The global methane budget and the organic carbon cycle

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Biogenic methane processes

Shallow fluid systems

Shallow fluid systems are widespread natural processes that result from the interaction of geological, biological and chemical processes. They are identified from characteristic geological features that are widely distributed along near-shore, continental slope and deep oceanic environments and include shallow gas accumulations, pockmarks, seeps, mud-volcanoes, authigenic carbonate precipitations and gas hydrates (Jensen 1992; Römer et al. 2012; Skarke et al. 2014; Dupré et al. 2007; Pierre et al. 2017; Hovland et al. 2002; Judd et al. 2002).
Such shallow fluid systems represent both a benefit and a hazard for human activity. They may contain shallow energy resources but they also play an important role in shaping and stability of the seabed and therefore may present a risk for infrastructure installed at the seabed. They also affect marine biological processes as they can alter locally the composition of the ocean water.
A major component of such fluids is methane which is associated with water or other gases such as carbon dioxide (CO2) or heavier gaseous (ethane, propane) or liquid hydrocarbons (HCs). Methane (CH4) is the smallest HC component and can have different origins:
• Methane can derive from the decomposition of Sedimentary Organic Matter (SOM) during the early stages of diagenesis through bacterial activity (Whiticar et al. 1986; Schulz et al. 2015). Such methane is called primary biogenic gas. In the framework of this PhD thesis, we are interested in the analysis of this type of gas and we refer to it as microbial gas or biogenic gas.
• Methane and other HC components can also have thermogenic origins as the product of thermal cracking of the organic matter during catagenesis (Tissot and Welte 1984; Jones et al. 2008).
• Methane can also be the result of the bacterial degradation of oil accumulated in reservoirs (Milkov 2011) which is called secondary biogenic gas.
• A different type of methane is represented by abiotic methane (resulting from an abiotic source of carbon). It is generally interpreted as related to reaction of H2 with CO2. In marine environments, abiotic methane is the result of reactions at high-temperatures along mid-oceanic ridges (Welhan and Craig 1979, Charlou et al. 2002; Proskurowski et al. 2008). Onshore, it has been found in different settings, like ophiolitic massifs, intracontinental cratons or volcanic/hydrothermal gas systems (Etiope et al. 2011). It is found much more rarely compared to thermogenic or biogenic gas.
The scientific community has always been interested in methane in the subsurface (Rice and Claypool 1981; Rice 1992; Whiticar 1994) as it is ubiquitous (from continental margins to mid-oceanic ridges) and represents a major energy source that has a lower carbon footprint than liquid HC or coal. In addition, the quantification of natural methane sources and sinks, both at present day and in the geological past, are of interest to the scientific community working on present and future global climate change (Regnier et al. 2011; Saunois et al. 2016). Despite this increasing interest, methane generation and biological degradation mechanisms are still not well understood.
In addition to the OM composition, methane generation is controlled by several other factors such as surface temperature, eustatic sea level variation and temperature changes at the sea bottom (Judd et al. 2002), as well as the type of microorganisms that mediate the reactions of gas generation (Boetius et al. 2000). Therefore, an accurate characterization of the microbial gas generation process is necessary in order to understand the evolution and the importance of this potential energy source.
