Controls on overpressure evolution during the gravitacional collapse of the Amazon deepsea fan

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OVERPRESSURE MECHANISMS

Abnormally high pore pressures within sedimentary successions are the result of fluid retention within a formation. In general, fluid is retained in the pores, but when it does not find a path for expulsion and pressure dissipation, it generates a pressure imbalance. This imbalance is a consequence of different generating mechanisms (Fig 2.1), among which the most common are:
Stresses related to disequilibrium compaction (undercompaction) and tectonics stress (lateral compressive stress) (Osborne and Swarbrick, 1997; Burrus, 1998; Law and Spencer, 1998; Nashaat, 1998; Swarbrick and Osborne, 1998; Grauls, 1999; Zoback, 2007);
Fluid volume increase during temperature increases (aquathermal), mineral transformation often release water (Hower et al., 1976; Law et al., 1983; Bruce, 1984; Hall, 1993; Burrus, 1998; Swarbrick and Osborne, 1998, Bowers, 2001), hydrocarbon fluids can be released due to primary or secondary cracking (Law and Spencer, 1998; Swarbrick and Osborne, 1998; Grauls, 1999).
Fluid movement and buoyancy (Osborne and Swarbrick, 1997; Swarbrick and Osborne, 1998).
Among the mechanisms mentioned above, the principal ones resulting in large overpressures are disequilibrium compaction and fluid volume expansion during gas generation (Swarbrick and Osborne, 1998). It is emphasized that disequilibrium compaction is often associated in basins with a high sedimentation rate, including Tertiary deltas and some intracratonic basins.
For studies on the pressure evolution (normal or abnormal) of sedimentary basins, it is important to identify the main generating mechanism, or set of mechanisms. These mechanisms can be tested from hypothetical scenarios, in numerical or analog models.

Disequilibrium compaction

This mechanism, also defined as vertical loading stress, is most common in deltaic depositional environments. The sediments are initially unconsolidated and remain in a suspended load with the carrying fluid until they are deposited at the base of a water body to be further compacted as the pressure exerted by younger sediments progressively increases. Overpressure generated by undercompaction is common in basins with a high sedimentation rate, including Tertiary deltas and some intracratonic basins (Swarbrick and Osborne, 1998). This process occurs as high sedimentation rates causes increasing burial of low-permeability and hence undercompacted sediments. As burial continues, the pore fluids start to support part of the weight of overlying sediments, instead of the grain-to-grain contacts, and the fluid becomes overpressured.
Undercompaction can preserve porosity, when sediments with low permeability inhibit the escape of pore fluid (fluid retention) as rapidly as the pore space can compact. Fluid retention is controlled by rapid sedimentation and by the low intrinsic permeability of rocks such as shales or well-cemented carbonates, formations are also referred to as ‘seals’. Measuring the permeability of seals is difficult and prone to large inaccuracies.
According to McBride et al (1991), sandstones compact at a relatively slow rate, from a starting point of approximately 40-50% porosity to as little as 5-10% porosity, mainly due to the rearrangement of sand grains, as well as due to dissolution at grain contacts. Low rates of compaction together with high rates of fluid expulsion (due to high permeability) lead to reduced porosity with depth. In contrast, clays have a typical porosity of approximately 65-80% at the point of deposition, and compact more quickly than sands. In addition, clay compaction is largely controlled by mudstone lithology, and in certain conditions, the sediments are not fully compacted relative to their overburden.

Tectonic stress

Tectonic stress may increase the total stress state and cause overpressures. This situation is common in strike-slip and reverse-fault stress regimes known to be under high lateral stresses, such as parts of California (Berry, 1969). According to Swarbrick and Osborne (1998), lateral compression can increase pore pressures in the same way as vertical stress causing over pressuring through disequilibrium compaction. The stress variation can result from variations of intraplate stress due to plate tectonic processes or from lithospheric flexure during deglaciation (Swarbrick and Osborne, 1998; Zoback, 2007). Recent rapid sedimentation and upsurges in horizontal compression that drive compaction can increase pore pressure.

