A 3D SEISMIC CASE STUDY from OFFSHORE ANGOLA Kehinde

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DEPOSITIONAL ARCHITECTURE AND PROCESSES OF SEDIMENT

GRAVITY FLOWS: A 3D SEISMIC CASE STUDY FROM OFFSHORE

ANGOLA

Kehinde Olafiranye

Department of Earth Science and Engineering

Imperial College London

Thesis submitted in accordance with the requirements for

the degree of Master of Philosophy of Imperial College London

October 2013 Declaration of originality

I declare that this thesis is entirely my own work carried out in the Department of Earth Science and Engineering at Imperial College London under the supervision of Drs Christopher Jackson (Imperial

College), David Hodgson (University of Leeds), and Gary Hampson (Imperial College). Full citation and adequate referencing have been made where any material could be construed as the work of others. No part of this work has been previously submitted to this or any other academic institution for a degree or diploma, or any other qualification.

Kehinde Olafiranye

Department of Earth Science and Engineering

Imperial College London

‘The copyright of this thesis rests with the author and is made available under a Creative Commons

Attribution Non-Commercial No Derivatives licence. Researchers are free to copy, distribute or transmit the thesis on the condition that they attribute it, that they do not use it for commercial purposes and that they do not alter, transform or build upon it. For any reuse or redistribution, researchers must make clear to others the licence terms of this work’

Abstract

Deepwater environments are characterised by the deposits of mass flows (e.g. debrites, slumps, slides), sediment density flows (turbidites), and background hemipelagic and pelagic suspension fallout. ‘Mass transport’ is a general term used for the failure and downslope movement of sediment under the influence of gravity in both subaerial and subaqueous environments, the products of which are called Mass

Transport Deposits (MTDs).

The aims of this study are: (i) to investigate the temporal and spatial development of sediment gravity flows, particularly the coupling of MTDs, turbidites and hemipelagites (ii) to characterise and interpret the seismic facies present within a well-imaged deep-water succession, offshore Angola, with particular focus on the external morphology and distribution of MTDs and associated sediment gravity flows; (iii) to document the geometry, scale, distribution and kinematic importance of structures within the MTDs and relate these to the emplacement of these deposits, in order to understand the impact of seafloor bathymetry on the distribution of the deep-water deposits; and (iv) to characterise and classify the development of pockmarks that are formed contemporaneously or successive to the emplacement of the

MTDs.

Key results of this study are that: (1) the grade of the submarine slope is a primary intra-basinal control on the emplacement and depositional patterns of MTDs and associated sediment gravity flows, (2) the emplacement of MTDs in upper slope environments promotes the capturing and trapping of sediments,

(3) strain indicators genetically linked to MTDs record kinematic information related to the initiation, translation and ultimate deposition of the MTDs, (4) pockmarks may develop from the dewatering of

MTDs due to their rapid transportation and consequent entrapment of fluid, and (5) the distribution of pockmarks in slope sediments may be haphazard, but pockmarks may also form preferentially along fault traces and along the evacuation planes of MTDs.

Table of contents

Declaration of originality ...... 2

Abstract ...... 4

Table of contents ...... 5

List of figures ...... 9

Chapter One ...... 14

Introduction ...... 14

1.1 Rationale and motivation ...... 14

1.2 Aims and objectives ...... 16

1.3 Study area ...... 17

1.3.1 Tectono-stratigraphic evolution ...... 17

1.4 Research methodology ...... 20

1.5 Structure of the thesis ...... 20

Chapter Two...... 22

Overview of sedimentation on submarine slopes and gravity-flow deposition ...... 22

2.1 Introduction ...... 22

2.2 Triggering of gravity flows ...... 23

2.3 Forms and characteristics of gravity flows ...... 24

2.3.1 Mass Transport Deposits (MTDs) ...... 24

2.3.2 Sediment gravity flows ...... 27

2.3.3 Flow translation and transformation ...... 28

2.4 Rheological behaviour and the classification of sediment gravity flows ...... 30

2.5 Impact of active structures on slope sedimentation ...... 31

2.5.1 The concept of slope grade and associated processes ...... 31

2.5.2 Type and impact of bathymetry on slope depositional patterns ...... 33

2.6 Internal architecture and strain indicators within MTDs ...... 36

2.6.1 Headwall Domain ...... 37

2.6.2 Translational Domain ...... 38

2.6.3 Toe Domain ...... 40

2.7 Seismic facies analysis ...... 40

2.8 Impact of MTDs and sediment gravity flows on fluid movement and pockmarks ...... 47

2.8.1 Generation and transfer of overpressure ...... 48

Chapter Three ...... 52

The role of tectonics and mass-transport complex emplacement on upper slope stratigraphic evolution: a 3D seismic case study from offshore Angola ...... 52

3.1 Abstract ...... 52

3.2 Introduction ...... 53

3.3 Geological setting ...... 57

3.4 Dataset and methodology ...... 61

3.5 3D seismic interpretation ...... 65

3.5.1 The Lower Seismic Package (LSP) ...... 65

3.5.2 Middle Seismic Package (MSP) ...... 74

3.5.2.1 Seismic Unit 3 (SU3) ...... 74

3.5.2.2 Seismic Unit 4 (SU4) ...... 74

3.5.2.3 Interpretation of the MSP ...... 74

3.5.3 Upper Seismic Package (USP) ...... 75

3.6 Discussion ...... 78

3.6.1 Stratigraphic organisation of deep-water successions and the role of MTDs ...... 78

3.6.2 Controls on accommodation development and the impact on upper slope depositional patterns 81

3.7 Conclusion ...... 83

Chapter Four ...... 85

3D seismic analysis of pockmark distribution and development within mid-Pliocene to modern strata, offshore Angola ...... 85

4.1 Abstract ...... 85

4.2 Introduction ...... 86

4.2.1 Sources of pockmark-forming fluids ...... 86

4.2.2 Fluid migration pathways for pockmark development and distribution ...... 89

4.2.3 Study aims ...... 90

4.3 Geological setting ...... 91

4.4 Dataset and methodology ...... 96

4.5 Results ...... 100

4.5.1 Density and distribution of pockmarks ...... 100

4.5.2 Morphology of pockmarks ...... 104

4.5.3 Thickness of Seismic Units ...... 107

4.5.4 Fault distribution ...... 110

4.5.5 MTD distribution in unit SU2 ...... 112

4.6 Discussion: Sources of fluid and conduits for the development of pockmarks ...... 115

4.6.1 Water expelled from MTDs and associated sediment gravity flows ...... 116

4.6.2 Hydrocarbons from deep-seated source rocks and breached reservoirs ...... 117

4.7 Conclusions ...... 119

Chapter Five ...... 121

Conclusions and future work ...... 121

5.1 Conclusions ...... 121

5.1.1 Seismic facies, stratigraphic architecture and bathymetric controls on upper slope depositional patterns ...... 121

5.1.2 Spatial distribution and development of pockmarks and controls on their evolution ...... 122

5.2 Future Work ...... 123

5.2.1 Seismic inversion and forward modelling ...... 123

5.2.2 Linking sediment gravity flows stratigraphic succession and eustatic sea-level curve ...... 123

5.2.3 Field analysis of relevant analogues ...... 124

5.2.4 Geochemical analysis of fluid samples ...... 124

Acknowledgements ...... 125

References ...... 126

Appendices ...... 163

List of figures

Figure 1.1 : (a) Map of southwest African margin (b) Map of tectonic framework of the Lower Congo...14

Figure 1.2: Generalized stratigraphic column for the Lower Congo Basin……………………………....16

Figure 1.3: Regional geologic transect, offshore Angola...... 18

Figure 2.1: Submarine flow transformation from initial slide failure to final...... 23

Figure 2.2: Rheological behaviour (stress-strain) relationship...... 24

Figure 2.3: Schematic cross-section illustrating the styles of accommodation developed on continental margins...... 30

Figure 2.4: Schematic representation of a MTD...... 34

Figure 2.5: (A) A density property image which allows comparison of input data with seismic models displayed in B and D. (B) A synthetic seismogram with 65Hz frequency content (C) The 45Hz synthetic seismogram. (D) The 32Hz synthetic seismogram of Ainsa II outcrop...... 41

Figure 2.6: (A) A seismic section of the channel complex in Angola, oblique to depositional strike. (B) A suggested interpretation of the Angolan system...... 42

Figure 2.7: Synthetic seismic-reflection profiles. A) Profile created with 120 Hz frequency, zero-phase

Ricker wavelet. B) 60 Hz profile. C) 30 Hz profile. D) 15 Hz profile...... 44

Figure 2.8: Shallow subsurface seismic-reflection profile from the western Gulf of Mexico……………45

Figure 2.9: Subsurface pressure profiles...... 47

Figure 3.1: Schematic cross-section illustrating the styles of accommodation developed on continental margins...... 52

Figure 3.2: (a) Map of SW African margin showing the location of the Lower Congo Basin in relation to main bounding hinge zones. (b) Map of tectonic framework of the Lower Congo Basin...... 54

Figure 3.3: Regional geoseismic section across the Angolan margin...... 54

Figure 3.4(a): Time-structure map of the base Middle Pliocene illustrating the structure of the study area...... 56

Figure 3.4(b): Broadly strike-orientated (i.e. NW-SE), uninterpreted and geoseismic sections illustrating the structural style of the study area, and the key seismic horizons and units...... 57

Figure 3.5: Stratigraphy, lithologies, depositional elements, and mapped seismic horizons and seismic units within the stratigraphic interval of interest...... 59

Figure 3.6: Broadly strike-orientated (i.e. NW-SE), uninterpreted and geoseismic sections through the relatively proximal part of the eastern depocentre...... 60

Figure 3.7: Frequency range extracted from (a) the entire 3D seismic cube and (b) within the interval of interest...... 61

Figure 3.8: Seismic and geoseismic sections through the relatively distal part of the eastern depocentre illustrating the external form of, and the internal seismic facies and structures in SU1–5...... 63

Figure 3.9: Morphology of Horizon A...... 64

Figure 3.10: Thickness and complimentary palaeogeographic interpretation maps of SU1-6...... 66

Figure 3.11: (a) Broadly strike-orientated (i.e. NW-SE) seismic section through the relatively proximal part of the eastern depocentre. (b) Broadly dip-orientated (i.e. ESE-WSW) seismic section through the eastern and western depocentre. (c) Instantaneous root-mean square (RMS) amplitude map illustrating seismic expression of the top surface of SU2...... 67

Figure 3.12: (a) Time-structure map of a reflection event within the megaclast within the SU2 MTD. (b)

Instantaneous dip map of a reflection event within a feature within the SU2 MTD...... 69

Figure 3.13: Seismic expression, morphology, scale and distribution of slope gullies in the upper part of the studied succession...... 75

Figure 3.14: Schematic cross-sections illustrating the styles and evolution of relief and accommodation types on continental margins based on our observations from offshore Angola...... 79

Figure 3.15: Chart illustrating the styles, dimensions and orientations (relative to slope) of different styles of tectonic and deposition relief and accommodation on continental margins...... 81

Figure 4.1: (a) Map of southwest African margin showing the location of the Lower Congo Basin in relation to main bounding hinge zones. (b) Map of tectonic framework of the Lower Congo Basin...... 90

Figure 4.2: Regional geologic transect offshore Angola...... 91

Figure 4.3: (a) Structural map illustrating the setting of the study area. (b) Oblique-dip oriented geoseismic profile, across the lower slope, showing the analysed seismic units (SUs) and structures in the study area...... 93

Figure 4.4: Strike oriented vertical profiles of the analysed seismic units (SUs) and structures in the study area...... 94

Figure 4.5: Description of the six mapped seismic units (SUs) and their constituent seismic facies...... 95

Figure 4.6 : Illustration of stratal architecture in the analyzed seismic units (SU1-SU6)...... 96

Figure 4.7: Classifications of (a) pockmark type based on their morphological and reflector characteristics, and (b) pockmark spatial distributions...... 98

Figure 4.8: Variance maps of the top surfaces of (a-f) seismic units SU1-SU6...... 99

Figure 4.9: Maps showing the spatial distribution and density of pockmarks at the top surfaces of (a-d) seismic units SU3-SU6...... 100

Figure 4.10: Detailed variance maps illustrating the distribution of pockmarks developed at the (a) top-

SU3 and (b) top-SU4 surfaces. (c) Uninterpreted seismic and (d) interpreted geoseismic sections along a string of pockmarks...... 101

Figure 4.11: Detailed variance map illustrating the distribution of pockmarks at the (a) top-SU5 surface.

(b) Uninterpreted seismic and (c) interpreted geoseismic section showing fault systems in relation to pockmarks...... 102

Figure 4.12: Detailed variance map illustrating the distribution of pockmarks at the (a) top-SU6 surface.

(b) Uninterpreted seismic and (c) interpreted geoseismic section showing fault systems in relation to pockmarks...... 102

Figure 4.13: Graphs of long axis versus short axis for pockmarks developed at the top surfaces of seismic units (a-d) SU3, SU4, SU5, and SU6...... 104

Figure 4.14: Graphs of pockmark heights versus average diameter for the pockmarks developed at the top surfaces of seismic units (a-d) SU3, SU4, SU5, and SU6...... 105

Figure 4.15: Isochore maps illustrating the thickness of seismic units (a-f) SU1 to SU6...... 107

Figure 4.16: Maps of (a-c) the chaos seismic attribute, and (d-f) interpreted deposits from iso- proportional slices taken at three levels within seismic unit SU2: (a, d) lower slice, taken at 25% thickness from the base of the unit; (b, e) middle slice, taken at 50% thickness from the base of the unit; and (c, f) lower slice, taken at 75% thickness from the base of the unit...... 112

Figure 4.17: (a) Uninterpreted seismic and (b) interpreted geoseismic sections of seismic facies in the area upslope of the eastern depocentre aligned along depositional dip...... 113

Figure 4.18: (a) Uninterpreted seismic and (b) interpreted geoseismic sections of seismic facies in the area upslope of the eastern depocentre aligned along depositional strike...... 114

Figure 4.19: Model for the development of pockmarks in the Mid-Pliocene to Holocene succession in

Block 4, Lower Congo Basin, offshore Angola...... 118

List of tables

Table 2.1: Classification of sediment gravity flows on the basis of rheology……………………………31

Chapter One

Introduction

1.1 Rationale and motivation

Deepwater environments are characterised by the deposits of mass flows (e.g. debrites, slumps, slides), sediment density flows (turbidites), background hemipelagic and pelagic suspension fallout, and contour currents. Mass flow events are common to all continental margins, but the processes priming and triggering the initial slope failure remain poorly understood (Harders et al, 2010; Mosher et al., 2010).

The predominance of deposits of mass flow on most continental margins around the world indicates that they are an important element of any deep-water stratigraphic succession (Moscardelli and Wood, 2008).

