Stable carbon isotope compositions during the thermal alteration of organic matter

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Stable carbon isotope compositions during the thermal alteration of organic matter

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Title:
Stable carbon isotope compositions during the thermal alteration of organic matter
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Conkright, Margarita E.
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Tampa, Florida
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University of South Florida
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English
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xii, 131 leaves : ill. ; 29 cm

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Carbon -- Isotopes -- Analysis ( lcsh )
Kerogen ( lcsh )
Methane ( lcsh )
Dissertations, Academic -- Marine science -- Doctoral -- USF ( FTS )

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Thesis (Ph. D.)--University of South Florida, 1989. Includes bibliographical references.

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University of South Florida
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University of South Florida
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All applicable rights reserved by the source institution and holding location.
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025637981 ( ALEPH )
21356460 ( OCLC )
F51-00174 ( USFLDC DOI )
f51.174 ( USFLDC Handle )

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STABLE CARBON ISOTOPE COMPOSITIONS DURING THE THERMAL ALTERATION OF ORGANIC MATTER by Margarita E. Conkright A dissertation submitted in partial fulfillment of the requirements of the degree of Doctor of Philosophy in the Department of Marine Science in the University of South Florida April 1989 Major Professor: William M. Sackett, Ph.D.

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Graduate Council University of South Florida Tampa, Florida CERTIFICATE OF APPROVAL Ph.D. Dissertation This is to certify that the Ph.D. Dissertation of Margarita E. Conkright with a major in Marine Science has been approved by the Examining Committee on January, 1989 as satisfactory for the dissertation requirement for the Ph.D. degree. Examining Committee: Major Professor:0W .M: v ( Member: E.S. Van Vleet Mem}?e'l: J t6. Walsh --I -\<"\..,.. Member: E.w.' Baker

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ACKNOWLEDGEMENTS I would like to especially thank Dr. W.M. Sackett for his teaching, support, and invaluable assistance through out this project. I would also like to acknowledge Drs. E.S. Van Vleet and E. W. Baker for their comments and time spent on this dissertation as well as Drs. R.H. Byrne and J.J. Walsh. Dr. R.M. Garrels provided great insight into some of the research problems and encouragement when needed. I will really miss him. I am grateful to T. Barber, R. Burke and G. Pauly for discussions about this work. I would like to thank Ken Peters for providing the samples and M. DeFlaun for analyzing the organic carbon content in my samples. Marion Kooker was great in allowing the use of her typewriter and for the encouragement she gave. Thanks very much to W.W. Gregg for his infinite patience throughout this time. This research was supported by grants from Chevron Oil Field Research Company, Sohio Oil Company, the Department of Natural Resources, the Department of Marine Science at the University of South Florida and the Patricia Harris Minority Fellowship. ii

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TABLE OF CONTENTS LIST OF TABLES v LIST OF FIGURES vii ABSTRACT x CHAPTER 1. INTRODUCTION 1 carbon Isotopic Compositions in Nature 3 Carbon Isotope Compositions of Methane 7 Biogenic Methane 7 Thermogenic Methane 9 Abiotic Methane 16 Objectives 18 CHAPTER 2. MATERIALS AND METHODS 20 Samples Analyzed 20 Organic Carbon Combustion 21 Pyrolysis 25 Methane Separation and Combustion 25 Calcium carbonate Analysis 26 Methane-co2 Exchange Experiments 26 Mass Spectrometer Analysis 27 Gas Chromatography Analysis 27 CHAPTER 3. STUDY OF MODEL COMPOUNDS 29 Earlier Work 29 Pyrolysis of n-octadecane at 600"C 33 Pyrolysis of Decacyclene 36 Discussion 40 CHAPTER 4. APPLICATION OF THE PYROLYSIS CARBON ISOTOPE METHOD 44 Earlier Work 44 Pyrolysis of Bakken Shales of North Dakota 52 Organic Carbon 55 Formation of Methane 56 Isotopic Composition of Methane 58 Formation and Isotopic Composition of C02 62 Time Series Experiments 64 Formation of methane 64 Formation and isotopic composition of C02 70 Isotopic composition of the organic residue 71 Discussion 72 iii

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Cape Verde Rise Kerogens Organic Carbon Content Time Series Experiment Formation and Isotopic Composition of C02 Formation of Methane Isotopic Composition of Methane Formation and Isotopic Composition of the Organic Residue Summary and conclusions CHAPTER 5. CARBON ISOTOPE EXCHANGE AT 600C Earlier Work Results and Discussion CHAPTER 6. CONCLUSIONS CHAPTER 7. LITERATURE CITED APPENDICES APPENDIX 1. RESULTS FOR THE LN2CF FOR THE PYROLYSIS OF 76 82 86 88 90 92 94 96 98 98 101 105 111 122 THE BAKKEN SHALES 123 APPENDIX 2. RESULTS FOR THE TIME SERIES ANALYSIS OF BAKKEN SHALE SAMPLE 38423-15. 124 APPENDIX 3. CARBON MOLE RATIO RESULTS FOR THE TIME SERIES PYROLYSIS OF BAKKEN SAMPLE 38423-15 124 APPENDIX 4. RESULTS FOR THE TIME SERIES PYROLYSIS OF BAKKEN SAMPLE 36650-1 125 APPENDIX 5. CARBON MoLE RATIO FOR THE TIME SERIES PYROLYSIS OF BAKKEN SHALE SAMPLE 36650-1 126 APPENDIX 6. GAS CHROMATOGRAPHY ANALYSIS OF THE TIME SERIES PYROLYSIS OF KEROGEN SAMPLE 31 127 APPENDIX 7. RESULTS FOR THE TIME SERIES ANALYSIS OF KEROGEN SAMPLE 31 (613C OF THE PARENT 0 CARBON= -26.4 / o o) 128 APPENDIX 8. RESULTS FOR THE CAPE VERDE RISE KEROGEN LIQUID NITROGEN CONDENSABLE FRACTION (LN2CF) 129 APPENDIX 9. RESULTS FOR THE CAPE VERDE RISE KEROGEN RESIDUAL ORGANIC MATTER 130 APPENDIX 10.EXCHANGE BETWEEN EQUAL AMOUNTS OF METHANE AND CARBON DIOXIDE (2ml OF METHANE AND DIOXIDE USED) 131 iv

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Table 1. Table 2. Table 3. Table 4. Table 5. LIST OF TABLES Interlaboratory comparison of organic carbon percent determinations. Samples used are the Bakken shales of North Dakota provided by Chevron Oil Field Research Company Pyrolysis of n-octadecane at 600C Pyrolysis of decacyclene at 500C (CMR x 100 represents % conversion of carbon to methane) Pyrolysis of decacyclene at 600C (CMR x 100 represents % conversion of carbon to methane). Determination of the isotope fractionation factors for methane formation from the pyrolysis of model compounds at various temperatures. Methane values used are those of initially formed methane (determined from the y-intercept in figures 6 and 7) 24 33 3 8 38 4 2 Table 6. Data for Bakken shales from North Dakota (provided by Chevron Oil Field Research Company) 51 Table 7. Table 8. Table 9. Table 10. Pyrolysis, at 600C (for 120 hrs.), of North Dakota Bakken shale samples of varying maturities (ranked from high to low H / C ratios) 53 Results for the time series pyrolysis of Bakken shale sample 38423-15 65 Results for the time series pyrolysis of Bakken shale sample 36650-1 65 Data for the analysis of the Cape Verde Rise kerogens from DSDP Leg 41, Site 368 (after Peters et al., 1983) 8 1 v

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Table 11. Table 12. Results from the pyrolysis, at 600 C (for 120 hr.), of the Cape Verde Rise kerogens Carbon isotope exchange between equal amounts of methane and carbon dioxide at 600C vi 84 103

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Figure 1. Figure 2. Figure 3. Figure 4. Figure 5. Figure 6. Figure 7. Figure 8. Figure 9. LIST OF FIGURES Carbon isotope compositions in nature (after Fuex, 1977) 6 Formation and degradation of kerogen (after Tissot and Welte, 1984) 10 Elementary compositions of humic acids and kerogens (after Tissot and Welte, 1984) 1 2 Schematic of the system used for the preparation of samples (TC= thermocouple; R= reservoir; TP= Toepler pump; M= manometer; LN= liquid nitrogen trap; ID= water trap; F= furnace) 22 Fractionation of methane from model compounds at various temperatures. Data from Sackett et al., 1968; Frank and Sackett, 1969;and Frank et al., 1974 (after Frank et al., 1974) 3 1 Carbon isotope fractionation vs. the carbon mole ratio of accumulated methane to n-C18H38 for pyrolysis experiments at various temperatures 35 Carbon isotope fractionation vs. the carbon mole ratio of accumulated methane to C36H18 for pyrolysis of decacyclene at 500 and 600 C 39 General scheme of kerogen evolution from diagenesis to metagenesis in the van Krevelen diagram (after Tissot and Welte, 1984) 46 Representation of the shift of methane precursor carbon (MPC) composition with natural maturation and artificial pyroly s i s (Sackett, 1984) 4 9 vii

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Figure 10. Figure 11. Figure 12. Figure 13. Figure 14. Figure 15. Figure 16. Figure 17. Location of the Bakken shales Relation between Carbon Mole Ratio (CMR) and H/C ratios from the pyrolysis of Bakken shales at 600C Relation between d13C and H/C ratios for the pyrolysis of Bakken shales at 600o c Change in the isotopic composition, with time, of carbon dioxide and methane for Bakken sample 38423-15 Change in the isotopic composition with time of co2 CH4 and the organic residue for Bakken sample 36650-1 Change in the composition with time of Bakken samples 36650-1 and 38423-15 with H/C ratios of 1.20 and 0.68 respectively. Numbers by points indicate pyrolysis time in hours Relation between carbon mole ratio (CMR) and d 13c ( <513C-CH4 <513c-parent carbon) for the Bakken shale. Numbers by points indicate the H/C ratios Location of the Cape Verde Rise samples from DSDP Leg 41, Site 368 Figure 18. Lithologic column showing the location of samples 58, 59. 60 and 62 in the Cretaceous black shale sequence from Site 368, DSDP 54 57 59 66 6 7 69 7 5 7 7 Leg 41 (after Simoneit et al., 1981) so Figure 19. Change in the o13C-TOC with proximity to the sill of the Cape Verde Rise kerogens. Hatched area represents the sill. Results include data from Simoneit et al., 1981 83 Figure 20. Gas chromatography results from the pyrolysis of Cape Verde Rise kerogen sample 31 8 7 Figure 21. Change with time of the isotopic composition of methane from the pyrolysis of Cape Verde Rise kerogen sample 31 8 9 Figure 22. Correlation between the carbon mole ratio (CMR) and H / C ratios from the pyrolysis of Cape Verde Rise kerogens 91 viii

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Figure 23. Change in the 613c of the organic carbon and organic residue of the Cape Verde Rise kerogens with proximity to the sill 95 Figure 24. Carbon isotope exchange between equal amounts of methane and carbon dioxide at 60ooc 102 ix

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STABLE CARBON ISOTOPE COMPOSITIONS DURING THE THERMAL ALTERATION OF ORGANIC MATTER by Margarita E. Conkright An Abstract Of a dissertation submitted in partial fulfillment of the requirements for the degree of Doctor of Philosophy in the Department of Marine Science in the University of South Florida April 1989 Major Professor: William M. Sackett, Ph.D. X

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The use of the amount and carbon isotopic composition of methane as a maturation index was tested by pyrolysis of sedimentary organic carbon (kerogen) at 600C. The parameters used to describe the maturity are CMR (CH4-Cjkerogen carbon) and the A 13c ( r13c r13c ) u CH4-0 OC With increasing maturities, smaller amounts of methane are generated and there is a decrease in the fractionation between methane and the parent carbon. This procedure is attractive because it can be used on whole rock samples and on kerogens which do not contain vitrinite or palynomorphs. The pyrolysis of Bakken shale samples, with varying maturities, show high correlation coefficients between the CMR and h13C vs. the atomic H/C ratios (r = +0.91 and -0.89 respectively) which indicates that each of these parameters, independently, can be used as a maturity index. The Bakken shale pyrolysis experiments also show that methane generated from the most thermally altered samples is up to 2joo heavier than the parent carbon. In addition, methane-co2 exchange experiments, at 600C, show a shift toward heavier methane values after heating of CH4 and C02 for 504 hrs. The isotopic composition of methane formed under high temperature regimes, such as from highly metagenic and subducted rocks, may be determined by exchange reactions if xi

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any C02 is present. For these reasons, it becomes difficult to use carbon isotope compositions of methane to distinguish between thermogenic and mantle methane without any other supporting evidence. The effect of metagenesis on the isotopic composition of organic carbon was determined for a suite of kerogen samples from the Cape Verde Rise, DSDP Leg 41, Site 368. This site is characterized by intrusions of hot diabase sills. The isotopic composition of the organic carbon is thermallycontrolled. With increasing maturities, the c513c-oc values are heavier due to a loss of lighter carbon in the form of methane. This is shown by a decrease in the carbon mole ratio, with increasing maturities. Abstract approved: xii MajorProfessor:W.M. Sackett Graduate Research Professor Department of Marine Science t v-Date of Approval

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CHAPTER 1 INTRODUCTION Throughout the history of mankind, three primary combustion sources of energy can be identified: wood, coal, and oil and natural gas. The use of wood as fuel dates back to prehistoric times and was the primary energy source until it was replaced by coal. Mining of coal began in the British Isles during the reign of Queen Elizabeth I and continued to grow in importance with the advent of the Industrial Revolution. Crude oil was first sold in large scale in the 1860's, but did not become a major fuel source until the development of the internal combustion engine. The shift toward natural gas began in the 1930's as technology was developed for long-distance, low cost gas pipelines. Presently, oil and gas account for three fourths of the energy consumed in the United States. The Arab oil embargo in 1973 created an international energy crisis which was just a presage of a more permanent crisis as oil and gas supplies dwindle. As a result, much emphasis is placed on understanding of the formation of oil and gas, and the characteristics of surrounding sediments which may provide clues as to the location and magnitude of

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2 potential oil and gas accumulations. Much current research focuses on the development of techniques to analyze and describe the characteristics of oil-producing rocks in an attempt to predict their hydrocarbon producing potential. Commercial accumulations of methane are generally believed to originate from either the microbial decomposition of organic matter in anoxic environments (biogenic gas) or thermal decomposition of organic matter (thermogenic gas). Recently, it has been advanced that commercial accumulations of nonbiogenic methane may exist in the deep crust and mantle (Gold and Soter, 1980). Stable carbon isotopes have proven to be useful in characterizing methane sources and history of available organic matter. Carbon has two stable isotopes; mass-12 with an abundance of 98.89%, and mass-13 with an abundance of 1.11% (Nier, 1950). The difference in chemical reactivity of isotopes, due to the mass difference, results in significant variations in the ratios of stable carbon isotopes in organic matter produced during physical and biological processes (Urey, 1947). Stable carbon isotopes have been used in studies of the origin and evolution of pre-biological organic matter in meteorites and terrestrial material, studies on biosythetic pathways, flow of carbon through food webs, and studies on the source, formation and migration of petroleum and natural gas. The isotopic composition of carbon is expressed using the del notation:

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[ 13c;12csample -1 13c;12cstandard ] 3 X 1000 Results are expressed as parts per thousand difference from a standard material. If the 613c is negative, this signifies the sample is depleted in carbon-13 relative to the standard, and the sample is referred to as "light": if the 613c is positive, the sample is enriched in carbon-13 relative to the standard, and the sample is referred to as "heavy" The standard used is Chicago PDB, a carbonate, Belemnitella americana, from the Cretaceous PeDee formation in South Carolina. A measure of the difference in the isotopic composition of two substances is the coefficient a, which is called the fractionation (separation) coefficient (Hoefs, 1980; Galinov, 1985) and defined as: where R is the ratio of any two isotopes in the chemical compound. In terms of the 6 values, a is expressed as: (6A + 1000) ( 68 + 1000) Carbon Isotopic Compositions in Nature Variations in the ratios of carbon-13 to carbon-12 are a result of two different reaction mechanisms; a kinetic isotope

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4 effect and a thermodynamic exchange effect. The isotope effect depends on mass difference of the isotopes, differences in the activation energies and temperature. The kinetic isotope effect is a result of different reaction rates due to the mass difference of the isotopes (Urey, 194 7; Hoefs, 1980) Chemical bonds formed by a heavier isotope are stronger than bonds formed with a lighter isotope therefore the activation energies for a reaction are different for the lighter and heavier isotopes. Generally, the products of a reaction are enriched in the lighter isotope, and the residue enriched in the heavier isotope (Galinov, 1985). The kinetic isotope effect decreases with increasing temperature. The thermodynamic exchange effect results in distribution changes of the isotopes in chemical substances or molecules (Urey, 1947; Bigeleisen and Mayer, 1947; Hoefs, 1980). Isotope exchange reactions can be written as: 12cH + 13co <===> 4 2 The equilibrium constant K, for this reaction is K = The isotopic equilibrium constant K, is temperature dependent. Fractionations are 1 at very high temperatures, and approach 0 at 0K which represents complete isotope separation (Hoefs, 1980).

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5 Nier and Gulbransen ( 1939) used the variations in the ratios of carbon-13 to carbon-12 to define two major reservoirs; an inorganic and an organic reservoir. The inorganic reservoir was found to be 2% enriched in carbon-13 relative to the organic one. Craig (1953) surveyed a large variety of carbon compounds and found atmospheric C02 to be depleted in carbon-13 relative to oceanic bicarbonate. Shown in Figure 1 is a representation of the stable carbon isotope distributions in nature (after Fuex, 1977). Oceanic plankton has variations of up to 15j oo (Sackett et al., 1974; Wong and Sackett; 1978) with values ranging from -30joo for cold water plankton, to about -16/oo for plankton in tropical and temperate waters. Degens et al. (1968b) analyzed the major biological constituents of marine plankton and found hemicellulose, proteins and pectins to be depleted in carbon-13 by 17joo relative to oceanic bicarbonate, and the extractable lipid fraction depleted by about 3 0 / o o Marine plants and animals are enriched in carbon-13 relative to terrestrial plants. Terrestrial plants which follow the Calvin-Benson or C-3 pathway are generally more depleted in carbon-13 (-32 to -22joo) than those which follow the Hatch-Slack or C-4 dicarboxylic acid pathaway (-23 to -12 0joo; Deines, 1980; Smith and Epstein, 1971). The o13c of organic carbon in surface sediments is the same as that of the organisms living in the environment of deposition (Sackett, 1964; Sackett et al., 1965; Sackett et al., 1974). This range is -28 to -16/oo depending on the

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-90 ocll 1 N PERMIL PDB 80 -70 -60 -50 -30 -20 INORGANIC CARBON AT'IIIOSP>RBON CAF180111 IN LIVING ORGANISMS 'IIIAR I NE PLANTS VERTEBRATES & II\IVEFITEBRATES 'IIIAR I NE PLANOC:TON l----------l LIPID FRACTION OF 'IIIARINE PL .. NTS LAND PLANTS L I P I D FRACTION O F LAND PLANTS H HYDROCARBON SOURCE 'lilA TER IALS RECE"H SEDIMENTS { MARINE ORGANIC CAH8 (JN 1N ORGANIC CARBON CLASTIC ROCt
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7 relative amounts of marine and terrestrial carbon input {Sackett and Thompson, 1963; Peters et al., 1978; Conkright and Sackett, 1986). The total organic carbon is not significantly altered during maturation processes therefore stable carbon isotopes can be used to study the depositional history of the sediments (Redding et al., 1980; Rigby et al., 1981; Yeh and Epstein, 1981; Schoell, 1984). Coals, formed from the thermal alteration of plant matter, have a distribution similar to modern land plants {-25 to -23/oo; Wickman, 1952; Craig, 1953; Colombo et al. 1969) and are generally carbon-13 'enriched relative to crude oils (around -30/oo to -26/oo). Methane produced from the decomposition of organic matter generally has a range from -50 to -25/oo, whereas biogenic methane is generally more 13c-depleted (-90 to -50/oo; Fuex, 1977). The o13c of mantle methane emitted from the East Pacific Rise ranges from -17.6 ojoo to -15/oo (Welhan, 1981). Carbon Isotope Compositions of Methane Biogenic Methane Microbial decomposition of organic matter in anoxic sediments may result in the production of biogenic methane. Methane producing bacteria are a specialized group of anaerobes which rely on hydrogen, carbon dioxide, acetate, methanol or methylamine substrates (Zehnder et al., 1982).