Clay diagenesis

Diagenesis consists of physical, chemical, and/or biological changes that alter sediments during lithification or post-depositional compaction. Diagenetic processes are driven by increasing pressure and temperature with depth, and may result in the formation of new minerals, recrystallization and lithification. Several diagenetic mineral transformations common in sedimentary successions lead to dehydration, or the release of bound water.
During diagenesis, bound water becomes free water. If the increased amount of free water cannot escape, then the pore pressure becomes abnormally pressured. Dehydration reactions include gypsum-to-anhydrite in evaporitic sediments, and coalification (Law et al., 1983). The most common transformation involves the dehydration of smectite, a multi-layered, mixed-layer clay commonly found in mud-rocks. Smectite transforms to a new mineral, illite, along with the release of water.
In many mud-dominated basins, a gradual and systematic downward change from smectite to illite is observed in stratigraphic successions, broadly coincident with a transition to high amounts of overpressure. The transition occurs over a temperature range of 70-150°C (Bruce, 1984; Schneider et al., 2003) and appears to be independent of sediment age and burial depth.

Hydrocarbon generation

This mechanism is driven by temperature increase with depth, and depends on time, the presence and richness of source rocks, facies distributions and permeability. Even though hydrocarbon generation may promote the development of overpressure, and affect compaction disequilibrium, it is difficult to quantify its effects on pore pressure distribution (Grauls, 1999) Meisser (1978) and Du Rochet (1978) observed primary migration in source rocks and interpreted the transformation from kerogen to oil, or primary cracking, to occur at temperatures from 100-120°C.
Progressively, after hydrocarbon expulsion from source rocks, a pressure transference occurs due to secondary hydrocarbon migration, also called differential fluid density. In this overpressure mechanism, the formation contains a fluid with a different density than water. This dynamic transfer occurs along structures with significant dip, such as active faults, leakage of seal and through the permeability of rocks with a significant dip.
Secondary oil cracking can also occur during this oil-to-gas transformation, at temperatures exceeding 175-180°C (Hedberg, 1974, Mackenzie and Quigley, 1988). In some cases, at depths of burial typically between 3-5 km, the compressibility of gas must be considered as well as its solubility in brine. Large volume increases are still likely, and can potentially generate overpressures.