‘Mass transport’ is a general term used for the failure and downslope movement of sediment under the influence of gravity in both subaerial and subaqueous environments (Shanmugam, 2006) - the product of which is called Mass Transport Deposits (MTDs). As a result of water depth and due to their complex internal character and large scale, MTDs are frequently investigated using geophysical methods (e.g. high-resolution seismic reflection or sidescan-sonar data; Masson et al., 2006). MTDs are typically characterised on seismic reflection data by sharp-based and sharp-topped packages of chaotic seismic reflections of variable amplitude and frequency (Weimer and Shipp, 2004). The emplacement of MTDs commonly creates bathymetric relief that may influence subsequent sedimentation patterns in deep-water settings (Shanmugam, 2000; Marr et al., 2001; Posamentier and Kolla, 2003; Moscardelli et al., 2004;

Pickering and Corregidor, 2005; Armitage et al., 2009; Jackson and Johnson, 2009). While the occurrences of MTDs above an erosive surface is common knowledge (various references herein), the influence of varying lithology on the magnitude of translation has rarely been discussed. MTDs vary in thickness from several metres to several tens of metres, and they may vary in areal extent from a few tens to hundreds km 2 (e.g. Dan et al., 2010; Gee et al., 2006; Moscardelli et al 2006) to several thousands to tens of thousands of km 2 (e.g. Bryn et al., 2005; Van der Merve et al., 2009).

There are a number of economic reasons to understand the triggers for and range of processes involved in the initiation, transport and arrest of MTDs.

1. Geohazards: Due to large volume of material transported within them and the velocity at which

they move across the seabed (propagated at about 90 km/h and reaching close to 450 km/h before

impact; modelling simulation study of Harbitz 1992), processes of MTDs may represent a

geohazard and pose a threat to offshore installations (e.g. platform instability and seafloor cables)

and coastal communities (Mulder and Cochonat, 1996; Frey-Martinez, 2010). For example, the

deposits of large, recent submarine slope failures are observed at the modern seabed and near

seafloor in numerous oil provinces (e.g. Israel (Almagor and Garfunkel, 1979), Norway Storegga

tsunami (Bryn et al., 2005), the Niger Delta (Nissen et al., 1999), and Gulf of Mexico (Sawyer et

al., 2009)).

2. Hydrocarbon exploration: MTDs can be critical for the hydrocarbon prospectivity of sedimentary

basins as they represent a highly effective mechanism by which large volumes of sediments are

transferred from shallow to deep water (Frey-Martinez, 2010). Furthermore, their emplacement

can influence basin bathymetry, the routing of turbidity currents and the deposition of reservoir-

quality sandstones (Prather et al., 1998; Armitage et al., 2009; Jackson & Johnson, 2009).

3. Hydrocarbon production: The deposition of MTDs can alter the original textural properties and

external and internal architecture of reservoir intervals, generate traps, and influence migration

pathways. When mass flow occurs, the original stratification is modified, the lithologic

compositions are mixed, and the primary volumes of sediment are altered (Frey-Martinez, 2010).

This reworking may decrease net-to-gross, reduce porosity, and degrade the permeability of the

deposits. In addition, the occurrence of large-scale MTDs can trigger seal failure and vertical

migration of hydrocarbons due to a rapid drop in lithostatic pressure (Haflidason et al., 2002,

2004).

4. Drilling operations: The lithological character of MTDs and their bounding surfaces may impact

the operational performance of jetted conductors and suction anchors pile design (Newlin 2003;

Shipp, et al., 2004). For example, the deposition of MTDs is commonly associated with the

expulsion of pore waters (Piper et al., 1997) and, as a consequence, the shear zones at the base of

MTDs in the near subsurface are commonly overcompacted when compared to unfailed

sediments. It has been demonstrated that sediment overcompaction can significantly reduce the

rate of conductor penetration during deepwater jetting operations (Shipp et al., 2004).

1.2 Aims and objectives

Few studies have investigated the entire length of an MTD, from its source area to the distal limit of the deposits. Also, few analyses have looked at the internal kinematics and the nature of the basal and upper surfaces of MTDs. This project uses 760 km 2 of high quality 3D seismic data from Block 4 (Lower

Congo Basin) within the continental slope offshore Angola. The key aims and objectives of the research project are to:

1. Understand controls and processes governing MTD emplacement

2. Deduce the source of the fluids responsible for the development of pockmarks

3. Investigate the temporal and spatial development of sediment gravity flow deposits particularly

the coupling of MTDs, turbidites and hemipelagites.

4. Analyse the impact and implications of slope topography and bathymetry on the distribution and

emplacement of MTDs.

5. Characterise the style and controls of the topography developed along the upper surface of

MTDs.

6. Map and document the relevant kinematic indicators (such as striations, grooves, ramps and flats,

faults, and folds) and their distribution within MTDs and their embedded megaclasts.

7. Document and characterise the morphology and distribution of pockmarks and assess the controls

of sediment gravity flows on the formation of pockmarks.

1.3 Study area

The study area is located within the Lower Congo Basin Block 4 in water depth 225-700 m offshore Angola. It lies between 7°30’S and 7 °45’S which is approximately 40-100 km from the present- day coastline (Fig1.1).

Fig. 1.1: (a) Map of southwest African margin showing the location of the Lower Congo Basin in relation to main bounding hinge zones. (b) Map of tectonic framework of the Lower Congo Basin indicating in red box the location of Block 4 in relation to other Blocks. Modified from Anderson et al., (2000).

1.3.1 Tectono-stratigraphic evolution

The Lower Congo Basin (Fig. 1.1) along with several other major sub-basins in the passive continental margin of West Africa was generated during the Early Cretaceous in response to the opening of the South

Atlantic Ocean (Brice et al., 1982; Marton et al., 2000). The main source rocks for the Lower Congo

Basin are the bituminous shales of the Lower Cretaceous Bucomazi Formation deposited during the rift phase (Burwood, 1999; Cole et al., 2000), which expelled oil and gas from the early Miocene onward

(Cole et al. 2000) (Fig. 1.2). After the rift phase, a thick evaporite accumulation, the Loeme Formation, was deposited (Fig. 1.2) (Brice et al., 1982; Uchupi, 1992), which later facilitated gravitational sliding of

17 the postrift sequence, by acting as a decollement layer, resulting in the formation of numerous salt structures and salt minibasins (Duval et al., 1992; Lundin, 1992; Liro and Coen, 1995; Valle et al., 2001;

Fort et al., 2004). The postrift sequence is thus divided into an upper slope extensional domain characterized by grabens and rafts and a downslope contractional domain characterized by salt diapirs, canopies, and walls (Fig. 1.3) (Lundin, 1992; Marton et al., 2000; Rouby et al., 2002). The study area which comprises salt-induced grabens and rafts is located at the transition between the salt pillow and turtle back province and the unrafted Pinda shelf (Fig. 1.3). From Late Cretaceous to Early Oligocene time, an aggradational carbonate/siliciclastic ramp developed in response to low-amplitude/low-frequency sealevel changes and continuously hot climate (greenhouse period) (Séranne, 1999). This interval includes the Tertiary source rocks producing thermogenic oils and gas by cracking of organic matter

(Burwood, 1999). From Early Oligocene to present time, the sedimentation is dominated by the progradation of a terrigenous wedge that reflects high-amplitude/high-frequency sea-level changes and alternating dry and wet climatic conditions. The high-frequency climate variations triggered deep incision and erosion and deposition of large amounts of terrigenous material in a turbiditic fan off Congo and

Angola (Brice et al., 1982; Uchupi, 1992; Droz et al., 1996). The important spreading of the Congo plume over the canyon edge has led to the deposition of mud-dominated sediments on the slope, whereas turbidite currents are transferred to the deep-sea fan within the Zaire canyon (Bentahila et al., 2005). The

Pliocene to Present sediments constitute a seal above the Oligocene-Miocene turbiditic interval

(Burwood, 1999). Thermogenic hydrocarbons, produced in the pre-salt lacustrine sequence and in the clastic sediments and marls of Albian–Cenomanian Moita Seca source rocks (Cole et al., 2000), have migrated upward and are now preferentially trapped within the Oligo–Miocene silty-sandy channels (Gay et al., 2006b), constituting the most prolific hydrocarbon reservoirs in this area (Burwood, 1990). From early Pliocene time to present, the shelf and the slope were deeply incised by a canyon of the proto-Congo

River (Gay et al 2007).

Fig. 1.2: Generalized stratigraphic column for the Lower Congo Basin. Taken from Valle et al., 2001.

Fig. 1.3: Regional geologic transect offshore Angola. The major structures and depositional sequences are highlighted. COB is the continent– ocean boundary. Note the occurrence of the salt layer within the study area, indicated by the red box. The location of the transect is indicated in Figure 1.1b. Modified from Lentini et al., (2010).

1.4 Research methodology

This study uses a single volume of 3D seismic dataset covering an area of 760 km 2 and more than five seconds two-way travel time (TWT) vertical interval. Velocity data needed to make time-depth conversion was obtained as an analogue from published studies that used similar seismic dataset offshore

Angola. The study concentrated on the topmost 1.5 seconds of the data. Seismic facies were defined and surfaces (tops and bases) were mapped for seismic units identified within the study interval. Isopach maps were generated to analyse thickness variability in each unit. Seismic attribute extractions (dip, RMS amplitude, and variance) were carried out on the interpreted surfaces to investigate the physical geological features identified. Manual counting of pockmarks was carried out to understand their density and link their distribution to geological controls. The detailed methodologies are outlined in Chapter 2 and further described in Chapters 3 and 4.

1.5 Structure of the thesis

This thesis is made up of five chapters. Chapter One gives an introduction of the research and comprises the main motivation and rationale of the research, including an overview of the subsurface geological setting. The chapter also presents highlights of the approach and methodology of data analysis.

Chapter Two provides the reviews of the relevant literatures on sedimentation on submarine slopes and gravity-flow deposition. This involves factors that cause the triggering of gravity flows, their subsequent

20 translation and emplacement, the different forms and characteristics of gravity flows, and the link between gravity flows and rheology. It also outlines the impact of active structures on slope sedimentation in the context of slope grade and slope bathymetry and their resultant depositional patterns. The chapter discusses the internal architecture and strain indicators in the tripartite domains of MTDs, summarizes the type seismic facies that identify gravity flow deposits and underlines the impacts of gravity flows on fluid movement and the development of pockmarks.

Chapters Three and Four are written in the form of a paper as they discuss the analysis and results of the main research. Chapter Three ( www.sciencedirect.com/science/article/pii/S0264817213000536 -

Marine and Petroleum Geology 2013) presents a 3D seismic analysis of the role of tectonics and mass- transport complex emplacement on the upper slope stratigraphic evolution with the major highlight being a model illustrating the styles and evolution of relief and accommodation types on continental margins.

Chapter Four describes the controls of the spatial distribution and development of pockmarks, and their significance in fluid migration within Pliocene to modern strata, offshore Angola, West Africa.

In Chapter Five a synthesis of the major findings of the research is presented in relation to the stated aims and objectives (section 1.2).

Chapter Two.

Overview of sedimentation on submarine slopes and gravity-flow deposition

This chapter provides a review of background literature that underpins the independent research presented in Chapters 3 and 4. The chapter is structured as follows: 1) General introduction highlighting slope morphologies and their associated re-sedimentation processes; 2) Initiation processes including some factors which cause the triggering of gravity flows; 3) Classification of sediment gravity flows linked to their rheological behaviour and characteristics; 4) Control and influence of topo-structures on slope sedimentation of gravity flows; 5) Development of internal architecture and strain indicators within

MTDs and their significance; 6) Seismic facies analysis relevant to the identification of gravity flows: 7)

Influence of gravity flows on fluid movement and their relationship to the development of pockmarks.

2.1 Introduction

Submarine slopes are considered to be the inclined part of the sea floor between the shelf-slope break and the basin floor (Shanmugam, 2006). Modem continental slopes around the world average about

4o (Heezen et al., 1959), but slopes can range from less than 1 o to greater than 40 o. Slopes on active margins (e.g. Califomia and Oregon) are generally steeper than those on passive margins (e.g. Louisiana)

(Pratson and Haxby, 1996).

Prior et al. (1984) and Prior and Bornhold (1988) pointed out the importance of slope instability on the morphology and sedimentary processes of British Columbia fjord delta fans where most of the slope failures have been attributed to liquefaction of delta-front sediments because of excessive regional sedimentation. In the Norwegian–Barents Sea continental margins, Elverhøi et al. (1997) and Laberg and

Vorren (2000) suggested that the extensive failure which develops in thick sequences of clay-rich glaciogenic sediments on the upper slope is the causative agent for the observed downslope sediment movement. Piper et al. (1985) and Papatheodorou and Ferentinos (1993) linked sediment failures in the

Gulf of Corinth trench, and on the Grand Banks continental slope to the generation of earthquakes. The

22 complex interaction of subsurface structuration and slope-instability processes is the main factor controlling the geomorphological and sedimentological evolution of the intra-slope basins of the northwest Gulf of Mexico continental slope (Lee, 1990). In addition to a variety of depositional processes that played major roles in controlling facies distribution in intra-slope basins (Shanmugam, 2006), some of the principal factors that controlled the sea-floor topography in intra-slope mini-basins include salt tectonics, compression, and strike slip along thin-skinned transfer zones (Apps et al., 1994) in which salt withdrawal and diapirism are particularly important. These two factors, salt withdrawal and diapirism, resulted in sea-floor topography of modem continental slopes that is quite complex in terms of shape and relief (Shanmugam, 2006).

Sedimentation on submarine slopes is dominated by re-sedimentation processes arising from slope instability. For example, the oversteepening of the flanks by the seaward mobilization of underlying salt masses resulted in the abundance of slumps, slides, and debris flows in the Beaumont Basin

(Tripsanas et al., 2004). Gardner et al. (2002) indicated that active shedding of large debris blocks on slope and base-of-slope environments is the result of underwater landslides. Generally, deposits of muddy slumps and debris flows and of pelagic-hemipelagic settling are the dominant depositional facies in slope environments (Bryant et al., 1990).

2.2 Triggering of gravity flows

Landslides are initiated when the shear stress oriented downslope (driving shear stress) exceeds the shear strength (resisting stress) of the slope-forming material (Hampton et al., 1996). The initiation of mass transport deposits (MTDs) has been attributed to several factors, such as: salt movement in intra- slope basins (Tripsanas et al., 2004); high sedimentation rates on the fronts of large deltas (Coleman and

Prior, 1982); earthquakes and seismic activity (Radbruch-hall and Varnes, 1976; Fine et al. 2005); hurricanes or cyclic loading by storms (Prior and Coleman, 1982); loading or over-steepening of slopes

(Assier-Rzadkiewicz et al., 2000); glacial deposit loading and reloading (Mosher et al., 2004); under- consolidation or generation of overpressure in the substrate (Prior and Coleman, 1982); precipitation (i.e. where landslides have a subaerial extension; Longva et al., 1996); slope-parallel, mechanically-weak

23 layers in bedded sedimentary sequences (Haflidason et al., 2003; Bryn et al., 2005); gas hydrate dissociation (Sultan et al., 2003); sea-level change (Weaver and Kuijpers, 1983); emplacement of overlying MTDs (Van der Merwe et al., 2009); volcanic activity (Masson et al., 2002); and meteorite impacts (e.g. the Chicxulub impact at K-T boundary in the Yucatan, Mexico; Claeys et al., 2002).