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8 Methanogenesis is limited by the presence of oxygen or dissolved sulphate concentrations in excess of 1. OmM (Whi ticar and Faber, 1986}. Two pathways are used by methanogens; acetate fermentation, the primary pathway in freshwater sediments, or C02 reduction occurring mainly in marine sediments. The range of o13c for biogenic methane formed by co2 reduction is -110 to -60/ o o, and -65 to -50/ o o by acetate reduction (Whiticar and Faber, 1986}. Carbon isotope fractionation during methanogenesis is determined by kinetic isotope effects rather than isotope exchange between CH4 and C02 (Whiticar et al., 1986} since the rate of carbon isotope exchange in shallow sediments is too slow to permit isotope re-equilibration of the methane that has formed with the C02 reservoir (Sackett and Chung, 1979; Giggenbach, 1982}. Previous studies by Galinov (1969}, Alekseyev et al. (1972) and Nakai et al. (1974} had postulated an isotope exchange equilibrium mechanism catalyzed by bacteria. In the initial zone of methanogenesis, methane is most carbon-12 enriched and becomes heavier with increasing depth and time as the C02 reservoir becomes depleted in carbon-12. A sharp decrease in methane concentrations from saturation levels to trace amounts in marine and certain freshwater sediments may be due to aerobic or anaerobic methane oxidation (Martens and Berner, 197 4} Methane consumption primarily takes place in a thin layer below the sulfate-reducing zone (Reeburgh, 1976, 1980}. Microbial oxidation of methane is

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9 associated with a kinetic isotope effect where the heavier isotope is concentrated in the residual methane, unlike methanogenesis. Thermogenic Methane Oils form in basins where thick accumulations of sediments occur (Hunt, 1979). Organic matter (plants, plankton and marine organisms) of marine origin deposited in sedimentary basins, is degraded by bacteria to form gases (ie. C02 and methane) and nitrogen, oxygen and sulfur containing humus complexes (Stahl, 1977). There is an enrichment in carbon-12 during the reactions which form the humic complex (Tissot and Welte, 1984). During diagenesis (T= 25-60C), nitrogen, and carbon monoxide and dioxide are expelled and these humic complexes are converted to kerogen. Kerogen is defined as the insoluble organic carbon fraction in sediments and b elieved to be the oil source material (Durand, 1980). Figure 2 is a schematic of the decomposition of organic matter and the formation of kerogen. The thermal degradation of kerogen to produce hydrocarbons occurs in the temperature range of 50-200C (catagenesis). Catagenesis can be described as a process of disproportionation whereby hydrocarbons with increase d hydrogen content are generated, and the residual kerogen becomes depleted in hydrogen (Mciver, 1967; Tissot et al., 197 4)

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Liwiat .,, ..... Zlu If Ill ltrltill Figure 2. Formation and degradation of kerogen (after Tissot and Welte, 1984) 10

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11 The thermal history of the rock is the most important factor in oil formation (Hunt, 1977). The formation of petroleum from organic matter is dependent on source material, time and temperature. Petroleum is formed at temperatures between 65-150C, defined as the "oil window" (Hunt, 1979). The thermal destruction of organic matter occurs in the temperature range of 150-250C (Price, 1983). Above 200C, heavy hydrocarbons are thermally unstable and graphite and methane are the only products (Stahl, 1974; Hunt, 1975; Barker, 1977). The alteration of kerogen to petroleum is dependent on the hydrogen content and structure of the kerogen (Dow, 1978). On the basis of H/C and 0/C atomic ratios, van Krevelan (1963) defined three types of kerogen and their evolution paths as shown in figure 3. Type I is produced from organic matter deposited in lacustrine environments where algal material is the principal form of organic matter. It is the most oil-prone, and least common type of kerogen. Type II kerogens are the most common and predominantly produced in reducing environments where autochthonous (phytoplankton, zooplankton, other detritus) material is the principal form of organic matter. Type III kerogens are more important in gas rather than oil formation if buried deep enough. The main source of organic matter for Type III kerogens is terrestrial plants. Aliphatic hydrocarbons, with high hydrogen content, are abundant in types I and II kerogens whereas type III kerogens are high in aromatics with a reduced hydrogen content. These

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u ...... X u 2 0 050 010 020 030 -----ATOWIC 0/C----- PRINCIPAL PRODUCTS OF KEROGEN EVOLUTION t::l CH4 [:=J OIL -GAS RESIDUAL ORGANIC MATTER (no potent1ol for 011 or gas) Figure 3. Elementary compositions of humic acids and kerogens (after Tissot and Welte, 1984) ...... IV

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13 differences in chemical compositions result in different maturation pathways, as well as different products (van Krevelen, 1963). During catagenesis and metagenesis, cracking reactions generate light hydrocarbons from kerogen as well as from previously formed crude oil (Stahl, 1977). Initially formed methane is enriched in carbon-12 due to a kinetic isotope effect. The carbon isotope fractionation of methane is associated with the cleavage of carbon bonds to form a methyl free radical which reacts with hydrogen to form methane (Sackett, 1978). Less energy is required to rupture 12c-12c bonds than 13c-12c bonds. Based on data from the cracking of propane-1-C13, at 500-550C (Stevenson et al., 1948) the rupture of 12c-12c bonds is about 4% more frequent than the rupture of 13c-12c bonds. Field and laboratory experiments have shown heavier methane values with increasing maturity of the source rock (Stahl, 1977; Schoell, 1983; Sackett, 1984). The difference between the isotopic compositions of the methane and source carbon decreases as a result of the temperature dependence of the kinetic isotope effect (Sackett et al., 1968) andjor methane formation from aromatic groups enriched in carbon-13 (Galinov, 197 4) The isotopic value of methane, in commercial deposits, is primarily a function of the isotopic composition of the source carbon and the fractionation which occurs during its formation (Fuex, 1977). Gas from terrestrial environments is 10-15joo

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14 lighter than methane from sapropelic environments (Stahl, 1975). On the basis of its isotopic composition, methane has been characterized by Galinov (1969), Stahl (1975), Schoell (1980) and Rice and Claypool (1981) as: a) shallow, dry gas: methane associated with bacterial production during diagenesis. The isotopic range of this methane is -90 to -55/oo and the amount of higher hydrocarbons is negligible. b) gas associated with oil generation: methane generated during catagenesis has an isotopic range of -58 to -40/oo. c) deep, dry gas: gas from oil source rocks past the oil generating stage or from cracked petroleum has a range of -40 to -20/oo. Several investigators have shown there is little or no isotopic alteration of the organic matter during the diagenetic and catagenetic stages, therefore the values of the total organic carbon reflect the source of organic matter (Redding et al., 1980; Rigby et al., 1981; Yeh and Epstein, 1981; Schoell, 1984). There is some discrepancy as to whether thermal processes alter the isotopic composition when the kerogen reaches the metamorphic stage. Barker and Friedman (1969) found an enrichment of carbon-13 in low organic carbon metamorphosed rocks relative to high organic carbon non-metamorphosed rocks. They attributed these differences to either loss of isotopically light organic compounds in metamorphosed rocks, or the presence of two types of organic matter with different isotopic signatures and degradation

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characteristics in non-metamorphosed rocks. 15 Baker and Claypool ( 197 o} in a study of incipient metamorphism on organic matter, found differences between metamorphosed and non-metamorphosed rocks only when the isotopic composition of the sources differ. McKirdy and Powell (1974} and Hoefs and Frey (1976} found an enrichment in 13c in metamorphosed vs. unmetamorphosed samples of up to 10joo. Hoefs and Frey (1976} suggested the 13c enrichment found in the metamorphosed rocks as due to kinetic isotope effects during methane formation, equilibrium exchange reactions between calcite and graphite, or preferential aqueous oxidation of 12c in graphite to C02 Laboratory pyrolysis experiments by Chung and Sackett (1979} from four shale samples found only a small enrichment in the residual fraction relative to the kerogen composition. Similar results were reported by Peters et al. ( 1981} for sapropelic and humic sediments, and Simoneit et al. (1981) for Cretaceous black shales altered by diabase intrusions (DSDP Site 41-368). Their results show a maximum enrichment in the residual fraction of 2joo. However, two studies have shown kerogen becomes more negative with increasing thermal stress. Jenden et al. (1982} studied protokerogens in the Guaymas Basin, Gulf of California (DSDP Sites 477, 478 and 481} which had been thermally altered by dolerite sills. Protokerogens at Sites 477 and 481 were 1.5joo more negative relative to the original carbon material. Peters et al. (1981) found residues from sapropelic and humic sediments to be initially more negative and become

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16 systematically heavier with increasing pyrolysis temperature. He proposed this trend was a result of initial loss of isotopically heavy functional groups, such as carboxyls, followed by loss of light methane leaving the residual carbon enriched in carbon-13. Abiotic Methane The association of methane with mantle-derived gases emitting from hydrothermal vents has led to increased interest into the source of this deep methane. Welhan and Craig (1983) measured methane in the East Pacific Rise hydrothermal vents to be -17.6 to -15/oo which is 10-20/oo heavier than organically derived thermogenic methane. Based on the methane values, the absence of any organic sources near the vents, and the absence of other low molecular weight hydrocarbons, they suggested an inorganic origin for this gas. They were uncertain whether the heavy methane values were the original isotopic composition of methane in the upper mantle or the result of high temperature exchange between CH4 and co2 within the basalt and extracted from the basalts by the circulation of seawater (Welhan, 1981). Several possible origins for methane from the mantle have been proposed. Gold and Soter (1980) suggested that large amounts of hydrocarbons, formed at the time of earth's accretion under high temperature and pressures, would liberate enough methane when degraded to account for large

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17 accumulations of gas. Some of the evidence they presented in favor of this argument was the presence of large concentrations of dissolved methane in waters overlying plate boundaries and rift areas (ie. the Red Sea brines or hydrocarbon plumes in the East Pacific Rise) The correlation between major oil and gas regions and the principal zones of past and present seismic activity, and phenomema associated with earthquakes such as flames, explosions and bubbling bodies of water can be attributed to the presence of a combustible gas such as CH4 or H2 Abiotic methane can be formed from the reduction of C02 and co by hydrogen gas (Fischer-Tropsch reaction) at temperatures of at least 250C and in the presence of a catalyst (Hunt, 1979). The Fischer-Tropsch reaction can account for the formation of organics in carbonaceous chondrites and bitumens found in igneous rocks (Galinov, 1985). This reaction can result in an initial large fractio-nation between CH4 and C02 ( 81 I o o ) due to kinetic isotope effects (Lancet and Anders, 1970). As methane becomes the dominant product with increasing time, its isotopic composition is determined by exchange between co and co2 approaching the equilibrium fractionation values predicted between CH4 C02 and CO (Craig, 1953; Bottinga, 1969; Richet et al., 1977). Studies of geothermal gases have been carried out in an attempt to evaluate the origin, history and composition of the volatile phases associated with the magmas. These studies

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18 have concentrated on isotope exchange reactions between CH4 a nd co2 and on geothermometry. Several conflicting results have arisen from these studies. Hulston and McCabe (1962) found that isotopic equilibrium between CH4 and co2 accounted for their results in New Zealand geothermal gases. Craig ( 1963) found that carbon collected from lavas agrees with carbon found in basalts (-24 to -18/o o) while geothermal methane was in the range of -30 to -25/oo characteristic of organic methane (Craig, 1953; Hulston and McCabe, 1962; Wasserburg et al., 1963). Gunter and Musgrave (1971) agree that a chemical reaction is not resposible for the methane values observed in hydrothermal systems. Giggenbach (1980) concluded that the difference in the isotopic compositions of coexisting CH4 and C02 is based on the initial compositions and cooling rates of the rising geothermal fluids. DesMarais et al. ( 1981) proposed that methane isotopic values in Cerro Prieto, Mexico, were independent from the co2 values. Lyons and Hulston (1984) further concluded that at present it is not possible to distinguish between mantle methane and thermogenic methane since thermogenic breakdown of organic matter in sedimentary rocks undergoing subduction and metamorphism may also contribute to deep gas deposits. Objectives 1. To undertake a systematic study of the changes in carbon isotopic compositions of total organic carbon from immature

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19 through over-mature whole rocks and kerogens from the same suite of rock. This should yield definite information on the effect of metamorphism on the o13c of organic carbon by resolving the discrepancy as to whether organic matter becomes lighter or heavier in the metagenic stage of maturation. 2. To utilize the amounts and isotopic composition of CH4 vs. the parent carbon compositions to further examine Chung and Sackett's (1979) and Sackett's (1984) maturation model. This model will be described in Chapter 4. The potential results from this research would be: a) a firmly established universal and objective maturation index, and b) data on the maximum amounts of thermally generated methane that can be expected from subducted rocks. 3. To set limits of metagenically derived methane in an attempt to distinguish between mantle methane and thermogenic methane. This would help in determining the importance of the upper crust and deep mantle as a source of methane. 4 To determine whether carbon isotope exchange between methane and carbon dioxide, at high temperatures (600C), controls the isotopic composition of the methane. This will be accomplished by measuring the changes in the isotopic composition of methane and carbon dioxide in laboratory exchange experiments.

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20 CHAPTER 2 MATERIALS AND METHODS Samples Analyzed A detailed description of the sedimentary samples analyzed, Bakken shales from North Dakota and kerogens from Cape Verde Rise, are presented before each pertinent section. The Bakken shales are organic rich black shales of Mississippian age, identified as the source rock of the Mississippian oils found in the Williston basin (Williams, 1974; Zumberge, 1983). They have a wide range of thermal maturities and contains primarily amorphous kerogen of aquatic origin (Leenheer, 1983). These samples, provided by Chevron Oil Research Field Company (CORF}, were ground and extracted with methylene chloride. The Cretaceous Black shales from the Cape Verde Rise, Eastern Atlantic are core samples from DSDP Leg 41-Site 368. These are organic-rich, Cretaceous black shales which were intruded by hot diabase sills during the Miocene (Lancelot et al., 1977). As a results of these intrusions, the changes in the organic matter are a result of thermal alteration rather than organic matter type. These kerogens were separated by CORF. The procedure for concentrating the kerogen from the sedimentary rock is destruction of the

PAGE 34

21 mineral matrix with 6N HCl and 60% HF. The residues are dried ih a vacuum oven at 60C, rinsed with methylene chloride and re-dried (Peters et al., 1983). The model compounds used for the characterization of methane formation from aliphatic and aromatic hydrocarbons were n-octadecane ( c18H38 ) and decacyclene ( c36H18 ) These were obtained through Aldrich Chemical Co.; n-octadecane 97% b.p. 317C; decacyclene m.p. > 325C. Gases used for the exchange experiments and gas chromatography standards were Scotty II Analyzed gases; co2 can mix 105 with 99.8% purity and methane can mix 109 with 99% purity. Organic Carbon Combustion Samples were weighed, wrapped in aluminum foil to prevent any loss of sample during the introduction of material into the combustion system, and placed in pre-combusted clay boats. The organic carbon was combusted using a circulation technique similar to that described by Craig (1953). Referring to Figure 4, the boats were placed in a quartz combustion tube half filled with cupric oxide and the system was pumped down and pressurized to about 20 torr with oxygen gas. The gases were circulated, using a Toepler pump (TP} 15 minutes through the furnace (F2; around 850C), 15 minutes through the furnace and a water trap containing a liquid nitrogen/isopropyl slush (ID; temperature about -80C) and then 15 minutes through the furnace, water trap and a liquid

PAGE 35

TO DIFFUSION PUMP ._ ____ LN2 10 VACWM PUMP F2 eooc LNl R TO SAMPLE BULB FOR MASS SPEC TP M Figure 4. Schematic of the system used for the preparation of samples (TC= thermocouple; R= reservoir; TP= Toepler pump; M= manometer; LN= liquid nitrogen trap; ID= water trap; F= furnace) N N

PAGE 36

nitrogen trap (LNl) in which co2 was frozen. 23 co2 gas was measured in a calibrated manometer before being transferred to an evacuated sample bulb. The residual organic carbon fraction was combusted in this same manner. The standard deviation for the results using this method is .2/oo. Organic carbon content can also be determined in a CHN analyzer by placing the sample in a quartz tube kept at 1030C and adding pure oxygen to the helium stream. The mixture of gases is then passed over Cr2o3 to ensure complete combustion, then over copper at 650C to remove excess oxygen and reduce the nitrogen oxides, and then through a Porapak QS chromatographic column heated to approximately 100C. The individual components are separated and eluted as N2-co2-H20 and measured in a thermal conductivity detector. The organic carbon content was determined by CORF using a Carlo Elba CHN analyzer and compared to results measured at USF, by M. DeFlaun, using a similar analyzer. These results are presented in Table 1. There is disagreement between the percent organic carbon measured in this system and the values measured by CORF. Generally, the combustion data was similar to either the Chevron or USF CHN carbon content results. The values reported for carbon content in the text are those measured from the combustion of organic matter in the Craig-type combustion system.