Microbial methanogenesis is the result of complex degradation processes of the OM that take place at different stages of diagenesis (Whiticar et al. 1986; Floodgate and Judd 1992; 15 Whiticar 1999; Schulz and Zabel 2006). Degradation can be assimilated to an anaerobic respiration resulting first in methane (CH4) and subsequently, in moderate quantities, in ethane (C2H6) production. The efficiency of this process is determined by the quality and quantity of the OM and the temperature conditions (Schulz and Zabel 2006). Poor-OM layers (TOC ˂ 0.5%) can however generate important volumes of microbial gas (Clayton 1992). Therefore, large accumulations of biogenic gas can be found in offshore environments when conditions at the seafloor are compatible with an anaerobic respiration. For instance, in deltaic settings such as the Amazon cone (Arning et al. 2013) or the Carupano Basin (Schneider et al. 2012), generation and accumulation of biogenic gas is occurring because the dispersed OM is transported and deposited with relative high sedimentation rates at low geothermal gradients which are key conditions for biogenic gas generation (Clayton 1992). With increasing burial and temperature, the activity of the micro-organisms decreases (Clayton and Hay 1994). Above 60 to 80°C, the generation of microbial gas slows down considerably, and progressively thermal cracking of the OM takes place (Zeikus and Winfrey 1976; Zeikus 1997; Whiticar et al. 1986; Whiticar et al. 1999).
Pathways of OM degradation by microbial activity are linked to the type of microorganisms, their metabolic pathways, and to the type of carbon source. Metabolisms of methanogens include three main pathways: methylotrophy, hydrogenotrophy and acetotrophy. As methyl compounds are easily degraded, methyl fermentation is restricted to the uppermost part of the sediments (i.e. first meters). The main mechanisms are carbon dioxide reduction and acetate fermentation (Whiticar et al.; 1986). These different pathways may coexist but the reduction of H2 and CO2 from microbial activity is still considered the most efficient process to produce large methane quantities in marine environments. Indeed, notably in the framework of the PAMELA (Passive Margin Exploration Laboratories) project, it has been shown that acetotrophs are absent in marine environments suggesting a competition with sulfate reducing bacterias (Odobel 2017).
The biodegradation of OM depends on its reactivity to microbial activity. The OM is usually subdivided into a refractory part and a biodegradable part (Burdige 2007, 2006). A part of the biodegradable fraction (called “labile” such as proteins or carbohydrates) is rapidly degraded into methane whereas a more refractory part (called “thermo-labile” such as vitrinite and lipids) is remobilized as temperature increases and only then becomes available for micro-organisms to produce additional methane (Burdige 2011; Kamga 2016). Whereas the refractory TOC content is easily defined through geochemical analysis (e.g. Rock-Eval technique or LECO) (Espitalié et al. 1977, 1984; Lafargue et al. 1998; Behar et al. 2001), it is not the same when we try to reconstruct the biodegradable compounds of the initial TOC content. It is still challenging to characterize the biodegradable OC as it is linked to OM preservation and mineralization processes (Burdige 2007). Different types of OM have different reactivities and, depending on their quality, they can also undergo different selective preservation and/or re-mineralization processes (Burdige 2007, 2011). Therefore, it is important to consider the effect of the quality and quantity of the OM on methane production, especially with regards to the effect of its initial labile and thermo-labile fractions. This study therefore aims to understand the effect of the OM quality on CH4 production, using geochemical analysis that is integrated with basin modelling.