Overall, these factors can be divided into two main types; 1) those related to the intrinsic geological characteristics of the landslide material (e.g. overpressure due to rapid deposition or the presence of a weak layer) and 2) those driven by transient external events (e.g. earthquakes or climate change). In fact, for many examples, the initial slope failure event results from a combination of the factors listed above

(Masson et al., 2006). However, the pre-failure conditions and mechanics of flow of MTDs are rarely known and have to be inferred from the architecture and kinematic indicators within, below or above the associated deposit, or within the upslope source area (Mosher et al., 2010). Numerical modelling (e.g.

Mazantti, 2010) and flume experiments (e.g. Zakeri et al., 2010) provide additional insights and constraints into the processes that may trigger gravity flows.

2.3 Forms and characteristics of gravity flows

Gravity-driven processes in deep-water environments are broadly classified into two types: mass transport and sediment gravity flows.

2.3.1 Mass Transport Deposits (MTDs)

‘Mass transport’ is a general term used for the failure and en masse downslope movement of sediments under the influence of gravity in both subaerial and subaqueous environments (Shanmugam,

2006), the products of which are called Mass Transport Deposits (MTDs). Mass-transport processes are composed mainly of slides, slumps, and debris flows (Dingle, 1977; Varnes, 1978; Hampton et al., 1996), which are important gravity-driven processes for transporting deep-water sand and gravel (Shanmugam,

2006).

2.3.1.1 Slides

Slides are movements of coherent masses of sediments bounded on all sides by distinct failure planes where the internal structure of the mobilized material is largely undisturbed, such that internal deformation during their translational movement and displacement is limited (Mulder and Cochonat,

1996; Shanmugam, 2006) (Fig. 2.1).

Fig. 2.1: Submarine flow transformation from initial slide failure to final plume or hemipelagic deposits. (After Mulder and Cochonat, 1996).

Submarine slides can travel hundreds of kilometres, and run-out distances of between 50 – 100 km are common (Hampton et al., 1996). However, longer run-out distances can be attained in some situations

(e.g the Nuuanu Slide, offshore Hawaii; Normark et al., 1993). Slide development can be either simple or complex. In a simple slide, the main slide body does not generate other significant failures, although some secondary smaller slides may form at the scar, and there is no significant retrogression of these smaller slides (Prior and Suhayda, 1979). In complex slides, the motion of the sedimentary body causes instability in adjoining sediments, and the volume of such induced secondary slides is similar to the volume of the initial slide or slump in a process termed retrograde failure (Prior and Suhayda, 1979). Pull-apart features may be present particularly near the head of the slide (Lewis 1971). Key diagnostic characteristics of slides are their composition of gravel to mud lithofacies, lateral extent of up to 100s km, transported

25 shallow-water blocks or megaclasts encased in deep-water strata (e.g Twenhofel, 1932), primary basal glide plane, basal zone of shearing, secondary internal glide planes, and upslope tensional faults (e.g

Booth et al., 1993).

2.3.1.2 Slumps

Slumps are also coherent masses of sediment that move on a concave-up glide plane and undergo rotational movements causing some degree of internal deformation. In practice, distinguishing slides from slumps can be challenging (Shanmugam, 2006). In slumps, large folds are commonly overturned in the direction of movement; however, folds and faults can have any orientation and may even be parallel to the dip of slope (Thompson 1973). The general characteristics for the identification of slump deposits include their composition of gravel to mud lithofacies, basal zone of shearing, upslope tensional faults, downslope compressional folding or thrusting (i.e. toe thrusts), slump folds overlain by undeformed strata, contorted layers interbedded with uncontorted layers at core scale, and irregular upper contact (Shanmugam, 2006).

2.3.1.3 Debris flows

The addition of fluid to the slumped or sliding sediment mass may transform it into a debris flow , which transports sediment as an incoherent body in which inter-granular movements predominate over shear-surface movements (Shanmugam, 2006) (Fig. 2.1). In debris flows, the sediment-fluid mixture moves en masse (Dasgupta, 2003) with characteristic Bingham plastic behaviour (Fig. 2.2).

Fig. 2.2: Rheological behaviour (stress-strain) relationship showing a range of deformations with two end-members, Bingham plastic and Newtonian fluids. The graphs show that there is a fundamental rheological difference between debris flows (Bingham plastics) and turbidity currents (Newtonian fluids). Modified from Dott, 1963 and Shanmugam, 1997a.

Debris flow deposits show little consistent preferred fabric (Nardin et al., 1979) within layers of varying thickness consisting of a relatively fine-grained 'matrix' with fragments of original bedding and laminae in lens-like angular to well rounded pieces (Pickering et al. 1986). Grading and fabric are generally poorly

26 developed. Fabrics that develop can have a-axis patterns that are parallel to flow or random with imbrications oriented either up or down flow (Harms 1974). The beds in debris flows are generally massive and unstratified and a basal zone of inverse grading with imbrication may occur as a consequence of shearing (Nardin et al., 1979). Internal flow structures can occur at any level within the deposit

(Hampton, 1975). In some cases, large clasts may project above the bed level (Walker 1975; Middleton and Hampton, 1976).

2.3.2 Sediment gravity flows

2.3.2.1 Turbidites

A turbidity current is a gravity-driven flow in which fluid turbulence maintains grain dispersion in large part of the flow (Bagnold, 1962; Shanmugam et al., 2006). A turbidity current is characterized by the flow of a sediment–fluid mixture within an ambient fluid body, and the density contrast between the flow and the ambient fluid body is caused by high concentration of sediments within the flowing mass

(Simpson, 1982; Allen, 1985). The duration of many turbidity currents, especially the largest ignitive flows and the hyperpycnal plumes, from many ten of hours up to several days, allows most of them to be considered as quasi-steady or steady flows (Normark and Piper, 1991). Turbidity currents are Newtonian in rheology (Fig. 2.2), and do not have the strength to carry coarse-grained sand and gravel (Dott, 1963;

Lowe, 1979), except that isolated outsized clasts can be transported at some height within the flow and subsequently deposited in such a position (Postma et al., 1988). When a turbidity current results from the transformation of a slide or a slump (Fig. 2.1), the scarp in the source area can give a rough idea of the original sediment volume, but this volume can be strongly increased by erosion during translation of the precursor slide or slump (Shanmugam 2006). On seismic profiles, turbidity current deposits are generally characterized by continuous strong and multiple parallel reflectors indicating stratified sandy and muddy turbidites (Mulder and Cochonat, 1996).

2.3.2.2 Pelagites and hemipelagites

Pelagic sediments accumulate in the open ocean, and primarily consist of the skeletal parts of planktonic organisms, together with some minor amounts of very fine silt and clay, much of which has reached the open ocean by aeolian transport. These sediments are deposited mainly through very slow accumulation by settling, with dissolution or dilution of biogenic material (Stow and Piper, 1984c;

Pickering et al. 1986). The proportion of terrigenous material in pelagites may be increased by the preferential dissolution of the biogenic fractions, such that a true record of the primary depositional composition is not preserved (Stow and Piper, 1984c; Dean et. al., 1985). Hemipelagites are also fine- grained sediments comprising mixtures of biogenic and terrigenous material that have been deposited by a combination of vertical settling and slow lateral advection, and they are widespread in most deep-sea environments, especially those close to continental margins and in other settings near to terrigenous sediment sources (Pickering et al. 1986; Stow and Tabrez, 1998).

Both pelagites and hemipelagites have similar depositional dynamics characterised by slow accumulation of waning energy, late-stage flows in quiescent deepwater environments (Fig. 2.1). Pelagites and hemipelagites are generally characterised by mudstone lithofacies, parallel lamination, faint normal grading, bioturbation, deep-marine body fossils and trace fossils, and sheet-like geometry conformable to underlying sea-floor topography (Shanmugam, 2006).

2.3.3 Flow translation and transformation

Subaqueous gravity-driven processes such as slides, slumps, debris flows, and turbidity currents, are important agents for transporting sediments downslope into deep-marine environments (Shanmugam,

2006). After failure of a submarine slope, these flow processes may transform dynamically as part of a continuum (Dott, 1963; Varnes, 1978; Martinsen, 1994; Hampton et al., 1996; Mulder and Cochonat,

1996; Bull et al., 2009) (Fig. 2.1). Instability of the earlier sediments is the pre-requisite for generation of flows involved in resedimentation processes. Such instability normally comes into existence either due to over-steepening of the parent deposit (Fairbridge, 1947; Banerjee, 1966; Shepard et al., 1977; Prior et al.,

1986, 1987) or through some process of liquidization including seismogenic slumping and storm wave-

28 induced landslides (Heezen and Ewing, 1952; Morgenstern, 1967; Piper et al., 1985; Prior et al., 1989;

Weaver et al., 1992; Garcia and Hull, 1994; Lee and Chough, 2001). Remobilization of sediments through subaqueous sliding and slumping has also been reported by Syvitski and Farrow (1983), Prior et al. (1984,

1986), Friedman et al. (1992), and Lee et al. (2002). Friedman et al. (1992) specifically pointed out that subaqueous slumping is broadly of two types: (1) coherent slumping and (2) incoherent slumping. In coherent slumping, the strata within the displaced mass remain intact and may be deformed, while in incoherent slumping, the entire body of sediments is thoroughly mixed and churned, and almost all traces of stratification are obliterated. The entrainment of ambient fluid (usually sea water) reduces the strength and increases the fluid behaviour of the debris, thereby causing it to flow rather than slide (Hampton,

1972). For example, the incorporation of only a few percent of water typically decreases the strength of landslide debris by a factor of two or more (Shanmugam, 2006), thereby making landslide debris commonly fluidized with incorporation of a small amount of water. But after initial movement, if the downslope component of the gravitational pull exceeds the yield strength of the moving mass, the flow continues as a debris flow. Under subaqueous conditions, the moving sediment-fluid mixture undergoes body transformation (Fisher, 1983) on reaching a steeper slope. Under such conditions, the moving mass is accelerated (Laval et al., 1988) due to enhanced gravitational pull, and the flow tends to become turbulent. This further disturbs the relative tranquillity of the adjacent fluid due to inertia contrast, so that a part of the ambient fluid becomes entrained in the sediment-fluid mixture, and the flow becomes a turbidity current due to a decrease in grain/water ratio of the flowing mass. Hampton (1972) postulated that dilution and stripping of surface materials might cause generation of turbidity currents from subaqueous debris flows. Piper et al. (1985) also predicted transformation of subaqueous debris flows into turbidity currents.

Flow types involved in resedimentation processes are commonly generated through transformation of flows due to gradual change in grain/water ratio of the flowing system (Dasgupta,

2003).

2.4 Rheological behaviour and the classification of sediment gravity flows

Subaqueous sliding and debris flows represent relatively cohesive flow types, and should be distinguished from relatively dilute, sediment gravity flows such as turbidity currents (Dott, 1963). The rheological behaviour of MTDs, as expressed by the relationship between applied shear stress and shear strain (Fig. 2.2), is an important property to classify gravity flows (Gani, 2004; Shanmugam, 2006) in terms of plastic and Newtonian rheologies. Plastic deformation occurs at low stresses and proceeds as elastic deformation (Allaby and Allaby, 2003). Plastic deformation occurs in rocks under conditions of high pressure and temperature, producing a permanent alteration of their shape, but without failure by material rupture. In contrast, Newtonian behaviour is the manner in which a fluid flows if it exhibits ideal viscous strain (Allaby and Allaby, 2003). Newtonian fluids (i.e. fluids with no yield strength), like water, will deform instantly and permanently when shear stress is applied and the deformation is linear (Fig. 2.2)

(Shanmugam, 2006). Cohesion and friction between clastic grains lend a degree of stability to sediments, while gravity-induced, downslope shear stress acts in opposition to this (i.e. it causes distortion of the sediment-fluid mix) (Dott, 1963). A distinction occurs if the cohesive strength is exceeded and more-or- less a coherent plastic mass (i.e. a debris flow) passes to a turbid, viscous suspension (i.e. turbidity current). Debris flows have been modelled as flow of a plastico-viscous substance (Hampton, 1972), or a non-Newtonian (Coussot and Meunier, 1996) or plastic rheology (Dott, 1963; Johnson, 1970), which has yield strength and deforms viscously at stresses greater than the yield strength (Fig. 2.2) This is analogous to the classification originally proposed by Dott (1963). In a non-Newtonian Bingham plastic, a critical value of shear stress (the yield stress) has to be crossed before there is any deformation, after which the deformation is linear i.e., a Bingham plastic is a combination of an ideal plastic and a Newtonian fluid

(Gani 2004). In a non-Newtonian dilatant fluid there is no yield strength, but the deformation is nonlinear with applied stress in such a way that it becomes progressively harder to deform the fluid (Fig. 2.2).

Turbidity currents exhibit Newtonian rheology, and debris flows have a non-Newtonian rheology. Debris

30 flows can be divided further into ‘cohesive debris flows’ (non-Newtonian Bingham plastics), and ‘non- cohesive debris flows’ (non-Newtonian dilatant fluids). Table 1 is a classification by Gani (2004) which encompasses MTDs and their associated sediment gravity flows.

Table 2.1: Classification of sediment gravity flows with a simplified nomenclature for the flow types as well as their deposits. Flow rheology is the basis of this classification. Direction of increasing fluid content is roughly analogous to the down-slope evolution of sediment gravity flows. Flow states and sediment support mechanisms are incorporated to give a comprehensive picture of the nature of these flows (after Gani, 2004).

2.5 Impact of active structures on slope sedimentation

2.5.1 The concept of slope grade and associated processes

Upon dislodgement from the shelf edge, sea floor gradient and confinement are the two primary intra-basinal controls on flow distance, location of deposition, and the mechanisms and types of sediment gravity flows (Pirmez et al., 2000). Sea floor gradient or slope grade is essentially governed by slope angle (Prather et al., 1998; Prather 2000, 2003), which invariably creates various types of accommodation on the slope. Slope grade adjusts across the stratigraphic succession between two end-members; ‘graded slope’ and ’above-grade slope’ (e.g. Prather et al., 1998; Prather 2000, 2003). A graded slope lacks significant topography whereas an above-grade slope is characterized by well-developed enclosed intra- slope basins or subtle changes in depositional gradient which result in low-relief stepped or terraced

31 topography. The type and distribution of accommodation and the relationship of the depositional surface to an idealized or hypothetical graded profile (Fig. 2.3) differentiate between above-grade and graded slope types (Prather 2000, 2003; Steffens, 2003), where the hypothetical profile is defined by the long- term morphodynamic equilibrium between depositional and erosional processes on the slope (Ross, 1989;

Swift and Thorne, 1991).

