = o is the time of the basin initiation, S(t) is the total thickness of sedimentary cover, Pa and Pw are the asthenosphere and water densities, respectively, Zw(t) is the paleodepth of water column at the time t, and ps(t) is the average density in the sediment column:
s(t) f ps(Z,t)dZ ps(t ) = 0
(6.14)
s(t)
The porosityp(Z,t) and the density ps(Z,t) of rocks within the sedimentary column and the weight of the column are computed for every time of basin evolution to determine the value of the average density in Eq. 6.14. The statement about equality of the weight of column A1A1 to the weight of colu m n B1B1 in Fig. 6.1o leads to expression for ZTb: z r b ( t ) - ZWb(O) -
G(t) -
G(O) Pa g
(6.15)
6.i
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Here, G is the weight of the basement column of fixed length to and g is the gravity acceleration: lo
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The weight G is calculated during modeling for every time step. The density of lithospheric rocks t~(Z,t) is a function of temperature r(z,t) and pressure P(Z,t):
pl(Z, t) = po(Z, t) [1 - a T ( Z , t) + tiP(Z, t)]
(6.17)
where po(Z,t) is the density at the standard conditions P = 1 atm and T = 2o °C, a is the thermal expansion coefficient and/3 is the isothermal compressibility of the rock. The parameter Po reflects the variations in the rock type vs. depth (crust, mantle, granitic or basaltic rock, the garnet peridotite to pyroxene peridotite, and pyroxene peridotite to plagioclase peridotite compositional transition in the mantle (Forsyth and Press 1971), and the changes in the density distribution within the basement due to basement stretching (Figs. 6.z, 6.3). The effect of change in the depth of the compositional transition boundary (plagioctase peridotite into garnet) on the tectonic subsidence amounts to the first tens of meters. Another compositional transition (pyroxene peridotite to plagioclase peridotite) commonly is not taken into account because, according to pressure-temperature conditions (Forsyth and Press 1971), the transition occurs within the crust, but not in the mantle. The bases of columns AA, BB, AIA1, B1B1 (Fig, 6.10) are located at the isostasy depth, Zi. In our model, Zi coincides with the lower boundary of the computed domain, Zlow, which was determined previously. The rather great depth for Zlowassumes that any small stress difference must be relaxed, and steady pressure must be restored, because the mantle rocks are rheologically weak at great: depths Z =Zi. In the frame of the local isostatic approach, Eqs. 6.13, 6.15, and 6.16 describe the main processes that can contribute to changes in the tectonic subsidence amplitude. The tectonic curve (Eq. 6.13), computed by removing of the water and the sediment load (solid line in Fig. 6.zc), must coincide with the subsidence (Eq. 6.15) determined by variations of temperature and pressure in the lithosphere (dashed line in Fig. 6.zc). The comparison of these tectonic curves allows additional control of the sequence of tectonic and thermal events in the lithosphere, and these variations are assumed in our basin modeling. This control, however, has a relative, rather than absolute, character. This isostatic approach assumes a local isostatic response of the lithosphere to the load. The stretching of the lithosphere or its cooling from a warmer state is believed to account for the subsidence of the basement surface, whereas the thermal activation of the basement is assumed to result in raising of the basement surface. The high relief near the Red Sea is an example of a thermal activity action; the Afar region is a result of simultaneous action of stretching and thermal activation of the basement. In dynamically active regions (young mountain belts, accretion prisms), the basement sinking can result from collision of the plate margins with overthrusting of
z3o
Chapter 6 • Burial History and Kinetic Modeling for Hydrocarbon Generation
neighboring massifs, nappes, etc. (as in the South Caspian Basin), whereas uplift of the basement surface also can result from the dynamic compression of lithosphere boundaries, such as that taking place on the front slopes of accretion prisms of island arcs or within the Alpine-Himalayan mountain belt; however, all of these processes are essentially nonisostatic and are characterized by free-air gravity anomalies exceeding loo mgal (Ushakov et al. 1979). In such areas, the dynamic correction to tectonic subsidence must be introduced into basin modeling. Another isostatic mechanism, explaining basement subsidence and related to the compositional change of light basalt to heavy eclogite, was suggested in Artyushkov et al. (1979, 198z). This mechanism, however, has no support in the observations of oceanic lithosphere, and particularly in the processes of jumping the midoceanic ridge axis (Galushkin and Dubinin 1991, 1992).
6,1.5 Conclusions Our computer program simulates the burial and thermal histories and petroleum potential in sedimentary basins and considers the thermal regime in the sedimentary blanket and the underlying lithosphere. Consequently, modeled calculations of basement tectonic subsidence can be used to control the sequence of tectonic and thermal events in the lithosphere. These computations assume that the lithosphere has local isostatic response to load; therefore, applications to areas in dynamic active belts having anomalously high values of free-air gravity require corrections for tectonic subsidence. Our program integrates the geological thermal stage of the source obtained from Rock-Eval pyrolysis data into the fitting procedure to determine kinetic reaction parameters. This results in better estimates of the low-energy part of the kinetic spectrum for source rocks having a vitrinite reflectance of o.5-o.8%. Neglecting the geological stage of organic matter maturation leads to a shift in the energies spectrum toward high energies and, consequently, results in underestimating hydrocarbon output. The unambiguous solution of kinetic spectra restoration requires detailed investigation of source rocks of the same kerogen type, but with different maturation levels. 6.2
Applying the Model to Saharan Basins The major goal of this chapter is to describe applying the galo basin-modeling program to hydrocarbon generation in provinces having different tectonic and sedimentation histories. The choice of North Africa to demonstrate the proposed system was conditioned by the following factors: (1) This province is one of the world's most important oil- and gas-producing regions. The oil- and gas-bearing rocks formed under different conditions and exhibit different properties that are closely associated with specific features of their evolution; (2) two periods of geologic evolution of the region are divided by the Hercynian orogeny, which resulted in significant uplift and erosion and in thermal pause, (3) the potential source beds in Saharan basins, in particular graptolitic Silurian shales, are considered among the world's richest; (4) although there
6.2 • Applying the Model to Saharan Basins
231
are a m p l e g e o c h e m i c a l data for t h e N o r t h Africa basins, t h e s e data are d i s p a r a t e a n d p o o r l y p r o c e s s e d . A s y s t e m a t i c effort to i n t e g r a t e t h e s e d a t a w i t h b a s i n h i s t o r y is needed. I n t h i s study, we u s e d a vast a m o u n t o f g e o c h e m i c a l data, i n c l u d i n g the results o f pyrolysis, c h r o m a t o g r a p h i c a n d optical analysis, a n d soluble o r g a n i c m a t t e r extractions. Pyrolysis m e a s u r e m e n t s were o b t a i n e d in a n o p e n Rock-Eval s y s t e m at differe n t h e a t i n g rates. T h e m o d e l i n g h a s b e e n c a r r i e d o u t for t e n areas ( T a k h o u k h t (TKT), A k f a d o u (AKF), M e r e k s e n (MRK), E 1 - B o u r m a (ELB), Keskessa (KA), Sbaa (SBAA), O u e d Saret (OS), O u e d K e r r a n e (ODK), Fort Lalleland (FLD), a n d A g u e m o u r (GM)) l o c a t e d i n e i g h t b a s i n s (Fig. 6.11). T h e T r i a s s i c p r o v i n c e b a s i n s ( O u e d e l - M y a , G h a d a m e s , Trias, a n d N o r t h Illizi) h a v e b e e n s t u d i e d in the greatest detail. The results o b t a i n e d for o t h e r b a s i n s are u s e d in the s y n t h e s i s o f t h e data into a r e g i o n a l f r a m e w o r k a n d are n o t r e p o r t e d here. Oil-source r o c k c o r r e l a t i o n s are b a s e d u p o n the dist r i b u t i o n o f s a t u r a t e d Clo a n d a r o m a t i c Cs h y d r o c a r b o n s , h y d r o c a r b o n C4-C 7, gaso-
Fig. 6.11. Location map of principal Saharan basins and studied fields. * indicates fields for which full computer simulation (burial, thermal, and hydrocarbon generation histories) was conducted. Fields studied by classic geochemical methods are specified in italic, AKF = Akfadou, ALR = Mrar, AR = E1-Agreb, AT = Ait Kheir, BKH= Bou Khechba, BST= Bst, DECH= Dech, D]HN= Djebel Heiran, DRT = Draa Tamra, DL = Edjeleh, El(= Ektaia, EKR = E 1-Khtir, ELB = E1-Bourma, FLD = Fort Lalletand, FZ = Fogaret ez-Zoua, GBC = Gour Bouchareb, GLA = Guellala, GM = Aguemour, GSL = Gsl, GTT = Gtt, HAD = Haid, HFR = Hfr, HMD = Hassi Messaoud, HR = Hassi R'Mel, IC = Ioucha, KA = Keskessa, KB = Krechba, MGD = Megadine, MDR = Messdar, M R K = Mereksen, NEK = Nekhila, OCT = Oued Chouicat, ODK= Oued Kerrane, ODZ = Odz, OK] = Benkahla, OKP = Berkaoui, ONR = Oued elNoumer, OS = Oued Saret, OTLA = Otla, RB = Rhourde el-Baguel, RE = Bit Rebaa, RG = Reggane, RN = Rhourd Nouss, RYB = Rhourde Yakoub, SBAA = Sbaa, SED = Sedoukane, STAH = Stah, TFT = Tin Fouye-Tabankort, TG = Tiguentourine, TKT = Takhoukht, T M T N = Timedratine, TOT = Toat, W T = Wadi-Teh, ZAR = Zar, ZES = Zemlet el-Nouss, ZM = Zemlet Mederba, ZR = Zarzaitin
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Chapter 6 - BurialHistoryand KineticModeling for HydrocarbonGeneration
line compositions, hydrocarbon C5-Cs oil compositions, and oil-source rock thermolysis products (soluble organic matter neglected for biodegraded oils), as well as effective source rock distribution. Carbon isotope composition (~13C) also was correlated for five oil and bituminoid fractions (hexane, hexane-benzene, benzene, benzenemethanol, asphaltene). Three regions in northern Africa were considered (Fig. 6.tl): the Takhoukht area of the northern part of the Oued el-Mya Basin where there were contrasting Paleozoic and Mesozoic tectonic and sedimentation histories and Hercynian erosion was at a maximum; the Akfadou area of the Ghadames Basin where there was moderate Hercynian uplift and erosion; and the Mereksen structure of the Illizi Basin where there was Hercynian erosion. Hydrocarbon sources include Ordovician and Silurian shales in the Takhoukht area, and Middle and Late Devonian shales in the Akfadou and Mereksen areas. These source shales have similar organic matter abundance and types. Both the Oued el-Mya and Illizi Basins are mature exploration areas; however, the Ghadames Basin is poorly explored due to the high cost of drilling and to technical issues. Also, the promising Paleozoic and Triassic layers are at great depths; available geological and geochemical data, although modest, suggest high hydrocarbon potential. Thus, any new effort in this field comes to the foreground of interest. The southern and western basins (Tindour, Reggane, Ahnet, Mouydir; see Fig. 6.11)are also poorly known. Nevertheless, available data from them have been integrated to describe the main trends of hydrocarbon occurrences. For brevity; the details of the tectonic and thermal histories are given only for the northern Oued el-Mya Basin. Directly relevant thermal features for hydrocarbon generation are reported for the Ghadames and Illizi Basins. Geochemical and modeling data acquired for areas in other basins have been compiled and integrated into the regional synthesis and conclusions. Approximately 23o samples of representative source shales were selected for detailed study from a large quantity of analyzed rock samples from throughout the Saharan basins. Cores were obtained from source intervals that show significant gas- and oilgeneration potential; that is, the Gothlandian, Givetian, Frasnian, Famennian, Strunian, and Tournaisian shales (Table 6.3). These intervals represent the main sources in the basins. Average relevant data are given when describing source formations and in discussions. Oils representing the major producing fields were studied as part of a separate oil-source rock correlation effort, and these data are reported here in the framework of regional synthesis. A large quantity of geochemical data, including the results of pyrolysis, chromatographic and optical analysis (Ro, TAI), and soluble organic matter extractions (OME), have been considered in this study. These data constrained the modeling procedures and ensured the reasonable estimates of hydrocarbon potential. Pyrolysis experiments were used to study bulk hydrocarbon generation and associated changes in composition of the kerogen. Open pyrolysis experiments were used to determine kinetic parameters and to quantitatively describe organic maturation reactions via a computer model based on simulation of tectonic, thermal, and geochemical histories. The thermal history model is integrated with available geochemical data from basins having widely varying geological conditions to assess hydrocarbon generation, migration, and accumulation.