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24 Table 1. Interlaboratory comparison of organic carbon percent determinations. Samples used are the Bakken shales of North Dakota provided by Chevron Oil Research Field Company. SAMPLE CORFC CHN1 USF CHN2 COMBUSTION3 38421-4 10.0 10.4 11.1 36650-1 9.4 11.0 11.7 38423-15 11.0 10. 4 9.4 36662-1 12.8 11.3 10.8 38422-8 12.1 13. 6 12.0 36652-1 14.2 11.0 12.1 36653-1 17.8 18.1 18.8 36658-1 13.7 11.9 13.6 38422-21 19.2 12.9 19.6 36659-1 14.0 14.3 12.2 36651-1 12.5 12.8 11.0 36660-1 12.2 9.6 9.1 38420-39 9.9 20.1 18.8 36654-1 15.4 13.5 12.1 36649-1 9.6 8.1 8.7 Chevron Oil Research Field Company (CORFC) Carlo Elba CHN analyzer 2 3 University of South Florida Carlo Elba CHN analyzer Craig-type combustion of organic matter

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25 Pyrolysis Pyrex tubes were pre-combusted at 500C before use to remove any traces of adsorbed hydrocarbons and other gases. Solid samples were weighed, placed in the pre-combusted pyrex tubes, and sealed under vacuum. These samples were pyrolyzed 120 hrs. at 600C or for varying time periods for the time series experiments. N-octadecane was placed in small glass capillary tubes and weighed. These capillary tubes were then placed in the pyrex tubes and sealed under vacuum. Methane Separation and Combustion The separation of methane from other pyrolysis products was performed in the high vacuum system illustrated in Figure 4. The samples were initially cracked into the evacuated system with a device similar to that described by DesMarais and Hayes ( 19 7 6) The hydrocarbon gases were cooled in a liquid nitrogen trap (LN1) for 30 minutes, prior to the quantitative transfer of methane into the larger reservoir (R), via the Toepler pump (TP). Methane, which has a vapor pressure of 10mm at LN2 temperatures, was not frozen out in this trap. Carbon monoxide also has an appreciable vapor pressure at this temperature but none was detected using gas chromatography analysis of the products. The trapped hydrocarbons, referred to as the liquid nitrogen condensable fraction (LN2CF) were manometrically measured and transferred

PAGE 39

to an evacuated sample bulb. 26 In the exhaustive pyrolysis experiments, the only gas present in this fraction was C02 Methane was combusted using the same procedure as the combustion of organic matter except the circulation time was decreased to 5 minutes for each stage. The co2 gas generated from the combustion of methane was measured and transferred to an evacuated sample bulb. The LN2CF was combusted, to remove traces of water vapor, by cycling for 5 minutes through the furnace, 5 minutes through the water trap, and finally freezing for 5 minutes in a liquid nitrogen trap. Calcium Carbonate Analysis Inorganic carbon was determine d by reacting 6ml of 8 5 % phosphoric acid with the sediment sample in a sample bulb. The acid was mixed with the sample, after evacuation of the sample bulb, and the generated gases were allowed to pass through a trap cooled to about -80C (LN2 + isopropyl) before being frozen in another trap cooled to liquid nitrogen temperatures. The amount of gas produced was manometrically measured. This procedure was repeated until no more co2 gas was produced. Methane-C02 Exchange Experiments Methane and C02 were frozen in the apparatus shown in Figure (4). A pyrex tube was attached to the system and

PAGE 40

27 evacuated while the methane freezing trap (LN2) was pre-cooled with LN2 As soon as the lowest vacuum was attained, as measured by a thermocouple gauge {TCl), the breakseal was placed in the trap and pre-cooled. A vacuum was then pulled on the methane freez_ing flask and 2cc's of CH4 and C02 each were injected into the system. The gases were frozen for 30 minutes and then sealed while pulling a vacuum to eliminate the presence of any other gases such as 02 The tubes containing the gases were then pyrolyzed for varying periods of time and the methane and C02 separated as described in the previous sections. Duplicate samples were prepared for GC analysis and the only products detected were CH4 and C02 Mass Spectrometer Analysis o13c values were determined in a triple collector Mat 250 Finigan Isotope Ratio Mass Spectrometer (IRMS). All samples were measured relative to a Nori t working standard ( o13c = -24.8/oo vs. PDB). Gas Chromatography Analysis The composition of hydrocarbons gases was determined in a Hewlett Packard 5710A gas chromatograph {GC) equipped with a flame ionization detector. The column used was 3rnrn i .d. x 1.5m o.d. stainless steel columns packed with Porapak Q, mesh size 80-100 and programmed from 80-150 C at 16/minute. For

PAGE 41

28 the exchange experiments, the GC used was a Hewlett Packard 5710A equipped with a thermal conductivity detector. The helium carrier gas flow rate was 50mljmin and the oven temperature was held constant at 90C. The column used was 6nun OD stainless steel column with 3m grade 12 silica gel which allowed for baseline separation of H2 02 N2 CH4 and C02

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29 CHAPTER 3 STUDY OF MODEL COMPOUNDS Earlier Work The isotopic fractionation of methane from the thermal cracking of carbonaceous materials and simple model compounds has been studied by Sackett and co-workers using laboratory simulation studies (Sackett et al. 1968; Sackett, 1968; Frank and Sackett, 1969; Frank et al., 1974; Sackett, 1978). Laboratory pyrolysis studies have been used to explain the formation of petroleum and natural gases, and the potential of sediments to produce oil and gas (Baker, 1974; Martin, 1975). High temperature pyrolysis (500C and higher) has been used to generate information on the evolution and structural changes of kerogen (Giraud, 1970), and on model compounds to define mechanisms of hydrocarbon generation (Sackett et al., 1968; Sackett, 1968; Frank and Sackett, 1969; Sackett, 1978). The use of high temperature pyrolysis as a simulator of natural processes has been questioned (Bernard, 1978; Schoell, 1980; Hoering, 1984) since one of the problems associated with the pyrolysis procedure is the temperature necessary to offset the longer time requirements of natural systems. High

PAGE 43

30 temperatures may not substitute for long time (Chung, 1976; Bernard, 1978). Furthermore, high temperature dry pyrolysis may ignore some hydrogen transfer reactions in which wate' r plays an important role under natural conditions (Hoering, 1984). In the metamorphic stage of kerogens, water may affect the amounts of methane and C02 generated (Hoefs and Frey, 1976). However, Schoell (1980) found that results from pyrolysis experiments by Sackett and co-workers were comparable to natural methane compositions and bore a close resemblance to the formation of methane in nature. Furthermore, Evans and Felbeck (1983a) suggest that high temperature pyrolysis of organic material can be a valid simulator of the processes of petroleum formation and maturation. Sackett (1968) concluded that time, temperature and the nature of the terminal carbon-carbon bonds determine the fractionation of methane generated from thermal cracking. The nature of the carbon-carbon bonds is related to the type of organic matter, ie., terrestrial material contains a higher proportion of aromatic groups and marine and algal material are more enriched in aliphatic groups. The importance of terminal carbon-carbon bonds was determined from the cracking of propane, n-butane, n-heptane and n-octadecane at 500C as shown in figure 5 (data from Sackett et al., 1968; Frank and Sackett, 1969; and Frank et al., 1974). Fractionation of methane from these compounds shows a decrease with increasing bond strengths of the terminal methyl groups (Sackett et al.,

PAGE 44

z w 0 w > c ..J w cz .. % (,) u t() -30 -28 t6 A ; OCT A()[ CAN[ &ooc a !IO()C SACK[ T T et a l o ooc :s!IOc '3 j I ...j I -2t I OCTAOECANE a cooc THIS MPER J I !-22r-r--zl -r -r-Cl tSOBUTFJIE .... 0 ...j 1 ____ : o Figure 5. PROPANE 0.10 0.20 0.30 MOLES CM4 lt.40LE CARBON Fractionation of methane from model compounds at various tempe+atures. Data from Sackett et al., 1968, Frank and Sackett 1969 and Frank et al., 1974 (after Frank et al., 1974) 31

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32 1968). Bond dissociation energies decrease with increasing carbon number. Methane from n-octadecane ( c18H38 ) shows the highest fractionation, -25joo at 500C. As large molecules are cracked, decomposition products form with different carbon-carbon bond strengths and therefore different fractionations. These results show the energy differences between 12c-12c and 13c-12c are more significant in organic compounds with low bond dissociation energies. The nature of the c-c bonds of the organic molecule will determine the fractionation of methane from these compounds. The kinetic isotope effect is temperature dependent and decreases with increasing temperature (Urey, 1947). Sackett (1968) found a temperature dependence in the formation of methane from n-octadecane between 350 and 600C. The temperature dependence of the kinetic isotope effect was found to be unimportant in methane formation from neopentane between 500 and 600C (Frank et al., 1974) but a definite factor in the cracking of n-octadecane between 400 and 500C (Sackett, 1978). The carbon isotope fractionations for n-octadecane, at 400 and 500C, are -27.9 and -25.4/oo respectively. This study further investigates methane fractionation from n-octadecane at 600C. and from a highly aromatic model compound, decacyclene (C36H18 ) at 500 and 600C. These two compounds are assumed to be representative of the aromatic and aliphatic carbon found in kerogen.

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33 Pyrolysis of n-octadecane at 600C The cracking of n-octadecane to form methane should yield 9. 5 mmoles of methane if all the hydrogen is used in the production of methane (based on the hydrogen to carbon ratios) -------------> 9.5 CH4 +8.5C The carbon mole ratio (CMR) is defined as the ratio of accumulated methane carbon to parent carbon. It is a measure of the amount of carbon convertable to methane. For the cracking of n-octadecane, the theoretical maximum of methane was generated after 48 hrs. (theoretical maximum is 52.8% or CMR = 0. 528 as regulated by the H/C ratios) All of the available hydrogen is used up in the formation of methane. Table 2 shows the results for the pyrolysis of n-octadecane at 600C from 12 to 160 hrs Table 2. Pyrolysis of n-octadecane at 600C TIME 613C-CH CMR 1113c (hours) 0 4 (CH4/C) 13 < 1 o o) 6 C(CH4-c18H38 ) 12 -37.2 0.450 -3.8 24 -36.3 0.504 -2.9 48 -35.0 0.534 -1.6 96 -34.7 0.540 -1.3 160 -34.7 0.540 -1.3 n-octadecane 13 ( CPDB = -33.4/oo)

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34 The isotopic composition of methane generated from n-octadecane becomes more carbon-13 enriched reaching a plateau after 96 hrs. The isotopic composition of the methane remains 1.3/oo lighter than the parent carbon. Figure 6 is a plot of the carbon mole ratio (CHJC18H38 ) and the !J.13 c ( o13ccH4-o13cc18 H 38 ) for n-octadecane data from Sackett (1978) and for these results. The true fractionation during methane formation occurs when methane generation approaches zero. The y-intercept represents the isotopic difference between the initially-formed methane and the parent compound. The fractionation of methane at 600C is -16.5/ o o which compares to results by Sackett (1978) for CMR >0.1 at 500C. These results suggest the absence of a temperature effect in the fractionation of methane from n-octadecane between 500 and 600C, which is comparable to the results found by Frank et al. (1974) for the cracking of nee-pentane. The two curves shown in figure 6, for the 500C temperature range, represent a break in the fractionation of methane at CMR = 0.1 (Sackett, 1978). Initially formed methane is generated from the cleavage of terminal methyl groups and is more carbon-13 depleted due to the low bonding dissociation energies of n-octadecane. As the compound is cracked, smaller hydrocarbons form and methane is generated with smaller fractionations (Sackett, 1978). This process occurs at CMR >0.11. Therefore the largest fractionation is observed in the initial methane, which becomes progressively heavier as more methane is generated from smaller

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""'"' co (f) Z co 0 (..) (f) Co I z 00 (') ... co ...., 0 (') ... -28 -21 -24 -22 -20 -18 -18 -14 -12 -10 -a -e -4 -2 400(SACKETT, 1878) Y = 148X-27. 8 r = 0.847 SOO <<0.1)(SACKETT, 1978) Y:133X-25.4 r =0.987 500 <>0.1)(SACKETT, 1978) Y: 38X-16.8 r= 0.949 eoo Y = 28X-18.S r = 0.97 o. 1 0.2 0 3 0 4 0.5 0.6 CH41c H (mole ratio) 18 38 35 Figure 6. Carbon isotope fractionation vs. the carbon mole ratio of accumulated methane to n-c18H38 for pyrolysis experiments at various temperatures

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36 hydrocarbons . A problem associated with the use of 600C is that reactions occur quite rapidly and we are unable to study early methane formation. This curve represents methane formation from smaller hydrocarbons. Further analysis of the curves shown in figure 6, show a temperature effect for initially formed methane of 2.5/oo between 400 and 500C, and no temperature effect for subsequently formed methane between 500 (CMR > 0.1) and 600C. In the same way we are unable to study initially formed methane at 600C due to the speed of the reaction, the curve at 400C only represents initial methane since cracking reactions are slower at this temperature. Pyrolysis of Decacyclene The model compound decacyclene (C36H18 ) was chosen for these pyrolysis experiments because it is thought to be a stable and highly aromatic molecule which could be representative of the aromatic fraction in kerogen. This experiment was designed to determine the amount and isotopic composition of methane generated.from an aromatic compound. Decacyclene (C36H18 ) was pyrolyzed for different time periods at 500 and 600C. The theoretical maximum amount of methane that can be formed from decacyclene is 12. 5% (as regulated by the hydrogen to carbon ratio). ------------> 4.5 CH4 + 31.5C

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37 Results for the pyrolysis of decacyclene are shown in Tables 3 and 4. Pyrolysis of decacyclene at 500C for 84 hrs. generated only 17.6% of the theoretical maximum of methane. The CMR ranged from 0.015 after 11 hrs. to 0.022 after 84 hrs as shown in Table 3. This means only 1.5 to 2.2% of the carbon was converted to methane. Methane values range from -31.2 to 0 0 Methane values by 1.5 joo and were 5.1 joo lighter than the parent carbon after 84 hours. By contrast, pyrolysis of decacyclene at 600C generated 41.6% the theoretical amount of methane after 48 hrs. The carbon mole ratio increases from 0.041 after 12 hrs. to 0.052 after 96 hrs. (refer to Table 4). This is a 5.2% conversion of carbon to methane. The CMR remains constant after 48 hours implying no more carbon and hydrogen groups are available for methane formation which points to the stability of the compound. These results are also meaningful in illustrating the importance of the carbon mole ratio as a measure of the "reactive" carbon and hydrogen available for methane formation. The H/C ratio for decacyclene is 0.5 but less than 5% of decacyclene carbon is converted to methane. The isotopic composition of generated methane ranges from -28.2 to -26.3/oo. After g6 hrs., methane is 0 -1.7/oo relative to the parent material. These values are heavjer than those generated from decacyclene at 500C. Figure 7 is a plot of the CMR vs. A13c for these results. Methane fractionation is -9.7/oo at 500C and -10.7joo at

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Table 3. Pyrolysis of decacyclene at 500C (CMR x 100 represents % conversion of carbon to methane) TIME (hours) 11 24 39 84 decacyclene 13 ( CPDB = o13C-CH 0 4 ( 1 o o) -31.2 0.015 -6.6 -30.7 0.017 -6.1 -30.4 0.018 -5.8 -29.7 0.022 -5.1 -24.6loo) Table 4. Pyrolysis of decacyclene at 600C (CMR x 100 represents % conversion of carbon to methane) TIME (hours) 12 24 48 72 96 decacyclene 613 ( CPDB = o13C-CH 0 4 ( 1 o o) -28.2 0.041 -3.6 -27.7 0.046 -3.1 -26.4 0.051 -1.9 -26.3 0.051 -1.9 -26.3 0.052 -1.7 -24.6loo) 38

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. -7. 0 ,..... co =fa -6.0 ('t) 0 0 ('t) -5.0 I J: -4. 0 0 0 C") Co -3.0 0 -2. 0 Y=211X-9.7 r =0.99 500C Y = 169X-10. 7 r = 0.98 0 .01 0 .02 0.03 0 .04 0.05 0.06 CH 4/ c 36 H 18 (mole ratio) Figure 7. Carbon isotope fractionation vs. the carbon mole ratio of accumulated methane to c36H18 for the pyrolysis of decacyclene at soo and 6ooc 39

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40 600C. Temperature dependence of the kinetic isotope effect results in a decrease in fractionation with increasing temperature. Since we find an increase by 1 o o with an increase in temperature, we can assume the differences in these results are not due to a temperature effect between 500 and 600C. Results from this study for n-octadecane, and neo-pentane from Frank and Sackett (1969) show no temperature dependence between 500 and 600C. Further analysis of the data shows that methane values exhibit less fractionation at the higher temperature even though methane at t = 0 (zero-intercept) is 1joo lighter for the higher temperature. This difference can be due either to differences in cracking products with different bond dissociation energies or the fractionation of methane from decacyclene is not linear. Discussion To demonstrate the dichotomy between the aliphatic and aromatic carbon compounds, two representative pure compounds were chosen: n-octadecane, representing the aliphatic carbon pool, and decacyclene, representing the aromatic carbon pool. These results show that pyrolysis of n-octadecane yields the theoretical amount after 48hrs. (CMR = 0.53) and pyrolysis of decacyclene, at 600C, only yields 41% the amount after 48 hours, and remains constant for the remainder of the experiment. These experiments suggest that although the aliphatic-type carbon pool is quantitatively converted t o

PAGE 54

41 methane (controlled by the hydrogen content), the aromatictype carbon pool is more resistant to thermal degradation. The CMR is a measure of the "reactive" hydrogen and carbon converted to methane as shown by the decacyclene results. Aromatic carbon is more carbon-13 enriched than the aliphatic generated carbon which results in heavier methane from aromatics. Methane generated from being both compounds approaches the isotopic composition of the parent molecule with increasing pyrolysis time. This is more apparent at 600 than at 500C. Table 5 shows the isotope fractionation factors for the formation of methane, at various temperatures, from these two compounds (the definition and computation of the isotope fractionation factor is found in pg. 3 of this text). There are two factors involved in determining the fractionation of methane: temperature dependence of the kinetic isotope effect and carbon-carbon bond types. For n-octadecane, there appears to be a temperature effect in the region between 400 and 500C. Even though the data for 400C only refers to initially formed methane, there is a 2.5/oo difference between 400C and 500C (CMR <0.1). There is no difference between 500C (CMR > 0 .1) and 600C. The fractionation factors shown in Table (5) are smaller for decacyclene than for n-octadecane. Since there is no apparent temperature effect for decacyclene in the temperature range measured, these difference must be due to differences in the carbon-carbon bond types.

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42 Table 5. Determination of the isotope fractionation factors for methane formation from the pyrolysis of model -compounds at various temperatures. Methane values used are those of initially formed methane (determined from the yintercept in figures 6 and 7). TEMPERATURE COMPOUND ( o C) 400 n-octadecane 500 (CMR<0.1) n-octadecane 500 (CMR>0.1) n-octadecane 500 decacyclene 600 n-octadecane 600 decacyclene *a = 1 + ( o13ccH4 + 1000) 13 1 + ( 0 cparent carbon + 1000) A 13c o 13C-CH (o I o o) 0 4 ( 1 o o > a -27.9 -61.3 0.9711 -25.4 -58.8 0.9737 -16.8 -50.2 0.9826 -9.7 -34.3 0.9900 -16.5 -49.9 0.9829 -10,7 -35.3 0.9890

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43 These results agree with those from Frank et al. (1974) for nee-pentane. The fractionation of methane at these temperatures, 500 and 600C, is regulated by the nature of the carbon-carbon bonds in the parent molecule and the subsequent cracked products and not on temperature. These results can be applied to fractionation observed in methane generation from kerogen. Most methane formation from kerogen will occur from aliphatic-type carbon compounds. With increasing thermal stress this methane will approach the isotopic composition of the parent compound. As temperature continues to rise, some methane will be generated from aromatic-type carbon compounds and will have much heavier isotopic compositions. A careful measure of methane generation from kerogen should produce a break in the curve due to a depletion of aliphatic-type groups and a shift toward formation of methane from aromatic groups with heavier isotopic compositions. The amounts of methane generated will decrease due to a depletion of carbon available for methane formation as the thermal stability of the kerogen increases with time. The CMR should prove to be a useful parameter for determining the amounts of hydrogen and carbon which are actually available for methane generation.