Fig. 2.3: (a) Schematic cross-section illustrating the styles of accommodation developed on continental margins. Note the development of ponded accommodation on the slope due to normal faulting, salt flow and thrusting, which are all types of ‘tectonic accommodation’. Modified from Prather (2003) and Steffens et al. (2003). The inset diagram illustrates the styles of depositional accommodation that can develop due to the emplacement of MTDs. In this case, relief and accommodation can form in response to the emplacement of ‘megaclasts’ (left-hand side of diagram) or deposit ‘buckling’ (right-hand side of diagram), which may be related to shortening of the deposit during abrupt flow freezing. Modified from Armitage et al. (2009) and Jackson and Johnson (2009).

Deposition of MTDs by downslope translation which may be accompanied by sub-aqueous background sedimentation usually takes place on an above-grade slope characterised by mid-slope structures (e.g. salt- or thrust-cored ridges; Beaubouef and Friedman, 2000; Prather, 2003). The trigger and emplacement of MTDs contribute to the evolution of above-grade slope as a result of the development of rugose topography at the top of MTDs (Kneller and McCaffrey, 1999; Frey-Martinez et al., 2006). During the deposition of turbidites, the relief of antecedent topography is often diminished due to the filling of ponded accommodation, which is accompanied by a decrease in gradient that leads to flow deceleration, reduction in flow competence, and eventual deposition (Prather et al., 1998; Prather, 2000, 2003; Sprague et al., 2005). Where minor relief is still present during the flow of turbidity currents, hydraulic jump and

32 flow-stripping represent a viable mode of sediment transportation across topographic highs into adjacent depocentres (e.g. Komar, 1971; Piper and Normark, 1983; Hodgson et al, 1992). Sediment transport through regions of topographic relief may also be achieved through shunting, fill and spill, switching and diversion (e.g. Hodgson et al, 1992).

Overall, an above-grade slope causes inefficient delivery and distribution of sediment gavity flows by concentrating potential coarse grained, sandy sediments around the upslope edge of local topographic highs (Mutti, 1985; Bouma, 2004; Sprague et al., 2005). Once topographic highs have been erased, depositional patterns are characterized by graded slopes which may reduce the potential to trigger mass movement (e.g. Galloway, 1989). Deposition at this stage takes place through slope bypass and is unconfined (Prather et al., 1998; Prather 2000, 2003). Eventually, slope instability may occur through slope oversteepening (Prather 2000, 2003).

The aggregate result of the interaction of all the processes of slope destruction and construction is to evolve from an initial hypothetical graded profile at a deep stratigraphic level to a new hypothetical graded profile at a shallower stratigraphic level achieved through the healing of both inherited and contemporaneous topography (Fig. 2.3).

2.5.2 Type and impact of bathymetry on slope depositional patterns

Sea floor bathymetry is a major, first order control on the deposition and spatial distribution of sediment gravity flows on submarine slopes. The development of seafloor bathymetry invariably creates accommodation space for sediment gravity flows to accumulate. Two types of bathymetric controls are responsible for the development of depositional patterns; structural bathymetry and depositional bathymetry.

2.5.2.1 Structural bathymetry and its impact on depositional patterns

Structural bathymetry is the primary relief of the palaeo-seafloor which pre-exists prior to the deposition of sediment gravity flows. This bathymetry may also form synchronously when the seafloor reacts to loading by sediment gravity flows (e.g. Vendeville and Cobbold, 1987; Cobbold et al., 1989;

Vendeville 2005). Pre-existing structural bathymetry is created by faulted and folded seafloors as the result of basin-scale tectonic deformation, whereas synchronous structural bathymetry is the consequence of localized withdrawal, collapse, and evacuation of a mobile salt or shale substratum (e.g. Prather et al.,

1998; Prather, 2000). The influence of the diapiric- and fault-controlled continental slope on the rapidly prograding shelf edge-to-upper slope area often results in a topography consisting of highs and lows with a variety of subbottom characteristics (Bouma et al., 1990). In addition, contemporaneous growth faulting can also create synchronous structural seafloor bathymetry. Various architectural and morphological occurrences of palaeo-seafloor structural bathymetry have been documented by Beaubouef and Friedmann (2000), Prather et al. (1998), Prather, (2000, 2003) and Steffens et al, (2003), leading to the identification of four types of related accommodation space: ponded accommodation, healed-slope accomodation, incised submarine valley accommodation, and slope accommodation.

Ponded accommodation space is characterized by a bowl-shaped, three-dimensional depression that is formed by tectonic movements associated with salt withdrawal in the distal portion of intra-slope basins, but may also be related to shale diapirism or complex faulting patterns (Booth et al., 2000). If a blockage or a sill at the rim of the depression is low, a density current may be able to move a portion of its sediment load across it; otherwise, the downdrip basin(s) will become starved of sandy sediment, and pelagic and hemipelagic deposits will provide the only fill (Bouma et al., 1990). Healed-slope accommodation is the volume difference between the ‘total’ and the ponded accommodation, the ‘total ’ (Fig. 2.3) being determined by locating the points of maximum positive curvature (the ‘hills’) on the slope in 3D and then draping a 3D convex hull that passes through these points (Steffens et al. 2003). Incised submarine valley accommodation represents the space created by submarine erosion (Prather, 2003). Slope accommodation is a three-dimensional space between the top of ponded space and the graded slope profile (Booth et al.,

2000; Fig. 2.3).

It follows from these definitions that upon filling, ponded accommodation inherently passes onto healed-slope accommodation, which in turn develops into slope accommodation. Incised submarine

34 valley accommodation may occur subsequent to the evolution of ponded, healed-slope or slope accommodation space.

2.5.2.2 Depositional bathymetry and its impact on depositional patterns

Depositional or stratigraphic bathymetry is a secondary relief developed on top of MTDs as a result of the coherent freezing of debris flows, and has been termed MTD-related topography (e.g. Mulder and Alexander, 2001; Armitage et al., 2009; Jackson and Johnson, 2009). The development of this bathymetry may be further enhanced by the emplacement of megaclasts, which may be transported within the parent flow of the MTDs (Jackson, 2011). MTDs commonly develop initial bathymetry that later influences sedimentation in deep-water settings (Shanmugam, 2000; Marr et al., 2001; Posamentier and

Kolla, 2003; Moscardelli et al., 2004) where they control, amongst other things, the routing and architecture of successive sediment gravity flows. They have depositional relief of several metres to hundred of metres, as indicated by baselapping of overlying turbidites on the upper surfaces of field-scale debrites (Pickering and Corregidor, 2005; Amerman et al., 2007; Armitage et al., 2009; Jackson and

Johnson, 2009) as well as in seismic reflection datasets (Deptuck et al., 2003; Moscardelli et al., 2006; and Wood, 2008).

Depositional topography creates secondary debrite-related mini-basins and results in the internal partitioning of overlying high frequency laterally continuous sediments with depositional patterns analogous to the deposition within ponded accommodation (Beaubouef and Friedmann, 2000; Diegel et al., 1995; Prather et al., 1998; Prather, 2000, 2003; Steffens et al., 2003), as the available space is filled quickly by sheet sands at the base of each cycle (Booth et al., 2000). However, due to lack of any significant, superimposed overall slope gradient on the MTD-related topography to support the tilting required to cause a spill, this MTD-related topography cannot support the sediment spill phase that characterises ponded accommodation space in an intra-slope basin. Additionally, erosional processes may be capable of levelling the top of the topography, thereby cancelling any appreciable relief. As a result, the ponded phase filling of MTD-related topography is typically followed by the emplacement of blanketing or widespread turbidites (e.g Bouma et al., 1990).

2.6 Internal architecture and strain indicators within MTDs

During the initiation, translation and cessation of transportation, kinematic indicators are formed and ultimately preserved at the base of, within and along the upper surface of MTDs. These geological structures record information related to the behaviour and nature of component material, and the magnitude and direction of motion at the time of translation and emplacement (Bull et al., 2009). The geometry and internal deformation of MTDs are consequences of the mechanism of failure and the morphology of the slope over which translation occurs (Strachan, 2002a; Lucente and Pini, 2003). These deformational elements are constrained by the rheology of the transported material, which itself is dependent on several factors including lithology and strain magnitude and rate (Cosgrove, 2007).

Modelling of kinematic structures by characterising and analysing the external geometry and internal distribution of deformational structures is useful in order to unravel the strain history and delineation of

MTDs (Strachan, 2002a; Lucente and Pini, 2003; Jackson, 2011). Bull et al., (2009) provided the first attempt at characterizing the various types of kinematic indicators (Fig. 2.4) that can be identified within

MTDs using seismic reflection data.

Fig. 2.4: Schematic representation of a MTD and the likely occurrence and associations of kinematic indicators relative to the various domains: (1) headwall scarp, (2) extensional ridges and blocks, (3) lateral margins, (4) basal shear surface ramps and flats, (5) basal shear surface grooves, (6) basal shear surface striations, (7) remnant blocks, (8) translated blocks, (9) outrunner blocks, (10) folds, (11) longitudinal shears indicatingfirst order flow fabric, (12) longitudinal shears indicating second order flow fabric, (13) pressure ridges, (14) fold and thrust systems. Taken from Bull et al., (2009).

Lewis (1971) suggested that a systematic distribution of strain is exhibited within many MTDs.

This is typically characterised by extensional structures within the upslope region, and contractional structures within the downslope region. However, many MTDs identified in ancient sedimentary successions do not fit this general model, and it does not account for the huge variety and association of structures observed (Martinsen, 1989; Trincardi and Agnani, 1990; Strachan, 2002a, b; Lucente and Pini,

2003; Van der Merwe et al. 2009).

2.6.1 Headwall Domain

The headwall domain (Fig. 2.4) encompasses the upslope, extension-dominated region of the

MTD. Two important kinematic indicators found in this domain are headwall scarps, and extensional ridges and blocks (Bull et al., 2009). The headwall scarp essentially represents an extensional failure surface, and therefore forms in the same way as an extensional fault. They typically define the updip boundary of an MTD, where the basal shear surface ramps up to cut stratigraphically higher and younger

37 strata, and intersect the palaeo-seafloor. Headwalls are often conspicuous features that can be easily identified in both cross-section (e.g. seismic sections) and in planform (e.g. seismic horizon maps). In plan view, headwall scarps appear as continuous scarps that may comprise one or more salients and are typified by an arcuate geometry (Prior et al., 1984; Bøe et al., 2000; Imbo et al., 2003). The headwall scarp typically propagates along-strike, perpendicular to the direction of the minimum compressive stress, which is generally oriented parallel to the slope due to the effect of gravity on the sediment mass. Thus, the orientation of the headwall scarp can yield important kinematic information, as it may reveal the initial direction of bulk material movement. The headwall may be completely evacuated, with the basal shear plane being fully exposed (Haflidason et al., 2004). Alternatively, material may have undergone limited translation, and the headwall may therefore represent only a minor topographic feature (Solheim et al., 2005).

2.6.2 Translational Domain

This domain comprises the main translated body of the MTD between the updip and downslope limits of the deposit. The movement of the failed material downslope across the basal shear surface ( sensu

Bull et al. 2009) can lead to intense deformation and the formation of a variety of features that can provide important information regarding the kinematic development of the flow and associated deposit

(Martinsen, 1994; Strachan, 2002a). Kinematic indicators within the translational domain include structures along the lateral margins of the MTD, the basal shear surface itself, which may be characterised by ramps and flats, grooves and striations and folds, internal body features such as folds and faults, and structures developed along the top surface of the MTD. The lateral margins are the dip-parallel, lateral boundaries of an MTD and they form parallel to the gross flow direction; these features are a primary constraint on the overall MTD transport direction. Depending on data coverage, it may be possible to trace the lateral margins of an MTD from the headwall to the toe (Gawthorpe and Clemmey, 1985).

The basal shear surface represents the primary surface along which the MTD is translated and it may yield important information about the transport direction(s) and the magnitude of sediment

38 entrainment at the base of the flow (Varnes, 1978; Martinsen, 1994; Gee et al., 2005; Bull et al., 2009).

The basal shear surface may be laterally continuous over great distances and is typically concordant to bedding (Varnes, 1978). However, the basal shear surface may cut up or down to a new stratigraphic level across a feature known as a ‘ramp’. A ramp is defined as a segment of the basal shear surface that cuts discordantly across bedding, whereas a ‘flat’ is a section that is bedding-parallel (Trincardi and Argnani,

1990). On seismic data, ramps are easily identifiable in a similar way to faults or the headwall scarp. In planform, ramps can be seen as scarps or linear features that mark sharp topographical variations of the basal shear surface. Trincardi and Argnani (1990) and Gawthorpe and Clemmey (1985) report that most ramps trend perpendicular to the main flow or movement direction.

Grooves are linear or sinuous features that occur on the basal shear surface. They are typically ‘v’ shaped in cross-section, continuous over many tens of kilometres, are generally orientated downslope and may diverge seawards (Posamentier and Kolla, 2003; Gee et al., 2008). These features are interpreted to be the result of the erosional action of coherent blocks of translated material at the base of a debris-like flow (Posamentier and Kolla, 2003; Gee et al., 2008). In seismic data, grooves are most obvious from attribute maps (e.g. amplitude, variance, and dip) of the basal shear surface of an MTD. Grooves may be difficult to identify in seismic section without reference to a map, as they are typically narrow and shallow (Bull et al., 2009). Striations have a similar origin to grooves, but they are wider than the latter and have a flatter base; in addition, striations typically do not diverge downslope (Gee et al., 2005). In general, the combination of ramps and flats, grooves and striations delimit deformation features and zones of mechanical shearing (Fortuin, 1978; Varnes, 1978; Gawthorpe and Clemmey, 1985; Strachan, 2002a;

Bull et al., 2009; Van der Merwe et al., 2009; Alves and Lourenço, 2010), and detachment of weak inter- layers between the destabilized and churned body of rock and its more stable structurally intact surrounding.

A variety of fold styles has been observed within MTDs at outcrop. These structures are traditionally used to determine the palaeoslope direction (Farrell and Eaton, 1988; Webb and Cooper,

1988; Eva and Maltman, 1994; Strachan, 2002a, b, 2006). Farrell and Eaton (1987, 1988) proposed that

MTDs which have undergone relatively little translation will contain upright, buckle-type folds.

Progressively more deformed and more translated MTDs will contain tight, isoclinal and recumbent folds.

Folds in the translated body of a MTD can therefore be useful in identifying the transport and palaeoslope direction, and possibly in determining the degree of translation.