Methanogenesis

Microbial gas, or primary biogenic methane, can be distinguished from thermogenic gas based on composition and isotopic signature. Biogenic gas is usually extremely dry (C1/(C2+C3) > 100) and characterized by a light isotopic signature (δ13C-CH4 between -90‰ and -50‰) (Bernard et al. 1976; Milkov and Etiope, 2018). In contrast, secondary biogenic gas displays a more variable isotopic signature, sometimes close to the signature of thermogenic gas (δ13C > – 55‰) (Katz, 2011).
Microbial gas is the result of the respiration by micro-organisms. This process first tends to use molecular oxygen as long as oxygen is available in the environment. When O2 is depleted, microbial respiration proceeds consuming electron acceptors in the order NO3−, Mg2+, Fe2+, then SO42− (Archer, 2007; Clayton and Hay, 1994). Once the sulfate reduction ends, methanogenesis starts (Fig. 3).
Methanogenesis requires anoxic conditions, low sulfate content, low temperatures, presence of OM, and sufficient pore space to host micro-organisms (Rice 1992). Depending on the amount of available SO42−, methanogenesis follows different generation pathways (Whiticar and Schoell, 1986) (Fig. 4). As mentioned before, in seawater the preferential pathway is reduction of CO2 by molecular hydrogen (Eq. 1) (Archer 2007). In contrast, in fresh water environment, the preferential pathway is fermentation of acetate (Eq. 2) (Hoehler et al. 1999).
CO2 + 4H2 => CH4 + 2H2O (1)
CH3COOH => CH4 + CO2 (2)
Figure 3. Distribution of the main elements along the sedimentary column (modified after Clayton and Hay 1994).
Despite the fact that methane can be generated either in fresh or marine environment, the relative proportion of each pathway and the final amount of generated CH4, are strongly controlled by the availability of substrate and hydrogen for the methanogens, which can be limited by the competition with sulfate reducing bacteria (Whiticar et al. 1986). The rate of methane generation is influenced by several parameters where temperature (Belyaev et al. 1983) represents one of the main controlling factor. The microbial activity is classified by increasing temperature resistance: Psychrophiles, Mesophiles, Thermophiles and Hyperthermophiles (Fig. 5). The peak of CH4 generation is around ~40°C corresponding to the mesophiles level, whereas neither psychrophiles nor thermophiles are capable of generating important microbial gas volumes (Katz 2011).
Figure 4. OM degradation pathways (acetate fermentation vs CO2 reduction) in freshwater and marine settings (Whiticar et al.;1986; Whiticar 1999).
Figure 5. Evolution of the bacterial activity levels as a function of increasing temperature (Katz 2011).
Methanogenesis is linked also to the salinity of pore water (Waldron et al. 2007) since microbial activity may be inhibited when salinity is too high (Zindler 1993; Schoell 1980). Sufficient pore space is also required to allow growth of microbial population (Chapelle and Lovley 1990). However, as observed by Chapelle and Lovley (1990), the abundance of substrate can impact the overall amount of methane but it doesn’t seem to be a controlling factor of the rate of methane generation.
Clayton (1992) determined that the key controlling factors of microbial gas generation and accumulation are thermal gradient and burial rate of the source rock. If the sedimentation rate is too low, the metabolizable organic carbon may be consumed in the aerobic or sulphate reduction zone and only a small fraction can reach the methanogenesis zone, resulting in a dispersion of the generated methane to the surface. If sedimentation is too fast, the sedimentary organic matter (SOM) may by-pass the methanogenesis window too quickly to generate large amounts of gas. Figure 6 shows the optimum heating rate (the product of geothermal gradient and sedimentation rate) for microbial gas production (Clayton 1992). This diagram was constructed with data from geological settings where extensive methanogenesis is observed such as in California, Po Valley, Offshore Peru and Offshore Louisiana (Clayton 1992 and references therein). The upper boundary of the diagram is equivalent to a heating rate of 18°C/My while the lower boundary is equivalent to a heating rate of 7°C/My (Clayton 1992). The optimum balance between production and accumulation with normal thermal gradients (20-40 °C/km) occurs at burial rates of 200-1000 m/My and yields heating rates located between 7 and 18 C/My.
Figure 6. Sedimentation rate vs geothermal gradient converted to heating rate for methanogenesis (Clayton 1992).
The microbial gas production is likely when the curve is comprised between 18°C/Ma and 7°C/ Ma.

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Evidence of fluid systems at the seafloor

When CH4 generation occurs at shallow depths and hydrostatic pressures, in un-consolidated sediments (Fig. 7) (Verweij et al. 2018) there may be no conditions for trapping of gas. Potential seals cannot hold even moderate gas columns as they easily exceed the low capillary entry pressures at these shallow depths (Fig. 7). This lack of effective sealing results in continuous fluid emanations at the ocean floor. These continuous emanations contrast with features of focused fluid escape structures such as pockmarks which are commonly associated with physico-chemical manifestations such as mineral precipitations of authigenic carbonates or hydrates formation.
Figure 7. Pore-fluid pressures vs depth for Plio-Pleistocene and Paleogene sequences in the Southern North Sea delta (modified after Verweij et al. 2018). Pore-fluid pressures are hydrostatic or close to hydrostatic in the first 1000 m of depth.