6.2 o Applying the Model to Saharan Basins
z35
6.2.1 Geological Framework
The Saharan Platform is believed to have been a single cratonic entity from the Cambrian until the Holocene (Burollet 1967b). Epeirogenic warping and locally significant faulting affected sedimentation in some basins. The formation of the Saharan basins is believed to have been affected by repeated continental collisions during the Phanerozoic, resulting in faulted-block terrain (Furon 1963).Although subsequent neo-tectonics influenced the structural features of the basins, the orientation of principal structural elements is inherited in part from the tonic trends of the Precambrian basement (Burollet 1967b). The post-pan-African structural picture of the Saharan Platform, including the structural effects of the Caledonian and Hercynian genies, is obtained from the Phanerozoic structural evolution of northeast Africa presented by Klitzsch (1971,1981,1986), Klitzsch and Wycisk (1987), and Schandelmeier et al. (1987). The anticlinal arches and their complicating highlands are confined, as a rule, to buried projections (horsts) of the Precambrian basement and are represented by Paleozoic sediments of reduced thickness with frequent interruptions and scours, discordantly overlapped (in the northern areas) with Mesozoic and Cenozoic sediments (Figs. 6.12, 6.13). The synclines are filled with Paleozoic and Mesozoic-Cenozoic sediments. The stratigraphic column is more complete and thicker than those of the anticlines, especially in the western part of the Saharan Platform (Figs. 6.12, 6.13).The central and southern Saharan basins outcrop along a north-northwest-south-southeast uplifted zone that runs along the Ougarta range, the Hoggar and Reguibat massifs, and also at the south Illizi homocline, with the Tassili Najjer among the largest Paleozoic exposures (Deynoux 1983). The western Saharan syncline was subjected to rapid subsidence during the Paleozoic and early Mesozoic. This resulted in the formation of a large basin (7-8 km deep) of Paleozoic sediments. The central and eastern Saharan synclines in the Paleozoic were areas of slower subsidence (in individual periods of the Paleozoic), and marine sediments are comparatively thin (from z to 4 km) (Fig. 6.13). During the early Mesozoic, the eastern Saharan synctine, including the Trias, Oued el-Mya, and Ghadames Basins (Figs. 6.1z; 6.13a,b), subsided rapidly. This resulted in the accumulation of up to 4 km of marine and, in places, continental Triassic, Jurassic, and Cretaceous sediments. Basement faults are extensive in the Saharan Platform. These faults influenced the formation of various tectonic features, including vast depressions and basins (synclinoria, step grabens). Rejuvenation of basement faults along the northern border of the Floggar massif (shield) created horsts and grabens with folds in Paleozoic strata. In the northern Ghadames and Oued el-lHya Basins, which were uplifted at the end of the Early Carboniferous, Paleozoic strata were stripped off before a postHercynian transgression (Burollet 1989). In Tunisia and northern Libya, a set of faults created tilted blocks and steps along the southern margin of a proto-Tethys Ocean (Klitzsch 1971),causing the deposition of Carboniferous and Permian shallow-marine carbonate tithofacies. A brief tectonic and sedimentation history of the Saharan Platform is summarized in Table 6.4. At the beginning of the Paleozoic, the South Pole was located just north of Africa in the lapetus Ocean (Petters 1991). Quartz-rich sandstones began to accumulate in
236
Chapter 6 . Burial History and Kinetic Modeling for Hydrocarbon Generation
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the Cambrian along North Africa when there was a progressive marine transgression, which continued into the Ordovician (Burollet 1989; Klitzsch 199o). Uplifts, such as the Hoggar and Reguibat, did not exist at that time; they rose after the Ordovician. Meanwhile, the Tindouf, Reggane, Ahnet, Mouydir, Ghadames, Tilizi, Murzuk, and Kufra (in Libya to the east) intracratonic basins subsided (Petters 1991). Gondwana drifted over the South Pole during the Ordovician (Neugebauer 1989), and by the Late Ordovician the South Pole was located far inland in northwestern Africa, leading to widespread continental glaciation. By the Middle Ordovician, interlayered shallow-marine sandstones and argillites had been deposited. Dark argillites, consisting of micaceous graptolitic and trilobitic clays, were deposited during an extensive marine transgression presumably due to melting of the Saharan glacial cap (Beuf et al. 1971; Rognon 1971). No substantial interruption in deposition between the Ordovician and Silurian is observed. Repeated marine transgressions during the Early Silurian resulted in the deposition of thick, dark graptolitic clays, argillites, and sandstones. The Ghadames and Oued el-Mya Basins, which have high potential for hydrocarbon generation, were major argillite depocenters at this time. Early Devonian continental sediments with plant remains discordantly overlie the Silurian strata. Alternating sequences of clays and sandstones reflect an extended cycle of alternating marine transgressive-regressive deposition during the Devonian. Devonian shales, particularly the Middle and Upper Devonian, and Silurian shales are considered to be the principal source rocks in the Saharan basins. Carboniferous shales are also considered to be good source rocks. General uplift resulting from the Hercynian orogeny led to a major withdrawal of the sea (Bishop 1975). The dominant feature of post-Hercynian erosion on the Saharan Platform is a T-shaped anticlinorium that extends from Algeria into Tunisia. To the east, a projection of the Hercynian Nefusa uplift of Libya extends westward and connects with this anticlinorium. The absence of Permian sediments in the Algerian region suggests that this area remained uplifted. Marine transgression taking place at this time resulted in the deposition of thick Permian marine sediments in Tunisia and Libya to the east. These Permian shales form seals to Silurian sandstone reservoirs in the Libyan oil fields. Restriction of the western Tethys basin and post-Hercynian subsidence along the margin of the African landmass led to a new cycle of sediment deposition, which included a thick series of Triassic and Liassic evaporates. This two-stage history influenced source rock and reservoir rock diagenesis. Triassic sediments are widespread in the northeastern part of the Saharan Platform, namely in the Ghadames, Oued elMya, and Trias basins, as well as in a part of the northern flank of the Illizi Basin. Triassic fluvial and shallow-marine sandstones commonly overlay the surface of the Hercynian unconformity. Triassic andesitic and basaltic flow's are abundant in the Triassic section and commonly overlie Cambrian and Ordovician sandstones above the Hercynian unconformity, forming a good seat. Volcanic activity at this time suggests a thinned crust and thermal activation related to the Hercynian orogeny. The Middle Jurassic and younger section on the stable Saharan Platform is dominated by relatively thin lagoonal dolomites, evaporates, and shales. Cretaceous sediments consist of alternating evaporates, limestones, dolomites, and thin layers of sandstone. Aptian-Albian nearshore carbonate facies are oil-bearing in Tunisia. Tertiary sedimentation is particularly present in Tunisia and farther over a wide scale of thickness (up to 7 ooo m). These sediments occur from the Paleocene to the Pliocene in the east, and
240
Chapter 6 • Burial History and Kinetic Modeling for Hydrocarbon Generation
in the nearshore areas of Gabes Gulf on the Tunisian Mediterranean coast. The Tertiary Alpine orogeny uplifted the unstable part of the platform and formed a multitude of folds and complex structures. The Atlasides folded belt forms the northern province. 6.2.2 Oued eI-Mya Basin 6.2.2.1
Tectonic Subsidence and Thermal History The burial and thermal histories of the northern Oued el-Mya Basin are presented in the model. Two methods are used to calculate the relative change in tectonic subsidence that determines the sequence of thermal and stretching events in the lithosphere (Fig. 6.2). A brief description of this sequence follows. Slight variations in the amplitude of tectonic subsidence from 6oo to 48o Ma indicate only moderate variations in heat flow during this time. This reflects slow cooling of the basement lithosphere from a thermal state having an initial heat flow of about 52 m W m -a. Basement subsidence from about 4oo to 35o Ma accompanied the deposition of about 2 50o m of clays and sands, and involved basement stretching with an amplitude of about 1.2 for 95 Ma (Fig. 6.2). Slow stretching rates resulted in a Moho depth change rather than a change in isotherm depths. Sinking of the isotherms at about 49o Ma was due to climate cooling, which continued up to the Early Carboniferous. The subsequent rise of the isotherms at 49o-35o Ma was due to the transition from low-temperature gradients in the basement (high thermal conductivity) to higher temperature gradients in the sedimentary cover (low thermal conductivity). Devonian sedimentation was followed by an interruption that lasted throughout the Carboniferous. The subsequent Hercynian orogeny resulted in uplift and erosion of the northeastern part of the basin, including the Takhoukht region. We estimate that about a 2oo m of Devonian and Silurian sediments were eroded. Thermal activation of the lithosphere in the northern Oued el-Mya Basin began in the Late Carboniferous (28o Ma) (Fig. 6.2). Thermal diapir uplift occurred at an average rate of about 5-5 km Ma-1 for a period of lo Ma Diapirs remained immobile for 35 Ma at a depth of less than 3o km. Surface heat flow reached 9o mW m -2, which is close to the values observed in present-day continental rifts (Smirnov 198o). The presence of relatively thick Triassic volcanics in the Oued el-Mya Basin is evidence of high thermal gradients in the Permian-Triassic. Subsidence of the basement in the Middle Triassic was a consequence of rapid cooling of the anomalously warm basement. Rapid deposition of salts and anhydrides, with their high thermal conductivity, also contributed to the sinking of isotherms in the Jurassic and the Cretaceous. In the Early Cretaceous, deposition was accompanied by stretching of the lithosphere (stretching amplitude about 1.2), which lasted to the end of the Cenomanian (Fig. 6.2). This second stretching phase accounts for the subsidence of the top of the basement during the last thermal activation of the lithosphere, which began in the Berriassic (145 Ma). The last thermal activation was accompanied by uplifting of the thermal diapir's roof at an average rate of 1 km Ma -1 for approximately 2o Ma in the Aptian and Albian. This roof remained at a fixed depth of about
6.2 • A p p l y i n g t h e M o d e l t o Saharan Basins
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60 km from the Albian to the present. The last thermal event explains the rather high temperature gradients in the present-day sedimentary section, which contains thick evaporates and a relatively significant level of maturation in the Lower Silurian rocks (Ro= o.7o-o.8o%). The rising of isotherms in the Lower Cretaceous is related to this heating event, as well as to the deposition of low-conductivity sediments; however, the deposition of 8oo m of sak during the Albian and Cenomanian resulted in short-term sinking of the isotherms. Slow sedimentation during the Cenozoic resulted in only minimal variation in the depths of isotherm and heat flow. The relatively high value of this heat flow (about 6o mW m -2) is in accordance with the high value of the presentday thermal gradient in the salt-bearing sediments of the northern Oued el-Mya Basin. The calculated present-day temperatures correlate well with the temperatures measured at 3 739, 3 785, and 3 989 m in boreholes. 6.2.2.2
Source Rocks
The principal hydrocarbon source rocks of the Oued el-Mya Basin are Silurian (Gothlandian) and Devonian (Ernsian, Givetian, Frasnian, Famennian) shales and, to a lesser degree, Ordovician shales (E1-Gassi and Azzel formations). The present areal distribution of Paleozoic (Ordovician, Silurian, and Lower Devonian) shales is a function of their initial distribution and the extent of Hercynian erosion. Maximum initial thicknesses were south, southwest, and west of the basin. Present-day thicknesses range from 6oo to 7oo m in the south to 28o to 66o m in the west and 22o to 46o m in the center of the basin. The Takhoukht section has about 4oo m of shales. The Ordovician shales contain mainly sapropelic, or mixed, organic matter and have an average TOC (total organic carbon) of o.9%. In our thermal model for the Takhoukht region, the base of the Ordovician section reached the main stage for oil generation (Ro = o.7%; TTI (time-temperature index) = 9o) at the end of the Cretaceous (Table 6.5). In the southern part of the basin, the main stage for oil generation could have
Table 6.5. Computed characteristics of the main source formations, Takhoukht area, Oued el-Mya Basin a
Layer
Depth (m)
T (°C)
Ro (%)
TTI
Early Carboniferous ( - 3 6 0 Ma) Ordovician shales Lower Silurian shales
2 438 - 2 635 2 359 -- 2438
84 82,5 -
90 84
0.490 - 0,526 0.480 - 0.490
4 3-
6 4
70 65 -
90 70
End o f the Mesozoic (~65 Ma) Ordovician shales Lower Silurian shales
3 743 - 3 922 3672 - 3 743
107.3 - 102.7 101 102.7
0.670 - 0.700 0.660 - 0,670
3924-4t00 3854-3924
103.2 - 108 101.2 - 103.2
0,735 - 0 . 7 6 7 0.723 - 0 . 7 3 5
Present Ordovician shales Lower Silurian shales
131 - 159 1 2 8 - 131
a T - t e m p e r a t u r e ; R0 - vitrinite reflectance calculated by kinetic model of vitrinite of Sweeney and Burnham (1990); TTI = time-temperature index (Lopatin 1971 ; Waples 1980).
242
Chapter 6 . Burial History and Kinetic Modeling for Hydrocarbon Generation
Fig. 6.14. Hydrocarbon (ttC) yields (solid line), rates of hydrocarbon generation (dashed line), and expulsion threshold in the geological history of the Silurian source shales of the Oued el-Mya Basin. Calculations used time-temperature history of the Silurian rocks and activation energy spectrum shown in Makhous et al. (1997). Two stages of hydrocarbon generation took place in the basin history: in the pre-erosion Carboniferous and during the Cretaceous-Cenezoic
400 300 O"
300
Time (Ma) 200
HC yield - - - Rate of HC generat on
0
100
Secons stage /L 6 ofHCgeneration / t v . ~
o
200-~
:t
,
,
3}
i l00
0
First stage of HC generation I "
0
been reached as early as the end of the Paleozoic. Today, in the northern part of the basin (Takhoukht area), organic matter at the base of the Ordovician is mature with respect to oil generation (R o = o.73-o.77%; TTI = 13o-16o; Table 6.5). In the southern part of the basin, Ordovician source rocks are in the gas window. Silurian and Early Devonian shales contain sapropelic, mixed and humic organic matter having a TOC range of 1.o to lO.O%. Lower Silurian radioactive shales in the central and northeastern parts of the basin contain up to 16% TOC. Our modeling shows that the onset of oil generation in the Lower Silurian and Early Devonian shales occurred at the beginning of the Cretaceous (Re = o.65%; TTI = 7); peak of oil generation (Re = 0.70%; TTI= 75) occurred as early as in the Albian (Figs. 6.14, 6.15). Today, the Silurian shales in the Takhoukht region are mature (Re = o.71-o.73%; TTI = lOO-13o), whereas in the south they are overmature (Re = 1.25-1.7o%), and peak oil generation occurred in the Paleozoic. The Middle Devonian, Late Devonian, and Carboniferous section in the southwest and northwest of the Oued el-Mya Basin is characterized by high-TOC shales (0.5-2.5%) and by mature to overmature kerogen (R o = o.7-1.5%). Carboniferous shales are considered potentially good sources for gas generation only because of their moderate organic matter contents and restricted occurrences within the basin. The Triassic, Jurassic, and Cretaceous sediments have low organic matter contents (generally TOC = O.l-O.3%) and low maturation level (Re = o.4o-o.5o%); consequently, they have poor hydrocarbon potential.