2.6.3 Toe Domain

The toe domain (Fig. 2.4) represents the downslope region of the MTD (Bull et al., 2009). The toe domain commonly has an overall convex-downslope morphology (Prior et al., 1984). The main kinematic indicators contained within the toe domain include pressure ridges, thrusts and folds. Pressure ridges are positive, parallel to subparallel, linear to arcuate ridges that are orientated broadly perpendicular to the overall flow direction (Prior et al., 1984; Masson et al., 1993; Martinsen, 1994; Bøe et al., 2000; Laberg and Vorren, 2000; Posamentier and Kolla, 2003; Frey Martinez et al., 2006;

Moscardelli et al., 2006). Although they are typically developed in the toe domain, pressure ridges can occur elsewhere due to localised topographical variations in the basal shear surface, or obstacles to flow

(Masson et al., 1993). In seismic data, pressure ridges are best viewed on maps of the MTD upper surface.

Thrust and fold systems are commonly described from the toe domains of MTDs (e.g. Lewis, 1971;

Trincardi and Argnani, 1990; Martinsen, 1989; Frey-Martinez et al., 2006), where they may affect sections up to 500 m thick (Gafeira et al., 2007); this places them on a similar scale to some tectonically- driven thrusts. Thrusts typically affect the entire thickness of the MTD body in the vicinity of the ramp above the basal shear surface, with the most likely transport direction being perpendicular to the strike of the thrust faults (Ramsey and Lisle, 2000).

2.7 Seismic facies analysis

Seismic facies analysis involves the delineation and interpretation of reflection geometry, continuity, amplitude, frequency, and interval velocity, as well as the external form and three-dimensional associations of groups of reflections where each of the seismic reflection parameters contains information

40 of stratigraphic significance (Sangree and Widmier, 1978). A seismic facies unit can be defined as a sedimentary unit that differs from adjacent units in terms of its seismic characteristics (Roksandic 1978).

Key descriptive seismic facies parameters include reflection amplitude, reflection continuity, reflection configuration, geometry of seismic facies unit, and its relationship with adjacent seismic facies (Pacht et al., 1993). An interpretation of the depositional processes and environments, and sediment transport direction, may ultimately be inferred based on a broader understanding of the regional setting of the specific study area (e.g. Sangree and Widimer, 1977; Feeley et al., 1985; McHargue and Webb, 1986;

Vail and Wornardt, 1991; Prather et al., 1998; Beaubouef et al., 2000). The first unified approach to seismic stratigraphic analysis was provided by Payton (1977) in which a threefold hierarchical division of seismic stratigraphic analysis was presented, consisting of seismic sequence analysis, seismic facies analysis, and seismic reflection character analysis. Most published studies (e.g Weimer, 1990; Beaubouef and Friedmann, 2000; Olafiranye et al., 2013) have concentrated on sequence and facies analysis because they provide a framework for regional interpretation of broad stratigraphic packages, palaeogeography, and geological history.

Progradational shelf-margin and slope environments typically contain thick marine sediments and have water depths sufficient for the development of complex arrangements of sigmoidal and oblique clinoform reflection patterns, and the basin slope and floor environments include a variety of additional seismic facies (Sangree and Widmier, 1978). Based on seismic stratigraphy and facies, Beaubouef and

Friedmann (2000) reported that the fill of slope basins exhibits vertical cyclicity reflecting alternating deposition of mass transport complexes (MTCs), distributary channel-lobe complexes, leveed-channel complexes, and hemipelagic drapes. Weimer (1990) and Beaubouef and Friedmann (2000) identified

MTCs as hummocky and mounded reflections with low amplitude, chaotic seismic facies units that can be interpreted as mud-rich complexes of slumps, slides, and debris flow deposits. The distributary channel- lobe complexes are sand-rich depositional units characterized by moderate to high reflection amplitude and continuity that exhibit onlapping, compensatory and locally erosional internal reflection geometries and fan shaped plan-view geometries that record deposition from highly concentrated sediment gravity

41 flows delivered across the basin floor through networks of relatively small distributary channels or as broad sheet-like flows. The leveed-channel complexes form a distinct seismic facies characterized by low amplitude, highly continuous reflection character and “gullwing” cross-sectional profiles that are considered to form from overbank deposition of low concentration turbidites and to contain low to moderate sand percentages. Hemipelagic drapes are thin, highly continuous seismic units that represent hemipelagic mudstones deposited during periods of abandonment and sediment starvation.

Seismic forward modelling, which is a method of quantitative analysis of seismic amplitudes through integration of petrophysical properties (velocity and density) with interpreted lithology distribution and reservoir architecture, can be carried out to ground-truth subsurface seismic facies (e.g.

Bakke et al, 2008; Armitage et al, 2010). The increase in exploration and development of depositionally complicated and heterogeneous petroleum reservoirs (e.g. deep-water turbiditic deposits) has led to a number of attempts to simulate their seismic-reflection response (e.g., Batzle and Gardner, 2000;

Campion et al., 2000; Coleman et al., 2000; Tomasso, et al., 2006; Schwab et al., 2007; Bakke et al.,

2008; Armitage et al, 2010).

Bakke et al. (2008) used the available rock properties (velocity and density logs) from a deepwater well from offshore Angola, in order to compare the Ainsa II turbidite system in Spain to turbidite systems buried in the subsurface. The properties in the modelled section are then assumed to have similar pressure and temperature conditions to the studied turbidite system in offshore Angola. The seismic modelling was carried out on depositional elements by integrating well data, 3D seismic data and analogue outcrop data. The 3D seismic interpretation and seismic modelling was performed using zero- phase wavelets with reversed polarity, where peak (black) and trough (red) represent negative and positive acoustic impedance, respectively (Fig. 2.5). In order to create synthetic seismic sections a 2D convolution with a Ricker wavelet was performed (Bakke et al., 2008), which simulates a 2D migrated stacked section with a 2D resolution parameter applied to the reflection series (e.g. Toxopeus et al.,

2003). The resulting synthetic seismogram is comparable to real seismic cross-sections.

Fig. 2.5: (A) A density property image which allows comparison of input data with seismic models displayed in B and D. (B) A synthetic seismogram with 65Hz frequency content. The different units are indicated by dotted green lines. Note that units 1 and 2 can be identiﬁed as two separate geobodies at this resolution. All seismic sections and seismic models are shown in time sections, but all scales are in metres. A black reﬂector represents a decrease in acoustic impedance. (C) The 45Hz synthetic seismogram, which possesses a similar resolution as the seismic data from offshore Angola. (D) The 32Hz synthetic seismogram of Ainsa II outcrop. This frequency content is common in seismic data of deep-water deposits in the Gulf of Mexico and in the North Sea. (From Bakke et al., 2008).

The Bakke et al. (2008) model (Fig. 2.5) provides a link between the architectural geometries observed in outcrop data and the subsurface 3D seismic data from offshore Angola (Fig. 2.6). Lateral changes from sandy channel axis to more shale rich channel margins are to some degree reflected in the model by a lateral degradation of amplitude strength (Bakke et al., 2008).

Fig. 2.6: (A) A seismic section of the channel complex in Angola, oblique to depositional strike. Note that the section intersects a narrow channel (Facies F2) at three locations, whereas the lower and broader part of the system is only intersected once. The clastic system contains three seismic facies. (B) A suggested interpretation of the Angolan system. Facies F1 is interpreted to represent in-channel deposits, with several depositional units (e.g. thick dotted lines). Facies F2 possibly contains conglomeratic channel ﬁll. The location of facies F3, which is interpreted to represent ﬁne-grained, channel. (From Bakke et al., 2008)

Armitage and Stright (2010) provided the first seismic forward modelling of mass transport deposits, using outcrop data from the Cretaceous Magallanes foreland basin located in the Ultima

Esperanza District of southern Chile. In their model, each lithofacies was assigned a synthetic velocity and density value from which an acoustic impedance model was generated. Random Gaussian noise of approximately 10% was added to the mass-transport deposits in the acoustic impedance model, so that the synthetic seismic-reflection profiles better replicate actual seismic data. A minimum bed thickness of 40 cm was modeled; however, only bundled groups of beds can be imaged in seismic-reflection profiles. The data-rich lithological model ensures that geological detail (e.g. lithofacies distribution) is not a limiting

44 factor in the production of seismic-reflection profiles. Synthetic seismic-reflection profiles were constructed from the impedance model using simple 1D, zero-offset acoustic modeling. Varying the peak frequency of a zero-phase Ricker wavelet by 120, 60, 30, and 15 Hz generated four different seismic- reflection profiles (Fig. 2.7). The synthetic seismic sections were constructed using a 2 ms sample rate and a shot/receiver spacing of 30 m, similar to conventional petroleum industry seismic-reflection data.

Fig. 2.7: Synthetic seismic-reflection profiles. A) Profile created with 120 Hz frequency, zero-phase Ricker wavelet. The following features are highlighted: i) reflection phase change due to lateral lithofacies changes, and ii) reflection truncation due to bed pinchout, but which might be interpreted as erosion. B) 60 Hz profile. The following features are highlighted: i) reflection phase changes due to lithofacies changes, and ii) high amplitude reflections due to tuning. C) 30 Hz profile. The following features are highlighted: i) reflection phase change due lithofacies changes, and ii) lateral increase in reflection amplitude due to a lithofacies change. D) 15 Hz profile. Unit boundaries and internal geometries in the sandy units are ambiguous. Vertical black lines on all profiles indicate the location of the synthetic well logs. (From Armitage and Stright, 2010).

The Armitage and Stright (2010) models showed that at the low frequencies of 15 and 30 Hz, both units B and C are poorly resolved (Fig. 2.7); their boundaries with adjacent MTDs are ambiguous. Although the top of Unit C exhibits a high-amplitude reflection the internal reflection-amplitudes are low, which is more characteristic of MTDs. At 30 and 60 Hz, the base of Unit B in the northern part of the outcrop has a channelized geometry, which, inferred by the high-amplitude response, is sand-rich. The interpretations made from the synthetic models when applied to a subsurface example from the Gulf of Mexico (Fig. 2.8;

Posamentier and Walker, 2006) shows alternating chaotic mass-transport deposits and sandstone deposits, as inferred from their high-amplitude reflection response. Subtle topography is resolved on the top surfaces of the MTDs (Fig. 2.8ii), similar in scale to the MTD topography of Unit 3 in the Armitage and

Stright (2010) lithological model (Fig. 2.7). This scale of topography would not be resolved at the lower frequencies (e.g., 30 Hz) associated with [subsurface] reservoir depths.

Fig. 2.8: Shallow subsurface seismic-reflection profile from the western Gulf of Mexico, acquired in approximately 2700 m water depth. The profile shows (inferred) alternating mass-transport and sandstone deposits. Maximum peak frequency is approximately 70 Hz (comparable to Fig. 2.7B). Rafted clasts are observed in the MTDs. i) Lateral changes in reflection amplitude interpreted as lithofacies changes. ii) MTD topography of a similar scale to that observed in the higher-resolution synthetic seismic-reflection profiles (Fig. 2.7A, B). Figure modified from Posamentier and Walker (2006).

2.8 Impact of MTDs and sediment gravity flows on fluid movement and pockmarks

The presence of fluid escape features like seafloor pockmarks (Hovland and Judd, 1988), and subsurface fluid-escape chimneys (Gay et al., 2006; Andressen et al., 2011) indicate localized zones of

47 high fluid pressures in subsurface sediments (Judd and Hovland, 2007). Overpressure (i.e., pore pressure in excess of hydrostatic pressure) may develop from rapid sediment loading in low-permeability environments (Gibson, 1958; Rubey and Hubbert, 1959). These overpressures can significantly influence fluid migration through porous sediments in stratigraphic layers (Harrington and Horseman 1999; Gay et al., 2007; Andressen et al., 2011) in areas of differential sediment loading (Kjeldstad et al., 2003), where pressure gradients may drive lateral and downdip flow and contribute to fluid venting and slope instability

(Dugan and Flemings, 2000, 2002). For example, several areas on the continental slope of the formerly glaciated Norwegian margin were subject to very high sediment accumulation rates during Pleistocene glaciations (Berg et al., 2005; Rise et al., 2005). Rapid loading occurred repeatedly in these proglacial depocenters due to the waxing and waning of channelized and fast flowing ice streams (Sejrup et al.,

2004, 2005). As a result, one of the largest known submarine slope failures is most likely associated with overpressure development due to high sedimentation rates during the emplacement of the Storegga Slide

(Leynaud et al., 2007). Numerous fluid escape features occur in the northern flank of the Storegga Slide,

Nyegga region, due to the generation of associated subsurface overpressure (Hustoft et al., 2009).

2.8.1 Generation and transfer of overpressure

Pore-water in sedimentary basins flows from areas of higher hydrodynamic potential to areas of lower potential (Bjorlykke, 1993). This inherent flow relies on the interplay of lithostatic and hydrostatic pressures (Fig. 2.9). The lithostatic pressure of a sedimentary regime, sometimes known as overburden, is the theoretical maximum possible in situ pore pressure in the absence of any tensile strength in the rock

(Mann and Mackenzie, 1990). Under 100 % lithostatic pressures, the pore fluids support all the weight of the overlying sediments including brine and grains (Mann and Mackenzie, 1990). Hydrostatic pressure is the pressure exerted by a fluid at equilibrium at a given point within the fluid, due to the force of gravity

(e.g. Mann and Mackenzie, 1990; Bjørlykke, 1993; Osborne and Swarbrick, 1997). Hydrostatic pressure increases in proportion to depth measured from the surface because of the increasing weight of fluid exerting downward force. The difference between pore pressure and hydrostatic pressure is called

48 overpressure. A rock is said to be overpressured when its pore pressure is significantly greater than hydrostatic pressure (Mann and Mackenzie, 1990).

Fig. 2.9: Subsurface pressure profiles of lithostatic, hydrostatic, overpressure/underpressure, and fracture pressure. Modified after Gressier et al., (2010).

Overpressures occur in rocks whose pore fluids support some of the weight of the overlying rock grains in addition to the weight of the overlying and displaced fluids (Mann and Mackenzie, 1990;

Bjorlykke, 1993). There are not enough grain-grain contacts in the framework of overpressured rocks to support all the weight of the overlying sediments because the rocks have not been sufficiently compacted.

Fracture pressure (Fig. 2.9) is the fluid pressure at which fractures propagate perpendicular to the direction of minimum compressive stress (Mann and Mackenzie, 1990). The pressure required to attain fracture pressure is between the hydrostatic and lithostatic pressure, typically 70-90% of the lithostatic pressure (Du Rouchet, 1981). When fracturing occurs, pore fluids and pressure leak into the fractures.

When a small volume of sediment receives a significant pressure charge from a much larger volume, then pore pressures may reach lithostatic pressure and trigger fluid flushing. For example, mud volcanoes occur when the pore pressure transiently exceeds lithostatic pressure (Mann and Mackenzie, 1990).