Gravity Tectonics

The study of gravity tectonics or thin-skinned extension has received special attention in the last decades since this form of deformation has structural consequences in important petroleum producing areas. These studies were based on case studies and analogue modeling (Vendeville and Cobbold, 1988; Vendeville, 1991; Nelson, 1991; Vendeville and Jackson, 1992a, 1992b; Jackson, 1994; Mohriak et al., 1995; Morley and Guerin, 1996).
Thin-skinned extension consists of the development of structural systems within a sedimentary cover due to its horizontal translation above a basal detachment surface. It may take place in mobile layers consisting of evaporites or clays, which can act as one or multiple detachment levels. On continental margins, the presence of siliciclastic and/or carbonate sedimentary successions overlying rheologically more ductile stratigraphic levels may allow a mechanism of thin-skinned extension to develop (Vendeville and Cobbold, 1987; Cobbold et al., 1989).
According to Mourgues (2003) based on analogue and numerical models, extensional deformation begins when the gravity potential is sufficient to overcome the internal resistance (friction) of the sediment load and the resistance to gliding along potential basal surfaces. The detachment levels can be of two types:
Ductile behavior – décollement – that accommodates the movement of the sediment cover through a continuous deformation. The shear strength in this case is a function of viscosity and shear rate. Mobile levels with rheological behavior similar to evaporites can deform with practically no resistance under conditions of low deformation rates.
Detachment surface, which corresponds to a discontinuity between two fragile levels, on which the translation or sliding of the sedimentary cover occurs. The development of such as a surface requires overpressures generated by interstitial fluids, which are responsible for reducing friction (resistance to movement) between the interfaces (for example, overpressured clays).
The gravity potential is directly proportional to the slope of the detachment surface and gradient slope. Active sedimentation may play an important role as a triggering agent, acting as the motor of gravity tectonics (Vendeville and Cobbold, 1987). As a result, gravity tectonics can be generated by combination of two inducing mechanisms:
Gravity gliding: involves the translation of a sedimentary cover on a sloping detachment surface towards the basin under the action of gravity, and is generally associated with the beginning of thermal subsidence of the oceanic lithosphere (Fig. 2.2) (Schultz-Ela, 2001; Rowan et al., 2004). The intensity of the sliding process is directly proportional to the slope of the bottom surface and to the slope of the mobile substrate, and can occur even on surfaces with soft slopes (less than 2°).
Gravity spreading: movement on ductile levels is induced by variations in the lithostatic pressure gradient generated by differential sediment load, for example due to the construction of prograding sedimentary prisms such as deltas, carbonate platforms and submarine fans (Ramberg, 1981; Gaullier and Vendeville, 2005). According to Rowan et al (2004), the erosion of the continental shelf and slope during lowered relative sea level, as well as bypass events leading to the deposition of large volumes in the slope and in the ocean basin, attenuate the gravity potential and, consequently, the sedimentary cover (Fig. 2.3).
Figure 2.2: Passive margin with basinward-dipping detachment and thus a significant component of gravity gliding (not to scale). Basinward tilting is enhanced by differential thermal subsidence and cratonic uplift and is reduced by proximal loading subsidence. Modified from Rowan et al. (2004).
Figure 2.3: Passive margin with failure dominated by gravity spreading (not to scale). (a) Progradational deposition along the outer shelf and upper slope increases the differential sediment load and thus drives spreading. (b) and (c) Upper-slope bypass and distal deposition on the lower slope and abyssal plain reduces the surficial slope and the gravity potential, slowing or stopping spreading. From Rowan et al. (2004).
Several authors have proposed that gravity tectonics is associated with distinct structural domains, based on the interaction between sedimentation and structuring, and on local morphological characteristics of the substrate. These analyses were obtained based on a combination of improvements in seismic data processing and analogue modeling. The mechanism of extensive surface deformation develops a structural system composed of a proximal extensive domain, mechanically connected to a distal compressive domain (Fig. 2.4; Vendeville & Cobbold, 1988; Cobbold et al., 1989; Nelson, 1991; Cobbold & Szatmari, 1991; Vendeville, 1991; Vendeville e Jackson, 1992a e 1992b; Jackson et al.,1994; Mohriak et al., 1995; Morley & Guerin, 1996). However, other studies recognize an intermediate structural domain, mechanically interconnecting the extensive and compressive domains, characterized by the rigid translation of the sedimentary cover (Fig. 2.4; Damuth, 1994; Hooper et al., 2002; Oliveira et al., 2005; Lecomte and Vendeville, 2008).
Figure 2.4: Schematic model representing a possible structural zonation generated by extensive surface deformation. Modified from Perovano (2008).
The extensive domain is structurally composed of synthetic antithetic faults, rollover anticlines and, in the case of an evaporite layer, salt rollers formed behind the fault plane. The listric faults are predominantly perpendicular to the direction of translation and horizontally in depth, anchored to levels of decollement, accommodating translation and tilting of the sedimentary blocks (Vendeville & Cobbold, 1987; Cobbold et al., 1989).
The compressive domain accommodates the extensional and the translation deformation of the sedimentary cover. The shortening of the sediment cover in this domain can take place in the form of different compressive structures, such as, fold belts, thrust faults, diapirs or walls of salt or shale. The compressive structures usually form at the base of the slope, close to the limit of effective influence of the levels of detachment. In the case of deformation associated with overpressured levels, this limit is defined by the decrease of the interstitial pressurization and consequent loss of mobility of the mobile level (Thomas, 2001; Rowan et al., 2004).
The tectonic evolution of passive continental margins related to occurrence of gravity tectonics with overpressured shales is directly controlled by the balance of the generation and dissipation rates of the interstitial fluid overpressure condition. According to Mourgues (2003), the maintenance of the overpressure state depends on the permeability (porosity) of the rocks and the velocity of generation of interstitial pressure, i.e. it depends on the ratio between the rate of fluid release and loss of pressurization, controlled by porosity and rock fracture, and the rate of pressure generation through physical, chemical and mineralogical mechanisms.