6.2.2.3 Maturation History and Hydrocarbon Generation in the Takhoukht Area Ordovician Sources (EI-Gassi Formation). Modeling results for the Ordovician shales are presented in Table 6.5. The Ordovician shales have an average hydrogen index (HI) value of ~95 mg HC / g TOC and an average TOC of o.78%. HIIOI (oxygen index) correlations show that the Ordovician source rock's kerogen is a mixture of type I kerogen with the initial potential of 7:o mg HC / g TOC (Espitali~ et al. 1988) and type II
6.2 • Applying the Model to Saharan Basins 0 0
•
1-
~
~ ~,
S
D
¥1
¥
N.~
z43
C i
f
P i
IT
T
J
¥1
¥
K i
¥1
CZ ~
¥
i
~
0.5
0.95 ~ 1.3
4-
1.8 ~ 5-
s90
s40
4;o
44o
3;o
340
2;0
Geological
240
1;0
140
;0
40
0
time (Ma)
Fig. 6.15. Generalized pattern of the main source rocks, burial, generation, and expulsion histories. Solid lines = the main variant with erosion; dashed lines = variant free from erosion. Long dashed lines
represent a stabilization of source subsidence in some areas beginning in the Late Cretaceous and during the Cenezoic.It stands to reason that the achieved maturation level during source subsidence is irreversible, and Ro values do not change in the course of uplifting
kerogen with the initial potential of 630 mg HC / g TOC. Using the paleotemperatures computed for the Ordovician section by our model, we obtain total hydrocarbon yields of 84 mg HC / g TOC for type I kerogen and 493 mg HC / g TOC for type II kerogen. The residual potentials are 626 (type II) and 137 (type I) mg HC / g TOC. We infer that the observed residual potential of z95 mg HC 1 g TOC represents a mixture of kerogen types (about 68% type II and 32% type I). Our calculations show that the hydrocarbon yield from the Ordovician sources was 4o-45 mg HC / g TOC by the end of the Carboniferous, and is 36o mg HC / g TOC today. Hydrocarbon yield during the first Paleozoic stage of maturation represents 12% of the final yield. Silurian Sources. Restoration of the activation energy spectrum for hydrocarbon generation in our program used an example from the Gothlandian source shales in the Takhoukht area (Fig. 6.2). The initial hydrocarbon potential was determined to be HIo = 630 mg HC / g TOC, which is typical for open-marine kerogen (type II) (Espitali~ et al. 1988; Ungerer et al. 199o). The present residual potential of the sources is about 45% of the initial potential, implying that about 55% was generated between about 350 Ma and the present (Fig. 6.14). According to our calculations, hydrocarbon yields were 96% liquids and 4% gas. Less than 0.2% of liquid hydrocarbons were subjected to secondary cracking. The calculated hydrocarbon history is shown on Fig. 6.14, The yield during the Carboniferous (from 360 to 286 Ma), prior to the Permian erosion, accounts for about 6% of the total generated hydrocarbons. A small local peak on the hydrocarbon yield rate curve between 360 and 286 Ma (Fig. 6.14) corresponds to this first stage of hydrocarbon generation. The comparatively low yield rates are associated with moderate tem-
z44
Chapter 6 - Burial History and Kinetic Modeling for Hydrocarbon Generation
peratures (82-85 °C). A second and final stage of hydrocarbon generation occurred during the Campanian (12o-9o Ma), when source temperatures exceeded loo °C. In this stage, the rate of hydrocarbon generation was one order of magnitude higher than the rate of generation during the Carboniferous (Fig. 6.14). The activation energy spectrum for hydrocarbon generation (Fig. 6.2) indicates that two reactions, the first having an/~i (activation energy) of 50 kcal mo1-1 and the second 52 kcal mo1-1, were the major contributors to hydrocarbon generation and account for about 86 and 13% of the yield, respectively. Expulsion from the source is assumed to commence when zo% of the free pore space is saturated with liquid hydrocarbons. For the most organic-rich Gothlandian source shales (TOC = 11.8%), ~o% pore saturation was achieved at the beginning of the Coniacian (about 88 Ma; Fig. 6.14), when the hydrocarbon generation was as high as about 35 mg HC 1 g TOC and source rock temperatures reached 9o-loo °C. The expulsion threshold for shales having TOC = ~4.4% was attained 3 Ma earlier. Mesozoic subsidence was accompanied by formation of new structures at the same time that pre-existing traps were completely or partially destroyed. Hydrocarbons migrated from the destroyed traps and accumulated in traps having Triassic and lurassic seals. Consequently, reservoirs in the basin are either Triassic (Berkaoui, Benkahia, Guellala) or Paleozoic (Hassi Messaoud, E1-Agreb) (Fig. 6.11). Further subsidence of the Paleozoic sources during the lurassic and the Cretaceous caused hydrocarbon generation, which was followed by hydrocarbon migration and accumulation in traps in the northern part of the basin. Both Triassic and Paleozoic reservoirs were filled with hydrocarbons generated in Paleozoic sources. Most of the hydrocarbons generated in Paleozoic source rocks during the ]urassic-Cretaceous and following lateral or, to a greater extent, vertical migration were trapped in Triassic reservoirs. Migration pathways along the Amguid-el-Biod/Hassi Messaoud axis were in a southward direction far from subsided zones to more uplifted zones. The migration of hydrocarbons generated in Devonian shales to more uplifted zones also took place in the east from the Ghadames Basin. 6.2.3 Ghadames and Illizi Basins 6.2.3.1
TectonicSubsidence and Thermal History The first major tectonic event affecting the Ghadames Basin was the Hercynian orogeny. Prior to its onset at the close of the Carboniferous, 2 8oo-3 9oo m of Paleozoic sediments had accumulated. The Hercynian orogeny resulted in uplift and the subsequent erosion of about 9oo m of Late Paleozoic sediments (Fig. 6.15). In the lurassicTriassic, the Ghadames Basin was the center of inversion tectonic movements, which caused subsidence of its northern, northwestern, and western parts; that is, in regions that experienced the most uplift at the close of the Pateozoic (Makhous et al. I995). Evaporates deposited at this time covered an area that stretched far beyond the boundaries of the depression and had thicknesses exceeding 9OO-lOOOm in the Messdar, E1-Khtir, Fort Lalleland, and Keskessa areas. By contrast, there was minimal subsidence in the Triassic-lurassic and an absence of evaporates in areas subjected to maxi-
6.2 - Applying the Model to Saharan Basins
245
mal subsidence in the Paleozoic (Rhourde Nouss, Ektaia, and the southeast area of the basin). Maximal total thickness of all Triassic-Jurassic formations, including detrital sediments and evaporates, is 19oo-190o m in the Wadi-Teh, Haid, and Bir Rebaa areas. In the Cretaceous, subsidence rates in the southeastern, western, and northwestern regions of the basin decreased, and the final configuration of the Ghadames Basin was formed. The Ghadames Basin is a Mesozoic tectonic structure whose closure took place in the Cretaceous. The sequence of tectonic events in the Akfadou area is derived from an analysis of the tectonic subsidence of the basement surface. The sequence basically repeats the sequence determined for the Takhoukht area: stretching through the OrdovicianDevonian, thermal activation during the Permian, and basement stretching in the Early Cretaceous, with thermal activation of the lithosphere from the beginning of the Cretaceous until the present. A high maturation level of organic matter is observed in the lower horizons of the sedimentary section: Ro = 1.oo-ln9% at depths of 3.0 3.9 km. Higher temperatures are characteristic of the sedimentary section of the Akfadou area as compared to the Takhoukht area. Temperatures of about 12o °C were reached in Early Devonian sediments. Our model calculations for the end of the Early Carboniferous (about 33o Ma), and before the Hercynian uplift and erosion, give temperatures and maturation levels for the Ordovician, Silurian, and Early Devonian shales consistent with early to main stages of hydrocarbon generation (Table 6.6). Toward the close of the Carboniferous (about 289 Ma), the sapropelic and humic organic matter of the Ordovician, Silurian, and probably Devonian source shales was sufficiently mature to generate both liquid and gaseous hydrocarbons. Silurian and Devonian shales in other areas of the basin are also characterized by considerable potential
Table 6.6. C o m p u t e d characteristics of the m a i n source formations, A k f a d o u area, G h a d a m e s Basin a
Layer
Depth (m)
T (°C)
Ro (%)
TTI
3220-3440 2 7 8 0 - 3220 2116 --2780 1280-2116
113 - 118 1 0 2 - 113 86 - 102 6286
0.723 - 0 . 7 5 6 0.640 - 0 . 7 2 3 0.510 - 0.640 0,367-0.510
4300 4500 3870-4300 3 270 .- 3 870 2 5 5 0 - - 3 270
t 39 144 128 - 139 114 - 128 96 114
1,111 - 1.t70 0,920 1.111 0,770 - 0.920 0,650 - 0.770
4650-4850 4240-4650 3650-4240 2 9 5 0 - 3 650
146136 121 104 -
1.267 1.082 0.861 0.709
Close of Early Carboniferous ( - 3 3 0 Ma) Ordovician shNes Silurian shales Lower Devonian shales Middle and Upper Devonian shales
42 17 5 0.8-
54 42 17 5
End of the Mesozoic ( - 6 5 Ma) Ordovician shales Silurian Shales Lower Devonian Shales Middle and Upper Devonian shales
920 425 165 42
-1300 - 920 - 425 - 165
] 923 866 323 88
-2775 - ] 923 - 866 - 323
Present Ordovician shales Silurian shales Lower Devonian shales Middle and Upper Devonian shales
152 146 136 121
-
1.378 1267 1.082 0.861
a T = temperature; R0 = vitrinite reflectance calculated by kinetic model o f vitrinite o f Sweeney and Burnham (1990); 3-1-1= t i m e - t e m p e r a t u r e index (Lopatin 1971 ; Waples 1980).
z46
Chapter 6 • Burial History and Kinetic Modeling for Hydrocarbon Generation
to generate liquid hydrocarbons. The relatively moderate Hercynian uplift in the region and consequent erosion slowed down organic matter maturation slightly. According to our modeling, the resulting temperature decrease ranges from ao °C for sediments at the base of the Ordovician to 16 °C for sediments at the base of the Carboniferous. Toward the end of the Mesozoic subsidence (about 69 Ma), organic matter in the Ordovician, Silurian, and Devonian shales was subjected to further maturation, and temperatures ranging from 114 to 144 °C and Ro values of o.8-1.2% were reached (Table 6.6). Ordovician and Silurian source shales realized most of their hydrocarbon potential in the Paleozoic. In the Mesozoic, hydrocarbon generation occurred throughout the Paleozoic section, including Carboniferous shales, in areas of maximal subsidence. During the entire history of sedimentation in the Illizi Basin, the granodiorite basement of the Mereksen area was a horst with an amplitude of some zoo m, but rapidly damping out in Cambrian-Ordovician sediments. In the Triassic and particularly in the Jurassic, tectonic movement caused horizontal stretching of the crust, resumption of movement on old faults, and the generation of new faults. The amplitude of faults within and bounding the horst decreased considerably with time. These processes were accompanied by subsidence of the basement. The Mereksen region is considered to be an old consedimentation structure with a tendency for disintegration, with subsidence of the basement in the Jurassic and at the beginning of the Cretaceous. Prob-
400
6OO
Time (Ma)
200
~2
E OJ
4 Sec bot
5
.......
Bas
.........
Jsol refl
. . . .
lsol
Fig. 6.16. Burial, thermal, and maturation histories of the sedimentary section of the Akfadou region, Ghadames Basin. The greater volume of sedimentation compared with the variant of Oued elMya Basin in Fig. z in Makhous et al. (1997) led to the greater level of maturation of organic matter in Silurian rocks, despite a moderate level of Hercynian uplift
6.2 - Applying the Model to Saharan Basins
247 Time (Ma)
310
510
II0
'\,
Tr-J
"X
oo"......................... A
2
o ~
Ds
S-""
¢-
¢1
..~.....
- o 5 0°/o 4..
"\',
s
.
.
.
.
.
3
° z -
Basement surface
.........
Isolines of vitrinite reflectance
.
"'~..~
"~'Z
,,,"
~ N °!° | a o C : ......]
Sedimemary layers boundaries
-
.......
.
/
.
.
Isotherms
I1//I / / #
1
I
Fig. 6.17. Burial, thermal, and maturation histories of the sedimentary section of the Iltizi Basin in the Mereksen region. Relatively higher maturation level is achieved in the absence of Hercynian uplift in the area
able source rocks in the Mereksen region are thinner than those in the Ghadames Basin due, in part, to slower sedimentation rates (Figs. 6.16, 6.17). Toward the end of the Carboniferous, Paleozoic basal rocks were at a depth of about 3 2oo m. Interrupted sedimentation was the only effect of the Hercynian orogeny. The region has generally higher thermal gradients than the Ghadames Basin (Makhous et al. 1995). The burial history and thermal model for this basin use an analysis of tectonic subsidence vs. time, and are in good agreement with the present temperature gradient and maturity levels estimated from measured vitrinite reflectance. The sequence of tectonic events that correlate with variations in tectonic subsidence of the basement surface includes periods of minor basement stretching and thermal activation commencing in the Early Cretaceous and continuing to the present. This thermal activation in the basement helps to explain the relatively high temperatures observed in the present sedimentary section. Temperatures of lO5-1o8 °C, measured at a depth of 2 776 m in a Mereksen borehole, are in good agreement with the calculated value of lO7 °C at the base of the Middle Devonian layer at a depth of 2 779 m (Fig. 6.17). The relatively high stage of organic matter maturation (Ro = o.7-1.2%) measured in Devonian shales occurring at depths of 2.5-3.1 km in the Mereksen area exceeds the results of our numerical simulation (Ro = o.8% at depth of about 2.9 kin) (Table 6.7). This high stage of maturation could be attributed to water infiltration into outcropped basin strata and to the thermal effect of rising groundwater flows; however, this problem requires an additional detailed hydrologic investigation.
248
Chapter 6 . Burial H i s t o r y and Kinetic M o d e l i n g f o r H y d r o c a r b o n G e n e r a t i o n
Table 6.7. Computedcharacteristics of the main source formations,Mereksen area, Illizi Basina Layer
Depth (m)
R0(%)
T(°C)
TTI
Close o f Carboniferous (~288 Ma) Ordovician shales Silurian shales Lower Devonian shales Middle and Upper Devonian shales
2 493 - 3 177 2213 - 2 4 9 3 2060-2213 1 741 - 2060
97 - 112 90 - 97 8 6 - 90 78 - 86
0,600 .- 0,707 0.535 - 0.600 0,511-0.535 0.455 - 0.511
14 7.7 5.6 2.9
-
47 14 7,7 5,6
End o f the Mesozoic (~65 Ma) Ordovicianshales Silurian shales Lower Devonian shales Middle and Upper Devonian shales
3179-3832 2918-3179 2777-2918 2485 - 2 7 7 7
115 108105 98 -
129 115 108 105
0,814 1,020 0#50 --0.814 0.730 - 0 . 7 5 0 0,670 - 0.730
3179 - 3 832 2918-3179 2 777 - 2 918 2485 - 2777
116 t09106 99 -
130 1t6 109 t06
0,862 0.790 0,760 0.712
266
-
710
160
-
266
125
-
160
75
-
125
Present Ordovician shales Silurian shales Lower Devonian shales Middle and Upper Devonian shales
.- 1,087 -0.862 - 0.790 - 0.760
391
--1209
229
-
391
191
-
229
109
-
191
a T = t e m p e r a t u r e ; R0 = vitrinite reflectance calculated by kinetic model of vitrinite of Sweeney and 8u rnham (1990); TTt = time-temperatu re index (Lopatin 1971; Waples 1980).