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44 CHAPTER 4 APPLICATION OF THE PYROLYSIS CARBON ISOTOPE METHOD Earlier Work Petroleum and natural gases are formed through the alteration and maturation of organic matter during burial. Study of the thermal history of sedimentary rocks is important in hydrocarbon exploration since it makes possible the prediction of a) the level of maturation; b) the potential for petroleum and natural gas occurrence and c) the type of organic matter. The potential for petroleum generation can b e determined by a sediment's thermal history and source/type of organic matter. Immature kerogen is characterized as a metastable, polycondensed structure (Tissot and Welte, 1984). Low temperature chemical reactions are characterized by elimination of heteroatomic compounds such as C02 and H20, and the formation of alkyl and cyclic structures. This is followed by increasing rupture of c-c bonds and a subsequent production of various hydrocarbons. Catagenesis can be described as a process of disproportionation whereb y hydrocarbons with increasing hydrogen content are generated

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45 and the residual kerogen becomes depleted in hydrogen (Mciver, 1967; Tis sot et al. 197 4) Methane is the predominant product past the oil-generation stage. A general scheme for kerogen evolution is described by the van Krevelan diagram (figure 8) which shows a decrease in H/C and 0/C atomic ratios with increasing maturity for different kerogen types as a result of decarboxylation reactions (loss of oxygen) and aromatization and polymerization reactions (loss of hydrogen). H/C ratios are an established technique for determining the thermal history of sediments (Tissot et al., 1987). These ratios are measured on the kerogen which is separated from the source rock by strong acid digestion of the associated minerals. This method is very time consuming and laborious and may result in the loss of hydrogen and carbon-containing functional groups during the strong acid treatment. Kerogen isolation procedures are also susceptible to fractionation and contaminations risks (Gilmour and Pillinger, 1985). Vitrinite reflectance is another well established method for determining the maturity of sediment samples. Vitrinite is the most common maceral group found in humic coals and most sedimentary kerogens. It is formed after sedimentation in the diagenetic environment by humification of lignin and cellulose of plant cells (Dow and O'Connor, 1982). With increasing rank, there is a systematic increase in hardness and reflectivity (as shown in Figure 8). One of the problems associated with the use of vitrinite reflectance is the inability to

PAGE 59

.2 e J c 0 0 .05 0 .10 0 .15 0 .20 030 ---------------------. Atemie ratiO OIC Vitrinite reflectance i z 0 4 ,_1 Apptoaomote 0111 .. ol ..tronote refiKIOIIct --ol t ile f i eld ol ltroqlfl Eoluti oll pall'lt of til llfiiiC o pol 'YII" of roQell Figure 8. General scheme of kerogen evolution from diagenesis to metagenesis in the van Krevelen diagram (after Tissot and Welte, 1984) 46

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47 distinguish between primary and recycled vitrinite, and between vitrinite and other materials which may resemble it, such as solid bitumens (Dow and O'Connor, 1982). Chung and Sackett (1979) found that the amounts and stable carbon isotopic compositions of methane relative to the total organic carbon (TOC) can be used as measures of the maturity of the TOC. Based on the results from nineteen coals and twelve sol vent-extracted shale samples with different maturities (as determined by other parameters), Sackett (1984) proposed the pyrolysis-carbon isotope method (PCM) for determining the maturity of coal and shale samples. The parameters used to describe the indices are derived from the isotopic difference between the total accumulated methane, produced by exhaustive pyrolysis, and the total carbon (A 13c = o13ccH 4 o13croc> andjor the difference in mole ratio of methane (total produced by exhaustive pyrolysis) to parent carbon (CMR = CH4-C/KEROGEN-C). The PCM model is based on the concept of methane formation from a thermally induced, non-renewable pool of methane precursor carbon (MPC), such as alkyl side chains and linking groups, following the microbial alteration of organics. The initially formed methane is isotopically the lightest (since 12c-1 2c bonds are easier to rupture than 13c-12c bonds) which leaves the residual pool isotopically heavier. With increasing time and temperature, the carbon-carbon bonding in the pool becomes stronger and more 12c depleted which results in a reduction of amount and carbon-12 content in subsequently generated

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48 methane. This process is illustrated in Figure 9 by the shift from MPCinitial to MPCtoday The pyrolysis experiments artificially mature the sediments forcing a shift from MPCod t ay to MPCpost-pyrolysis The extent to which this shift occurs is dependent on how much the sediments have already matured by natural processes. "Thus, the more methane produced relative to the MPC pool and the larger the fractionation we obtain by pyrolysis, the closer is the MPCtoday to the MPCinitial pool and the more immature is the kerogen. Conversely, the smaller the amount of methane and isotope fractionation, the closer is the pool to its "final" mature state" (Sackett, 1984). The amount of methane generated will be controlled by the amount of hydrogen available in the MPC pool for methane production. The isotopic composition of the methane will be controlled by the kinetic effects during carbon-carbon breakages in the MPC pool. The characteristics of the MPC pool are mainly source-controlled. This procedure is attractive for the following reasons: 1. It depends on well established physical (ie. preferential cleavage of 12c-12c bonds over 13c-12c bonds) and chemical laws and objective chemical and isotopic measurements (Sackett, 1984) 2. It does not rely on the presence of palynomorphs, vitrinite or petrographic measurements. 3. It can be performed on whole rock samples and does not require kerogen isolation with its potential for loss or change in some of the carbon matrix. It only requires solvent

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4'3C (%o) o 5 "tb .... '\((; ,.).'\. g..t-" '\ o(c .... '\((; 0 o" .05 0.1 0.15 CH4 -C/KEROGEN-C (MOLE RATIO) Figure 9. Representation of the shift of methane precursor carbon {MPC) composition with natural maturation and artificial pyrolysis (Sackett, 1984) \0

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50 extraction of the organic matter (Conkright et al. 1986). This study was undertaken to provide further verification of the pyrolysis-carbon isotope maturation procedure as outlined above. The potential results from this research would be: a) a firmly established universal and objective maturation index, b) data on the maximum amounts of thermally generated methane that can be expected from subducted rocks and c) isotopic criteria for differentiating between methane produced from subducted kerogen and that produced by other reactions deep within the earth. The PCM was applied to a suite of Bakken shale samples provided by the Chevron Oil Field Research Company. These samples were of various maturities as inferred from atomic H/C values, and data generated from Rock-Eval pyrolysis such (temperature of maximum hydrocarbon generation), P.I. or production index ( s1;s1+s2 where s1 is the amount of hydrocarbons already present in the rock and s2 the hydrocarbons generated from the pyrolysis of the kerogen), and H.I. (hydrogen index which is the ratio of s2 to organic carbon) (these data are presented in Table 6). and the production index both increase with increasing thermal maturity; H/C atomic ratios and the hydrogen index decrease with depth as reactive hydrogen becomes less available. The PCM was al8o applied to kerogen samples extracted from Cretaceous black shales from the Cape Verde Rise, Eastern Atlantic supplied by the Chevron Oil Field Research Company and are described in the next section. These samples were

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51 Table 6. Data for Bakken shales from North Dakota (provided by Chevron Oil Field Research Company) EXTRACTED SAMPLE ROCK WT. DEPTH Tin ax HI8 Pib No. (g) (ft) (oC) 38420-39 29.0 7545 439 340 0.12 36650-1 39. 0 7562-8 426 573 0.03 36653-1 7.7 7577 431 543 0.06 38421-4 43.4 9899 457 150 0.24 38422-8 31.5 10002 439 870 0.08 36654-1 13.0 10006-9 435 333 0.11 38422-21 27.8 10054 432 750 0.09 36651-1 12. 5 10253-8 454 149 0.26 36660-1 15.8 10332-50 453 117 0.20 36659-1 5.2 10638 452 95 0.23 36658-1 7.5 10796-9 448 104 0.22 36652-1 11.0 11000-26 445 178 0.16 36662-1 2.5 11139-43 454 88 0.27 38423-15 29.0 11259 455 110 0.42 36649-1 10.0 11262-5 450 96 0.38 8Hydrogen Index from Rock-Eval pyrolysis (mg HC/ g TOC} bProduction Index, S1/(S1+S2), from Rock-Eval pyrolysis

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52 obtained by DSDP Leg 41-Site 368 and are of various maturities as determined by H/C ratios and vitrinite reflectance percentage. Pyrolysis of Bakken Shales of North Dakota The Bakken shale is an organic rich black shale found as a thin clastic unit in the subsurface of the Williston Basin which covers most of North Dakota in the United states (Williams, 1974). It is of Mississippian age and identified as the source rock of the Mississippian oils found in this basin (Williams, 1974; Zumberge, 1983). The kerogen is primarily amorphous (70-95%) with an algal origin as inferred from the high H/C ratios (Webster, 1984) Webster also identified herbaceous kerogens (0-20%) and coaly kerogens {<30%). Samples varied in depth between 7545 and 11262 feet with varying maturities as inferred from the H/C ratios. Data for these samples are presented in Tables 6 and 7 and location of samples in Figure 10. Pyrolysis of model compounds and kerogen samples, for a period of several hours to several days, show that no more methane is generated after 96 hours. To insure complete pyrolysis, the Bakken shale samples were pyrolyzed for 120 hours at 600C

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53 Table 7. Pyrolysis, at 600C (for 120 hrs.), of North Dakota Bakken shale samples of varying maturities (ranked from high to low H/C ratios). SAMPLE OC CMR H/C No. % 36650-1 11.7 -26.1 -27.6 -1.5 0.222 1.20 36653-1 18.8 -29.0 -31.6 -2.6 0.183 1.15 36654-1 12.1 -28.4 -30.6 -2.2 0.203 1.10 38422-8 12.0 -28.2 -30.7 -2.5 0.213 1. 09 38422-21 19.6 -28.8 -30.1 -1.3 0.180 1.03 38420-39 18.8 -28.6 -30.5 -1.9 0.181 0.90 38421-4 11.1 -28.4 -28.3 +0.1 0.107 0.83 36651-1 11.0 -28.4 -26.6 +1.8 0.089 0.82 36652-1 12.1 -26.7 -25.3 +1.4 0.132 0.81 36659-1 12.2 -28.4 -27.0 +1.4 0.152 0.80 36660-1 9.1 -28.2 -28.2 0.0 0.122 0.76 36662-1 10.8 -28.6 -27.7 +0.9 0.073 0.75 36658-1 13.6 -27.8 -28.2 +0.6 0.080 0.73 36649-1 8.7 -28.4 -27.0 +1.4 0.099 0.70 38423-15 9.4 -26.9 -24.7 +2.2 0.110 0.68 *H/C ratios provided by Chevron Oil Field Research Companydetermined on pre-extracted shale samples

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Saskatchewan 31152-1, .. 3MI2-1 I 31420-1 31652-1-. 3..51-1- ._ .... _, 3M5._1 ..... 3M51-1Figure 10. Location of the Bakken shales 3M50-1 .,., Manitoba Blamarck I UM12571 Ot" "'"

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55 Organic Carbon The organic matter (OC) in sediments is a composite of that photosynthesized in the overlying waters and that believed to be transported from other synthesis sites by currents (Sackett, 1964; Sackett et al., 1965; Degens, 1969; Conkright and Sackett, 1986). Several investigators (Redding et al., 1980; Rigby et al., 1981; Lewan, 1983) have shown there is little or no isotopic alteration of the organic matter during the diagenetic and catagenetic stages, therefore, the cS13c values of the total organic carbon reflects the type of organic carbon (if there are no losses of carbon). The organic carbon (OC) values are those for the samples after extraction with methylene chloride and represent the organic carbon that was actually combusted (see methods section for determination of organic carbon content). The carbon content varies between 8. 7 and 19.6% and the cS13C values range from -26.1 to -29.0/oo with a mean value of -28.1/oo (see Table 7). In general, these light values and high oc suggest a highly restricted-anoxic depositional environment (resulting in high preservation of organic matter) with a dynamic cycling of C02 between the atmosphere and epicontinental sea, fixation by marine algae and bio-oxidation resulting in a 12c enrichment in dissolved C02 and a concomitant production of organic matter with cS13c values of about -28/oo (Conkright et al., 1986). There is no relationship between the maturity of the samples (as measured

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56 by the H/C ratios) and the o13C0c which supports the above observations that isotopic variations in the oc are source-controlled rather than maturity controlled. Formation of Methane The methane to carbon mole ratio (CMR) varies from 0.073 to 0.222 (a 7.3 to 22.2% conversion of carbon to methane; see Table 7) As shown in figure 11, the highest amounts of methane are generated by the least mature samples (high H/C ratios) and decrease as the hydrogen to carbon ratio decreases, as predicted by the model. The high correlation between the carbon mole ratio and H/C ratios (r= 0.91) shows how the carbon mole ratio is a measurement of the hydrogen available for methane production. Methane generation was highest for samples with H/C ratios higher than 0.9 (18-22% conversion) Chung and Sackett (1979) show that higher carbon content (%) is correlated with the volume of methane generated from the pyrolysis of shale samples at 500C. This correlation \vas due to the dependance of methane production on the amount of carbon available. Such correlation is not found in the pyrolysis of the Bakken shale samples, which shows that not all of the carbon, in the Bakken shale samples, is available for methane generation. This supports the premise that the amount of methane generated is dependent on type of organic matter and not just on carbon content. This conclusion is

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57 o. 0 6 0.7 0.8 0 9 1 0 1 1 1. 2 H/C RATIOS Figure 11. Relation between carbon mole ratio (CMR) and H/C ratios for the pyrolysis of the Bakken shales at 600"C

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58 further supported by the decacyclene results in the previous chapter. Isotopic Composition of Methane As shown by the model, the isotopic composition of the methane reflects the extent to which the organic matter has been altered. The 613 CCH4 values range from -31.6 to 0 -24.7 joo. The heaviest values coincide with the most thermally altered samples and conversely, the lightest values with the least mature samples. With increasing temperature, there is an increase in the rupture of 12c-12c bonds and a subsequent depletion of 12c in the carbon pool, which results in an enrichment of carbon-13 in the MPC pool. Methane becomes heavier than the parent carbon in samples with H/C ratios lower than 0.9 (shown in figure 12). Displa-cement toward methane isotope values heavier than the total organic carbon at higher temperatures have been found by Chung (1976) in the pyrolysis of anthracite at 500C and by Arneth and Matzigkeit (1986) in the pyrolysis of sediments from the Williston Basin at 500 and 600C. One posible reason for this displacement toward heavier values is intramolecular isotopic differences o r 13c heterogeneity. Abelson and Hoering (1961) found that carboxyl groups in amino were enriched in carbon-13 relative to the total amino acids. The lipid fraction is enriched in carbon-12 by about 8joo relative to the whole plant (Abelson and Hoering, 1961; Park and Epstein,

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+3 +2 -r =-0. 87 ; + 1 0 I 0 0 C') (f) 0 T"" .... I (f) T"" co -1 -2 -3 0.6 0.7 0 8 0 9 H/C 1 0 1 1 1.2 Figure 12. Relation between b.13c and H/C ratios for the pyrolysis of Bakken shales at 600C 59

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60 1961; Smith and Epstein, 1971 and DeNiro and Epstein, 1978). Galinov (1974) describes this intramolecular isotopic heterogeneity as a result of thermodynamic equilibrium during biosynthesis. These results have also been explained as a result of kinetic isotope effects during the synthesis of biological molecules (DeNiro and Epstein, 1978; Monson and Hayes, 1982). Chung (1976) concluded that carbon-13 heterogeneity is better explained by kinetic isotope effects rather than by thermodynamic equilibrium. Fractionation in methane is mainly due to the lower energy required to break 12c-12c bonds than 13c-12c bonds. With increasing maturity, new carbon-13 distributions develop in the kerogen leaving some methane precursor functional groups highly enriched in carbon-13. This latter view is more consistent with the results found in these pyrolysi. s experiments and with the carbon pool concept. Carbon-13 heterogeneity can become an important factor in determining the isotopic composition of methane in samples with H/C ratios lower than 0.9. A similar break in the curve may occur in figure 11 at H/C = 0.9. This break could be fortuitous or due to a shift in methane generation from aromatic rather than aliphatic type carbon groups. Increasing time and temperature results in an increase in aromaticity and polymerization and a concomitant decrease in methane precursor carbon. The enrichment of methane relative to the parent carbon can also be a result of isotope exchange between methane and

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61 co2 Several investigators (Sackett and Chung, 1979; Harting and Maas, 1980; Giggenbach, 1982; and Lyon and Hulston, 1984) have shown that isotope exchange between CH4 and C02 is extremely slow at temperatures below 500C. At higher temperatures, Harting and Maas (1980) showed that some exchange takes place. The possibility of this reaction occurring at 600C can not be discarded but seems unlikely due to the high correlation between b.13ccH4 and the H/C ratios. This correlation would not exist if the methane isotopic values were controlled by the C02 fraction. The correlation between b. 13c values and H/C ratios is shown in figure 12. Again, this is a parameter that relates the fractionation of methane to increasing maturity. The b.13c values are positive with H/C ratios lower than 0.9 and reach a maximum value of +2. 2 for an H/C ratio of 0. 68. This demonstrates that methane can be heavier than its parent carbon. Heavy methane values were found by Welhan and Craig ( 1983) in the East Pacific Rise hydrothermal vents. An abiotic source was proposed due to the heavy methane values (-15 to -17.6/oo), the absence of any organic sources near the vents and the absence of other low molecular-weight hydrocarbons. The Bakken shale methane results show that heavy methane values can be generated from highly altered sediments. It therefore becomes difficult to use carbon isotope compositions of methane to distinguish between thermogenic and mantle methane without other supporting evidence.

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62 Formation and Isotopic Composition of co2 Prior to the pyrolysis of the Bakken shales, all samples were tested for the presence of carbonates (this procedure is described in the methods section). Samples 38422-8 and 38421-4 were analyzed for caco3 content since bubbles were detected upon the addition of acid. Sample 38422-8 contained 0. 34% carbonate while sample 38421-4 contained 3.7%. The isotopic compositions were -5.9 and -2.1/oo respectively. The possibility of production of C02 from the carbonates is remote since the decomposition of caco3 generally does not occur at temperatures below 800C unless impurities such as Sio2 are present. caco3 can react with sio2 to form co2 and casio3 Production of co2 from caco3 can also occur between 400 and 700C in the presence of oxygen (Gilmour and Pillinger, 1985; Swart et al., 1982) but since pyrolysis involves heating in the absence of oxygen, caco3 has been disregarded as a possible contaminant. The amounts of carbonate present were not significant enough to warrant acid digestion of the samples which could have resulted in the loss of some of its carbon and hydrogen containing functional groups. The amounts and isotopic composition of the C02 were highly variable and showed no relationship with increasing maturity, amounts and isotopic composition of methane, or carbon content (Appendix 1). Intuitively, one would expect the amount of co2 to decrease with increasing thermal stress

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63 since the source of C02 is decarboxylation reactions which result in a decrease in the availability of oxygen groups. Immature sediments contain a higher proportion of thermally labile c-o functional groups which decrease with increasing temperatures. The C02 data suggests that even in the most mature samples there may still be enough carbonyl functional groups available to produce co2 This would suggest a thermal stability in the C-O groups found in the Bakken shales. Arneth and Matzigkeit (1986) have speculated on the possibility of a temperature-dependent release of different types of oxygen functional groups of various stabilities and isotopic compositions to explain a 13c enrichment of pyrolitically-derived C02 Oxygen release can occur not only from the carbonyl and carboxyl groups but also from ether bonds, phenols or heterocyclic structures. Bond dissociation energies and enrichment in either carbon-12 or carbon-13 in these groups will determine the isotopic composition of the C02 The lack of correlation between C02 and CH4 supports the absence of isotope exchange between the two fractions. Isotope exchange reactions lead to a continual enrichment of carbon-13 in the CH4 fraction and a concomitant depletion of carbon-13 in the C02 fraction until equilibrium is attained (Craig, 1953; Bottinga, 1969; Richet et al., 1977; and Lyon and Hulston, 1984).