Overpressures can be generated through mineralogical reactions (e.g. smectite to illite) (Osborne and Swarbrick, 1997), hydrocarbon maturation (Ungerer et al., 1983), thermal expansion of water

(Barker, 1972), and the failure of sediments to compact on burial to a porosity where the framework of rock grains can support all the weight of the overlying sediments. This last mechanism is by far the predominant cause of most large overpressures observed in sediments (Mann and Mackenzie, 1990).

Overpressures can significantly influence fluid migration in areas of differential sediment loading

(Flemings et al., 2002; Kjeldstad et al., 2003), where pressure gradients may drive lateral and downdip flow and contribute to fluid venting and slope instability (Dugan and Flemings, 2000; Dugan and

Flemings, 2002). Overpressuring can be transferred through the rocks by fluid movement in three dimensions (Osborne and Swarbrick, 1997). Where sandstones are isolated within shale sequences, flow is predominantly vertical (Mann and Mackenzie, 1990). However, if sandstones are areally extensive and interconnected, fluid and pressure may be transferred laterally (Mann and Mackenzie, 1990). Pressure may also be gradually redistributed below an impermeable barrier so that fluid pressure increases

50 immediately beneath the seal, but decreases at the base of the pressure compartment (Swarbrick and

Osborne, 1996).

Chapter Three

The role of tectonics and mass-transport complex emplacement on upper slope stratigraphic

evolution: a 3D seismic case study from offshore Angola

3.1 Abstract

Three dimensional seismic-reflection data of the mid-Pliocene-to-Holocene upper slope succession, offshore Angola, provide an opportunity to constrain the stratigraphic context, distribution, external morphology and internal strain within mass-transport deposits (MTDs). These data also allow an assessment of the impact that erosion and relief associated with MTDs have on upper slope stratigraphy and depositional patterns, and the role that MTDs play in achieving ‘grade’ on submarine slopes. The study area is dissected by a series of NW-SE-striking, thin-skinned, salt-detached normal faults, which bound a slope-perpendicular, intra-slope horst that divides the study area into two depocentres.

Three main seismic packages and their six constituent units have been mapped across the study area and reveal that, during the initial stages of deposition, a series of MTDs were emplaced, the thickness and distribution of which are controlled by the intra-slope horst. Substantial volumes of substrate were removed and entrained into the parent flow, and significant and irregular relief (150 m) was developed along MTDs upper surface. This MTD-rich package is interpreted to document a time when the slope was above grade, degradational processes dominated and sediment was trapped on the upper slope due to tectonic accommodation. Subsequent deposition was from either turbidity currents or and suspension fallout, at a time when the slope had begun to achieve ‘grade’ and depositional processes dominated. The associated depositional units display only minimal thickness variations with respect to the intra-slope horst, which had been ‘healed’ by this time; however, the unit displays pronounced and abrupt changes in thickness due to infilling of relief at the top of the preceding MTDs. The uppermost strata document a time when the slope was at grade and constructional process (i.e. aggradation and progradation) dominated. Deposition at this time was characterised by progradation of a mudstone-dominated, gullied slope system.

This study highlights the role that tectonically- and mass transport-driven changes in bathymetry can have on upper slope accommodation and sediment dispersal. From a hydrocarbon exploration perspective this is critical, because tectonic and depositional accommodation provide a mechanism for capturing and trapping clastic sediments in an upper slope setting, which is otherwise typically associated with coarse- grained sediment bypass.

3.2 Introduction

Accommodation on submarine slopes is the gap between the sediment surface (the background slope surface) and the equilibrium profile (the slope profile of no net erosion or deposition) (Pirmez et al. 2000;

Kneller 2003). The equilibrium profile can be perturbed by changes in submarine slope bathymetry through time and space (Fig. 3.1). For example, faulting and folding (e.g. Bryant, 1991; Hodgson and

Haughton, 2004; Kane et al. 2010), and the flow of salt and mud (e.g. Hodgson, 1992; Prather et al.,

1998; Prather 2000; Prather 2003; Steffens et al, 2003) can lead to the formation of intra-slope bathymetric relief and disturbance of the grade of the slope (Fig. 3.1).

Fig. 3.1 : (a) Schematic cross-section illustrating the styles of accommodation developed on continental margins. Note the development of ponded accommodation on the slope due to normal faulting, salt flow and thrusting, which are all types of ‘tectonic accommodation.’ Modified from Prather (2003) and Steffens et al. (2003). The inset diagram illustrates the styles of depositional accommodation that can develop due to the emplacement of MTDs. In this case, relief and accommodation can form in response to the emplacement of ‘megaclasts’ (left-hand side of diagram) or deposit ‘buckling’ (right-hand side of diagram), which may be related to shortening of the deposit during abrupt flow freezing. Modified from Armitage et al. (2009) and Jackson and Johnson (2009).

A graded slope is typically characterised by a smooth, possibly sigmoidal, shelf-to-basin profile, which contrasts with the more irregular profiles that commonly characterise out-of-grade or above-grade slopes

(e.g. O’Grady, 2000, 2001; Smith 2004; Prather et al., 2003; Steffans et al., 2003; Pyles et al., 2011).

MTDs may typically form at the base of an out-of-grade slope, as the system evolves towards equilibrium; slope progradation by accretion may then occur. When a slope is significantly above ‘grade’ a series of predominantly erosional or ‘degradational’ processes may act to eventually return it to a condition of ‘grade’; conversely, when a slope is below grade, depositional (i.e. progradational or aggradational) processes occur and return the slope to a graded profile (e.g. Ross et al., 1994, 1995;

Prather 2003; Steffens et al., 2003; Pyles et al. 2011). Changes in slope grade and equilibrium profile can therefore markedly influence the location and number of sediment fairways, the storage and dispersal patterns of sediment on the slope, and the behaviour of submarine sediment gravity flows (Haughton,

1994; Prather et al., 1998; Kneller and McCaffrey, 1999; Prather, 2000, 2003; Bouma 2004; Hodgson and

Haughton, 2004; Smith et al, 2004, Adeogba et al., 2005; Kane et al. 2010; Hodgson et al., 2011;

Moscardelli et al., 2012). From the point-of-view of hydrocarbon exploration and production, changes in slope accommodation directly influences the potential for sand-prone reservoir facies to be deposited in upper slope settings, which are typically associated with coarse-grained sediment bypass.

The role of primary depositional relief on controlling accommodation development on submarine slopes is less well understood. Primary depositional relief on submarine slopes is typically related to the deposition of mounded sedimentary bodies or the transport of large clasts in mass-transport deposits

(MTDs) (see inset in Fig. 3.1). Data from high-resolution, near-seabed, acoustic reflection profiles (e.g.

Hampton et al. 1996; Piper et al. 1997), industry-standard seismic-reflection datasets (e.g. (e.g. Collot et al., 2001; Deptuck et al., 2003; Faulkenberry 2004; Moscardelli et al., 2006; Frey-Martinez et al., 2006;

Moscardelli and Wood, 2008; Ashabranner et al., 2010; Mosher et al., 2010; Alves and Cartwright, 2009;

Gamberi et al., 2011; Amerman et al., 2011; Cossey et al., 2011; Dykstra et al., 2011), and outcrop studies

(e.g. Pickering and Corregidor, 2000; Pickering and Corregidor, 2005; Amerman et al., 2007; Armitage et al., 2009; Jackson and Johnson, 2009; Van der Merwe et al., 2009, Butler and McCaffrey, 2010) have

54 acknowledged the role that MTD-related depositional and erosional relief may have on local intra-slope and basin floor accommodation. However, the styles of relief associated with MTD-related accommodation, their three-dimensional form and areal extent, and their impact on subsequent sediment dispersal patterns remain poorly constrained. Furthermore, it is not clear over what timescales MTD- related accommodation may impact the stratigraphic development of deep-water successions. These uncertainties partly reflect the limited extent and two-dimensional nature of the majority of outcrops, and the lack of subsurface datasets that have used high-quality, 3D seismic reflection data.

This study focuses on the Mid-Pliocene-to-Holocene upper slope succession in Block 4, offshore

Angola (Figs 3.2 and 3.3). We aim to: (i) characterise and interpret the seismic facies present in this deep- water succession; (ii) describe and interpret the controls on the external morphology and distribution of

MTDs at any one stratigraphic level; (iii) establish the overall stratigraphic context of multiple MTDs and assess the role of the associated sediment gravity-flows in achieving grade on submarine slopes; (iv) document the geometry, scale and distribution of structures at the base of, within and at the top of the

MTDs, and relate these to the initiation, translation and emplacement of these deposits; and (v) assess the

Fig. 3.2: (a) Map of SW African margin showing the location of the Lower Congo Basin in relation to main bounding hinge zones. The continent-scale geographic setting of the margin is shown in the inset map. The location of the study area (Fig. 2b) is shown in the red box. (b) Map of tectonic framework of the Lower Congo Basin. Modified from Anderson et al. (2000) and Dutton and Trudgill (2009). The red box indicates the location of Block 4 and the study area (see also Fig. 4a). The study area lies at the southern end of a Miocene graben (see also Dutton and Trudgill, 2009). impact of tectonic structures, and depositional and erosional relief associated with the emplacement of

MTDs, on the stratigraphic architecture of the upper slope and subsequent sediment dispersal.

Fig. 3.3: Regional geoseismic section across the Angolan margin. Modified from Lentini et al. (2010). Major structures and depositional sequences are highlighted. COB=continent–ocean boundary. Note the occurrence of a thick, highly-mobile salt layer (Loeme Formation) across the entire width of the margin. The study area lies within the updip extensional domain (see red box). The location of transect is shown in Fig. 3.2b.

3.3 Geological setting

During the Jurassic and Early Cretaceous, a series of rift basins formed along the West African margin in response to basement-involved, upper crustal extension. The Lower Congo Basin, which lies between 7 °30’S and 7 °45’S, is one of these basins (Fig. 3.2a). After the main period of rifting, the West

African passive margin underwent thermally-induced subsidence during the Late Cretaceous to Present.

However, the presence of a thick, lower Cretaceous evaporite layer (Loeme Formation) and westward tilting of the margin in response to subsidence of the Atlantic Ocean and uplift of the African craton, resulted in thin-skinned, gravity-driven extension, which initiated in the Albian and continues to the present day (e.g. Duval et al., 1992; Lundin, 1992; Richert, et al., 1992; Spathopoulos, 1996; Eichenseer et al., 1999). Gravity-driven extension resulted in the breakup of platform carbonates (Pinda Formation) and the formation of a series of normal fault-bound raft blocks (Figs. 3.2b, 3.3 and 3.4) (Dutton and

Trudgill, 2009). During the Miocene, a period of continental uplift caused renewed thin-skinned extension and an influx of coarse clastic material (Hempton et al., 1990; Duval et al., 1992). In the western part of

Block 4, a major thin-skinned extensional system developed, which was characterised by a series of normal fault-bounded graben and half-graben (Figs. 3.2b, 3.3 and 3.4) (Dutton and Trudgill, 2009). In the upper- to mid-slope area a fault-bound graben, which is located above a salt weld between two Pinda

Formation raft blocks, passes downslope into fault-bound horst, which directly overlies a large raft block

(Fig. 3.4). The horst subdivides the upper continental slope into two main depocentres, which are herein termed the eastern and western depocentres (Fig. 3.4). These depocentres are located in the SW corner of

Block 4, about 16 km from the present-day shelf edge and in ca. 500 m water depth (Fig. 3.3). In this paper we focus on the stratigraphic development of the middle Pliocene-to-Holocene interval of clastic- dominated Malembo Formation, which is represented in the study area by ca. 620 m high clinoforms that dip towards the SE (Figs. 3.4b, 3.5 and 3.6). This unit was deposited during thin-skinned extension and comprises a succession of sandstone-dominated, submarine channel-fill deposits, mudstone-dominated

MTDs and slope mudstones.

Fig. 3.4(a): Time-structure map of the mid Pliocene illustrating the structure of the study area (Horizon A; see Figs. 4b and 5). The key normal faults, which initially developed in response to Miocene-age, thin-skinned, salt-detached extension, are marked (see Fig. 3.3 for basin-scale context). In the SE corner of the study area two depocentres (i.e. eastern and western depocentres) are developed due the presence of a normal fault-bound horst (intra-slope horst). The locations of seismic sections shown in Figs. 3.4b, 3.6 and 3.8 are shown. See text for full discussion.

Fig. 3.4(b): Broadly strike-orientated (i.e. NW-SE), uninterpreted (above) and geoseismic (below) section illustrating the structural style of the study area, and the key seismic horizons and units (SUs; see also Fig. 3.5). g=gullies. The location of the seismic section is shown in Fig. 3.4a. Note the intra-slope horst that divides the study area into an eastern and western depocentre.

Fig. 3.5: Stratigraphy, lithologies, depositional elements, and mapped seismic horizons and seismic units within the stratigraphic interval of interest. SU=seismic unit; LSP=lower seismic package; MSP=middle seismic package; USP=upper seismic package. Descriptions and interpretation of the key seismic facies types are also provided. The stratigraphic position of the base mid- Pliocene Unconformity is based on data in Dutton and Trudgill (2009).

Fig. 3.6: Broadly strike-orientated (i.e. NW-SE), uninterpreted (above) and geoseismic (below) section through the relatively proximal part of the eastern depocentre; this illustrates the structure of the study area, and the mapped seismic horizons and units (see also Fig. 3.5). Note the prominent incision at the base of SU2 and the pronounced relief along the upper surface of this unit, which is onlapped by and controls thickness variations in SU3. Note also the upwards transition from chaotic to more continuous seismic facies; this documents an overall upward transition from MTD-dominated stratigraphy to the development of gullied, mudstone-dominated slope system. See text for full discussion. The location of the seismic section is shown in Fig. 3.4a. Note that SU1 is not present on this line, and horizons A and B form a composite surface. See text for full discussion.

3.4 Dataset and methodology

This study uses a volume of 3D seismic-reflection data that covers 760 km 2 of the upper part of the continental slope (Figs. 3.2-3.4). The water depth within the survey area is between 225-700 m and the data are generally of very good-to-excellent quality. The seismic survey has a vertical sample interval of 4 ms TWT (milliseconds two-way time; TWT) and an inline and crossline spacing of 12.5 m. The seismic volume is time-migrated but not depth-migrated, thus true depth and stratal thicknesses are estimated using a clastic acoustic velocity of 2000 m/s (e.g. 100 ms TWT equals 100 m). Based on a dominant frequency of 45 Hz in the relatively shallow interval of interest, and given the velocity in the interval of interest, the vertical resolution of the data is estimated to be ca . 11 m (Fig. 3.7). The polarity of

61 the seismic data is SEG normal; e.g. a downward increase in acoustic impedance is represented by a black reflection and a downward decrease in acoustic impedance is represented by a red reflection.

Fig. 3.7: Frequency range extracted from (a) the entire 3D seismic cube (i.e. 0-4600 ms TWT) and (b) within the interval of interest (i.e. 0-1500 ms TWT). Note that the peak frequency in the interval of interest (45 Hz) is higher than the entire 3D seismic cube (17 Hz). See text for full discussion.