Slow, continuous sedimentation during the period beginning in the Permian resulted in minimal variation in isotherm depths and in the depths of Ro isolines (Fig. 6.17) in the Mereksen area. Our calculations of temperature and organic matter maturity in Ordovician, Silurian, and Devonian rocks at the close of the Carboniferous (approximately 288 Ma) are shown in Table 6.7. These calculations show that at the beginning of the Permian, organic matter in Ordovician and Silurian shales was in the lower part of the oil window (Fig. 6.17). Further subsidence contributed to a temperature rise and further maturation of organic matter. At the close of the Mesozoic, temperatures and maturation indicators in the Ordovician, Silurian, and Devonian shales were loo-13o °C, Ro = o.7o-l.oo%, and TTI= 7o-7oo (Table 6.7). These values suggest that at the close of the Mesozoic, no potential remained for liquid hydrocarbon generation in the Ordovician and some of the Silurian shales, whereas the Famennian shales were at peak generation conditions. Present-day temperatures, Ro, and TTI in the Ordovician, Silurian, and Devonian layers (Table 6.7) suggest that today Early Ordovician shales at a depth of about 3 8oo m are generating dry gas, and the top Ordovician shales are generating wet gas and condensate. The generation of liquid hydrocarbons in the Silurian, Early Devonian, and Middle Devonian shales to a large extent must be exhausted. 6.2.3.2
Maturation History and Hydrocarbon Generation in the Ghadames Basin The sediments in the Ghadames Basin in the Akfadou region were subjected to a higher thermal regime with reduced erosion amplitude and absence of evaporates in the geological section compared to the northern Oued el-Mya Basin in the Takhoukht region.
6.2 , Applying the Model to Saharan Basins
249
Ordovician shales (E1-Gassi and Azzei formations) in the Ghadames Basin and southern part of the Trias basin have a TOG range of o.5 to 1%. Ordovician shales in the north of Illizi Basin have TOC contents that average 1.3%. At present, the amorphous Ordovician organic matter is in the main phase of oil generation. Silurian (Gothlandian) shales in the Ghadames and Illizi Basins have TOC concentrations ranging from o.5 to 2.0%. The kerogen is amorphous and presently has no remaining oil potential. Areas with the higher TOC generally correspond to primary depocenters and to a moderate level of maturation. There are numerous areas where the Silurian shales have lower TOC and overmature organic matter. The maturation history of Silurian sources in the central part of the Ghadames Basin, including the Akfadou region, was only minimally affected by Hercynian erosion. Our paleotemperature calculations indicate that these sources generated petroleum as early as the Early Carboniferous (Fig. 6.16). Since the Hercynian uplift in the central part of the basin was minimal, the appropriate conditions for hydrocarbon generation were not interrupted (Fig. 6.16). Geochemical data show that these currently overmature Silurian source shales initially had considerable oil potential. Their estimated initial potential for hydrocarbon generation in the Ghadames Basin can be considerably higher than that estimated only on the basis of mean content of the present-day overmature organic matter, which is low. The Ghadames Basin contains thick beds of Devonian shales containing oil-prone amorphous organic matter and TOC concentrations ranging from o.5 to 5.o%. This organic matter is mature to overmature (R o = o.7-1.6%) in the Upper and Middle Devonian shales and often overmature (R o = o.8-2.0%) for Lower Devonian shales. These shales are considered to be excellent sources of liquid hydrocarbons. Middle and Late Devonian source shales in the central part of the Ghadames Basin are currently within the oil- or gas-condensate windows at temperatures of ~oo-~1o °C ('fable 6.5, Fig. 6.16). Our model calculations of the quantities of hydrocarbons generated and the generation rate of hydrocarbons in the Late and Middle Devonian shales of the Akfadou region are shown on Fig. 6.18. About 9o% of the initial hydrocarbon
Fig. 6.18. Hydrocarbon yields
(solid line), rates of hydrocarbon generation (dashed line), and expulsion threshold in the geologicalhistory of the Middie and Late Devonian source shales of the Ghadames Basin, Akfadou region
400 300
Time (Ma) 200
300
I
HCyield
]
8
/
- - - Rate o f HC ~
o
0
100
I ', i
.~
i
I I
I
6~
400
~o
U =1=
g U "r
~ 2oo 0
E
U
/
4"1" E
\
0
250
Chapter 6 - Burial History and Kinetic Modeling for Hydrocarbon Generation
potential has alreadybeen generated. The rate of hydrocarbon generation has a clearly defined peak in the Late Cretaceous, when formation temperatures reached 11o °C (Figs. 6.16,1.18). The decrease in the generation rates from 8o Ma to the present is because most of the generation is controlled by reactions having tow activation energies (50 and 52 kcal mol-1). Late and Middle Devonian source shales having an average TOC of 4-5% reached the expulsion threshold at the end of the Aptian (115 Ma). Gas products could account for less than 5% of total hydrocarbon output. For comparison, we give modeling results for a standard spectrum of kerogen type II with initial potential HIo = 63o mg HC/g TOC (Espitali~ et al. 1988). About 7o% of this potential was realized during the burial history of the formation. The expulsion threshold was reached in 7o Ma. The expulsion threshold for liquid hydrocarbons was attained in Campanian time (nearly 8o Ma) for Late Devonian source shales in the Mereksen region. During the Paleozoic, only some Ordovician and partially Silurian shales could realize their petroleum potential, whereas during the Mesozoic, the process of hydrocarbon generation occurred throughout the Paleozoic shales, including Carboniferous shales in the areas of their maximal subsidence (Fig.6.17). Carboniferous shales in the Ghadames and Illizi Basins contain gas-prone humic kerogen and have TOC ranging from i to 4%. Because the thickness of Carboniferous shales is considerable in this province (from 50o to 15oo m in its central part), they may be considered as a good source for hydrocarbon generation. Present-day vitrinite reflectance ranges from o.so to o.71% (TTI= 7-11o; Tables 6.6, 6.7). 6.2.3.3
Maturation History and Hydrocarbon Generation in the Illizi Basin Devonian shales in the Illizi Basin have properties similar to those in the Ghadames Basin but are thinner. According to our model, Middle and Late Devonian sources in
Fig. 6.19. Hydrocarbonyields (solid line), rates of hydrocarbon generation (dashed line), and expulsion threshold in the geological history of the Middle and Late Devonian source shales of the Itlizi Basin, Mereksen region
400
300
Time (Ma) 200
1 O0
4"g
~400
o~
I
~: 300
g
illi~lj
U O
0
200
I I I
O
E t00
-
0
6.2 • Applying the Model to Saharan Basins
251
the Mereksen area began to generate liquid hydrocarbons in the Permian (Fig. 6.17). Considering that the average TOC = 4.5% and the average measured S~= 7.5 mg HC / g rock, the residual potential should be H I = 167 mg HC / g TOC. This value differs from the modeled value by 6%. Total hydrocarbons output during the geological history of these source shales is nearly 75% of the initial potential (63o mg HC / g TOE; Fig. 6.19). In the Trias basin, the Devonian shales have low TOC and are not hydrocarbon sources. The initial and present-day TOC contents of the Devonian shales and their distribution in the Illizi and Ghadames Basins differ greatly from those of the Silurian. Late Devonian shales in the Ghadames and Illizi Basins contain the largest quantities of TOC (2-8%),which is considerably higher than in the corresponding Silurian shales (usually 2%); however, the organic carbon content in the Late Devonian shales decreases westward of the Saharan Platform. These variations are probably linked with the change of direction and the quality of detrital material in the Devonian as compared to Silurian shales. The Hoggar massif to the south was the main source of detrital material during the Early Silurian, whereas the Tihemboka-Zarzaitine-Alrar anticlinal system to the east was a dominant detrital source during the Late Silurian and Early Devonian. During the Ardenian orogeny, local highs originated in the Ghadames and Illizi Basins. These highs contributed detrital material to the basins during the Middle and Late Devonian. 6.2.4 Southern and Western Basins
The thermal histories of the Ahnet, Mouydir, Timimoune, Tindouf, Reggane, and other basins were also modeled. The Ahnet, Mouydir, and Timimoune basins had high hydrocarbon generation potential. In the Sbaa sub-basin of the southern Timimoune basin, the initial TOC content of Silurian shales (9%) far exceeds present-day values (3%). This area, characterized by a relatively moderate level of organic matter maturation (Ro = o.9-1.o%), generated substantial oil. Devonian shales, particularly Middle Devonian shales, have hydrocarbon potential similar to the Silurian shales, Silurian source shales (in the Ahnet, Mouydir, and North Timimoune basins) have relatively high TOC (2%) concentrations; however, the maturation level is high (Ro = 1.2-1.6%), and these shales are likely to be generating gas. Measured TOC ranges from 2 to 8% in the Middle Devonian shales, and from 1 to 5% in the Late Devonian shales, but with maximum decreasing concentrations westward from 1.5 to 5.0% in the Illizi Basin, 1.o to 3.5% in the Mouydir basin, and 1.o to 1.8% in the Timimoune basin. These variations are likely linked to changes in transport direction and in the provenances of detritus as compared to Silurian shales. The level of organic matter maturation in the Devonian sources in these basins is higher (Ro = I-4%) than it is in the equivalent sources of the Triassic province (Oued el-Mya, Ghadames, and Trias basins) due to differing sedimentation, burial, and tectonic histories. Consequently, gas generation is currently expected to occur in the south and west of the Sahara, with the exception of the Sbaa sub-basin, which is at a lower maturation level.
25z
Chapter 6 - Burial History and Kinetic Modeling for Hydrocarbon Generation
6.3
Summary and Conclusions Our modeling has shown that, in a number of basins - in particular, the Ghadames, southern Oued el-Mya, and Sbaa - the initial total organic carbon values for the present-day overmature kerogen (which is beyond the hydrocarbon generation maximum) exceed appreciably the present-day average content of residual total organic carbon (5% in the Upper Devonian shales of the Ghadames and Illizi Basins, 2.5% in Silurian shales in the northern Ghadames, and about 2.5-3% in the Silurian shales of the Sbaa sub-basin). Presumably, these areas were abundant as hydrocarbon generators. The Silurian source shales, despite their present occurrence mostly in the gas window, were not heated sufficiently high in the north Oued el-Mya Basin and in certain areas of the Ghadames Basin. The Devonian source shales of the Ghadames Basin also may be mentioned in this connection. This apparent inconsistency is attributable to erosion of a significant part of the Paleozoic sediments as a consequence of the Hercynian uplift. In the Ghadames Basin, Hercynian uplift amplitude and the extent of erosion are appreciably louver than in the Oued el-Mya Basin, and for this reason the apparent disagreement between the present-day temperatures and the organic matter maturation level is less contrasting. In the south and west Illizi Basin, the deepest subsidence occurred prior to the Hercynian uplift. The moderate Hercynian uplift in these areas entailed no substantial temperature drop; consequently, the organic matter maturation proceeded, although at a slower rate. The resultant effect is that the measured maturation level is higher than expected for the present-day temperatures. Consequently, regional average estimation of the initial total organic carbon requires evaluation of the Hercynian uplift amplitude and the extent of erosion of the Paleozoic sediments for each particular area - and their effect on the kerogen maturation. It would be of interest to determine also the original depocenters, because areas with overmature organic matter exhibit a lower content of total organic carbon in consequence to depletion of the major kerogen potential. High geothermal gradients, in the large majority of cases, are associated with zones of uplifted basement, most notably in the Hoggar massif and Ougarta chain adjacent to the Triassic province's southern and western basins. High geothermal gradients increase along the Amguid el-Biod ridge; relatively low geothermal gradients are typical in the northern Oued el-Mya and Ghadames Basins, where thick Mesozoic evaporates were deposited. In the Pateozoic, favorable conditions for hydrocarbon generation and accumulation occurred mainly in the south and in the southwest of the province. As to the central and northern regions, the generation of hydrocarbons would have taken place preferentially in the Mesozoic (Fig. 6.2o). Promising traps or advantageous structures are those located close to the subsidence zones, where the Silurian and Devonian source shales escaped uplifting and, consequently, the thermal pause, as well as the erosion. In particular, the Ghadames and Illizi Basins constituted a favorable province because they were active during the course of both the Paleozoic and Mesozoic. At the end of the Cretaceous, the eastern Sahara started generating gas.
6.3 • Summary and Conclusions
z53
r
Ir Traps formed 11 I1NorthOuedMya li after migration Ii basin (areaBST)
Cenozoic-Mesozoic (J-K) oil (gas) II generation ~.
Traps formed [_before m grat on ~ Berkaoui Ben Kahla EastOued Mya
T Traps buried
Tilrhmet Hassi R'Met
Hercynian orogeny (unconformity)
/ Destroyed U~ fields /J
I~ -Traps-h~y l! /
~
/ Paleozoic (D3-C) 11 lJ oil (gas) generation I~ "~
.