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64 Time Series Experiments Time series experiments were conducted on samples 36650-1 and 38423-15, the samples with the highest (1.20) and lowest (0.68) H/C ratios (results are shown in Tables 8 and 9 and Appendices 2, 3, 4 and 5). These samples were pyrolyzed, at 600C, for a period from 12 to 160 hrs. and and from 12 to 120 hrs. for samples 36650-1 and 38423-15, respectively. By artificially maturing these samples, the changes in amounts and isotopic composition of the gases can be more closely analyzed. In the case of sample 36650-1, the change in the isotopic composition of the organic residue was also determined. Formation of methane The isotopic composition of the methane shows a gradual enrichment in carbon-13 with increasing time for both sets of data (see figures 13 and 14). The range for the least mature sample, 36650-1, was -30.8 to -27.5/oo for a period from 12 to 160 hrs. The range for the most mature sample, 38423-15, was -26.8 to -24.7joo for a period from 12 to 120 hrs. For sample 36650-1, the methane was depleted in carbon-13 relative to the parent carbon by -1.4joo at 120 hrs. All methane values for sample 38423-15 were enriched in carbon-13 relative to the total organic carbon.

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65 Table 8. Results for the time series pyrolysis of Bakken shale sample 38423-15. TIME CMR 13 13 f:.13C-CH 6 Cco2 6 CCH4 4 (hr) (parent-CH4 ) 12 0.078 -2.2 -26.8 0.1 24 0.072 -5.2 -25.8 1.1 48 0.077 -6.1 -25.8 1.1 120 0.110 -6.4 -24.7 2.2 parent carbon = -26.9/oo Table 9. Results for the time series pyrolysis of Bakken sample 36650-1. TIME 613c CMR 613C CMR &13C (hr) (C02 ) (CH4 ) (OCR*) 12 -----30.8 0.185 -26.4 24 -7.7 0.087 -30.3 0.189 -25.8 48 -8.7 0.092 -29.5 0.186 -27.5 72 -10.5 0.116 -30.1 0.193 -27.2 96 -9.4 0.102 -29.2 0. 214 -27.6 120 0.130 -27.5 0.240 -28.1 160 -6.6 0.160 -28.1 0.276 -29.0 parent carbon = -2 6 1 j o o OCR = organic carbon residue

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0 -5 8 0"' -10 C02 .._... 0 -15 ('t) ..-Gc::> -20 -25 ___.--.. -30 0 24 48 72 96 120 144 TIME (hours) Figure 13. Change in the isotopic composition, with time, of carbon dioxide and methane for Bakken sample 38423-15 0'1 0'1

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-8 ....... 0 -0 r-GC) 0 -5 -10 -15 -20 -25 t- t -30 -35 0 24 ,._, 48 72 96 120 144 168 TIME (hot.rs) C02 -BCH4 Organic residJe Figure 14. Change in the isotopic composition, with time, of C02 CH4 and the organic residue for Bakken sample 36650-1 0\ -..l

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68 The methane to carbon ratio increases with time reaching a maximum conversion of 27.6% for sample 36650-1 at 160 hrs. and only 11% for sample 38423-15 for 120 hr (CMR =0 .110). Sample 36650-1 had a methanejcarbon ratio of 0.24 at 120 hrs. almost double that of sample 38423-15. The generation of methane halved from a sample with an H/C ratio of 1.20 to one with an H/C ratio of 0.68. Methane generation is not "exhaustive" as previously assumed since there is additional methane generation after 120 hrs. for sample 36650-1. The CMR increases from 0.240 at 120 hrs. to o. 276 at 160 hrs. The use of 12 0 hrs. for pyrolysis time was based on GC time series analysis of kerogen samples (described in the following section) and those reported in the literature by Sackett and co-workers. Even though the maximum amount of methane was not generated for this sample at 120 hrs., the high correlation between CMR and H/C ratios in figure 11 points to either "exhaustive" methane in the other samples analyzed, or exhaustive pyrolysis is not a pre-requisite for the model to work. I tend to support the first assumption because of the high correlation found in figure 11 which would not be found if the amount of methane generated was still rapidly changing. The change, with time, in CMR and is illustrated for both samples in figure 15. The two sets of data are considerably displaced from one another with no overlap in either parameter for the pyrolysis times that were used.

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eu 38423-15 (H/C:O.e8) r:0.744 Y 5 1 2 e x -3 o 1 0 .100 69 euo 3eeso-1 (H/C: 1.20) r :0.850 Ya2i.43X-9.4i 0 .300 CH 4 -C/KEROGEN-C (MC?LE RATIO) Figure 15. Change in the composition with time of Bakken samples 36650-1 and 38423-15 with H/C ratios of 1.2 and 0.68 respectively. Numbers by points indicate pyrolysi s time in hours

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70 Formation and isotopic composition of C02 The volume and isotopic composition of carbon dioxide generated from sample 36650-1 was highly variable with no trend with time. However, results for sample 38423-15 show the volume to be relatively constant between 12 and 120 hrs. and the isotopic composition constant from 48 to 120 hrs. o13Cc02 ranges between -2. 2 and -6.4/oo. The isotopic composition of C02 decreases in carbon-13 content with increasing pyrolysis time. Several investigators have reported a carbon-13 enrichment in C02 with increasing thermal stress (Arneth and Matzigkeit, 1986; Sackett, 1978; Chung, 1976) contrary to the results from this experiment. During initial pyrolysis of these samples, the LN2CF consists of C02 and c2-c6 hydrocarbons (Sackett, 1978). These gases have more positive values than co2 and disappear around 48 hours. The increase in carbon-12 content with time, measured in these experiments, can be accounted for by a decrease and ultimate disappearance of other hydrocarbons. The co2 value for sample 38423-15 appears to be around -6.2joo. The possibility of isotope exchange between CH4 and C02 needs to be considered for sample 38423-15 since CH4 becomes heavier as the co2 becomes lighter. Isotope exchange equilibrium between CH4 and co2 at 600C, as predicted by Bottinga (1969), Richet et al. (1977) and Lyon and Hulston (1984) would give co2 compositions about +11.5/oo relative to the CH4

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71 values. At 120 hr., the difference between the co2 and CH4 isotopic values is 18.3. Further evidence for the existence of isotopic exchange is the correlation found between o13c-co2 and o13C-CH4 (r = -0.88). This correlation could be fortuitous since previous results from the pyrolysis of the Bakken shales show no indication that exchange is taking place. Also, the high correlation between A 13c and the H/C ratios is not consistent with an exchange mechanism. If the isotopic composition of the methane were determined by exchange with other carbonaceous species present, no correlation would be expected between methane and the maturity of the samples. Thermodynamic exchange is dependent on temperature and not on the maturity of the samples. Further evidence is the variable C02 compositions for the time series pyrolysis of sample 36650-1 and the lack of correlation with methane isotopic values. Isotopic composition of the organic residue Several investigators have shown that organic matter is not significantly altered during the diagenic and catagenic stages which result in the formation of kerogen (Degens, 1969; Johnson and Calder, 1973; Redding et al., 1980; Rigby et al., 1981; and Lewan, 1983). Sowever, there is some discrepancy as to whether thermal processes alter the isotopic composition when the kerogen reaches the metamorphic stage (McKirdy and Powell,1974; Hoefs and Frey, 1976; Chung and Sackett, 1979;

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72 Peters et al., 1981; Simoneit et al, 1981 and Jenden et al., 1982) The isotopic composition of the organic residue for sample 36650-1 was analyzed by pyrolyzing it from 12 to 160 hours. These results, presented in Table 9 and Appendix 4, show the residue becomes systematically lighter with increasing time. After 160 hrs., the residue is 2.9/oo lighter than the parent carbon. The loss of carbon with heavy o13c values ranges from 27.6% after 12 hours, to 43.6% after 160 hours. The enrichment in carbon-12 in the residual carbon fraction may be due to an initial loss of isotopically heavy higher molecular weight hydrocarbons and a isotopically heavy C02 Discussion loss of The time series experiments, on the least and most mature samples, have cast some doubt on the mechanism which determines the isotopic composition of methane in these maturation studies. The PCM model is based on kinetic isotope effects as the mechanism which determines the isotopic methane values, therefore, the possibility of exchange reactions would negate the use of the A 13c parameter as a maturity indicator. Several explanations have already been advanced negating the presence of an exchange mechanism. To reiterate, a high correlation between A13c and H/C ratios refutes the existance of exchange between methane and carbon dioxide. If exchange

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73 occurs, the methane composition would be affected by the isotopic composition of the co2 which the time series results showed to be fairly constant once other hydrocarbons were consumed. Exchange between methane and the residual organic carbon is highly unlikely since the residual carbon is so complex that any exchange reactions between these two fractions would be extremely slow. The time series experiments produced some unusual results which do not need to be explained in terms of an exchange mechanism. Kinetic isotope effects and isotope heterogeneity are the likely processes which determine the methane carbon values. The high correlation coefficients between CMR and h. 13c with the H/C ratios indicate that each of these factors independently can be used as maturity indices. The CMR is an indicator of the amounts of hydrogen and carbon functional groups available for methane production and their decrease in availability with increasing thermal stress. This is shown by the shift in the MPC pool toward lower carbon mole ratios. The CMR is also a better measure of the H/C ratios since it measures the actual amounts of hydrogen and carbon used in methane production. The h.13c is an indicator of the decrease in fractionation with increasing time and temperature. With increasing maturity, the MPC pool becomes depleted in carbon-12 and fractionation decreases as the c-c bonding becomes stronger. Another factor which became evident is the carbon-13 heterogeneity found in the more thermally altered

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74 samples. Carbon-13 heterogeneity results in methane values being heavier than the initial total carbon. When the CMR and !J. 13c are plotted against each other (figure 16) we find a correlation of -0.82 (99% confidence limit); a decrease in the amount of methane to carbon is accompanied by an increase in carbon-13 content. The regression line represents the H/C ratios along the maturity pathway of this suite of samples. There appears to be a break in the data; samples with a high CMR and lighter values and H/C ratios higher than 1.0, and samples with a low CMR, methane values as heavy or heavier than the parent carbon, and H/C ratios lower than 0.9. This same break is shown in figures 11 and 12. Referring to figure 8, H/C ratios between 1.5 to 0.5 fall in the catagenic stage for type II kerogen evolution. Catagenic reactions result in aromatization and polycondensation of the organics. Previous results from an aromatic and aliphatic model compound (see previous chapter) showed formation of less and heavier methane from the aromatic fraction than from the aliphatic compound. The observed break in the data could be due to a depletion in the amount of methane formed from aliphatic-type functional groups which results in a greater importance of methane from aromatic-type groups. This break can also be a result of a lack of 8amples with H/C ratios in the range from 0.9 to 0.8. Based on the Bakken shale results shown in figure 16, we can illustrate the relationship between H/C ratios and the model parameters, and the potential of these samples for

PAGE 88

-3 -2 1 0 tw) 0 -
PAGE 89

76 methane generation given sufficient time and temperature. Referring to figure 16, a sample with an H/C = 1. 2 would represent the MPCtoday pool. Fitting the values for CMR and ll13C for a sample with H/C = 1.2 we can have an indication of its potential for methane generation. This model measures the structural changes of the kerogen with increasing thermal stress and gives us information into the changes in the elemental composition of the kerogen with increasing maturation. Cape Verde Rise Kerogens Kerogens from organic-rich Cretaceous black shales from the Cape Verde Rise, DSDP Leg 41,. Site 368, were analyzed in this study first to determine the effects of metamorphism on the o13c of organic carbon and second to provide additional confirmation of the PCM method. This site is characterized by organic-rich Cretaceous black shales which have been intruded by hot diabase sills during the Miocene. Kerogen samples were isolated by Chevron Oil Research Company using the standard procedures (see Peters et al., 1983). A general description of the samples is presented based on the site reports of Lancelot et al. (1977). The Cape Verde Rise is located at 17.30.4'N, 21.2'W and a water depth of 3366m (see figure 17). Site 368 was drilled on the southwest flank of the Cape Verde Rise which

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35 25 I N . . . . . . ..... . : : I .... AFRICA 77 Figure 17. Location of the Cape Verde Rise samples from DSDP Leg 41, Site 368

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78 is located northeast of the Cape Verde Islands. Seismic data suggests this rise is an accumulation of pelagic sediments during most or all of the Tertiary and Late Cretaceous. The only thermally mature samples at this site, are those associated with the igneous intrusive rocks. Two thin, fine-grained basic intrusive sills occur in cores 60-2 (0-10cm) and 60-4 (75-85cm); one coarse-grained diabase sill occurs from sections 60-5 to 62-3 for a total thickness of 12.5 meters. The contact between the uppermost thin sills and black shales is sharp with little apparent effect due to baking. The sills intruded these shales during the Miocene when there was maximum eruptive volcanic activity on the Cape Verde Islands. The intrusion of the sill has a strong metamorphic effect on adjacent sediments (Kendrick et al., 1977). The thermal effect of the intrusives decreases rapidly with distance away from the sills as noted in the % C, HC and Ro distributions (Baker et al., 1977; Kvenvolden, 1977). A lower thermal gradient is found above than below the sill (Peters et al., 1983). The Cretaceous black shales have been described as high potential source rocks and identified as aliphatic and alicyclic-rich Type II kerogens commonly found in marine organic matter deposited in reducing environments (Deroo et al., 1977; Dow, 1977). The low content of humic compounds suggests low original content of continental organic matter. Additional evidence for a marine origin of these shales was found by Erdman and Schorno (1977) based on the lipid c513C

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79 compositions (range -28.6 to -23.7joo). Interstitial gases in the sediments near the intrusives are derived from thermal decomposition of organic matter and biogenic sources (Baker et al., 1977). Methane is the predominant gas and becomes isotopically heavier with depth (-65 to -55joo). The c1-c4 hydrocarbon concentrations increase with proximity to the sill (Baker et al., 1977; Doose et al., 1977). The samples analyzed include sub-sections from cores 58, 59, 60 and 62 shown in figure 18. Cores 58 and 59 are cyclic interbeds of green and olive silty claystone of late Cretaceous age. Cores 60 and 62 are dark gray to black shales with interbedded diabase sills and a high content of organic debris. The samples analyzed have a narrow range of maturities as indicated by the atomic H / C ratios and vitrinite reflectances (Table 10) than 0 5 and vitrinite Samples with atomic H/C ratios lower reflectances higher than 2 % a r e classified as metamorphic (Tissot and Welte, 1 984). These t w o parameters negatively correlate very well with each other (r= -0.94). The reflectance values shown in T able 10 are somewha t higher than those reported by Peters et al. (1978) and Simoneit et al. (1981) but within the range of those r eported by Dow (1977).

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58 so 62 94SALT ; N TRUSION BLACI( SHALE S ANDS CARBONATE O>YRITE f't Figure 18. Lithologic column showing the location of samples 58, 59,60 and 62 in the Cretaceous black shale sequence from Site 368, DSDP Leg 41 (after Simoneit et al., 1981) 80

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81 Table 10. Data for the analysis of the Cape Verde Rise kerogens from DSDP Leg 41, Site 368 (after Peters et al., "1983) CORE SECTION INTERVAL (em) Above major sill 58-5 ( 1) 13-18 59-3(31) 106-107 60-1 ( 2) 104-110 Minor sill 60-2(47) 127-128 60-3(33) 17-18 60-3 ( 41) 75-76 60-3(5) 89-93 60-4 (35) 39-40 60-4(39) 70-71 Minor sill 60-4(44) 110-111 60-5(34) 24-25 Below major sill 62-3(57) 100-102 62-4 (52) 60-61 DEPTH BELOW SEABED (m) 927.5 947.1 950.5 952.2 952.6 953.2 953.4 954.4 954.7 955.1 955.8 972.6 973.7 DISTANCE FROM SILL (m) +29.1 +9.5 +6.1 +4.4 +4.0 +3.4 +3.2 +2.2 +1.9 +1.5 +0.8 -0.7 -1.8 ELEMENTAL ANALYSIS %C %H 37.2 3.56 71.8 7.30 29.4 1.60 58.2 2.70 61.7 2 .06 48.2 1.53 65.0 1.76 42.7 1.23 71.1 1.36 43. 4 0.94 81.3 0.82 65.6 3.67 ATOMIC RATIO H/C R o % 1.15 0.63 1.21 0.8 7 0.65 1.57 0.56 2.12 0.40 2.78 0.38 3.04 0.33 3.12 0.35 3.84 0.23 4.16 0.26 4.26 0.12 4.16 3.33 0.67 1.43

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82 Organic carbon Content The organic carbon content is high and variable ranging from 22.3 to 74.9%. The large range in oc values is a result of depositional variations rather than a thermal effect due to the intrusives (Dow, 1977; Simoneit et al., 1981). There is some discrepancy as to whether thermal processes alter the isotopic composition when the kerogen reaches the metamorphic stage (Barker and Friedman, 1969; Baker and Claypool, 1970; McKirdy and Powell, 1974; Hoefs and Frey, 1976; Chung and Sackett, 1979; Simoneit et al., 1981; Peters et al., 1981 and Jenden et al., 1982). Shown in figure 19 and Table 11 is the change in the kerogen o13c-oc values with depth. Figure 19 includes o13c-oc values from Simoneit et al. ( 1981) The o13c0c values range from -29.5 to -20.9joo With increasing proximity to the sill, the o13c-oc values become more carbon-13 enriched: therefore, the o13c-oc values may be thermally-controlled rather than source-controlled. The heavier values with increased metamorphism can be due to preferential loss of carbon-12 due to a kinetic isotope effect (Hoefs and Frey, 1973; Simoneit et al., 1981), or the thermally labile fraction of the kerogen m a y be composed of lighter carbon (Simoneit et al., 1981). No correlation was found between %C and o13c-oc which would not be expected since the organic carbon content is a factor of the depositional environment, and the o13c-oc values

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0 CD E :I: li: w c 920 930 940 950 960 970 980 83 !:::. Sinonelt et aJ. 1981 * This paper *6 ** -28 -27 -26 -25 -24 -23 -22. -21 -20 013 C-ORGANIC CARBON Figure 19. Change in c513C-TOC with proximity to the sill of the Cape Verde Rise kerogens. Hatched area represents the sill. Results include data from Simoneit et al., 1981

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84 Table 11. Results from the pyrolysis, at 600C (for 120 hrs.) of the Cape Verde Rise kerogens. SAMPLE DISTANCE oc 613C 13 CMR HMR 6 CCH4 % oc Above major sill 58-5(1) 29.1 31.4 -24.2 -26.3 0.083 0.243 -2.1 59-3 (31) 9.1 74.2 -26.8 -28.5 0.207 0.678 -1.7 60-1(2) 6.1 22.3 -24.9 -29.2 0.022 0.102 -4.3 Minor sill 60-2(47) 4.4 49.9 -24.6 -24.9 0.022 0.087 -0.3 60-3(33) 4.0 45.7 -25.3 -28.5 0.026 0.195 -3.2 60-3(41) 3.4 36.7 -23.9 -28.0 0.012 0.094 -4.1 60-3(5) 3.2 44.8 -24.3 -28.8 0.013 0.114 -4.5 60-4(35) 2.2 40.8 -23.2 60-4(39) 1.9 55.7 -22.9 -24.7 0.006 0.080 -1.8 Minor sill 60-4(44) 1.5 50.4 -20.9 60-5(34) 0.8 74.9 -24.0 Below major sill 62-3 (57) -0.7 29.3 -26.8 -28.2 0.016 0.232 -1.4 62-3 (52) -1.8 54.0 -27.5