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The Foz do Amazonas Basin

The Foz do Amazonas Basin represents the northernmost basin on the Brazilian Equatorial Margin (Fig. 3.1). It is located off the coasts of the State of Amapá and the northwest coast of the State of Pará, and extends over an area of 268,000 km2 (Brandão and Feijó, 1994). The basin includes a major seafloor depositional feature, the Amazon Fan, one of the world’s largest deep-sea fans, which extends from the continental shelf to abyssal depths of 4800 meters (Silva et al., 1999).
The basins of the Brazilian Equatorial Margin developed as a consequence of the separation of the South American and African plates in a transform tectonic context characterized by multiple phases of subsidence and complex structural styles (Matos, 2000). According to Szatimari et al. (1987) and Matos (2000), rifting of the Brazilian Equatorial Margin occurred in several evolutionary stages between the Aptian and Cenomanian. Matos (2000) recognized three stages in the tectonic evolution of the Equatorial Atlantic: pre-transform, syn-transform and post-transform.

Table of contents :

I. INTRODUCTION
I.1 Geopressure: Problematics and open questions
I.2 The Amazon Margin as a study area
I.3 Objectives of the thesis
II. OVERPRESSURE PROCESSES AND PRODUCTS
II.1 Overpressure mechanims
II.1.1 Desequilibrium compaction
II.1.2 Tectonic stress
II.1.3 Clay Diagenesis
II.1.4 Hydrocarbon generation
II.2 Gravity Tectonics
III. GEOLOGICAL SETTING
III.1 The Foz do Amazonas Basin
III.2 The Amazon Fan gravity tectonics
III.3 Thermal History and Hydrocarbon Potential
III.4 Summary
IV. DATA AND METHODS
IV.1 Dataset
IV.2 Methods
V. Controls on overpressure evolution during the gravitacional collapse of the Amazon deepsea fan
V.1 Introduction
V.2 Geological Setting
V.3 Data and Methods
V.3.1 Well data
V.3.2 Composite seismic reflection profile
V.3.3 Time-depth conversion of interpreted section
V.3.4 2D Structural Restoration
V.3.5 Pore pressure estimation from 1D geomechanical models
V.3.6 2D basin modeling of pore pressure and temperature
V.3.7 2D mechanical modeling
V.4 Model set-up and calibration from well data
V.4.1 2D seismic interpretation (TWTT)
V.4.2 2D lithofacies distribution in depth..
V.4.3 2D Basin modeling set-up
V.4.3.1 Porosity-permebaility relation
V.4.3.2 Boundary conditions
V.4.4 2D mechanical model set-up
V.5 Modeling results along profile A-B
V.5.1 2D restoration of SE Amazon Fan growth and deformation
V.5.2 2D basin modeling of pressure and temperature evolution
V.5.2.1 Present-day pore pressure
V.5.3 Temperature-driven fluid expasion mechanisms
V.5.3.1 Clay mineral transformation
V.5.3.2 Hydrocarbon gas genertaion
V.5.3.3 Present-day fluid expansion (unloading)
V.5.4 Evolution of deformation along the detachment and faults
V.6 Discussion
V.6.1 Pore pressure evolution in the southeast Amazon Fan
V.6.2 Overpressure mechanisms in the Amazon Fan
V.6.3 Implications for other deep-sea fans
V.7 Conclusions

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