\ \ \
~up fred . t~nnet, ~, • .' especa ~ n-7 yI~~-entre L. in Late rz_ -J North and ~ ~ EastSahara, [] Tramsformed +lnl-o,uye structuraltrend I I . . IJ . . .laDanKort, Amguid-Hassi DeTore mlgratlon Edjelen, ~o~.,A ~ [ ~ ~ Zarzaitin G'"~a~lame~s,
Daharregion
~I " ~ l ~ w I ~ I =!I. Irapslormect ~ ~
I~
[ I T r a p s moderately]] auellala, /flluplifted. especiallyll DraaTamra, [I inLatePZ ]IEIAg reb
Ahone;clirU bYtectonicinversion ~me~Un' - hypsometric denivellation Y - paleostructuresin ~ North-lllizi basin
Fig. 6.20. Generalizedmodelofpossibleoil-gasgeneration,migrationpathwaysandtrapping(specified examples)
Chapter 7
Degree of Preservation of Hydrocarbon Accumulation as Indicated by Carbon Isotope Analysis
In various regions of northeastern Algeria like Djebel Hamra (Ain-Rich Concession), S4tif, Mddjoun6s, Ain-Touta (Batna) and Djidjet indications of petroleum have been observed. Geochemical and tithological studies carried out previously (Makhous 1982) showed clear and convincing evidence for a secondary migration of hydrocarbons located in deposits within Upper Jurassic and Cretaceous (?) rocks. This region is characterized by an intense development of tectonic dislocations which served as pathways for the secondary migration mentioned, during which parts of these accumulations have been destroyed. As an unresolved problem we still have to delineate the extent of this secondary migration and to evaluate the volume of that part of the hydrocarbons that has been transferred through the zones of dislocation and then destroyed on the way. In other words, we have to find out to which degree the accumulations ultimately trapped in the study zone have been preserved. The analysis of the fractionation of the stable carbon isotopes 12C and 13C in carbonate rocks and organic matter extracted from them allows us to delineate the extent of the secondary processes connected to the migration of hydrocarbons and to undertake an attempt at defining the scale of destruction of potential hydrocarbon deposits. An additional objective of this investigation is to present an actualization of the isotope studies and of their applicability to prospecting for hydrocarbons. The importance and reliability of the information gleaned from carbon isotope studies will be outlined, in particular for tracing one of the prime aspects of petroleum research, viz. secondary migration. 7.1
Methods Employed The results of carbon isotope analyses are expressed as a ratio against the so-called PDB, i.e. calcite from a Cretaceous"Belemnitella americana" with a (13C/lZC)PDB= 88.99 (Craig 1957), according to the following formula: 13 12 13 12 ~3C = ( C/ C)sample- ( C/ C)pD B X103 (13C]12C)PDB
The oceans represent a basin connected to the sedimentation of carbonate rocks. If carbonate sedimentation in it takes place in equilibrium, then its carbon will be enriched in 13Cby 4%o against that of the bicarbonate dissolved in the same waters
2.56
Chapter 7. Degreeof Preservation of Hydrocarbon Accumulation
(Keeling 1968). The majority of the carbonate of marine origin exhibits a fi13C ranging from +6 to -3%o (Park 1960). The decomposition of organic matter leads to the liberation of relatively large amounts of"light" CO2 which will be mixed with the bicarbonates dissolved in water, leading to the formation of carbonates with a pronounced negative fi13C.~'Light" carbon in carbonate rocks may result from the oxidation of organic matter, the oxidation of methane, or the supply of the latter from deeper horizons. With such methane is associated the supply of carbon with fi13Cvalues ranging from -10 to -15%o. It has been noted that microbiological reduction of hydrocarbons by sulfate-reducing bacteria will lead to metabolic carbon dioxide enriched in 1~Cby 5-10%o (Thode 1958). As a consequence of the phenomena described above we will obtain "light" carbonates, the volume of which depends directly on the intensity of the reduction of sulfates or the oxidation of organic matter under participation of the migrated hydrocarbons. The presence of thick carbonate beds of chemical derivation enriched in light carbon isotopes serves as an indication of the oxidation by sulfate-reducing bacteria followed over time by the destruction of vast amounts of petroleum. The smaller the volume of such ~light" carbonates, the better the preservation of petroleum. In other words, a weak development of such carbonates will be a good indicator of the preservation of petroleum deposits (Makhous 1973).This criterion was employed by us during prospecting activities for petroleum and native sulfur in the Mesopotamian basin of Syria and Iraq.
7.2 General Data on Carbon Isotope Composition of Sedimentary Rocks (Carbonates) and Organic Matter from Northeastern Algeria 7.2.1 Carbon Isotope Composition of Carbonate Rocks The carbon isotope composition of carbonates at different levels at Djebel Hamra (Ain-Rich Concession), S~tif, M6djoun6s, Ain-Touta (Batna) and Djidjel was studied in detail. The deposits investigated are enclosed in rocks of the Upper Jurassic and Cretaceous at Djebel Hamra (HM-I) and in the Senonian and Paleozoic in the S4tif and M6djoun~s areas (Fig. 7.1). The mean isotopic compositions were calculated for micritic carbonate rocks not affected by secondary transformations for the various regions mentioned. The Senonian and Paleozoic micritic carbonates of the S6tif and M4djoun~s areas exhibit mean 6t3C values of -o.2 and -0.1%o respectively. The mean for the micritic carbonates of the Upper Jurassic at Djebel Hamra is +1.5%o. For the Ain-Touta and Djidjel areas the mean values of S'3C are -0.4 and -o.6%o respectively.
7.2.1.1 The Djebel Homra Region (HM- I) and AYn-Rich Concession We have already referred to the vertical migration of hydrocarbons within the rocks of the lower part of the succession in the Djebel Hamra region (HM-1) between 2 ooo and
7.2. General Data on Carbon Isotope Composition of Rocks and Organic Matter
E
~ ~~
~~
257
o
S
a o ©
o
8 E~
m
~
o
m~
0~.~
0.~
~ 8~o "~
o
o
~8 8~ ~ °8~ Z ~
Z
U
258
Chapter 7 • Degree o f Preservation of Hydrocarbon Accumulation
Table 7.1. Isotope composi-
t/on of carbon in carbonate rocks and in organic matter at Djebal Hamra region (HMq)
Depth (m)
6~C carbonates (%0)
6~3C organic matter (%0)
801.0 802,0 803,2 893.0
0.0 +2.8 +3.2 -46
-22.2
1334,4- 1 335.0 1429.0- 1430.3 1913.4- 1913.8
+4.5 -7.4 +0.6 -5.8
-21.9 -26.2 -22.3 -24.8
2411.0 -2414,0 2 683.2 - 2 686,4 2 753.0 --2 754.2 3 474.0 - 3 475.0
+1.5 + 1.7 +1.5 -6.1
-22,5 -22.2 -25.8
3526.0-3531.0
+2,3
36175 -3621.7 3 940.0 - 3 942.5 4 067.0 - 4 069.0
+1.5 +1.4 +0,9
800.0801,0 802.0892.0-
1493.0 - 1 494.3
-25.6
-24,0
4 069 m. The carbonate carbon of these rocks exhibits a truely homogeneous isotope composition of +o,o to 4.5%0 with the exception of three samples with 6'3C of -4.6, -6.1 and -7,4%o (Table 7.1, Fig. 7.2). Such an isotopic composition is characteristic of marine carbonates. We may thus suggest that during the unidirectional process of hydrocarbon migration the really quite limited recrystallization of the carbonates has not caused a substantial change of the carbonate carbon isotope composition of these rocks, although there are convincing indications for pathways of hydrocarbons in the sediments. In the upper part of the succession in drill hole HMq the carbonate carbon isotope composition fluctuates between well-defined limits of +4.5 and -7.4%0 (Table 7.1, Fig. 7.2). In this case, recrystallization of carbonates and neoformation of calcite or dolomite have taken place under participation of CO 2 which in certain instances contained a large portion of isotopically light carbon (12C)whereas in other instances there was a predominance of hea W carbon (13C). The results obtained, together with other geochemical data, show that in the zones in which the sedimentary rocks have been affected by infiltration waters the migrating hydrocarbons have been subjected to oxidation and other changes, to be discussed in detail below. 7.2.1.2
The $~tif and M~djoun~s Areas For samples taken up-dip of the Sdtif and M4djoun6s areas the 613C values range from +2.9 to -5.8%o (Tables 7.2, 7,3; Fig. 7.2). The "light" carbon makes itself felt in the carbonates where we find traces of recrystallization or of other mineralogical neoformations and in particular the appearance of mixed-layer clay minerals with perfectly ordered crystal structures. In the same samples microfissures are filled by secondary calcite or dolomite with detrital carbonate cement frequently being present between the crystals. In other cases we observe entire fields of neoformed minerals (dolomite
7.2 • G e n e r a l D a t a o n C a r b o n
-~9
-8
-7
~6
-4
-5
Isotope Composition of Rocks -3
-2
-1
+t
~ ~
and Organic Matter
+3
+2
~
+4
~ ~
+6
+5
~
z59
+7
.........
+8 8~C (%o)
DJEBELHAMRA
MEDJOUNES
AINTOUTA (SATNA)
.........DJIDJEL
.........MARINE
CARBONATES
. . . . . . . . . . FRESH~WATER CARBONATES
-9
--8
Fig. 7.2.
-7
q6 _ -S,
-4
~3
52
-I
+I
+2,
+3,
+4,
+5,
+6
+7, ~8
8~C
(%o)
Isotopic composition of carbonate carbon
Table 7.2. Isotope composition of carbon in carbonate rocks and in organic matter
(OM) in the Setif
region
No. of 613Ccarbosample nates (%0)
613Corganic Core(%) matter (%o)
No. of 61~Ccarbosample hates (%0)
6~3Corganic Corn(%) matter (%0)
6a 6b 7 8 9 10a 10c 10d l 1a 11b 11 c 11 d 12 20 21 22 23 24b
-25.01 - 22.2 -25.1 -
25 26c~ 26b 34 35 36 39a 39b 56 57 58 59a 59b 60 61 62 62a 62b
-24.4 -24.1 -
-5.6 2.0 2.0 0,1 2.9 -1.0 1.9 -0.6 -1.3 -1.5 - 1.8 -3.4 -0.9 1.8 2.0 0.1 0.1 -0,3
.-
- 23.8 -24.2
0.82 0.61 0.14 0.9 0.24 0,43 0.31 0.72 0.42 0.66 0.31 0.12
-0,9 -I A -1,2 2,1 1,5 0~1 -2.6 0.4 -0.6 -3.9 -5,2 -4.7 1,3 -1,7 0,5 -0,2 -0,9 -0A
-23.3 -23.8
0.86 1,46 0.75 0,09 0.10 0,09 0.10 0.11 0.12 0.14 0,64 0.80 0.27 0.13 0.18 0.25 025 0,31
260
Chapter 7 .
Degree of Preservation of Hydrocarbon Accumulation
T a b l e 7.3. Isotope composition of carbon in carbonate rocks and in organic matter
(OM) in the re-
gion of M~djounds
No. of 613Ccarbosample nates (%0)
613Corganic Corg(%) matter (%0)
No. of 613Ccarbosample nates {%o)
613c organic matter (%0)
Corg(%)
Io
0.I
-
1.08
30b
-0.3
-
Ib
-0.2
-
0.96
31
- 1.6
-
0.48
2a 2b 2c 3a 3b 4 5 13 14a 14b ] 4c
-3.2 0.8 -0.4 -0.9 0.0 -0,7 -0.4 -2.3 -1.7 -0.9 -0.2
-22.8 -24.4 -24.2 -
0.22 1.13 0.19 0.14 0.10 0.85 0.21 0.71 0.71
41 42a 42b 43 44 46a 46b 47a 47b 48a 48b
-0.7 0.2 0.5 0.1 -0.1 -0.3 -0.3 -2.2 -0,4 -4.5
-25,8 -24.2 -24.6
0.t8 0.15 0,16 0.17 026 0.] 8 0.21 0.29 0.13 0.22 O.12
t 5a
-2.0
-25.1
0.45
48c
0,1
-
0.21
15b 16 17 18 19a 19b 27a 27b 28a 28b 29a 29b 30a
0.6 0.3 -1.0 -0.3 -0.t -1.1 -18,9 0.7 0.3 0.1 -1J 0.3
-25.7 -22.0 -
0.63 0.71 0.22 1.05 0.26 0.20 0.20 0.19 020 0.49 0.26 0.64 0.37
49a 49b 50a
0.8 1,0 0.1 -1,5 2.2 -5.1 -5.4 -5.0 -1.7 -4.0 -5,8 05 2.0
-23.4 -26.3 --24.8 -25.9 -21.9
0.I 0 0.19 0.13 0.16 0.66 0.34 0.42 0.43 0.83 0.69 0.46 0.13 0.32
0.6
50b
63 65 65a 65b 66 67 68 70 71
0.21
o r calcite) w h e r e t h e c e m e n t i t s e l f e x h i b i t s a s e c o n d a r y ( n e o f o r m e d ) s t r u c t u r e . S u c h m i n e r a l o g i c a l n e o f o r m a t i o n s are g e n e r a l l y o b s e r v e d a l o n g t e c t o n i c f r a c t u r e s w h e r e relatively l o w 613C v a l u e s are a r e m a r k a b l e f e a t u r e (Fig. 7.3). The neoformed carbonate minerals are paragenetically associated with organic m a t t e r o f a n e p i g e n e t i c n a t u r e . T h e latter is l o c a t e d e i t h e r in t h e f i s s u r e s o f t h e r o c k o r in m a s s i v e f o r m o v e r t h e fields o f r e c r y s t a l l i z e d c a r b o n a t e s . U n d e r t h e s e c o n d i t i o n s s o m e t i m e s " f r e s h " p y r i t e s is e n c o u n t e r e d . T h e i s o t o p i c c o m p o s i t i o n o f t h e c a r b o n f r o m t h e r e c r y s t a l l i z e d c a r b o n a t e s in t h e s e f i s s u r e s i n v a r i a b l y e x h i b i t s 613C v a l u e s w i t h a p r e d o m i n a n c e o f t h e l i g h t e r i s o t o p e ( < - 2 . o % o ) . T h e i s o t o p e c o m p o s i t i o n o f t h e org a n i c c a r b o n f r o m t h e s a m e s a m p l e s is 2-3%o l i g h t e r t h a n t h e o r g a n i c c a r b o n f r o m m i c r i t i c r o c k s o r f r o m c a r b o n a t e s c o l l e c t e d away f r o m t e c t o n i c f r a c t u r e s , In o u r o p i n i o n , this v e r y case, like t h a t o f t h e u p p e r z o n e s i n t e r s e c t e d b y d r i l l h o l e HM-1, clearly s h o w s t h a t t h e r e h a s b e e n a n o x i d a t i o n o f t h e o r g a n i c m a t t e r t h a t h a s m i g r a t e d in t h e f o r m o f h y d r o c a r b o n s . T h e i n t e r a c t i o n o f t h e CO 2 t h e r e b y f o r m e d w i t h t h e r o c k s r e s u l t e d in a l i g h t e r i s o t o p i c c o m p o s i t i o n o f t h e total c a r b o n a t e c a r b o n . I n g e n e r a l , it c a n b e s a i d t h a t i s o t o p i c a l l y l i g h t c a r b o n a t e s are a s s o c i a t e d w i t h r o c k s rich in e p i g e n e t i c o r g a n i c m a t t e r .