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85 appear to be the result of the high thermal stress. Another aspect which becomes apparent is the differences in the correlations between H/C ratios and vitrinite reflectance with the o13c-oc. A better relation is found between o13c-oc and the H/C ratios than with vitrinite reflectance values. Peters et al. (1983) found that over short periods of time and high temperatures, vitrinite reflectance responds faster than the generation and alteration of bitumens. Rapid heating of organic matter results in changes which are not directly comparable to those of slow burial metamorphism (Peters et al. 1983). H/C ratios are dependent on the changes in the kerogens due to loss of hydrocarbons during maturation and may be a better maturity indicator for these data sets than vitrinite reflectance. Although, the high correlation between these two parameters indicates a close relationship between them, H/C ratios might be more sensitive to the rapid changes due to contact metamorphism. In general, the isotopic composition of the organic carbon appears to be thermally-controlled. Therefore, no conclusions as to its source can be based on its isotopic composition. The PCM can not be applied to these samples since the model presented in figure 9 assumes the methane precursor carbon pool is source-controlled. Data by Dow (1977), Baker et al. (1977) and Erdman and Schorno (1977) all point to a marine origin for the Cape Verde Rise kerogens. These shales were formed during the anoxic

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86 events in the middle Cretaceous (Tissot et al., 1980). These events were a result of the position of the continents, the warm equable climate, a large sea level transgression which resulted in warm, shallow seas with high productivity and the silled basin geometry which resulted in stagnation (Arthur and Schlanger, 1979). Time Series Experiment Time series analysis was conducted on sample 31, the most immature sample, to determine the pyrolysis time necessary for "exhaustive" methane formation. This sample was pyrolyzed for a period from 24 to 288 hours at 600C The products were determined by gas chromatography (GC) and the methane isotopic composition measured. The GC results are shown in figure 20 (additional data shown in Appendix 6 and 7) The amounts of ethane and ethylenejacetylene rapidly decrease after 48 hours of pyrolysis time. These hydrocarbons disappear after 120 hours, apparently degrading to methane and carbon. Based on these results, the pyrolysis time was set for 120 hours. The carbon mole ratio (CMR) remains constant after 96 hours indicating no more methane is formed relative to the parent carbon. The conversion of carbon in kerogen to methane peaks at 29.4%. The isotopic composition of the methane became increasingly heavier with pyrolysis time reaching a plateau

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'#. 40 0 4 8 12 16 20 24 28 32 4 t( 600C )-hrs 100 90 J: 0 '#. <( UJ 0: <( 80 70 50 20 40 80 120 160 200 At (600C)-hrs Figure 20. Gas chromatography results from the pyrolysis of Cape Verde Rise kerogen sampl e 31
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88 after 120 hours. Methane remains 1.6 lighter than the parent carbon as shown in figure 21. Formation and Isotopic Composition of C02 The amounts of carbon dioxide measured, presented in Appendix 8, were generally greater than the amounts of methane, contrary to expectations. With this high range in maturities, little or no C02 should form due to the absence of oxygen-containing functional groups. During the maturation of organic matter, oxygen is rapidly depleted (Tissot and Welte, 1984). Data by a myriad of investigators contradicts the premise that highly unreactive oxygen-containing functional groups would be present at these high levels of maturity. These samples are mainly powdery and graphitic in nature and have high absorptive properties (Sackett, pers.comm.). Exposure to the atmosphere could result in absorption of oxygen and even reaction with the carbon to form oxygen-containing functional groups (VanVleet, pers.comm.). Even though these samples are heated under vacuum prior to sealing, all oxygen may not have been driven off. Therefore, the amounts of C02 measured in this experiment may be invalid. The isotopic composition of the carbon dioxide was highly variable ranging from -28.3 to -23.6/oo with no relationship to maturity. Since the carbon dioxide appears not to be an inherent part of the kerogen, these isotope values will not be

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...,. lj 0 (t') LC) -34 -35 -36 -37 -38 12 24 48 96 120 160 TIME (hours) Figure 21. Change with time of the isotopic composition of methanefrom the pyrolysis of Cape Verde Rise kerogen 31 ():) \0

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90 further discussed. Formation of Methane The metamorphic stage of most of these samples, as indicated by the low atomic H/C ratios, signifies that not only does the methane precursor carbon pool become depleted with proximity to the sill, but less hydrogen is available for methane formation due to increase aromaticity of the kerogen molecule. Figure 22 shows a sharp decrease in the conversion of kerogen carbon to methane, with increasing maturity; from 20.7% at an H/C ratio of 1.21, to 0.6% when the H/C ratio is 0.23. A trend toward increased stability of the kerogen can be inferred from the decrease in the carbon mole ratio. In the metamorphic stage, less than 1% of the kerogen carbon is reactive. Based on the hydrogen content data presented in Table 10, the hydrogen mole ratio (HMR = CH4/H) or amount of hydrogen converted to methane, can be computed. The HMR values range from 0.753 for the least mature sample, to 0.080 for the most mature one. The carbon mole ratio and hydrogen mole ratio are indices to the amounts of "reactive" hydrogen and carbon available in the kerogen. These ratios decrease with increasing maturity due to a depletion in the precursor groups available and increased aromatization of the kerogen. Better correlations were found between the H/C ratios and CMR than between

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91 0.25 0 20 .. 0 15 0.1J .. 0.05 0.0 I ,. l I I I I I I I I t 0 0 2 0.4 0 6 0.8 1 0 1. 2 H/C Figure 22. Correlation between the carbon mole ratio ratio (CMR) and H / C ratios from the pyrolysis of Cape Verde Rise kerogens

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92 vitrinite reflectance and CMR. Isotopic Composition of Methane The isotopic composition of methane ranged from -29.2 to -24.7joo. There was no relationship between these values and the maturity of the samples. Generally, with increasing maturity, the isotopic composition of the methane becomes more carbon-13 enriched due to the kinetic isotope effect. Also, since the parent carbon o13c values decrease with proximity to the sill, the methane values would be expected to follow this trend. These kerogen methane results again cast some doubts on the fractionation mechanisms since a trend toward heavier values with increasing maturity should be observed. Previous results from the Bakken shales negate the presence of exchange reactions as the mechanism which determines the isotopic composition of methane at 600C (see the discussion in the Bakken shale section). The Bakken shale methane results were explained in terms of kinetic isotope effects and C-13 heterogeneity. The possibility of isotope exchange between methane and carbon dioxide in the Cape Verde Rise kerogens cannot be discarded since there is a trend of heavier methane values associated with the lighter C02 values and vice-versa. Exchange between these two gases would lead to an enrichment of carbon-13 in the methane and a subsequent depletion of carbon-13 in the co2 As the carbon dioxide predominates, its

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93 isotopic composition upon formation should be controlling. However, the kerogen time series experiment showed no evidence of an exchange mechanism since the methane values reached a plateau after 96 hours. Several reasons can be discarded as to the lack of a systematic change in the methane values with increasing maturation. The Bakken shale results were based on methane generation from the whole shale sample whereas the Cape Verde Rise kerogens have been subjected to intensive acid treatment which could have affected the carbon matrix and the subsequent methane isotopic compositions. However, there is a trend in the organic carbon isotopic composition toward heavier values, therefore the methane formed from this carbon should follow a similar trend. The presence of oxygen adsorbed onto the kerogen appears to react to form carbon dioxide but it may also result in oxidation of the methane to form additional carbon dioxide. However, if oxidation of methane were to occur, the CMR would also be affected whereas we see a high correlation between CMR and H/C ratios. Due to the small sample size, the methane isotopic measurements could not be repeated. In conclusion, these methane values may be a result of innacurate measurement of its isotopic composition or the effects of some other processes. These values will be disregarded throughout the rest of the discussion.

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94 Formation and Isotopic Composition of the Organic Residue The organic residue was combusted and analyzed for all the previous samples. The data are presented in figure 23 and Appendix 9. The isotopic composition of the residue was generally heavier with proximity to the sill, ranging from -2 7 6 to -2 o. 5 o o Generally, the isotopic composition of the residue was heavier than the parent carbon which points to a loss of light carbon in the organic carbon fraction. This difference decreases with increasing maturity. Since a decrease in the CMR is accompanied by a decrease in the difference between o13c-residue and o13c-oc, it is likely that the shift of c513c-oc toward heavier values with increasing maturity is due to a loss of isotopically lighter methane. This trend does not support the results found in the Bakken shale time series analysis. The Bakken shale time series showed the residue becoming more carbon-12 enriched, relative to the parent carbon, with increasing time. This was an immature sample, compared to these kerogens, with enough oxygen-containing functional groups present to generate carbon dioxide and therefore result in a loss of heavy carbon. The organic residue data for the Cape Verde Rise kerogens supports field and laboratory data which indicates the residual fraction becomes heavier with time. The CMR data supports the premise that this shift is due to loss of isotopically light carbon in the form of methane.

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......... (IJ '-(l) 1 :t: t-a. w c 920 930 940 950 960 970 980 ORGANIC CAF130N 6. ORGANC RESDl.E *6 6 *6 95 -28 -27 -26 -25 -24 -23 -22 -21 -20 Figure 23. Change in the cS13c of the organic carbon and organic residue of the Cape Verde Rise kerogens with proximity to the sill

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96 Summary and Conclusions The Cape Verde Rise Cretaceous black shales are metamorphosed, Type II kerogens, high in organic carbon content. The organic carbon content is source-controlled, but its isotopic composition appears to be thermally-controlled. With proximity to the sill, the o13c-oc values are heavier due to a loss of lighter carbon in the form of methane. This is shown by a decrease in the carbon mole ratio, with increasing maturities, associated with a decrease in the difference between the o13c-oc and the c513C-residue. The amount of reactive carbon available decreases sharply in the highly altered samples. The amount of reactive carbon in the metamorphosed samples is between 1 and 2%. reactive hydrogen is around 8%. The amount of The isotopic composition of the methane and the carbon dioxide is highly variable. The source of the carbon dioxide is probably adsoption of oxygen during handling, and not inherent in the kerogen as shown by the amounts of C02 measured. The isotopic composition of methane does not follow an expected trend of heavier values with proximity to the sill even though the carbon mole ratio decreases sharply, and the c513C-OC values become heavier. The problem appears to be in the isotopic determination of the methane values. Isotope exchange between methane and carbon dioxide is not a determining factor since time series analysis of kerogen 31 shows the methane isotopic composition reaching a plateau

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after 120 hours. The pyrolysis carbon isotope method is concept of a methane precursor carbon 97 based on the pool which characteristics are determined by the source of organic matter. The A13c parameter cannot be used in these samples. The CMR is a useful parameter and indicative of the amount of "reactive" kerogen carbon available and can be used as such for these samples. In addition, a hydrogen mole ratio can be calculated to measure the amount of "reactive" hydrogen available for methane formation. These two parameters, CMR and HMR, decrease with increasing maturities.

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98 CHAPTER 5 CARBON ISOTOPE EXCHANGE AT 600C Earlier Work Stable carbon isotopes have proven to be a successful technique in distinguishing between the different sources of methane and the mechanisms of formation. The isotopic composition of the methane gas will be determined by kinetic isotope effects, carbon isotope exchange between the methane and other carbon containing species present and by carbon-13 heterogeneity. Carbon isotope fractionation during methanogenesis is determined by kinetic isotope effects rather than isotope exchange between CH4 and C02 (Whiticar et al., 1986). The rate of carbon isotope exchange in shallow sediments is too slow to permit isotope re-equilibration of the methane that has formed with the C02 reservoir (Sackett and Chung, 1979; Giggenbach, 1982) The carbon isotopic composition of thermogenic methane will be determined by time, temperature of formation and the nature of the parent organic matter (Sackett et al., 1968). Kinetic isotope effects rather than isotope exchange

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99 equilibrium will determine the amount of isotopic fractionation of thermogenic methane formed at temperature below 500C. studies by Sackett and Chung (1979) and Lyon and Hulston (1984) have shown that exchange at temperatures below 500C is so slow that it can be considered negligible. However, Harting and Maas (1980) have shown exchange occurring between CH4 and C02 at temperatures above 500C. Therefore, sediments which have been subjected to high temperatures such as metamorphic or subducted sediments, or those that have been intruded by sills, could show some exchange between methane and other carbon bearing gases present. Results from the Bakken shale and the kerogen pyrolysis experiments have raised the question of whether isotope exchange reactions are controlling the isotopic compositions of CH4 co2 and the organic carbon at 600C. Isotopic equilibrium at 600C results in methane values 11.5joo lighter than C02 (Lyon and Hulston, 1984). In the time series experiment with Bakken sample 38423-15, results show CH4 becomes progressively lighter and C02 heavier with a difference in these two fractions, of 18.3/oo after 120 hours. This is equivalent to equilibrium temperatures between 370-400C (extrapolated value from Bottinga, 1969). The results for the kerogen time series analysis do not support the premise of exchange reactions since the methane values remain constant after 120 hours. The C02 fraction was not measured.

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100 The high correlation between A13c and H/C ratios for the Bakken shale samples appears to negate the presence of isotope reactions and point rather to a kinetic isotope effect. If exchange reactions do take place, the most abundant fraction will dominate the equilibrium reactions. In this case, the most dominant fraction is the organic residue and co2 the least abundant. The structure of the residual fraction is highly polymerized and stable and may not facilitate exchange, therefore methane may be the controlling gas. If this were the case, we should not find such a high correlation between CH4 and H/C ratios, since the exchange reactions are a factor of temperature, and not the maturity of a sample. A reason for postulating exchange reactions is the positive methane values found relative to the parent carbon in the more mature samples. This can be explained by exchange reactions or by a restructuring of the methane carbon pool with enrichment in carbon-13 as described in the model. The lack of correlation between A 13c and the H / C ratios for the cape Verde Rise kerogen samples can not be used as an indication of processes affecting the isotopic composition of methane. In order to determine if carbon isotopes can be used to study the origin and mechanism of methane which is generated from deeply buried sedimentary rocks, we must first be able to determine the mechanism of fractionation, effects andjor carbon isotope equilibrium, kinetic isotope so that we can distinguish between thermogenic methane generated from organic sedimentary rocks which have been subducted and deep source

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101 abiotic methane. If o13C-CH4 values are determined by isotope exchange equilibrium then these values can no longer be used to distinguish as to methane sources since they will be controlled by the more abundant carbonaceous species present. In order to determine the fractionation processes occurring at 600C, a series of experiments were conducted, to determine if exchange occurs between CH4 and C02 by heating equal amounts of methane and co2 for a period from 12 to 504 hrs. (see experimental section for description of procedures) Results and Discussion Figure 24 shows the shift in methane and C02 values (additional data shown in Appendix 10). Methane values were heavier by 1 .8joo and C02 values lighter by 1.4joo after 504 hrs (std. dev. .2joo). The shift in the isotopic values of CH4 and C02 was not noticeable until after 120 hours. Prior to this time period, the values remain fairly constant. A problem encountered in these experiments was the lack of 100% yield. The yields varied from 35-90% (mean= 75%). This can be due to a number of reasons such as loss of gas during the initial injection into the system, incomplete recovery of the gases injected during the freezing procedure, or innacurate reading of the amount of gas in the manometer. Results in Table 12 show reproducibility in the 613c values of 0. 2 o o even though the amount of gas recovered was not

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ro : ...... c: 0 C') ,.-c..o I Q) E 0 C') ,.-c..o ---0 C') ,.<1 6 4 I /). 13CcH4 2 I 0 I 11.5 %o I _!!!: 2 I /). 13Cco2 4 6 / 0 Figure 24. 100 200 300 400 500 00 TIME (hours) Carbon isotope exchan3e between equal amounts of methane and carbon dioxide at 600 C 1-' 0 t\)

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103 Table 12. carbon isotope exchange between equal amounts of methane and carbon dioxide at 6ooc. TIME (hrs) o13c-co 2 o13C-CH 4 aC02-CH4 dC02-CH4 12 -13.2 -37.7 1. 025 24.4 24 -13.3 -37.6 1.025 24.3 48 -13.1 -37.5 1.025 24.4 72 -13.1 -37.5 1.025 24.4 96 -13.2 -37.6 1.025 24.4 120 -13.3 -37.5 1.025 24.2 210 -13.4 -36.9 1.024 23.5 504 -14.0 -36.4 1. 023 22.4

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104 constant. This suggests the problem must occur after the gases have been heated since theory decrees that exchange reactions are controlled by the relative amounts of gas present. However, the differences in volume between duplicates might not be large enough to cause a difference in the results. A 13 d f d 13 13 u C 1s e 1ne as o Cc02 o CcH4 Heating of these two gases for 504 hours results in a Aco2-CH4 difference of 2 2 4 o o Based on the curves shown in figure 2 4 we can estimate the half-life of the reaction to be about 1080 hrs assuming an equilibrium value of 11.5/oo (based on the Tables by Lyon and Hulston, 1984). The shift in the CH4 and co2 values appears to occur after 120 hours which is the pyrolysis time for all the previous samples analyzed. Therefore, it appears that isotope exchange between CH4 and C02 does not affect the pyrolysis results. The large differences in the C02 and CH4 fractions in these pyrolysis experiments point to kinetic isotope reactions being the main process controlling the isotopic composition in the formation of these gases. These results suggest that carbon isotope values of methane, formed under high temperature regimes, may be determined by isotope exchange if any carbon dioxide is present. Isotope exchange would result in heavy methane values. This would make it difficult to differentiate between methane produced from subducted kerogen and that produced by reactions deep within the earth using stable carbon isotopes.

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105 CHAPTER 6 CONCLUSIONS The objectives of this research, delineated on page 18, are to evaluate the effects of metamorphism on the 613c of organic matter, test Sackett's {1984) maturation model, set limits of metagenically derived methane and determine exchange between methane and carbon dioxide at 600"C. These objectives were achieved by: 1. studying the amounts and isotopic composition of two model compounds, n-octadecane and decacyclene. 2. applying the pyrolysis carbon isotope method to two suites of samples with varying maturities, Bakken shales of North Dakota and kerogens from the Cape Verde Rise, and analyzing the various carbonaceous species present {CH4 C02 and the organic residue). 3. heating equal amounts of CH4 and C02 at 600"C for various time periods to measure exchange between these two fractions. The isotopic composition of organic carbon was studied to resolve the discrepancy as to whether the organic matter becomes lighter or heavier in the metagenic stage of maturation. This was accomplished by analyzing the Bakken shales which fall within the catagenic stage of alteration and

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106 the Cape Verde Rise kerogens which are mostly metamorphozed. The isotopic composition of the organic carbon is source-controlled for the Bakken shale samples since there is no trend in the c513c-oc with increasing maturation. The organic residue tends toward lighter carbon values with increasing heating time. This is probably due to the presence of sufficient oxygen-containing functional groups present in the shales to generate carbon dioxide and therefore result in a loss of heavy carbon. However, the c513c-oc values for the Cape Verde Rise kerogens appear to be thermally-controlled. Studies of the organic residue of the Cape Verde Rise metagenic kerogens, show an enrichment in carbon-13 with increasing maturation. The heavier residual values are due to a loss of isotopically light methane. With increasing maturation, the CMR decreases and the difference in isotopic compositions between the organic carbon and its residue decrease. The organic carbon itself becomes heavier with increasing metagenesis. The pyrolysis of model compounds yields information on the differences in amounts and isotopic composition of methane generated from aliphatic and aromatic-type compounds. Noctadecane and decacyclene are assumed to be representative of these carbon-types in kerogen. Generally, while most methane is generated from aliphatic-type compounds, methane from aromatic-type carbon is enriched in carbon-13. Methane from both compounds approached the isotopic composition of the parent carbon with increasing pyrolysis time.