Pockmarks and gas seepages

Pockmarks are morphological depressions in the soft seabed formed by fluid expulsions including gas (Hovland 1989). They can be cone- or saucer-shaped depressions with generally a size varying from 1 to more than a few 100 meters and height between 1 to more than 10 meters (Dimitrov and Woodside 2003). Pockmarks have been observed and described for the first time in the Nova Scotia continental shelf by King and McLean (1970) and, then, they have been recognized along many continental margins (Hovland and Judd 1988). These features are found widely distributed in the deep-water domain and in shallow environment (such as in the Lower Congo Basin or in the Norwegian offshore) (Gay et al. 2006; Rise et al. 2014) as the result of interactive processes between fluids and sediments.
It is now well accepted that pockmarks are the expressions related to fluid seepage at the seafloor (Hovland et al. 1984, Hovland et al. 2005) or of gas hydrate dissolution (Sultan et al. 2014), which can be originated by either thermogenic and biogenic gas emissions (Hovland and Judd 1988). They can be present as a cluster or as a string reflecting subsurface structures (e.g. fractures or faults) (Hovland et al. 2010; Pilcher and Argent 2007; Gay et al. 2007). Several studies (Sultan et al. 2010; Davy et al. 2010) tried to classify and differentiate pockmarks between classical fluid escapes with a conical shape (Type-I) and hydrate-bearing depressions developed over several kilometers (Type-II) (Fig. 8). Type-I pockmarks are characterized by a sub-circular shape whose diameters vary between 10 to 900 m, with a depth of 0.1 to 50 m (Riboulot et al. 2016). They are usually associated with a gas chimney or acoustic pipes and are commonly observed on continental margins. Type-II pockmarks show irregular morphology with diameters varying between 50 – 800 m and depths of about 10 m, and they are not associated to vertical chimneys (Riboulot et al. 2016).
Figure 8. Morphological and geometrical characterization of the two different types of pockmarks (Riboulot et al. 2016). a) Type-I pockmark with gas chimney association; b) Type-II pockmark with an irregular shape associated to hydrates occurrence.
According to Etiope et al. (2019), seepages can be classified into (1) onshore seeps in sedimentary basins (e.g. CH4 seeps and mud volcanoes), (2) offshore seeps where methane is usually released at shallow depths (200 to 400 mbsl) and CH4 can likely reach the atmosphere as a function of the amount of emitted gas and the water depth, (3) deep-sea seeps, (4) geothermal and volcanic manifestations, where CH4 and CO2 are usually associated.
The interest regarding these features is growing since they can be a key for understanding fluid flow in sedimentary basins (Chenrai and Huuse 2017). Pockmarks and paleo-pockmarks (Andresen et al. 2008) can be used as first indicators of active or paleo-fluid escapes to the seafloor, giving information on hydrocarbon and hydrate occurrences as well as on the paleo-fluid flow history. They can also contribute to the emissions of large quantities of gas directly into the ocean, with a possible transfer to the atmosphere, resulting in a negative impact on climate and environment. Distributions of pockmarks and seeps have been identified and mapped along the areas studied in this PhD thesis, both in the Offshore Zambezi (Deville et al. 2020) and in the offshore Aquitaine (Dupré et al. 2014, 2020; Michel et al. 2017).

Methane-Derived Authigenic Carbonate (MDAC)