This study is principally based on the seismic facies analysis of seismic reflection data (e.g.

Mitchum et al., 1977; Roksandi ć, 1978). A seismic facies unit can be defined as a sedimentary unit that differs from adjacent units in terms of its seismic characteristics. Key descriptive seismic facies parameters include: reflection amplitude, reflection continuity, reflection configuration, geometry of seismic facies unit, and their relationship with other seismic facies. An interpretation of the depositional processes and environments, and sediment transport direction, may ultimately be inferred based on a

62 broader understanding of the regional setting of the specific study area (e.g. Sangree and Widimer, 1977;

Feeley et al., 1985; McHargue & Webb, 1986; Vail & Wornardt, 1991; Prather et al., 1998; Beaubouef et al., 2000). In this study we use the seismic facies scheme of Prather et al. (1998) to describe and interpret the seismic facies in the middle Pliocene-to-Holocene succession in the Lower Congo Basin, offshore

Angola. This scheme was based on the analysis of 3D seismic reflection data from the Gulf of Mexico, and borehole cuttings and electrical log data were used to directly calibrate the lithology and depositional processes responsible for the various seismic facies. Furthermore, the stratigraphic succession encountered in the Gulf of Mexico comprises predominantly sandstone and mudstone deposits that are buried to an approximately equivalent burial depth to those encountered offshore Angola (i.e. <2 km). As a result, the acoustic and seismic response between the two locations is thought to be similar, although we recognise that different sediment accumulation and compaction rates, coupled with different styles and intensities of diagenesis, is likely to result in some differences in the physical rock properties and seismic responses.

Fig. 3.8: Seismic (above) and geoseismic (below) sections through the relatively distal part of the eastern depocentre illustrating the external form of, and the internal seismic facies and structures in SU1–5. Note the rugose top surface of the mass-transport deposits (MTDs) within SU2 and the strata-bound normal faults in SU2c. Note that the relief along the upper surface of SU2 is onlapped by and controls thickness variations in SU3. g=gullies. The location of the seismic section is shown in Fig. 3.4a.

Based on the recognition of surfaces of seismic-stratigraphic significance, such as onlap, truncation and downlap surfaces, and vertical variations in seismic facies, the studied interval was subdivided into six seismic units that are bound by seven mappable seismic reflections (SU1-6; Fig. 3.4b,

3.5, 3.6, 3.8 and 3.9). These six seismic units are grouped into three main seismic packages; the lower, middle and upper seismic packages (LSP, MSP, and USP; Fig. 3.5). The following section describes the seismic expression, geometry and distribution of these units, in addition to the characteristics of their bounding surfaces. Based on these characteristics, and by comparison to the seismic facies framework of

Prather et al. (2008), the depositional processes and overall depositional environment represented by each unit are interpreted.

Fig. 3.9: Morphology of Horizon A (where it forms a discrete surface outside of the eastern depocentre and where it forms a composite surface with Horizon A in the eastern depocentre; see text for full discussion). (a) Instantaneous root-mean square (RMS) amplitude map illustrating the development of striations (labelled ‘s’), and a frontal ramp (labelled ‘fr’) and lateral ramps (labelled ‘lr’) along the basal shear surfaces to MTDs in SU1 and 2. (b) Time-structure map illustrating the morphological expression of the map ramp in the SE corner of the eastern depocentre. (c) Instantaneous dip map illustrating the morphological expression of the map ramp in the SE corner of the eastern depocentre. (d) Broadly strike-orientated (i.e. NW-SE) seismic section through the relatively proximal part of the eastern depocentre. The key seismic horizons and units are labelled. Note the prominent incision at the base of SU2, and the pronounced relief along the upper surface of this unit, which is onlapped by and controls thickness variations in SU3. g=gullies; lr=lateral ramp. For location of the seismic section see Figs. 3.9a-c.

3.5 3D seismic interpretation

3.5.1 The Lower Seismic Package (LSP)

The Lower Seismic Package (LSP) is composed of Seismic Unit 1 and 2 (SU1 and 2), and is bound below and above by Horizons A and C respectively.

3.5.1.1 Horizon A

Across the majority of the study area, the base of the study interval and SU1 is defined by

Horizon A (Figs. 3.5, 3.6, 3.8 and 3.9d). The surface defines a marked change in seismic facies, from predominantly continuous, parallel, low-amplitude reflections below, to continuous to discontinuous, low-to-high amplitude reflections above, and is interpreted to define the base of the mid-Pliocene

65 succession (Figs 3.5, 3.6, 3.8 and 3.9d) (see Dutton and Trudgill, 2009). In the SE corner of the study area, to the east of the intra-slope horst, Horizon A is truncated by and merges with Horizon B; in this location SU1 is absent and SU2 directly overlies the composite surface (Fig. 3.8).

An amplitude map of Horizon A indicates that a series of linear amplitude anomalies are developed on this surface in the northern part of the eastern depocentre (marked ‘s’ in Fig. 3.9a). These anomalies are >2 km in length, ca. 30 m wide, spaced on average every 45 m, and trend NE-SW, subparallel to the regional dip of the slope and appear to terminate downslope at the updip end of SU1 (Fig.

3.9a). These anomalies can be clearly identified on the amplitude map (Fig. 3.9a), but they are very subtle and harder to see in cross-sections (e.g. Fig. 3.9d).

3.5.1.2 Seismic Unit 1 (SU1)

SU1 is bound below by Horizon A and above by Horizon B, and is thickest in the eastern depocentre (Figs 3.5, 3.8 and 3.10a). SU1 is dominated by chaotic reflections (ca. 75%) within the eastern depocentre; continuous, low-to-moderate amplitude seismic reflections characterise SU1 in the western depocentre (Fig. 3.11b).

Fig. 3.10: TWT-Thickness and complimentary palaeogeographic interpretation maps of SU1-6 (see Fig. 3.5). The traces of the key faults is only shown to allow the reader to easily relate the maps shown in D-F (i.e. SU4-6) with those shown in A-C (i.e. SU1-3); the presence of these fault traces does not imply that the faults were active during the specific time step. See text for full discussion.

The chaotic portion of SU1 covers 4.5 km 2 and it is thickest within the eastern depocentre, where it is up to 95 m thick (Fig. 3.10a). On seismic sections that are orientated parallel to the linear amplitude anomalies developed along Horizon A, the chaotic portion of SU1 thins downslope onto, and terminates on, the footwall of the horst, where it is only represented by a single seismic-cycle thick stratigraphic package (Fig. 3.11b).

Fig. 3.11: (a) Broadly strike-orientated (i.e. NW-SE) seismic section through the relatively proximal part of the eastern depocentre. Lateral ramps (labelled ‘lr’) developed along the prominent incision at the base of SU2 are well-illustrated on this section, as is the relief along the upper surface of SU2, which is onlapped by and controls thickness variations in SU3 and partly SU4. Gullies (labelled ‘g’) are well-developed at the base of SU5 (i.e. along Horizon E), within SU5 and within SU6. ‘K1’ and ‘K2’ are features referred to in the text. The location of the seismic section is shown in Fig. 3.11c-d. The location of the intersection with (b) is shown. (b) Broadly dip-orientated (i.e. ESE-WSW) seismic section through the eastern and western depocentre, crossing the intra-slope horst. The frontal ramp (labelled ‘fr’) developed along the base of SU2 is well-imaged on this section, in addition to moderate relief along the upper surface of SU2. Note the restriction of the chaotic, MTD portion of SU2 to the eastern depocentre, which is interpreted to be due to the presence of the intra-slope horst on the slope at this time. The location of the seismic section is shown in Fig. 11c-d. The location of the intersection with (a) is shown. See text for full discussion. (c) Instantaneous root-mean square (RMS) amplitude map illustrating seismic expression of the top surface of SU2 (i.e. Horizon C); note the rapid variability in amplitude, which is likely due to the rugosity of this surface and the variable acoustic properties of onlapping strata. The locations of Fig. 3.11a and b are shown. (d) Instantaneous dip map of the top surface of SU2 (i.e. Horizon C); note the rapid changes in dip associated with relief along this surface (see Fig. 3.11a-b). The locations of Fig. 3.11a and b are shown.

3.5.1.3 Seismic Unit 2 (SU2)

In the upslope part of the eastern depocentre SU2 is underlain by Horizon B, which in this location forms a composite surface with Horizon A; in the downslope part of the eastern depocentre and in the western depocentre, SU2 is underlain by Horizon B only (Figs. 3.5, 3.8, 3.11a and b). In the eastern depocentre a major incision surface is developed at the base of the unit along Horizon B (Figs. 3.6, 3.9 and 3.11). This incision surface has an arcuate map-view form and is composed of a SE-dipping, NE-SW- trending surface that displays c. 150 m of relief (labelled ‘lr’ in Fig. 3.9), a NE-dipping, NW-SE-trending surface that displays c. 90 m of relief (labelled ‘fr’ in Fig. 3.9), a NW-dipping, NE-SW-trending surface that displays c. 250 m of relief (labelled ‘lr’ in Fig. 3.9). The NW and SW components of this ‘scoop- shaped’ incision surface are relatively steep (>40°) and smooth at the scale of resolution afforded by the seismic data. The SE margin is also steep but has a stepped morphology in cross-section (Fig. 3.9). Relief along Horizon B results in SU2 being thickest in the eastern depocentre and the relatively thin in the western depocentre (Fig. 3.10b). Detailed mapping of locally-developed intra-SU2 seismic horizons allows SU2 to be sub-divided into three sub-units (SU2a, SU2b, and SU2c; Figs. 3.5, 3.6 and 3.8); the areal extent and seismic expression of these units differ, thus they are described separately below.

SU2a is characterised by moderate-amplitude, semi-continuous seismic reflections and is only locally developed in the northern part of the eastern depocentre (Fig. 3.8). SU2b overlies SU2a, and locally SU1, across Horizon B, which locally is strongly erosive and has up to 30 m of relief (Fig. 3.8).

SU2b is up to 140 m thick in the eastern depocentre due to relief developed along Horizon B, and thins and pinches out downslope to the SW (Fig. 3.11b). SU2b consists of moderate-amplitude, predominantly

(ca. 95%) moderately-continuous to chaotic seismic reflections and is largely confined to the NE side of the deep, NE-dipping, incision surface developed along its basal surface (Figs 3.10b and 3.11b).

However, downslope towards the SW (Fig. 3.11b) SU2b protrudes c. 80 m above and extends almost 11 km downdip of the SW margin of the incision surface; along-strike towards the NW (Fig. 3.11a), SU2b protrudes 125 m above and extends 2 km downslope of relief associated with the NW margin of the incision surface. The SE margin of SU2b appears to be largely confined within relief associated with the

69 basal incision surface (Figs. 3.6, 3.9d and 3.11a). In the eastern depocentre, immediately adjacent to the

NE-dipping erosion surface bounding the base of SU2b, a 200 m thick, 2.5 km wide, discrete package of continuous but folded, low-amplitude, parallel seismic reflections are observed (labelled ‘K2’ in Fig.

3.11a).

Fig. 3.12: (a) Time-structure map of a reflection event within the megaclast within the SU2 MTD (labelled ‘K2’ in Fig. 3.11a). (b) Instantaneous dip map of a reflection event within a feature within the SU2 MTD (labelled ‘K1’ in Fig. 3.11). For the location of both maps see Fig. 3.9c. See text for full discussion.

Mapping of an intra-package reflection over an areal extent of c. 6 km 2 indicates that the package is defined by a N-S-trending antiform (Fig. 3.12a), which is sub-parallel to the strike of adjacent erosion- related, base surface relief. Immediately adjacent to this package is another package of relatively

70 continuous, but highly-faulted reflections (marked ‘K1’ in Fig. 3.11a); the faults within this package strike broadly N-S-to-NE-SW, dip either to the E or SE and have relatively low displacements

(maximum of 18 m) (Fig. 3.12b). SU2c forms a N-S-elongate, wedge-shaped body of high-amplitude seismic reflections that is confined to and onlaps onto the NW margin of SU2b (Fig. 3.8). It is moderately-chaotic and contains several SE- or downslope-dipping, listric normal faults planes, which are overlain by rotated reflections (Fig. 3.8).

Horizon C is caps SU2 (Figs. 3.5, 3.6, 3.8 and 3.9). In the eastern depocentre the top surface of

SU2 is rugose and relief of up to 150 m is developed along Horizon C (Figs. 3.6, 3.8, 3.9 and 3.11). The majority of this relief is associated with the package of folded reflections located immediately to the NE of the NW-dipping surface developed in association with the basal incision (Fig. 3.9d). However, additional, lower-relief undulations, which do not display any specific trend in map-view, are observed in the eastern depocentre (Fig. 3.8, 3.11b and 3.11d).

3.5.1.4 Interpretation of the LSP

The chaotic seismic reflections within SU1 are suggestive of poorly-stratified and, therefore, highly-deformed strata. Based on the overall deep-water character of the studied succession and by analogy to similar deposits documented in reflection data from other deep-water successions, the chaotic portion of SU1 is interpreted as an MTD (cf. Prather et al., 1998; Armitage and Stright, 2010). Horizon A is therefore interpreted as the basal shear surface (BSS; sensu Bull et al., 2009) over which the MTD was translated and ultimately deposited (e.g. SU1 and SU2). The linear anomalies that trend broadly parallel to the dip of the slope are interpreted as basal grooves or striations developed along the BSS (cf. Nissen, et al., 1999; Gee., et al; 2005; Bull et al., 2009). It is likely that the MTD formed in response to a flow transformation from a precursor slide to a debris-flow, thus the locally-developed packages of more continuous reflections may represent more coherent, less disaggregated slide material, which may have only been translated a limited distance (e.g. Middleton and Hampton, 1976; Kneller and Buckee, 2000;

Mulder and Alexander, 2001; Gee et al., 2005; Bull et al., 2009). The striations along Horizon A are

71 broadly normal to the overall slope (cf. Fig. 3.9a and 3.9b), suggesting that the MTD in SU1 was sourced from the slope and translated to the W and WSW. A discrete headwall and zone of stratigraphic depletion cannot be identified in the current dataset, suggesting that the source area for the slide lies slightly outside of the study area to the N or NE. SU1 is thickest in the eastern depocentre and is almost fully restricted to the eastern depocentre on the NE side of the horst, suggesting that this structure was present on the seabed at the time of MTD emplacement. This tectonic feature clearly limited the run-out distance of the MTD and the thin interval of SU1 preserved on the footwall of the horst may represent a more dilute component of the MTD that out-ran the main, cohesive part of the flow and was deposited on top of the horst (Fig.

3.10a and 3.11b).