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107 The fractionations, measured as isotopic compositions of methane relative to the parent compound, for n-octadecane (600), decacyclene (5oooc), decacyclene {600C), Bakken shale 36650-1 {600C-least altered sample), and Bakken shale 38423-15 (most altered sample) are 0 -16.5joo, 0 -9.7 1 o o, -10.7joo, -9.5joo and -3.1/oo respectively. There are two factors involved in determining the fractionation of methane; temperature dependence of the kinetic isotope effect and carbon-carbon types. There appears to be no temperature dependence on the kinetic isotope effect between 500 and 6 0 0 o C. These fractionations decrease as the methane precursor pool decreases therefore these differences are due to differences in the carbon-carbon bond types. This becomes apparent in the fractionation of the most altered sample. The pyrolysis of the Bakken shales, shown in figures 11 and 12 as CMR and 11. 13c vs. H/C ratios, show a break in the curve around H/C = 0.9. This break is indicated by methane isotopic composition heavier than the parent carbon and a shift toward lower CMR values. This break could be fortuitous, or it could indicate a shift toward methane generation from aliphatic to aromatic-type carbon. The Bakken shale time series experiments on the least and most mature samples, and the positive Bakken shale methane values, have cast some doubt as to the mechanism which determines the isotopic composition of CH4 ; kinetic isotope effects or isotope exchange between methane, carbon dioxide and the organic residue. In the Bakken time series of sample

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108 35423-15, results show methane progressively lighter and carbon dioxide heavier, with a difference between these two fractions of 18.3/oo. This is equivalent to an equilibrium temperature between 370-4ooc. Isotope exchange reactions between methane, carbon dioxide and the organic carbon residue were discarded for the following reasons: 1. A high correlation between A 13c and H/C ratios for the Bakken shales since exchange reactions are dependent on temperature and not the maturity of the sample. 2. Exchange between methane and the residual carbon is highly unlikely since the residual carbon is so complex that any exchange between these two fractions would be extremely slow. 3. The time series experiment of the least mature cape Verde Rise kerogen sample showed the methane isotopic composition reaching a plateau after 120 hours. Also, the isotopic composition of the least mature Bakken shale sample was shown to be fairly constant once other hydrocarbons were consumed. 4. Experiments on the extent of exchange between methane and carbon dioxide showed that a shift in values occurs after 120 hours which is the.pyrolysis time for all samples analyzed (excluding the time series experiments). The half-life of the exchange reaction between CH4 and C02 at 6oo c is about 1080 hours assuming an equilibrium value of 11.5/ o o (based on Tables by Lyon and Hulston, 1984). Kinetic isotope effects and isotope heterogeneity are the processes which determine methane values. With increasing

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maturity, methane carbon-13 109 new carbon-13 distributions develop, leaving the precursor functional groups highly enriched in (Chung, 1976; Chung and Sackett, 1980). Regretfully, the kerogens which are mostly metamorphosed, did not yield reliable values and therefore no generalizations can be made as to the occurrence of 13c heterogeneity in other thermally altered samples. The Bakken shale methane results show that heavy methane values can be generated from highly altered sediments. In addition, the exchange experiments suggest that the isotopic composition of methane formed under high temperature regimes, such as from highly metamorphosed and subducted rocks, may be determined by isotope exchange if any C02 is present. This would result in heavy methane values. It therefore becomes difficult to use carbon isotope compositions of methane to distinguish between thermogenic and mantle methane without any other supporting evidence. The pyrolysis of the aliphatic carbon (n-octadecane) yielded the theoretical amount of methane after 48 hours (CMR = 0. 53) while the pyrolysis of the aromatic carbon only yielded 41% of the theoretical amount. The amount of kerogen carbon converted to methane, in the Bakken shales and the Cape Verde Rise kerogens, varied from 29% for the least thermally altered sample to 1-2% for the most thermally! altered samples. The CMR is an indicator of the amount of carbon and hydrogen functional groups available for methane production and their decrease in availability with increasing thermal

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110 stress. It is also a better measure of the H/C ratios since it measures the actual amounts of hydrogen and carbon used in methane production. The pyrolysis carbon isotope method measures the structural changes of the kerogen with increasing thermal stress and gives us information into the changes in the elemental composition and stability of the kerogen with increasing maturation. The CMR is an indicator of the amount of methane precursor functional groups available, and the a13c is an indicator of the decrease in fractionation with increasing time and temperature. The high correlation coefficient between CMR and A13c with the H/C ratios, for the Bakken shales, indicate that each of these factors independently can be used as maturity indices. This is not the case for the Cape Verde Rise kerogens since the PCM assumes the characteristics of the MPC pool are controlled by the source of organic matter. The CMR does yield information on the "reactivity" and relative maturity of these samples. This method is particularly attractive because it can be used on whole rock samples and on kerogens which do not contain vitrinite or palynomorphs.

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111 CHAPTER 7 LITERATURE CITED Abelson, P.H. and T.c. Hoering. 1961. Carbon isotope fractionation in formation of amino acids by photosynthetic organisms. Proc. Natl. Acad. Sci. 47: 623-632. Alekseyer, F.A., V.S. Lebeder and T.A. Krylova. 1972. Isotope composition of carbon in gaseous hydrocarbons and conditions for accumulations of natural gas. Internat. Geology Rev. 15: 300-308. Arneth, J.D. and u. Matzigkeit. 1986. Laboratory-simulated thermal maturation of different types of sediments from the Williston Basin, North America -Effects on the production rates, the isotopic and organo-geochemical composition of various pyrolysis products. Chem. Geol. 58:.339-360. Arthur, M.A. and s.o. Schlanger. 1979. Cretaceous "oceanic anoxic events" as causal factors in development of reefreservoired oilfields. Ainer. Assoc. Petrol. Geol. 63: 870-885. Baker, B.L. 1974. Generation of alkane and aromatic hydrocarbons from humic materials in Arctic marine sediment. Adv. org. Geochem. 6: 137-152. Baker, E.W., S.E. Palmer and W.Y. Huang. 1977. Intermediate and late diagenetic tetrapyrrole pigments, Leg 41: cape Verde Rise and Basin. In: Lancelot et al. ( eds. ) Initial Reports of Deep Sea Drilling Project Vol 41. pp. 825-838. Baker, E.W., W.Y. Huang, J.G. Rankin, J.R. Castano, J.R. Guinn, and A.N. Fuex. 1977. Electron paramagnetic resonance study of thermal alteration of kerogen in deepsea sediments by basal tic sill intrusion. In: Y. Lancelot, E. Seibold, et al.. Initial Repts. DSDP 41. u.s. Govt. Printing Office. pp. 839-847.

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112 Baker, D.R. and G.E. Claypool. 1970. Effects of incipient meta-morphism on organic matter in mudrock. Amer. Assoc. Petrol. Geol. Bull. 54:456-468. Baker, D.R. 1974. Pyrolysis techniques for source-rock evaluation. Am. Assoc. Petrol. Geol. Bull. 58(11): 2349-2361. Barker, F. and I. Friedman. 1969. Carbon isotopes in Pelites of the Preecambrian Uncompahgre Formation, Needle Mountains, Colorado. Geol. Soc. Am. Bull. 80: 1403-1408. Bernard, B.B. 1978. Light hydrocarbons in marine sediments. Ph.D. thesis. Texas A&M Univ. 144 pp. Bigeleisen, J. and M.G. Mayer. 1947. Calculation of equilibrium constants for isotope exchange reactions. J. Chem. Phys. 15: 257-261. Bottinga, Y. 1969. Calculated fractination factors for carbon and hydrogen isotope exchange in the system calcite-carbon dioxide-graphite-methane-hydrogen-water vapor. Geochim. Cosmochim. Acta 33(1): 49. Chung, H.M. and W.M. Sackett. 1980. during the pyrolytic formation carbonaceous materials. In: Carbon isotope effects of early methane from A. G. Douglas and J. R. Maxwell ( eds. ) Adv. Org. Press. Oxford. pp. 705-710. Geochem. 1979. Pergamon Chung, H.M. and W.M. Sackett. 1979. Use of stable carbon isotope compositions of pyrolytically derived methane as maturity indices for carbonaceous materials. Geochim. Cosmochim. Acta 43: 1979-1988. Chung, H.M. 1976. Isotope maturation of organic matter. University. College Station. fractionation during the Ph.D. Thesis. Texas A&M 162pp. Columbo, u., F. Gazzarrini, R. Gonfiantini, E. Tongiorgi, and L. Caflisch. 1969. Carbon isotope study of hydrocarbons in Italian natural gases. In: P.A. Schenck and I. Hevenaar (eds.). Adv. Org. Geochem. 1968. Pergamon Press. New York. pp. 499-516. Conkright, M.E. and W.M. Sackett. 1986. Stable carbon isotope evaluation of the contribution of terriginous carbon to the marine food web in Bayboro Harbor, Tampa Bay, Florida. Contrib. Marine Science 29: 131-139. Conkright, M.E., W.M. Sackett and K.E. Peters. 1986. Application of the pyrolysis-carbon isotope method for determining the maturity of kerogen in the Bakken shale.

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113 Advances in Organic Geochemistry 1985. Org. Geochem. 10: 1113-1117. Craig, H. 1953. The geochemistry of the stable carbon isotopes. Geochim. Cosmochim. Acta 3: 53-92. Craig, H. 1963. The isotopic geochemistry of water and carbon in geothermal areas. In: "Proc. Spoleto Conference on Nuclear Geology. E. Tongiorgi (ed.). Spoleto, Italy. pp. 17-53. Degens, E.T. 1969. Biogeochemistry of stable carbon isotopes. In: Organic Geochemistry. G. Eglinton and M.T.J. Murphy {eds.). Springer-Verlag. New York. pp. 304-329. Degens, E.T., M. Behrend, B. Gotthardt, and B. Reppmann. 1968b. Metabolic fractionation of carbon isotopes in marine plankton. II. Data on samples collected off the coast of Peru and Ecuador. Deep Sea Res. 15: 11-20. Deines, P. 1980. The isotopic composition of reduced organic carbon. In: P. Fritz and J. Fontes (eds.). Handbook of environmental isotope geochemistry. Vol 1. The terrestrial environments. A. Elsevier. pp. 329-406. DeNiro, M.J. and S. Epstein. 1978. distribution of carbon isotopes Cosmochim. Acta 42: 495-506. Influence of diet on the in animals. Geochim. Deroo, G., J.P. Herbin, J. Roucache, B. Tissot, P. Albrecht and J. Scheffle. 1977. Organic geochemistry of some Cretaceous black shales from sites 367 and 368, Leg 41, Eastern North Atlantic. In: Y. Lancelot, E. Seibold, et al . Initial Repts. DSDP 41. U.S. Govt. Printing Office. pp. 8 65-873. DesMarais, D.J., J.H. Donchin, N.L. Nehring and A.H. Truesdell. 1981a. Molecular evidence for the origin of geothermal hydrocarbons. Nature 292: 826. DesMarais, D.J. and J.M. Hayes. 1976. Tubecracker for opening glass-sealed ampoules under vacuum. Anal. Chern. 48:1651-1652. Doose, P.R., M.W. Sandstrom, R.Z. Jodele and I.R. Kaplan. 1977. Interstitial gas analysis of sediment samples from Site 368 and Hole 369A. In: Lancelot et al. (eds.). Initial Reports of Deep Sea Drilling Project Vol 41. pp. 861-864. Dow, W.G. and D.I. O'Connor. 1982. Kerogen maturity and type by reflected light microscopy applied to petroleum

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114 exploration. In: paleotemperatures. How to assess maturation and SEPM Short Course #7. pp. 133-158. Dow, W G. 1977. Kerogen studies and geological J. Geochem. Explor. 7: 79-99. interpretations. Dow, W.G. 1977. from leg 41: al. ( eds. ) Vol 41. pp. Contact metamorphism of kerogen in sediments Cape Verde Rise and Basin. In: Lancelot et Initial Reports of Deep Sea Drilling Project 815-816. Durand, B. 1980. Sedimentary organic matter and kerogen. Definition and quantitative importance of kerogen. In: B. Durand (ed). Kerogen. Technip. Paris. pp. 13-33. Erdman, J .G. and K.S. Schorno. 1977. Geochemistry of carbon: Deep Sea Drilling Project Leg 41. In: Lancelot et al. (eds.). Initial Reports of Deep Sea Drilling Project Vol 41. pp. 849-854. Evans, R.J. and G.T. Felbeck. 1983. High temperature simulation of petroleum formation. I. The pyrolysis of Green River Shale. Org. Geochem. 4: 135-144. Frank, D.J., J.R. Gormly and W.M. Sackett. 1974. Reevaluation of carbon isotope composition of natural methane. Bull. Amer. Assoc. Petrol. Geol. 58: 2319-2325. Frank, D.J. and W.M Sackett. 1969. in the thermal cracking of Cosmochim Acta 33: 811-820. Kinetic isotope effects neopentane. Geochim. Fuex, A.N. 1977. exploration. The use of stable isotopes in hydrocarbon J. Geochem. Explor. 7: 155-188. Galinov, E.M. isotopes. 1985. The biological fractionation Academic Press, Inc. Orlando. 261pp. of Galinov, E.M. 1974. Organic geochemistry of carbon isotopes. In: Advances in Organic Geochemistry 1973. B. Tissot and F. Bienner (eds.). Editions Technip. Paris. pp. 439-452. Galinov, E.M. 1969. Isotopic composition of carbon in gases of the crust. Internat. Geology Rev. 11: 1092-1104. Giggenbach, W.F. 1982. Carbon-13 exchange between C02 and CH4 under geothermal conditions. Geochim. Cosmochim. Acta 46: 159-165. Giggenbach, W.F. 1980. Geothermal gas equlibria. Geochim. Cosmochim. Acta 44: 2021-2032.

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115 Gilmour, I. and C.T. Fillinger. 1985. stable carbon isotopic analysis of sedimentary organic matter by stepped combustion. Org. Geochem. 8(6): 421-426. Giraud, A. 1970. chromatography to sedimentary rock. 439-455. Application of pyrolysis and gas chemical characterization of kerogen in Bull. Amer. Assoc. Petrol. Geol. 54: Gold, T and s. Soter. 1980. The deep earth-gas Scientific American 242:154-161. Gold, T. 1979. outgassing. Terrestrial sources of carbon and earthquake J. Petrol. Geol. 1:3. Gunter, B.D. and B C. Musgrave. 1971. New evidence on the origin of methane in hydrothermal gases. Geochim. Cosmochim. Acta 35:113. Harting, P. and I. Maass. 1980. Neve Ergebnisse zum kohlenstoff Isotope naustausch im system CH4 -C02 In: Mitteilungen zur 2. Arbeitstagung "Isotope in der Natur11 Leipzig. vol 2b: 13-24. Hoefs, J. Rocks. 1980. Stable isotope geochemistry. Minerals and Springer-Verlag. New York. 208pp. Hoefs, J. and M. Frey. 1976. The isotopic composition of carbonaceous matter in a metamorphic profile from the Swiss Alps. Geochim. Cosmochim. Acta 40: 945-951. Hoering, T.C. 1984. Thermal reactions of kerogen with added water, heavy water and pure organic substances. Org. Geochem. 5: 267-278. Hulston, J.R. and W.J. McCabe. 1962. Mass spectrometer measurements in the thermal areas of New Zealand, Part 2. Carbon isotope ratios. Geochim. Cosmochim. Acta 26: 399. Hunt, J .M. 1979. Petroleum geochemistry and geology. Freeman. San Francisco. 617pp. Hunt, J.M. and J.K. Whelan. 1977. Light hydrocarbons at site 367-370, Leg 41. In: Lancelot et al. (eds.). Initial Reports of Deep Sea Drilling Project Vol 41. u.s. Govt. Printing Office. Washington. pp. 859-860. Hunt, J.M. 1977. Distribution of carbon as hydrocarbons and asphaltic compounds in sedimentary rocks. Am. Assoc. Petrol. Geol. Bull. 61(1): 100-104.

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116 Hunt, J.M. 1975. hydrocarbons? Is there a geochmical depth limit for Petr. Eng. 47: 112-127. Jenden, P.D., B R .T. Simoneit and R.P. Philip. 1982. Hydrothermal effects of protokerogens of unconsolidated sediments from Guaymas Basin, Gulf of California: Elemental compositions, stable carbon isotope ratios and electron spin resonance spectra. In: Initial Reports of Deep Sea Drilling Project Vol LXIV, Part 2. u.s. Govt. Printing Office. Washington. pp. 905-912. Kendrick, J.W., A. Hood and J.R. Castano. 1977. Petroleumgenerating potential of sediments from Leg 41, Deep Sea Drilling Project. In: Lancelot et al. (eds.). Initial Reports of Deep Sea Drilling Project Vol 41. U.S. Govt. Printing Office. Washington. pp. 817-820. Krevelen, D. van. 1963. Geochemistry of coal. Geochemistry I. A. Breger ( ed. ) Pergamon York. pp.183-247. In: Organic Press. New Kuspert, W 1982. In: Cyclic and event stratification. A. Einsele and A. Seilacher (eds.). 482-501. Springer-Verlag. Berlin. Kvenvolden, K.A. 1977. Organic geochemistry Leg 41, Introduction and summary. In: Lancelot et al. (eds.). Initial Reports of Deep Sea Drilling Project Vol 41. pp. 815-816. Lancelot, Y., E. Siebold et al. 1977, Site 368: Cape Verde Rise Initial Reports of the Deep Sea Drilling Project, v.41, p.233-326. Lancet, M.S. and E. Anders. 1970. Carbon isotope fractionation in the Fischer-Tropsch synthesis and in meteorites. Science 170: 980-982. Leenheer, M.J. 1984. Mississippian Bakken and equivalent formations as source rocks in the Western Canadian Basin. In: Advances in Organic Geochemistry 1983. P.A. Schenck, J.W. DeLeeuw, and G.W.M. Lijmbach (eds.). Org. Geochem-6:521-532. Pergamon Press. Oxford. Lewan, M.D 1983. Effects of thermal maturation on stable carbon isotopes as determined by hydrous pyrolysis of woodford shale. Geochim. Cosmochim. Acta 47: 1471-1480. Lyon, G.L. and J.R. Hulston. 1984. Carbon and hydrogen isotopic compositions of New Zealand geothermal gases. Geochim. Cosmochim. Acta 48:1161-1171.