7.2 • General Data on Carbon Isotope Composition of Rocks and Organic Matter
261
o-
o
L
,=
~I
~
262 Table 7.4. Isotope composition of carbonate carbon and organic carbon from the AYn-
Touta and Djidjelregions
Chapter 7 • Degree of Preservation of Hydrocarbon Accumulation
No. of sample
6~3Ccarbonate (%0) 613Corganic (%0)
AT-1 AT-2 AT-3 AT-4 AT-5 DJ-1 DJ-2 DJ-3 DJ4 DJ-5
-0.4 -1.2 -3.0 2.4 -2.8 -0.3 -2.1 -1.6 -0.1 -5.4
-25.4 -25.1 -22.6 -25.4 -23.5 -23.8 -22.8 --25.6
7.2.1.3 The AYn-Touta (Batna) and Djidjel Areas The carbonate carbon isotope compositions of the Mn-Touta (Batna) and Djidjel areas are similar to each other (Table 7.4, Fig. 7.2), with the majority of the 6~3C in the samples being negative. They range from -o.1 to -5.4%o for Djidjel and from +2.4 to -3.o%0 for the A'/n-Touta region (Fig. 7.2). The regularities observed in the preceding regions of Djebel Hamra, S~tif and M~djounbs are encountered also here. The lowest 6~3Cvalues are notably found in recrystallized carbonates (calcite, dolomite) and especially in those associated with epigenetic organic matter. The paragenetic associations (recrystatlized carbonate + epigenetic bitumen) are located in the fissures where they sometimes occupy entire fields which still sometimes contain "islands" of not transformed micritic carbonate. The 6~3Cdata obtained in these two regions are, from a methodical point of view, of particular importance as clear indications of petroleum have been noted in the samples studied, viz. oil in a sample from Mn-Touta and gas in samples from Djidjel. This is proof of a concrete factor for the reliability of the indications furnished by the carbon isotopes for petroleum exploration in northeastern Algeria.
7.2.2 Carbon Isotope Composition of Organic Matter The carbon isotope compositions of organic matter from the five regions studied exhibit a narrow range of -21.9 to -26.3%0 (Tables 7.1, 7.z, 7.3, 7.4; Fig. 7.4), the lightest organic carbon coming from samples taken close to faults (Fig. 7.5). In those samples the organic carbon is a-3%o lighter than the organic carbon from micritic rocks or from carbonates derived from zones removed from the fault zones. There is, furthermore, the following trend: the lightest isotope composition in organic carbon was noted in samples distinguished by a high degree of recrystallization of the carbonate material which concurrently exhibit "lighter" carbonate carbon. A comparison of the 6uC of the carbonates with that of the organic carbon implies a direct link between them. The coefficient of correlation for all data appears to be fairly high, with the mean values being o.93 for Djebel Hamra, o.83 for S6tif, o.89 for M~djoun~s and o.97 for A~n-Touta and Djidjel. There is no correlation between the
7.3 . T h e M e c h a n i s m
(~13 C (O~o) -,32
-31
of Stable Carbon Isotope Fractionation
.30
~29
-28
-27
-26
-25
~63
-24
-23
-22
-21
-20
-19
-18
Isotopic composition of organic carbon Djebel Hamra . . . . . .
~ . ~ m
~
S~tif
~ r~ ~
r~r~r~__
M~cljoun~s
~
r ~
~ r ~
Ain Touta (Batna)
~
~
~
~
Djidjei
r~:
m
~..
_ _ ~ _
.....
. . . . . . . . IS 13
'13C'12C' ~ /~chant.
t
C= 61~C (%0) -32
-3~
-30
49
~ Oil
'~3C'120 / IPDI~
- t
if3CC/'~
3
I
N Carbonate carbon
--x10 2a
~ _ _
-27
-26
~ Organiccarbon -2s
-24
-2s
-22
-21
-20
-19
-18
Fig. 7.4. Variations in carbon isotope compositions
content of organic carbon and the isotopic composition either of the organic carbon or of the carbonate carbon. The data described above are further confirmation of the secondary migration of petroleum-type hydrocarbons across tectonic dislocations where oxidation of the latter led to secondary carbonates with varying degrees of "lighter" isotopic compositions ((Makhous 1979). As in the case of the carbonate carbon, the low variability of the absolute isotope values of the organic carbon (Fig. 7.4) and the low abundance of the lighter isotope over the area (Fig. 7.4, 7.5) indicate a rather limited migration of hydrocarbons through the succession studied as well as their ultimate destruction.
7.3 The Mechanism of Stable Carbon Isotope Fractionation (12C VS. 13C) and Regularities in Their Distribution in Jurassic and Cretaceous Deposits of Northeastern Algeria 7.3.1 Decomposition and Oxidation of Organic Matter As already pointed out, the available data indicate a rather narrow variation of (~I}C as well as a lower presence of 12C in the study area (Figs. 7.2, 7.6, 7.4, 7.5). This implies a mechanism of formation of 613C controlled by an equilibrium of the respective isotopes, viz. ~2C and ~3C of different origin which were encountered in the recrystallizational dnvironment of the carbonates. The carbon dioxide (C02) had to be formed through destruction of the organic matter, a process taking place during bacterial activity which produce CO2 by assimilation of organic matter and in particular of petroleum. Under aerobic conditions the
264
Chapter 7 • Degree of Preservation of Hydrocarbon Accumulation
L~
o
'/
~o
.Q
~6
o
r~ u~
o =
,¢ o
o
7.3 . The Mechanism of Stable Carbon Isotope Fractionation
a65
organic matter can be oxidized by free oxygen to furnish carbonic acid. The process of diagenesis of organic matter can take place along two distinctive pathways: 1. Decomposition: CxHyO~
> CO~ + C~_~HyO~_~ ,
z. Oxidation: C~HyOz+ O~
> C02 + Cx_lHyOz •
The intermolecular isotopic heterogeneity actually controls the differences in the formation of the isotopic composition of the carbon in the C02 resulting from the decomposition of organic matter (1) or from oxidation (2). The carbonic acid produced by the decomposition of organic matter and inheriting the carbon of the functional groups will be enriched in the heavier isotope because of the preferred concentration of ~3C in such groups. However, the carbonic acid formed by oxidation of organic matter possesses a carbon isotope composition corresponding to the mean isotope composition of the organic matter oxidized, i.e. it is characterized by the predominance of ~C over 13C. Process (I) takes place under anaerobic conditions and process (2) under aerobic conditions. It is obvious that under the specific geochemical conditions of northeastern Algeria, i.e. an oxidizing surface environment or penetration of atmospheric agents along faults, (aerobic) oxidation of hydrocarbons inevitably wilt lead to the formation of CO2 enriched with 1~C. The same carbonic gas with light carbon will, on encountering surrounding micrite carbonates, the isotope composition of which is characterized by a predominance of ~3C,produce a sort of leaching effect on the latter. We thus will have in solution bicarbonate of organic derivation together with carbonate ions from the dissolved initial carbonates: Ca~3CO3+ 1~CO~+ ttzO ------7Ca z+ + H13C03 + H~2C03 , CaMg(13CO~)2 + 21zC02 + 2H~O ..... > Ca -,++ Mg~+ + 2H13C0; + 2H12C0~ . The secondary carbonate recrystallized from these solutions may inherit to the same extent the isotope composition of the primary micrite carbonate as well as that of the oxidized organic matter. It is evident that the isotopic composition of the secondary carbonate will be controlled by the ratio of H12C03 to H13CO3in solution and thus by the degree of oxidation of the organic matter. This mechanism for the formation of a certain 613C is in perfect agreement with the results obtained in our study areas. However, the 613Cvalues observed in the lower portions of drill hole HM-1 (2 000-4 o69 m) as well as in certain isolated instances require further clarification. These essentially positive values could not have been established under participation of C02 resulting from the oxidation of organic matter as this should have led to a more or less pronounced predominance of 12C. Most of the samples collected over this interval and from certain other regions represent only non-transformed micrite carbonates. XRD analyses and microscopic observations only reveal very poorly crystallized limestones, the carbon isotope composition of which corresponds to that of carbonate rocks of marine origin. Even here,
266
Chapter 7 • D e g r e e
o f Preservation o f H y d r o c a r b o n A c c u m u l a t i o n
however, one sometimes observes fissures or fields filled with well-crystallized calcite or dolomite and associated with bitumens. The 613Cvalues of such neoformations appear to be rather elevated. This anomaly of positive values may be explained by the leaching of primary micritic rocks by CO~ derived from the decomposition of organic matter under anaerobic conditions by bacterial action. This process, as has been described above, leads to CO2 which inherits the '3C-enriched carbon of the functional groups.
7.3.2 Methanogenic Fermentation Carbon dioxide enriched with 13Cmay also form together with methane during methanogenic fermentation: CxHyOz ..... > COz + CH 4 + Cx_2gy_4Oz_ 2
Taking into account the biogenic isotope equilibrium, the COa and C H 4 resulting from the metabolisms should obtain different isotopic signatures, CO2 being enriched with I3C and CH4 with 12C. Traces of methane have actually been detected in the pores of certain samples from HM-1 with the aid of mass spectrometry of gas inclusions (Table 7.5). This CH 4 exhibits a rather light carbon isotope composition and could even have participated in the establishment of the 313C of the recrystallized carbonates in the upper parts of HM-L However, the observed traces of C H 4 are associated with greater concentrations of hydrogen (Ha) and nitrogen (N~) which would imply a primary origin of the methane (Table 7.5). The paragenetic association of CH4, H a and N~ is actually characteristic of brines of organic derivation. The joint occurrence of CHa, H~ and Na may be considered as supplementary evidence for the existence of secondary migration of hydrocarbons in the respective area.
7.3.3 Sulfate Reduction by Bacterial Activity We shall now discuss the rather interesting unique case of sample 27a which is made up of recrystallized isotopically very light carbonate. In the respective western part of
Table7.5. Composition (%) of gaseous phases occluded in rocks of drill hole HM-1
D e p t h (m) 1334.4- 1335.0
H2
N2
-
64.7
35.3
-
100.0
50
40
1429.0- 1430,3
Traces
2683.2 - 2686.4
10
3506.0- 3571.0
57
4067.0 - 4068.0
-
29 Traces
CO 2
Traces (very little gas)
CH a
14
7.4 - Conclusions
267
the M~djoun~s structure the carbonate carbon exhibits 313C of -18.9%o. Such a 8~3Cvalue corresponds to the reduction of sulfates like gypsum and anhydrite where the activity of sulfate-reducing bacteria, in the presence of hydrocarbons, forms H~S: CaSO 4 + 2H20 + CnHn+2
> CaCO3 + CO2 + Ha0 + H2S
In continuing the reaction, this hydrogen sulfide will be oxidized by oxidative infiltrating waters to result eventually in elemental (native) sulfur according to the following formula: 2H2S + O2
> 2S + 2H20
It is during this reaction corresponding to the model of sulfate reduction and oxidation of organic matter that the very"light" calcite is formed. It is clear that such a calcite wilt only inherit the isotopicatly light carbon coming from the hydrocarbons in the absence of micritic 13C-rich carbonates. Such examples of the association of isotopically very light calcite with native sulfur are also known from regions like the USA, Mexico, and the Middle East (Makhous 1974). To sum up, irrespective of the mechanisms leading to the isotopic compositions in the regions studied, the variations of (organic or mineral) 813C are very narrow. Furthermore, the abundance of the light isotopes is quite low (Tables 7.1, 7.2, 7.3, 7.4; Figs. 7.2, 7.4) and is restricted to the zones of secondary hydrocarbon migration. These observations may serve as indicators of a rather weak oxidation (destruction) of the final hydrocarbon accumulations in the province concerned. 7.4
Conclusions 1. The peculiarities of the carbon isotope composition in the study area are explained by a model of the formation of"light" carbonates by the intrusion of organic carbon consisting essentially of the light isotope ~2C.This intrusion took place after the oxidation of migrating hydrocarbons under aerobic conditions or under the influence of sulfate-reducing bacteria. Irrespective of the oxidative mechanism, the carbonic acid formed inherited the light isotope from the carbon of organic derivation. The final carbon isotope composition of the recrystaltized carbonates (calcite, dolomite) is controlled by the ratio in the solution of the H12CO3- ions derived from the oxidation of hydrocarbons to the H ~ 3 C O 3 - ions derived from the leaching of primary micritic carbonates enriched with 13C.Consequently, the carbon isotope composition (8~3C) and its variations may be used as a direct indicator of degree and even the extent of the oxidation or rather destruction of the hydrocarbon accumulations. 2. The fiuC values observed for the different regions indicate a lower abundance of the light isotope 12C on surface as well as at depth. Furthermore, the variations of the absolute 813Cvalues are also rather limited, the latter ranging from -7.5 to +4.5%o for the carbonate carbon and from -27 to -26.3%o for organic carbon. These data, on the whole, show that there is only a small volume of"light" carbonates which actually is in agreement with the insignificant amounts of petroleum destroyed
z68
3-
4.
5.
6.
7.