Methane-Derived Authigenic Carbonates (MDACs) are sedimentary features resulting from the Anaerobic Oxidation of Methane (AOM) (Jørgensen 1976, Boetius et al. 2000, Judd et al. 2007). MDAC precipitation takes place within the Sulphate-Methane Transition Zone (SMTZ) which corresponds to a boundary located mostly below the seafloor at variable depths (Boetius et al. 2000). MDACs are characterized by variable precipitation rates of several cm/ky to m/ky (ky = 1000 years) (Luff and Wallmann 2003).
During upward migration, the generated methane meets the SO42- of the downward diffusing seawater in the Sulphate-Methane Transition Zone (STMZ) where it is consumed by the activity of methanotrophic microorganisms with Sulfate Reducing Bacteria (SRB) in anoxic conditions (Boetius 2000; Conrad 2005; Thauer 2010; Lash 2015). The AOM redox reaction can be described as the net reaction between seawater sulfate and methane (Eq. 3):
CH4(aq) + SO42− → HCO3− + HS− + H2O(l) (3)
The dissolved inorganic carbon (bicarbonate) generated in the STMZ under anaerobic conditions increases alkalinity which promotes carbonate precipitation resulting in the formation of authigenic carbonates (Eq. 4) (Boetius 2000; Regnier et al. 2011; Lash 2015).
2HCO3- + Ca2+ CaCO3 + CO2 + H2O (4)
This process is a widespread diagenetic reaction along modern continental margins (Reeburgh 2007, Lash 2015). MDACs havean important role in climate control (Noble-James et al. 2020). Indeed, they represent a major sink for methane, with a consumption varying from 10 to 80% of the total up-ward migrating methane before it is released at the seafloor (Conrad 2009, Regnier et al. 2001).
MDAC can precipitate either in the form of isolated crystals dispersed in sediments, or chimneys or areal massive distributions when fluid emissions are important and sustained over time (Paull and Ussler, 2008). MDAC precipitation requires anoxic conditions which implies in most cases that these features precipitate in the sediments below the seafloor (Paull and Ussler, 2008). Therefore, according to the authors, the presence of MDAC at the sediment/water interface is usually the result of erosional processes which removed the uppermost sedimentary cover. Examples of such features are found along the Offshore Aquitaine (Dupré et al. 2014; Pierre et al. 2017; Dupré et al. 2020) and in the Offshore Zambezi (Deville et al. 2020) (Fig. 9).

Table of contents :

1. Introduction to biogenic methane
1.1 The global methane budget and the organic carbon cycle
1.2 Biogenic methane processes
1.2.1 Shallow fluid systems
1.2.2 Methanogenesis
1.3 Evidence of fluid systems at the seafloor
1.3.1 Pockmarks and gas seepages
1.3.2 Methane-Derived Authigenic Carbonate (MDAC)
1.3.3 Gas Hydrates
1.4 PhD thesis objectives as part of the PAMELA Research Project
1.4.1 The Mozambique Channel natural laboratory
1.4.2 The Offshore Aquitaine natural laboratory
1.4.3 PhD thesis outline
2. Natural sources of microbial gas
2.1 Biogenic source rock potential
2.2 Organic matter production
2.3 Organic matter composition
2.4 Preservation and burial of organic matter in marine settings
3. A case study from the Mozambique Basin (manuscript in submission process)
4. Modelling microbial gas processes at the basin scale
4.1 The OM reactivity model
4.1.1 The discrete model
4.1.2 The continuum model
4.1.3 Conceptual approach for biogenic CH4 processes at the basin scale
4.2 The TemisFlow basin model
4.3 Modelling biogenic gas migration
4.3.1 Modelling adsorption of biogenic gas
4.3.2 Modelling dissolution of methane in pore water
4.3.3 Modelling migration as a free gas phase
4.4 Trapping and sinks of biogenic gas
4.5 Critical parameters for biogenic gas generation and migration
5. A case study from the Offshore Aquitaine (Paper in Marine and Petroleum Geology, 2021)
6. Issues with the analytical procedure to characterize recent terrestrial OM
7. Conclusions
8. Future research on biogenic gas generation and migration
8.1 Hydrate stability zones as paleo-gas migration indicators
8.2 Additional research questions
9. References
10. Figures Table
11. Tables Table
12. Appendices
12.1 Publications
12.2 International Conferences
12.3 Internal PAMELA seminars
12.4 PhD thesis pitch
12.5 International Training and School
13. Summary
14. Résumé

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