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117 Martens, c.s. and R.A. Berner. 1974. Methane production in the interstitial waters of sulfate-depleted marine sediments. Science 185: 1067-1069. Martin, S .J. 1975. Thermally evolved hydrocarbons from the bitumen and kerogen constituents of whole rock. Adv. Org. Geochem 7. Mciver, R.D. 1967. Composition of kerogen-clue to its role in the or1g1n of petroleum. In: Proc. 7th World Petr. Congr. Mexico City. Elsevier. London. 2:25-36. McKirdy, D.M. and T.G. Powell. 1974. Metamorphic alteration of carbon isotopic composition in ancient sedimentary organic matter. New evidence from Australia and South Africa. Geology 2: 591-595. Monson, K. D. and J. M. Hayes. 1982. Carbon isotopic fractionation in the biosynthesis of bacterial fatty acids as a means of determining the intramolecular distribution of carbon isotopes. Geochim. Cosmochim. Acta 46: 139-149. Nakai, N., Y. Yoshida and N. Ando. 1974. Isotopic studies on oil and natural gas fields in Japan. Chikyu Kagaku 7/8: 87-98. Nier, A.O. and E.A. Gulbranson. 1939. Variations in the relative abundance of the carbon isotopes. J. Am. Chem Soc. 61: 697-698. Nier, A.O. 1950. A redetermination of the relative abundance of the isotopes of carbon, nitrogen, oxygen, argon and potassium. Phys. Rev. 77: 789. Park{ R. and s. Epstein. 2c and 13c in plants. 1961. Metabolic fractionation o f Plant Physiol. 36: 133. Peters, K.E., J.K. Whelan, J.M. Hunt, and M.E. Tarafa. 1983. Programmed pyrolysis of organic matter from thermally altered Cretaceous black shales. Bull. Amer. Assoc. Petrol. Geol. 67(11): 2137-2146. Peters, K.E., B.G. Rohrback, Geochemistry of artificially sediments-I: Protokerogen. 65: 688-705. and I.R. Kaplan. 1981. heated humic and sapropelic Am. Assoc. Pet. Geol. Bull. Peters, K.E., B.G. Rohrback, and I.R. Kaplan. 1978. Vitrinite reflectance-temperature determination for intruded cretaceous black shale in the eastern Atlantic. In: o .F. Oltz (ed.). Symposium on Geochemistry: low temperature metamorphism of kerogen and clay minerals.

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118 SEPM Pac. Sec. pp. 53-58. Price, L.C., T. Ging, T. Daws, A. Love, M. Pawlewicz and D. Anders. 1984. Organic metamorphism in the MississippianDevonian Bakken shale North Dakota portion of the Williston Basin. In: J. Woodward, F.F. Meissner and J.L. Clayton (eds.). Hydrocarbon source rocks of the Greater Rocky Mountain Region. Rocky Mountain Assoc. of Geol. Denver, Colorado. Price, L.C. 1983. Geologic time as a parameter in organic metamorphism and vitrinite reflectance as an absolute paleogeothermometer. Jour. Petrol. Geol. 6: 5-38. Redding, C.E., M. Schoell, J.C. Monin and B. Durand. 1980. Hydrogen and carbon isotopic composition of coals and kerogens. In: Advances in Organic Geochemistry 1979. A. G. Douglas and J .R. Maxwell (eds.). Pergamon Press. New York. Reeburgh, w.s. 1980. Anaerobic methane oxidation: rate depth distributions in Skan Bay sediments. Earth Planet. Sci. Lett. 47: 345-352. Reeburgh, w.s. 1976. Methane consumption in Cariaco Trench waters and sediments. Earth Planet. Sci. Lett. 28: 337-344. Rice, D.D. and G.E. Claypool. 1981. Generation, accumulation, and resource potential of biogenic gas. Am. Assoc. Petr. Geol. Bull. 65: 5-25. Richet, P., Y. Bottinga and M. Javoy. 1977. A review of hydrogen, carbon, nitrogen, oxygen, sulfur, chlorine stable isotope fractionation among gaseous molecules. Ann. Rev. Earth Planet. Sci. 5:65. Rigby, D., B.D. Batts and J.W. Smith. 1981. The effect of maturation on the isotopic composition of fossil fuels. Org. Geochem. 3: 29-36. Sackett, W.M. 1984. Determination of kerogen maturity by the pyrolysis-carbon isotope method. Org. Geochem. 6: 359-363. Sackett, W.M and H.M. Chung. 1979. Experimental confirmation of the lack of carbon isotope exchange between methane and carbon dioxide at high temperature. Geochim. Cosmochim. Acta 43: 273-276. Sackett, w .M. 1978. Carbon and hydrogen effects during thermocatalytic production of hydrocarbons in laboratory simulaton experiments. Geochim. Cosmochim. Acta 42: 571-

PAGE 132

119 580. Sackett, W.M, B.J. Eadie and M.E. Exner. 1974. Stable isotope composition of organic carbon in Recent Antarctic sediments. Adv. Org. Geochem. 6. France. pp. 662-671. Sackett, W.M, s. Nakaparksin and D. Dalrylpe. 1968. Carbon isotope effects in methane production by thermal cracking. In: Advances in Organic Geochemistry 1966. G.D. Hobson and G.C. Spears (eds.). Pergamon Press. New York. pp. 37-53. Sackett, W.M. 1968. Carbon isotope composition of natural methane occurrences. Bull. AAPG 52: 853-857. Sackett, W.M, W.R. Echelmann, M.L. Berner and A.W.H. Be. 1965. Temperature dependence of carbon isotope composition in marine plankton and sediments. Science 148: 235-237. Sackett, W.M. 1964. The depositional history and isotopic carbon composition of marine sediments. J. Mar. Geol. 2: 173-185. Sackett, W.M and R T Thompson. 1963. Isotopic organic carbon composition of recent continental derived clastic sediments of eastern Gulf coast, Gulf of Mexico. Bull. Am. Assoc. Petrol. Geol. 47: 525-531. Schoell, M. 1984. Stable isotopes in petroleum research. In: J. Brooks and D. Welte ( eds) Advances in Petroleum Geochemistry Volume 1. Pergamon Press. New York. pp. 215-246. Schoell, M. 1983. Genetic characterization of natural Am. Assoc. Pet. Geol. Bull. 67: 2225-2238. Schoell, M. 1980. The hydrogen and carbon isotopic composition of methane from natural gases of various origins. Geochim. Cosmochim. Acta 44:649. Simoneit, B.R.T., K .E. Peters and I.R. Kaplan. 1981. Thermal alteration of Cretaceous black shale by diabase intrusions in the Eastern Atlantic. II. Effects on bitumen and kerogen. Geochim. Cosmochim. Acta 45: 1581-1602. Smith, J.W., K.W. Gould, and D. Rigby. 1985. The stable isotope geochemistry of Australian coals. Org. Chem. 8(5): 341-347. Smith, B.N. and s. Epstein. 1971. Two categories of 13c/2c ratios for higher plants. Plant Physiol. 47: 380-384.

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120 Stahl, W. 1977. Carbon and nitrogen isotopes in hydrocarbon research and exploration. Chem. Geol. 20: 121-149. Stahl, W. and B.D. Carey. 1975. Source-rock identification by isotope analysis of natural gases from fields in the Val Verde and Delaware basins, west Texas. Chem. Geol. 16: 257-267. Stahl, W. 1974. carbon isotope fractionations in natural gases. Nature 251: 134-135. Stevenson, D.P., C.D. Wagner, o. Beeck and J.W. Otros. 1948. Isotope effect in thermal cracking of propane-1-c13 J.Chem. Phys. 16: 993-994. Swart, P.K., M.M. Grady and C.T. Fillinger. 1982. Isotopically distinguishable carbon components in the Allende meteorite. Nature 297: 381-383. Tissot, B.P., R. Pelet and P.H. Ungerer. 1987. Thermal history of sedimentary basins, maturation inndices, and kinetics of oil and gas generation. Am. Assoc. Petr. Geol. Bull. 71(12): 1445-1466. Tissot, B.P. and D.H. Welte. 1984. Petroleum formation and occurrence. Spriger-Verlag. New York. 699 pps. Tissot, B.P, G. Demaison, P. Masson, J.R. Delteil and A. Combaz. 1980. Paleoenvironment and petroleum potential of middle Cretaceous black shales in Atlantic basins. Amer. Assoc. Petrol. Geol. Bull. 64: 2051-2063. Tissot, B.P, B. Durand, J. Espitalie, and A. Combaz. 1974 . Amer. Assoc. Petrol. Geol. Bull: 55: 2177-2193. Tissot, B.P, B. Durand, J. Espitalie, and A. Combaz. 1974. Influence of the nature and diagenesis of organic matter in formation of petroleum. Am. Assoc. Petr. Geol. Bull. 58: 499-506. Urey, H.C. 1947. The thermodynamic properties of isotopic substances. J. Chem. Soc. 41: 562-581. Wasserburg, G.J., E. Mazor and R.E. Zartman. 1963. Isotopic and chemical composition of some terrestrial natural gases. In: J. Geiss and E.D. Goldberg (eds.). Earth Science and Meteorites. North. Holland. Chapter 12. Webster, R.L. 1984. Petroleum source rocks and stratigraphy of the Bakken formation in North Dakota. In: J. Woodward, F .F. Meissner and J.L. Clayton (eds.). Hydrocarbon source rocks of the Greater Rocky Mountain Region. Rocky Mountain Assoc. of Geol. Denver, Colorado.

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121 pp. 57-81. Welhan, J. A. and H. era ig. 19 8 3 Methane, hydrogen and helium in hydrothermal fluids at 21N on the East Pacific Rise. In: Hydrothermal processes at seafloor spreading centers. P.A. Rona, K. Bostrom, L. Laubier and K.L. Smith Jr. (Eds.). Plenum Press. New York. pp. 391-409. Welhan, J .A. 1981. Carbon and hydrogen gases in hydrothermal systems: the search for a mantle source. Ph.D. thesis. Univ. of Calif. at San Diego. 195pp. Welhan, J.A. and H. Craig. 1979. Methane and hydrogen in East Pacific Rise hydrothermal fluids. Geophys. Res. Lett. 6: 829-831. Williams, J.A. 1974. Characterization of oil types in Williston Basin. AAPG Bull. 58:1243-1252. Whiticar, M.J., E. Faber, and M. Schoell. 1986. Biogenic methane formation in marine and freshwater environments: C02 reduction vs. acetate fermentation-Isotope evidence. Geochim. Cosmochim. Acta 50: 693-709. Whi ticar, M. J. and E. Faber. 19 8 6. Carbon and hydrogen isotopes on gas samples from Leg 95, Sites 603D and 613. In: Initial Reports of Deep Sea Drilling Project. Vol 95. u.s. Govt. Printing Office. Washington. Wickman, F.E. 1953. Variations in the relative abundance of the carbon isotopes in plants. Geochim. Cosmochim. Acta 2: 243. Wong, W. W. L. and W .M. Sackett. 1978. Fractionation of stable carbon isotopes by marine phytoplankton. Geochim. Cosmochim. Acta 42: 1809-1815. Yeh, H. w. and s. Epstein. 1981. Hydrogen and carbon isotopes of petroleum and related organic matter. Geochim. Cosmochim. Acta 45: 753-762. Zehnder, A.J.B., K. Inguorsen and T. Marti. 1982. Microbiology of methane bacteria. In: D.E. Hughes, D.A. Stafford, B.I. Wheatley et al. (eds.). Anaerobic digestion 1981. Elsevier. Zumberge, J.E. 1983 Tricyclic diterpane distributions in the correlation of Paleozoic crude oils from the Williston Basin. In: Advances in Organic Geochemistry 1981. M. Bjoroy et al. (eds.) pp 738-745. Wiley Chichester.

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122 APPENDICES

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123 APPENDIX 1. RESULTS FOR THE LN2CF FROM THE PYROLYSIS OF THE BAKKEN SHALES SAMPLE o13C-LN CF 2 o13C-CH 4 ccCH4/LN2CF A13c 38422-23 -20.3 -31.2 2.1 +10.9 38422-8 -15.1 -30.7 31.7 +15.6 36653-1 -21.3 -31.6 25.5 +10.3 36660-1 -18.4 -27.8 4.3 +9.4 36666-1 -13.7 -30.6 10.6 +16.9 36658-1 -20.5 -28.3 1.9 +7.8

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124 APPENDIX 2. RESULTS FOR THE TIME SERIES PYROLYSIS OF BAKKEN SHALE SAMPLE 38423-15. TIME 13 A 13c-co 13 A 13C-CH Aco2-CH4 CC02 2 CCH4 4 (hr) {parent-co2 ) (parent-CH4 ) 12 -2.2 24.7 -26.8 0.1 24.6 24 -5.2 21.7 -25.8 1.1 20.6 48 -6.1 20.8 -25.8 1.1 19.7 120 -6.4 20.5 -24.7 2.2 18.3 parent carbon = -26.9joo APPENDIX 3. CARBON MOLE RATIO RESULTS FOR THE TIME SERIES PYROLYSIS OF BAKKEN SHALE SAMPLE 38423-15. TIME CMR-C0 2 (hr) CMR-CH4 CMR-OC 12 0.063 0.078 0.859 24 0.064 0.072 0.864 48 0.063 0.077 0.860 120 0.071 0.110 0.819 *The carbon mole ratio of the organic residue is determined by: 1 {CMR-C0 2 + CMR-CH4 )

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125 APPENDIX 4. RESULTS FOR THE TIME SERIES PYROLYSIS OF BAKKEN SAMPLE 36650-1 TIME o13c A13c-co o13c A13C-CH o13c f:.13C-OCR 2 4 (OCR1 ) (hr) (C02 ) (CH4 ) 12 -30. 8 -4.7 -26.4 -0.3 24 -7.7 18.4 -30.3 -4.2 -25.8 + 0.3 48 -8.7 17.4 -29.5 -3.4 -27.5 -1.4 72 -10.5 15.6 -30.1 -4.0 -27.2 -1.1 96 -9.4 16. 7 -29.2 -3.1 -27.6 -1.5 120 -27.5 -1.4 -28.1 -2.0 160 -6.6 19.5 -28.1 -2.0 -29.0 -2.9 210 -10.0 16.1 -23.4 +2.7 -----parent carbon = -26.1joo 10CR = organic carbon residue

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126 APPENDIX 5. CARBON MOLE RATIO FOR THE TIME SERIES ANALYSIS OF BAKKEN SHALE SAMPLE 36650-1 TIME (hr) 24 48 72 96 160 0.087 0.092 0.116 0.102 0.160 CMR-OC 0.189 0.724 0.189 0.722 0.193 0.691 0.214 0.684 0.276 0.564

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127 APPENDIX 6. GAS CHROMATOGRAPHY ANALYSIS OF THE TIME SERIES PYROLYSIS OF KEROGEN SAMPLE 31. AREA% TIME n METHANE ETHYLENE/ACETYLENE ETHANE 0.25 4 54.62 11.43 20.80 0.5 2 59.14 9.24 26.63 1 8 69.43 5.10 25.26 2 3 68.65 4.51 26.84 4 2 76.38 2.72 20.81 16 2 84.53 1.58 13.89 24 4 91.39 0.93 7.68 32 3 94.24 0.61 5.16 48 3 96.72 0 3.00 72 2 97.85 0 2.09 144 2 100 0 0 210 2 100 0 0

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128 APPENDIX 7. RESULTS FOR THE TIME SERIES ANALYSIS OF KEROGEN SAMPLE 31 (o13c OF THE PARENT CARBON= -26.4/oo) TIME wt (mg) mmC mlCH4 nunCH4 o13C-CH 4 CMR !113C 24 20.91 0.955 3.7 0.165 -31.1 0.173 -4.7 48 14.37 0.656 4.3 0.192 -29.6 0.293 -3.2 72 31.56 1.441 8.9 0.397 -29.4 0.276 -3.0 96 33.96 1. 551 9 9 0.442 -29.4 0.285 -3. 0 120 38.91 1.777 11.7 0.522 -28.9 0.294 -2.5 160 21.65 0.989 6.5 0.290 -28.0 0.294 -1.6 288 18.00 0.820 5.4 0.241 -28.0 0.294 -1.6

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129 APPENDIX 8. RESULTS FOR THE CAPE VERDE RISE KEROGEN LIQUID NITROGEN CONDENSABLE FRACTION (LN2CF) SAMPLE mlC02 13 Ccoz mmco2 CMR ccCHJccco2 Above major sill 58-5 ( 1) 1.3 -25.7 0.058 0.077 1. 08 59-3(311) 1.1 -23.6 0.051 0.022 8.94 60-1(2) 2.8 -24.5 0.125 0.193 0.11 Minor sill 60-2(47) 1.1 -25.8 0.048 0.031 3.94 60-3(33) 2.4 -25.2 0.108 0.091 0.29 60-3(41) 3.1 -23.7 0.136 0.188 0.06 60-3 (5) 2.7 -24.5 0.119 0.125 0.11 60-4(35) 0.6 -25.7 0.021 0.019 60-4(39) 0.5 -26.3 0.024 0.017 0.35 Minor sill 60-4(44) 0.3 -26.6 0.014 Below major sill 62-3(57) 0.4 -28.3 0.020 0.027 0.57

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130 APPENDIX 9. RESULTS FOR THE CAPE VERDE RISE KEROGEN RESIDUAL ORGANIC MATTER. SAMPLE DISTANCE oc o13c o13c A13 COC-OCR (%) oc OCR Above major sill 58-5{1) +29.1 31.4 -24.2 -23.4 +0.8 59-3 (31) +9.5 74.2 -26.8 -26.1 +0.7 60-1(2) +6.1 22.3 -24.9 -24.4 +0.5 Minor sill 60-2(47) +4.4 49.9 -24.6 -24.2 +0.4 60-3(33) +4.0 45.7 -25.3 -25.2 +0.1 60-3(5) +3.2 44.8 -24.3 -24.3 0.0 60-4(39) +1.9 55.7 -22.9 -23.0 -0.1 Minor sill 60-4(44) +1.5 -20.9 -20.5 +0.8 Below major sill 62-3(57) -0.7 29.3 -26.8 -27.6 -0.8

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131 APPENDIX 10. EXCHANGE BETWEEN EQUAL AMOUNTS OF METHANE AND CARBON DIOXIDE ( 2ml OF METHANE AND CARBON DIOXIDE USED). TIME mlC02 o13c-co 2 2 mlCH4 613C-CH 4 IJ.13C-CH 4 0 0 -12.6 0 0 -38.2 0 12 1.5 -13.1 -0.5 1.4 -37.7 0.5 1.3 -13.3 -0.7 0 0 0 24 1.8 -13.4 -0.8 0 0 0 1.3 -13.3 -0.7 1.6 -37.5 0.7 48 1.5 -13.2 -0.6 1.6 -37.5 0.7 1.6 -13.1 -0.5 1.8 -37.5 0.7 72 1.3 -13.1 -0.5 1.8 -37.5 0.7 96 2.1 -13.1 -0.5 1.8 -37.4 0.8 1.1 -13.3 -0.7 1.3 -37.7 0.5 120 0.7 -13.3 -0.7 1.6 -37.6 0.6 1.6 -13.2 -0.6 1.7 -37.3 0.9 210 1.6 -13.4 -0.8 1.6 -36.9 1.3 504 1.3 -14.0 -1.4 1.6 -36.4 1.8


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