Chapter 7 . Degree of Preservation of Hydrocarbon Accumulation
during its secondary migration from accumulations (or deposits) located in deeper underlying strata. In other words, the real hydrocarbon accumulations should have been well preserved. This conclusion may be considered as valid and may serve as an argument in favour of exploration for hydrocarbons in northeastern Algeria. A great similarity in the carbon isotope composition of 613C in samples collected in the different areas of northeastern Algeria has been noted. Taking into account this observation and the similarity of the geological situations in these areas, we may assume a uniformity of the geochemical processes which took place there and of the secondary migration, and, consequently, the existence of one large petroliferous province. In view of the great uniformity of the operation of the processes of migration and oxidation of organic fluids, relatively negligible variations in the carbon isotope composition should be expected. Any such variations are controlled by a variety of factors such as local geological and tectonic peculiarities, different quantities of organic fluids available, the mechanisms of oxidation, pH of the environment, etc. A regular trend for a heavier carbon isotope composition (organic and mineral) has been observed with increasing distance from organic accumulations and zones of hydrocarbon migration. This regularity underlines the fact that the presence of "light" carbonates due to oxidation of hydrocarbons may be considered as evidence, in addition to the geochemical and mineralogical criteria established previously by Makhous (1982),of the confirmation of zones of secondary migration. The greatest concentrations of recrystallized carbonates with the light isotope 12C directly indicate the zones of secondary migration themselves. The probablity of finding such carbonates with the light isotope 12C is much higher than observing traces of hydrocarbons or other paragenetic formations, considering the higher stability of the recrystallized carbonates. From the data obtained we may recommend the use of a 6~3Cvalue of -2%o or still lower in the carbonate carbon as an indicator of accumulations (deposits) of hydrocarbons in the carbonate rocks of northern Algeria. The exceptionally "light" carbonates (6~3C= 18.9%o) in the M6djounbs structure result from sulfate reduction by sulfate-reducing bacteria in the presence of hydrocarbons, leading eventually to the formation of elemental (native) sulfur. In view of the presence of thick layers of gypsum (CaSO4 x 2H20) and anhydrite (CaSO4) in the succession at the above-mentioned site (sample 27a), a potential for the presence of a native sulfur deposit is indicated. The practical importance of carbon isotope analyses lies in the possibility to distinguish carbonates formed (or transformed) during migration and oxidation of hydrocarbons from sedimentary or other carbonates even in the absence of real mineralogical differences.
Chapter8
Reconstruction of Temperatures from Organic and Mineral Diagenetic Criteria
8.1
Reconstruction of Temperatures from Degree of Structural Ordering in Mixed-Layer Minerals The mixed-layer minerals of the illite-montmorillonite type (I/M) in the shales as well as in the sandstones are characterized by an increase of the illite content with depth. It has been noted that the I/M mixed-layer minerals in shallow-buried sandstones contain much less illitic layers than the I/M coming from the shales. This difference in the illite content between sandstones and shales is controlled by the fact that the I/M in the sandstones are largely of authigenic origin and formed as cement in equilibrium with the physico-chemical conditions prevailing in the porous rock during their precipitation (Clauer et al. 1992,1994). As a consequence, the variation in the primary composition of the I/M is rather limited. It may be modified later during diagenesis when the environmental conditions change. In the shales the I/M are of detrital derivation and thus possess a varied composition. During burial and the concomitant increase of temperature the compositions of I/M in shales and sandstones become increasingly similar to each other. Ktibler (1993) has pointed out that the transformation of the swelling layers in mixed-layer clay minerals is controlled by reaction kinetics and permeability and thus an increase in temperature will speed up the reaction. An important role in the illitization of smectite is attributed to the compaction which reduces the permeability of the argillaceous and silty-argillaceous intercalations (Kiibler 1984). The proportion of the illitic layers in the I/M increases with depth and the unordered phases are transformed into ordered phases of the allevardite type (Fig. 8.1). This transition is easily recognized in diffractograms of samples saturated with ethylene-glycol by the disappearance of the 17-A peak and the appearance of a peak at 13-14 A indicative of a structural ordering in short chains (Fig. 8.2). Such a phase transformation has been initiated at various levels in the Oued el-Mya Basin at depths of 2 3oo-2 5oo m. The transformation of I/M of the allevardite type into those of the kalkbergite type with >8o% illitic layers in the lattice takes place at a depth of 4.3-4.8 km in the Paleozoic sediments. Corrensite, a mixed-layer mineral of the chlorite-montmorillonite type with an ordered structure, occurs at several levels within the Triassic Basin and in particular in the area of the Hassi R'Mel deposit (Plate 15). Corrensite is a highly useful geothermal indicator in sediments (Porrenga 1967; Ktibler 1973). In the area mentioned it starts to appear at a depth of 2.1 km and remains stable down to 2.3 km. The maximum temperatures reached were reconstructed on the basis of the appearance or disappearance of allevardite, kalkbergite and corrensite mixed-layer minerals (Fig. 8.2). Mineralogical and crystallochemical analyses of mixed-layer clay minerals reveal the pro-
270
Chapter 8 - Reconstruction o f Temperatures
<> Unordered phase <> <~ 4>
~ Well-ordered crystall structure
,..
.q.'.'.-'..'"
<> -oo-
C
~
-
x-
_x_.-x-,~ ~,,c~.x. * -x-_K_ ~ _
•c- -c>
-~-
~-
Illite degraded, pooriy crystallised 1 1O0
80
60 40 Smectite layers (%)
20
0
20
40 60 Illite layers (%)
80
Illite agraded, moderately crystallisect i
1O0
9
Illite degraded, poody crystallised
8 7 6 5 Half-width of peak at 10 A (mm)
4
3
Fig. 8.1 a. Proportion ofillite-smectite layers in crystal structure of mixed-layerminerals (
gressive structural ordering and an increase in the proportion of illitic and/or chloritic layers with depth and temperature (Huang et al. 1993) which enable us to use these criteria as geothermometers and for reconstructions of diagenesis. The temperatures reconstructed on the basis of crystallochemical peculiarities of the mixed-layer clay minerals like the appearance of kalkbergite and of other clay minerals (appearance of chlorite IIb) together with the reflectance of vitrinite show that certain Paleozoic petroliferous rocks of the Oued el-Mya Basin reached temperatures of up to 18o °C at a depth of 5.5 km. The uncorrected temperatures according to data on the diagraphy at this depth were around 14o-15o °C at a thermal gradient of 32 °C km -1. This shows that the maximum temperature was 30-40 °C higher than the actual uncorrected temperature. This obvious lowering of the temperature may be explained by the erosion of 1.5-1.7 km of the upper part of the Paleozoic succession, a conclusion that was confirmed by our kinetic and chemical modeling of the basin (cf. Chapter 6). The estimate of the amount of erosion also takes into account the R e which on surface attains a value of 0.22% with the thermal gradient remaining unchanged since maximum burial. The main uplift of the complex took place after maximum burial, the temperature indicating somewhere around the end of the Cretaceous and the start of the Paleogene. The slope of the paleotemperature curves constructed from the mineralogy of the (mixed-layer) indicator minerals is in good agreement with the present geothermal gradient at depths above 2.3 km (Fig. 8.2). The slight difference in slope is probably the result of the fact that the reaction of the I/M starts within the succession at slightly lower depths.
8.2 • Crystallographic Features of Clay Minerals as Thermal Indicators in Petroleum Geology
aT~
Temperature (°C) 1oo
Fig. 8.2. Paleotemperature gradients as determined by clay mineral geothermometers and present-day thermal gradient from uncorrected borehole temperatures. Claymineral temperatures were determined from changes in ordering and composition of mixed-layer illitemontmorillonite (I/M) clay and the appearance of corrensite. (Triassic province, except pyrophillite which was largely found in deeply buried Lower Paleozoic shales of Timimoune, Reggane and Tindouf basins). Depths are only of indicative signification
2~
\ \
\
(Randomly interstratified I/M, I-oh) I-I~ I/M compositionin sandstone I /-o-,/Mcompositioninshales I ~D/-~- C...... it. . . . position in sandstoneI ~5~ /~t~ Ch. . . . itei . . . . dstone I
\ ' X -!~ '~ -o-~ ~
*k-C> C ~
"~I~_a_-Q-~cL~- -0- _%-0r r e~, n s i t
e
~.~,%'~C~?
-o=o .
-~3
(Short-range o r d e r &
.-~TclayJ~_ ~ -o- ~
. . . .
\
\\-I -o-
(Long-range ordering) >85% illite layers in I/M
>300°C
Pyrophyllite 20 20 l
40
60
80
I00
40
60
80
100
n
r
I
Illite layers in I/M (%)
Chlorite layers in Ch/M (%)
- -
8.2
Crystallographic Features of Clay Minerals as Thermal Indicators in Petroleum Geology The kaolinite mineral species studied are kaolinite, kaolinite d (disordered kaolinite), dickite and nacrite. These polytypes have been described by Bailey (1963) on the basis of sense and degree of displacement of 1 : 1 layers and the position of vacant octahedral positions in the layer sequence. For the hydrated kaolinitic minerals, we have used the terminology of Keller and Johns (1976) which is based on endellite as the completely hydrated species and hatloysite as the partly or completely dehydrated species. The potytypes of chlorite have been described by Bailey and Brown (1962) and Hayes (197o). In Fig. 8.3b it is shown that montmorillonite, the mixed-layer clays and illite are located between pyrophyllite without interfoliar charge and the dioctahedral
272
Chapter 8
. Reconstruction of Temperatures
8.2 - Crystallographic Features of Clay Minerals as Thermal Indicators in Petroleum Geology
273
mica with a deficit of 1.o equivalent in the interfotiar charge for each 01o(0H)2radical (Lanson and Champion 1991). The montmorillonites possess a charge deficit in the range of 0.2-0.4 equivalents for each 01o(0H)2 radical in a layered structure of type 2 : 1,leading thereby to a swelling of the structure. Illite exhibits a charge of 0.8 equivalent for each 01o(0H)2 radical, a charge that is too high to provoke swelling of the structure when potassium occupies an interfoliar position. Illite thus resembles muscovite. Anyhow, two types of the mixed-layer minerals illite-montmorillonite may be distinguished in particular: (1) the allevardite (IMIMIMIM...) with an ordering of short chains usually encountered when the clays contain less than swelling layers (of montmorillonite) and (2) the kalkbergite (IMIIIMIII...) with ordering in long chains. Corrensite is characterized by an ordered alternation of the mixed layers of chlorite and montmorillonite.
Argillaceous Rocks. As a general rule, solid mineralogical criteria for the increasing degree of mineral maturity include: (1) the transformation of montmorillonite to illite via a sequence of mixed-layer minerals of the montmorillonite-iUite type, (2) the appearance of chlorite, and (3) the disappearance of potassic feldspars by decomposition. These mineral transformations may be described by the following reactions: montmorillonite + K-feldspar
) illite + chlorite + quartz
The proportion of illitic layers in the structural series illite-montmorillonite is the most sensitive indicator of the metamorphic degree shales. Furthermore, the illitemontmorillonite is characterized by a series of mineral species starting from disordered alternation, passing through ordering in short chains and then to ordering in long chains in the illite and montmorillonite layers. The transformation of illite to dioctahedral mica represents the culmination in the metamorphism of pelitic rocks prior to the green schist stage (Eber11993). In Fig. 8.3a we present the general mineral transformations observed in the Saharan basins as a function of the degree of metamorphism resulting from the burial of their argillaceous rocks.
Silty-Sandy Rocks. The diagenetic associations of clay minerals in the sandstones are more varied than in the shales which is certainiy due to the much greater permeability of the sandstones in comparison with that in the shales. The chemistry of the interstitial waters in the shales is largely controlled by the composition of their solid mineral matter and in particular by the decomposition of unstable detrital material, by the (potential) existence of solutions which are in equilibrium with the stable phases and with filtration processes through a membrance (Berry and Hanshaw 196o). The high permeability of the sandstones leads to a regime in which the interaction of the solution with the solid phase is of importance, the phases forming here being largely determined by the composition of the solution. The clay minerals encountered in the Saharan reservoir rocks and formed during diagenesis from the solutions include
Plate 15. Claymineralsof the chloritegroup,a, b, g, h Authigenicchlorite;c, d corrensite;e,f chamosite
274
Chapter 8 • Reconstruction
Fig. 8 . 3 . a Correlation of the temperature-dependant clay mineral assemblages in shales and sandstones, Saharan basins; b Distribution of montmorillonites, illites and mixed layers (I/M) within the compositional triangle pyrophytlite muscovite - celadonite (glauconite)
of Temperatures
i-- Montmorillonite .............
+
I-- (I/M) Random ~--+ ( I / M ) A l l e v a r d i t e
"~
--,--.--~ --. ( I / M ) K a l k b e r g .-- i
~--.lllite (IM)
2M Mica
Pyrophyllite Chlorite
t~ .....
~
..... ........ +. . . . . .
............
--
.... ~
Kaoline
K-Feldspar
~- - - 1 b d C h l o r i t e
......
~
~ ...... lib Chlorite
..... ................
Halloysite/kaolin
~ ......
~. . . . . . ,
t 0
d
Kaolinite r.....
-- + --- ~
lb Chlorite
......
+. . . . E. . . . . . .
m
Chamosite Dickite/nacrite Corrensite
IM lllite
a 0
,
T
50
100
T - -
150
+
i
-r-
200
250
300
Temperature
b
350
(°C)
Muscovite
3 Celadonite (glauconite)
Pyrophyllite
chlorite lb, ta dickite/nacrite and IM-illite (Fig. 8.3, 8.4; Table 8.1). The clay minerals identified as products of the metastabte decomposition of older minerals include kaolinite (from feldspars), corrensite (from ferromagnesian minerals), altevardite, kalkbergite and illite (from montmorillonite). From studies of Paleozoic sediments of the Russian Platform, Shutov et al. (197o) developed an hypothesis of the progressmg transformation of mineral phases of the kaolinite type with increasing subsidence. They concluded that the kaolinite mineral in little-compacted sediments is kaolinite d.
8,2 - C r y s t a l l o g r a p h i c
Features o f Clay M i n e r a l s as T h e r m a l I n d i c a t o r s in P e t r o l e u m G e o l o g y
~
c
°°.~:~ ~ ~
.~_~
~.~-~
.~_~7,:~
o
~E
u o
275
u Ill o ~ m c
~
8= ~_
o
o~
--
~
~
×~= -~.~
~.7
-~
E
g.
~
o
-~ ~ o ~
v ~.~
~_~.
k~ o
E
~
~ "E ~ .~
6 °
° u
~,
o
o
_
~.~
~
o
~-~
°
_i
',P
o,
I~
n
,., ,=,~
I
i i
>.
--
~
o
E % O~ ~,~ 0
~-~
+~ ~ - o ~ ' o v
~01
~'~_ O ~
,7+~7 " ~ '--- ~
-~ - 'o
o ~
. . . .
~Oo
,~
\.1
I
~,~
~'_--_--~¢ \ ,
-
.
b
"--
~
\
.~7 1
.....
\
/
/
,.,.=
~- u O ~" -
+~o
~ "" ;_~
,.~_~ :'~'0 "~
=
-o
.~
I ,I
"~~/
l~i-=~\i
t
t
0 u "~k...,A..~ v~
~-I"-1
o
I
-~oo
, o<~
~+"
~
u.-
o
~_~
E
~ u
t
\ ', I
~'-
o
0.~
~
" 1'-//°
O~
,
0
~"
+~
' •
\
,'/
~
/\
!
I
8o-0
°ill
276
Chapter 8
. Reconstruction o f Temperatures
Table 8.1. Structural features of Saharan day minerats
Mineral
Crystalline structure
Polytype
Parameter of crystalline celt
a Illite
Dioctahedric
1M
A
5.19
b 8.99
±0.02 Illite
Dioctahedric
2M~
Celadonite
Dioctahedfic
Chlorite
Trioctahedric
Ib
5.36
Chlorite
Trioctahedric
lib
5.35
5,18
_+0.03
•+0.02
5.36
Chamosite
TrioctahedrJc
5
Tficlinic
Dickil:e
Triclinic
2M~ (2M2)
_+0.06 +_0.06
90°00 ' _+0,03
+_0.02
_+0.04
106°50 , +_0.04
_+0.01
7.10 _+0.02
a=91°70 ' b = 104°50 '
7.05 _+0.04 7,05
8.90 _+0.01
5.16
8.96
_+0.02
+_0.05
96°8 , _+0.2
9.29
5.14
_+0.1
95°5' !0.05
_+0,02
•+0.01 Kaolinite
_+0.03
28.32
9.28
5.36
_+0.05 _+0.04 10.20 100°5,
_+0.02
+_0.02
.+_0.01 95°30 ,
14.20
9.30 _+0.02
C
] 01 °30'
_+0.01
9.30
fl
±0.02 20,20
9.02 _+0.01
Trioctahedric
A
10.31
8.98
5.21
c
.+_0.02
_+0.03
Chamosite
A
t4.25 _+0.01
96°50 ,
_+0.03
+_0.5
Subsidence leads to an ordering of the structure of this kaolinite d followed by the transformation to dickite or nacrite. According to these authors, nacrite is a"constraining" mineral forming in the same catagenetic stages as dickite, but as a vein-type mineral frequently encountered together with slickenside-minerals during brittle deformation. The respective evolutionary sequence is: kaolinite (detrital)
>kaolinite d ( in slightly compacted sediments) >kaolinite (in shallow-buried sediments) >dickite/nacrite (in deeply buried sediments).
In this sequence the authors ignored halloysite/endellite, important to dominant kaolinite minerals resulting from the alteration of feldspars in recent sediments. We propose the following diagenetic sequence: halloysite/kaolinite d
> kaolinite
> dickite/nacrite.
For the minerals of the chlorite group we arrive at the following sequence: 1. The species prior to the green schist stage are all chlorites of type I. z. Chlorite of type Ibd (Ib with disordered structure) may not be stabte above 7o-8o °C. 3- The transformation of type Ib into IIb catagenesis is contemporaneous with the lithological transformation of the pelites, i.e. it takes place in the sequence: shales >slates and illite >dioctahedral mica of type 2M (zoo °C) (Figs. 8.3, 8.4).
8.3 - Summary
277
Corrensite is present in several genetic types as described by Kiibler (1973). It has been observed in the Triassic dolomitic limestones associated with evaporites and in volcanic sandstones which experienced some burial. Correlations with other data like Ro or vitrinite reflectance and the temperatures measured in the holes show that the transformation of volcanic ferromagnesian minerals into corrensite takes place at a temperature of 9o-loo °C at which the haphazardly arranged layers of montmorillonite in the chlorite-montmorillonite structure largely disappear. Corrensite is everywhere encountered in complexes, the temperatures of which, as measured in the boreholes, went up to 148 °C. The results of our studies of the transformation of clay minerals and of the correlation between temperature and the different diagenetic minerals have been compiled as a model in Fig. 8.4. We want to stress that the differences in the nature of the monotonous mineral complexes in the shales and of the mineral associations in the sandstones are controlled to a great extent by the composition of the interstitial waters and, in the case of corrensite, they are indicators for detrital ferromagnesian minerals. Furthermore, the absence of a certain mineral does not imply that the respective sandstone has not been subjected to the corresponding diagenetic stage during which this mineral would have formed. In the sandstones it is the presence of such an indicator mineral that is of interest. 8.3
Summary The clay mineral associations may be used as indicators of temperature and degree of metamorphism reached by their enclosing rocks (Ktibler 1964, 1973). The chemical composition of the rock and of the interstitial fluid together with the mineralogy of the terrigenous rocks also furnishes characteristic indicators for the determination of the mineral associations formed at these temperatures. The association of the various clay minerals is also a function of time and of the duration of the reactions in the case of lower temperatures (Clauer et al. 1994). The mineral species most useful as geothermometers are: 1. in shales: illite-montmorillonite, allevardite, kalkbergite, illite, pyrophyllite and chlorite; 2. in sandstones: chlorite (chamosite), chlorite-montmorillonite (corrensite), dickite and illite; 3. in volcanic rocks: chlorite-montmorillonite and zeolites. In pelitic sediments the transformation of montmorillonite into illite and its subsequent recrystallizition into type IM-muscovite facilitates the subdivision of the respective rocks into zones of different temperature ranges. The mineralogy of the shales, i.e. appearance of certain structural polytypes, crystalline peculiarities, or ordering in mixed-layer clays, as well as vitrinite reflectance Ro in several parts of the Oued el-Mya basin (Allal and Hassi Messaoud-Agreb domes) and even in the Ghadames Basin confirm that the paleotemperatures were higher than the present ones which were determined in boreholes in Paleozoic sediments of the domes mentioned.
Chapter 9
General Conclusions
1. The porosity of many of the large hydrocarbon reservoirs is generally of a secondary nature, a fact that has become known only fairly recently. The favourable situation in this context is the formation of this secondary porosity prior to the hydrocarbon migration. It may be reduced considerably, but at greater depths it will be much better preserved than the primary porosity. 2. Taking into account solubility and stability boundaries of organic acids and of carbonic acid we may conclude that the organic acids are more efficient on a local scale for the dissolution of carbonate and silica cements. Where the source of the latter, i.e. the carbonic acid, is deep and further removed they may not increase the porosity by solution and removal of cement. The action of carbonic acid formed by decarboxylation on carbonate cement leads to the enrichment, in the dissolved carbonate, of the light carbon isotope inherited from the organic matter whereas the reaction of organic acids with the carbonate cement takes place under dissociation with the formation of secondary carbonate enriched with the heavy carbon isotope inherited from carbonates of mineral origin. 3. As a result of the activity of these acids enormous quantities of carbonates are transported upwards from diagenetically mature sandstones to become deposited in immature to semi-mature sandstones. Taking into account continued burial of terrigenous deposits, the carbonate cement is subjected to cyclic transformation and to upward transport, leading to the cementation of immature sandstones (at shallower depths) by carbonates and to their blockage. The primary migration of hydrocarbons generally follows on the formation of secondary porosity as during maturation of the organic matter the main phase of hydrocarbon generation takes place after the culmination of decarboxylation. This close association of hydrocarbon sources and reservoirs in time and space favours the accumulation of hydrocarbons in reservoir rocks with secondary porosity. 4. A balance calculation of secondary silica (determined with the aid of cathodoluminescence) indicates that in a number of sandstones more quartz is mobilized by pressure solution than exists actually in the form of silica cement. Intergranular pressure solution thus represents an important agent for the transfer of quartz in the sandstones. Some of the Saharan sandstones studied behave as silica "importers" during early stages of diagenesis and as "exporters" during the later stages. 5. Petrographic observations indicate that the largest portion of the silica cement could have been deposited prior to the main phase of pressure solution. In these cases the presence of early cements inhibited intergranular pressure solution and facilitated the preservation of a relatively large volume of the porosity. Intergranular
280
6.
7.
8.
9.
Chapter 9 - General Conclusions
porosity of sandstones refers to the part of the primary intergranular volume destroyed by the process of (mechanical and chemical) compaction as well as to the part of the intergranular volume remaining filled by the cement. A new mechanism for sandstone diagenesis is proposed here, a concept that implies that the transformation of silica in the sandstones takes place in its majority in the solid phase. It has been shown that dense monolithic sandstones as well as quartzites may form through development of low-energy crystal faces on the quartz grains and by twinning of crystals of the lapan- or Dauphin6-Brazilian type. This process is accompanied by the gradual reduction of porosity and permeability up to total fusion of the grains into a dense quartzitic rock. Aggregation of grains takes place by diffusion, sliding along screwn dislocations and face-to-face movement of large angles boundaries. This type of transformation of the rock in the solid phase is initiated by the reduction of the surface energy and in an open geological system the process of self-organization described above represents a comprehensive alternative mechanism to pressure solution. We were able to distinguish three phases of decompaction of an undulating nature overlapping in time and space: the first one corresponds to leaching mainly of the carbonate cement in the interval between 1.5 and 3.8 km depth and at 5o-15o °C, the second one entails leaching of carbonate, silica and alumosilicate cements between 2 and 3.5 km depth at temperatures of 8o-12o °C, and eventually the third one with intergranular pressure solution followed by removal of silica in an alkaline environment between 2.5 and 4.8 km depth and lOO-17o °C. On a geological scale these undulating phases of decompaction overlap each other. The main factors controlling the processes of compaction-decompaction of the reservoirs in the Saharan basins are: (a) development of secondary porosity, (b) establishment of an abnormal elevated formation pressure associated with the appearance of fractures, (c) formation of growth rims on quartz prior to the Mesozoic subsidence consolidating the matrix of the sandstones and making them resistant against ultimate compaction, (d) presence of Mesozoic evaporites leading to weak heating of the deposits below salts as a result of the low thermo-insulating properties of the salts and the low gravitational pressure exerted by them because of their low density, (e) early migration of the hydrocarbons which slows down and even stops the compaction of the reservoirs, (f) temperature and pressure, (g) thickness of the sandstones, (h) transformation of structural and textural features of the cement, and in particular of the argillaceous cement, and (i) tectonic processes and fracturing. The isotopic "portraits" (of carbon) observed in oil and bituminoid fractions extracted from Silurian and Devonian source rocks in the Triassic province are of special importance for correlations between oils and source rocks. Two different shapes of isotope curves have been observed in asphaltenes with light isotopes which, as shown before, are characteristic of marine source rocks. All groups of oils studied may be subdivided isotopically into two main groups which differ clearly in the isotopic composition of the five fractions of different polarity. The first group of oils closely reproduces the peak-shaped isotope curves characteristic of bituminoid fractions of Silurian age in which we also note a light isotopic composition of the asphaltenes and a narrow spread of ~'3C values. The isotopic"portraits" of the bituminoids from Devonian shales are notably different from those of the
GenerM Conclusions
281
Silurian source rocks and correspond to the isotopic "portraits" of the second group of oils. This is an important criterion in the search for oil. lo.The geochemical data on the argillaceous source rocks in organic matter of the Silurian and Devonian formations of the Triassic Basin and the Ghadames, Oued el-Mya, Illizi, Sbaa Basins, etc. reveal good correlations with the oils trapped in them. The correlation of the carbon isotope composition of the different oil fractions with the same fractions in bituminoids of potential source rock candidates, the distribution of saturated hydrocarbons (Cio), the hydrocarbon composition of gasoline (C4-C7) and the geological distribution of effective source rocks indicate that the oils from the northern and central parts of the Triassic Basin and from the south and west of the Illizi and Ghadames Basins essentially came from a Devonian source. The oils of the fields of Hassi R'Mel, Makouda, Air Kheir, Oued-Noumer, Diorf, Guellala, Berkaoui, Ben Khala and Hassi Messaoud originated essentially in the Silurian graptolite shales of the central region of the Triassic Basin. Migration probably took place along the Hercynian unconformity, filling Triassic and CambroOrdovician reservoirs on the way. 11. The majority of the oils in the southern Paleozoic province were formed during the Paleozoic when burial of the source rocks exceeded 2-3 km. However, the Paleozoic traps were destroyed by the Hercynian erosion and the hydrocarbons then were able to dissipate. During the Mesozoic, throughout the subsidence of the northern and northeastern parts of the platform, the burial depths again reached the conditions necessary for the generation of hydrocarbons. During the Cretaceous (postAptian) the depth of burial of the source rocks reached values of about 3 km which favoured mainly the formation of gas. 12. In the northern Mesozoic province the hydrocarbons were generated mostly during the Mesozoic. There are, however, numerous hydrocarbon accumulations of a Paleozoic age in the southwest of the eastern province. The most likely source rocks are of Silurian age although Devonian shales may also be considered as potential sources. The Middle to Upper Devonian shales representing source rocks are best developed in the Eastern Erg (Ghadames). However, the process of oil generation was interrupted by a reduction of the burial depth of the source rocks during the Hercynian orogeny and the subsequent erosion of the Paleozoic formations. Although hydrocarbons could have formed during the Paleozoic, very few favourable structures or traps existed at the time. In the eastern part of the Saharan Platform a large part of the hydrocarbons that formed during the Paleozoic disappeared as a result of the Hercynian erosion. 13. Crystalline features of the clay minerals can probably be used as geothermal indicators. The correlation of the temperatures calculated from the reaction of I/M and the appearance of chlorite IIb in the Paleozoic sediments of the southern Oued elMya Basin with the (uncorrected) values measured in the boreholes revealed that the enclosing sediments in the past had experienced temperatures some 3o-4o °C higher than those obtained from diagraphic data. I4. Modeling of the burial of the basin and of the formation of the hydrocarbons allows us to calculate changes in the thickness of the sedimentary successions and the thermal regime of the sedimentary cover of the platform and to evaluate the generation of hydrocarbons in the source rocks of the basin. Alternative methods for calculating the variations of the tectonic subsidence amplitude of the top of the
282
Chapter 9 - General Conclusions
basement within the framework of a local isotopic model were used for correcting the sequence of tectono-thermal events adapted to the model. 15. The attempt to locate on the Saharan Platform and interparticular in the Triassic Province traps of the non-structural type is of particular interest for prospecting in view of the exhaustion of the "inventory" of unexplored structural traps. The stratigraphic traps are more characteristic of the Paleozoic sedimentary complex because of the presence of wedging zones and angular unconformities in the respective basins. Lithological traps are developed essentially in the sandy-argillaceous formations of the Triassic and result from the facies variations characteristic of these sediments.
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