SU2a-c have a similar seismic facies expression to SU1, and are also interpreted as MTDs (cf. seismic facies A, B l, and B h of Prather et al., 1998; see also Beaubouef and Friedman, 2000). The erosive base to SU2 is therefore interpreted as a basal shear surface that was cut by the flow that deposited the overlying unit. The NE-dipping, upslope-facing margin to the incision surface is interpreted as a ramp at the front of the MTD (i.e. frontal ramp, e.g. Bull et al., 2009, or toe-wall, e.g. van der Merwe et al., 2011), whereas the SE- and NW-dipping margins are interpreted as ramps at the margins of the MTD (i.e. lateral ramps). This surface was associated with entrainment of a significant thickness (up to 250 m) of substrate into the parent flow (see Gee et al., 2005, 2006). Despite its scale, this surface only partly contained the overlying MTD, thereby allowing the genetically-related deposit to be classified as a ‘frontally-emergent mass-transport complex’ (sensu Frey-Martinez et al., 2006). Overall, the chaotic, MTD portion of SU2 does not extend into the western depocentre, suggesting that, during deposition of this unit, the intra-slope horst was present and formed a prominent bathymetric obstacle to the downslope transport of material within the parent sediment gravity-flow.

Locally-developed, semi-continuous reflections in SU2a and c may suggest that these deposit formed from precursor slides that underwent only partial disaggregation, in comparison to SU2b, which was more disaggregated. Furthermore, the broadly S- to SE-dipping, stratabound listric normal faults in

SU2c are interpreted as being related to downslope stretching of a slide, thereby implying that this unit

72 may have been deposited from the northern, S-dipping slope. The sediment source area for the sub-units of SU2a and c would, therefore, have been broadly similar to that for SU1. However, the orientation of the scoop-shaped basal erosion surface of SU2b and the sense of frontal emergence (i.e. to the SW and locally to the NW) strongly suggest that this unit was sourced from the E and transported downslope to the W. The composite MTC represented by SU2 therefore had two main sediment source areas; the S- dipping slope observed in the north of the study area and a broadly W-dipping slope that is inferred to exist to the E of the study area (Fig. 3.4a).

The package of continuous, contorted reflections that occur in SU2b in the SE of the study area is interpreted as a ‘megaclast’ that underwent relatively limited disaggregation and deformation in the parent debris-flow (e.g. MacDonald et al., 1993; Lee et al., 2004; Alves, 2010; Dunlap et al., 2010;

Jackson, 2011). The fold within the megaclast may have formed due to the impact of clast against the steep, NE-dipping frontal ramp. The fold orientation also supports an interpretation that the flow that deposited SU2b was sourced from outside of the study area to the E, and flowed towards the W-SW (i.e. perpendicular to the frontal ramp).

The relief along the top of SU2, which locally reaches up to 150 m, could have been caused by erosion of the top of the unit by subsequent sediment gravity flows. However, in the overlying stratigraphic interval (SU3), there is no evidence for depositional systems that may be associated with such large magnitudes of erosion (i.e. channel-confined channel systems). The rugose top of SU2 is therefore interpreted to have formed in response to ‘freezing’ and abrupt deposition of material in the parent debris-flow, possibly augmented by the development of intra-MTC thrusts (e.g. Shanmugam et al.,

1995; Frey-Martinez et al., 2006; Armitage et al., 2009; Jackson and Johnson, 2009; Bull et al., 2009).

The megaclast in SU2b clearly contributed to the development of local relief along the top of SU2, especially immediately updip of the frontal ramp defining the base of the unit. The relief along the top of

SU2 is only in the eastern depocentre in association with MTC material; along-strike of the margin, towards the NW, such relief is not observed, and the slope has a simple, smooth, basinward-dipping profile at the same stratigraphic level (Fig. 3.10b).

3.5.2 Middle Seismic Package (MSP)

The MSP is composed of SU3 and 4, and is bound below by Horizon C and above by Horizon E (Fig.

3.5).

3.5.2.1 Seismic Unit 3 (SU3)

SU3 is bound at the base and top by Horizons C and D respectively (Fig. 3.5), and is composed of laterally continuous, moderate-to-high amplitude reflections that onlap relief developed along the top of

SU2 (Figs. 3.6, 3.8, 3.9 and 3.11). The distribution of SU3 is, therefore, complicated because it is largely confined to a series of sub-circular, up to 16 km 2 (in areal extent), 150 m deep depocentres that are developed along the top of SU2 in the eastern depocentre (Fig. 3.10c).

3.5.2.2 Seismic Unit 4 (SU4)

The base and top of SU4 is defined by Horizons D and E respectively (Fig. 3.5, 3.6 and 3.11a).

The unit is thickest in the eastern depocentre and gradually thins downslope into the western depocentre

(Figs. 3.10d and 3.11b). The unit is composed of high-amplitude, laterally-continuous, parallel-to-subparallel seismic reflections, which are broadly parallel to and conformable with those in SU3 (Figs. 3.6 and 3.7).

3.5.2.3 Interpretation of the MSP

The seismic expression of SU3 and 4 are similar in terms of overall seismic response, if not reflection termination style, to seismic facies ‘C’, ‘D’ and ‘E’ of Prather et al. (1998). SU3 and 4 may therefore represent: (i) laterally continuous, sheet-like, sandstone-dominated bodies, which were deposited from turbidity currents; or (ii) sheet-like, mudstone-dominated strata, deposited by the fallout from suspension of pelagic and/or hemipelagic, fine-grained material. The thickness variations and distribution observed in SU3, in combination with pronounced stratal onlap onto the base-bounding surface (Horizon C), indicates that relief on the top of SU2 formed prominent relief during the early stages of deposition of SU3. However, with time, the amplitude of the depositional relief was reduced and

74 eventually healed in response to continued deposition; by the end of deposition of SU3 only minimal relief was present on the seabed upslope of the intra-slope horst. In contrast to SU3, SU4 displays only minimal thickness variations with respect to the intra-slope horst (Fig. 3.10d), suggesting that only minor tectonic relief was associated with this feature during its deposition.

3.5.3 Upper Seismic Package (USP)

The USP is composed of SU5 and 6, and is bound below and above by Horizons E and G respectively (Fig. 3.5).

3.5.3.1 Seismic Unit 5 (SU5)

SU5 is bound above and below by Horizons E and F respectively (Figs. 3.6 and 3.11). It is thickest in the eastern depocentre, and thins downslope to the W into the western depocentre and upslope to the E onto the SE-dipping clinoforms (Fig. 3.10e). A series of narrow (30–50 m wide), shallow (15 m deep), ‘v’ and ‘u’-shaped incisions, which are >4 km in length, are mapped at the base of SU5 (i.e. along

Horizon E). These incisions trend N-S to NE-SW and are best developed on the SW-facing slope that is located in the NE corner of the study area. The incisions merge downslope to the SW and eventually terminate on the eastern flank of the horst (Figs. 3.13c and d). The lower part of SU5 infills and onlaps incision-related relief (Figs. 3.9d, 3.11b and 3.13a). N-S-to-NE-SW-trending, ‘v’- or ‘u’-shaped incisions are also developed within SU5 and these are superimposed onto laterally continuous, parallel, high–to- moderate amplitude seismic reflections (Figs. 3.6, 3.8, 3.9d, 3.11a and 3.13). These incisions are 170–200 m wide, 20–30 m deep and can be mapped for up to 3 km, and they also appear to terminate on the eastern edge of the intra-slope horst (as exemplified by incisions along Horizon F that bounds the top of the unit; Fig. 3.13d).

Fig. 3.13: Seismic expression, morphology, scale and distribution of slope gullies in the upper part of the studied succession. (a) Seismic section from the eastern depocentre focused on gullies (labelled ‘g’) at the base of SU5 (i.e. Horizon E) and within SU5. See Fig. 3.13b for location of the seismic line. (b) Instantaneous dip map of Horizon E (i.e. base SU5). The locations of the close- up shown in Fig. 3.13c and the seismic section in Fig. 3.13a are shown. (c) Close-up of the instantaneous dip map of Horizon E (i.e. base SU5) shown in (a). (d) Instantaneous dip map of Horizon F (i.e. base SU6). (e) Instantaneous dip map of Horizon G (i.e. seabed). g=gullies; p=pockmarks.

3.5.3.2 Seismic Unit 6 (SU6)

SU6 is bound at the base by Horizon F and at the top by the seabed (Horizon G; Fig. 3.5 and 3.6).

SU6 comprises laterally-continuous, variable amplitude reflections (Figs. 3.6 and 3.11). The lower part of

SU6 infills and onlaps incision-related relief developed along the top of SU5 (Fig. 3.6). SU6 is thickest in the eastern depocentre and the unit gradually thins towards the west (Fig. 3.10f). Several prominent ‘v’- shaped incisions, which are 100-200 m wide and 20-30 m deep, are developed within and at the top of

SU6 (i.e. on the present seabed; Fig. 3.13e). These ‘v’-shaped incisions trend NE-SW, and extend from the eastern edge of the dataset, across the horst, and terminate in the western depocentre (Fig. 3.13e).

Within the western depocentre numerous circular depressions are developed. They are up to 70 m wide,

20-50 m deep and are developed over ca . 9 km 2 of the seabed (labelled ‘p’ in Fig. 3.13e).

3.5.3.3 Interpretation of the USP

Based on its overall seismic expression and comparison to SU3-5, SU6 is interpreted to be deposits of either (i) laterally continuous, sheet-like, sandstone-dominated bodies, which were deposited from turbidity currents; or (ii) sheet-like, mudstone-dominated strata, deposited by the fallout from suspension of pelagic and/or hemipelagic, fine-grained material. (e.g. Amery et al., 1978; Shaffer, 1990;

Pulham et al., 1993). Based on their map-view and cross-sectional geometries, their scale and their occurrence on a submarine slope, the subtle erosional features developed at the base of, within and at the top of SU5 and 6 are interpreted as slope gullies (e.g. Straub and Mohrig, 2009; Jobe et al., 2011; Sanchez et al., 2012). The origin of slope gullies is contentious, but they are thought to form in response to slope incision by thick, dilute, mud-rich turbidity currents, and are thought to characterise the low sediment- supply portions of passive margins (e.g. Jobe at al., 2011). Irrespective of their exact origin, the presence of gullies implies that SU5 and 6 are fine-grained and, therefore, characterised by mudstones rather than sheet-like, turbidite sandstones (i.e. seismic facies ‘D’ and ‘E’ of Prather et al. 1998). NNE-SSW to ENE-

WSW orientation of the gullies at the base of SU5, and the NE-SW trend of the gullies in SU5 and SU6

77 suggests southerly- to westerly-directed, downslope-directed gravity-flows and sediment transport at this time (i.e. broadly parallel to the overall dip of the slope). The gentle south-westwards thinning of SU5 and

6, and the extension of the slope gullies across the horst, implies that relief associated with this structure had been healed by this time. The circular depressions identified at the seabed within the western depocentre are interpreted as pockmarks formed in response to the focused upward flow of water, oil and/or gas (e.g. King and MacLean 1970; Hovland and Indreeide, 1980; Hovland 1981a, b; Moss and

Cartwright, 2010).

3.6 Discussion

3.6.1 Stratigraphic organisation of deep-water successions and the role of MTDs

Our seismic facies-based analysis of the Malembo Formation (middle Pliocene-to-Holocene), offshore Angola indicates that this deep-water succession has a distinct stratigraphic organisation, which is consistent with that predicted by the graded slope concept (e.g. Prather et al., 1998; Pirmez et al., 2000;

Pyles et al., 2011; Moscardelli et al., 2012). The lower package (i.e. LSP) is comprised of MTDs and associated deposits (e.g. debrites, slides), which are suggestive of slope degradation and deposition during a period when the slope was above grade. In contrast, the middle (i.e. MSP) and upper (i.e. USP) packages are characterised by gully-bearing, slope mudstone-dominated units that suggest the predominance of aggradational processes and slope accretion during the latter stages of margin development. We therefore interpret that the middle and upper packages document a time when the slope was at grade.

A key additional observation we make is that this long-term stratigraphic organisation of the upper slope was influenced by the evolving bathymetric template. First, deposition of SU1 and 2 at a time when the slope was above grade, in particular the distribution and thickness of the MTDs they contain was controlled by an intra-slope horst, which provided intra-upper slope accommodation (i.e. tectonic accommodation; Fig. 3.14). In contrast, during deposition of SU3-6, which document a progressive change to at grade conditions, intra-slope relief was minimal due to filling of accommodation by SU1 and

2 (Fig. 3.14). In summary, the Malembo Formation records an overall change in depositional processes from MTD emplacement and healing of tectonic relief associated with an out-of-grade slope, through healing of depositional relief (see below), to limited accommodation and the progradation of an at grade, gullied upper slope.

Fig. 3.14: Schematic cross-sections illustrating the styles and evolution of relief and accommodation types on continental margins based on our observations from offshore Angola. Key to all sections shown in (a). An overall upward transition from ponded to healed to slope accommodation is observed (i.e. Time 1 through 4), although higher-frequency, lower-amplitude variations in accommodation occurs due to formation and progressive filling of relief developed at the top of intra-slope MTDs (i.e. Time 2). In the inset in Time 2, (a), (b) and (c) represent stages in filling of ponded depositional accommodation on top of the intra-slope MTD. The relative roles of the various types of accommodation are shown on the chart insert on each time-step; the width of the triangle defines the relative importance of that style of accommodation compared to the other styles. See text for full discussion.

3.6.2 Controls on accommodation development and the impact on upper slope depositional patterns

Two main types of accommodation developed on the slope during deposition of the Malembo

Formation: (i) tectonic accommodation; and (ii) depositional accommodation (Figs. 3.1, 3.14 and 3.15).

Tectonic accommodation was related to the presence of a normal fault-bound horst, which developed in response to thin-skinned, supra-salt extension during tilting of the margin. The horst was oriented broadly perpendicular to the main sediment input direction and it partitioned the upper slope into an inner (i.e. eastern) depocentre and outer (i.e. western) depocentre, the latter being open to the main depocentre located downslope towards the W. Thickness variations in SU1-3 indicate that up to 250 m of relief was associated with the initial formation of the horst. Existing models, based largely on studies of the Gulf of

Mexico, suggest that the initial fill of ponded accommodation is dominated by layered and amalgamated, sheet-like sandstones (e.g. Mutti and Normark, 1987; Prather et al., 1998; Booth et al., 2000; Prather,

2003). However, in the example presented here, MTDs, which are inferred to be mud-dominated, dominated the initial period of intraslope basin filling; the development and dominance of MTDs at this specific stratigraphic level is interpreted to reflect the out-of-grade condition of the slope at this time, and the short transport distance that limited flow transformation. We suggest that the horst influenced the flow behaviour of the MTDs in SU2 by generating a steep frontal ramp, which limited the transport distance and degree of flow transformation.

Two types of depositional accommodation are observed: (i) upper surface relief and accommodation related to ‘buckling’ of the MTD upper surface during deposition; this type of relief can be up to a few tens of metres high and is equivalent to the Tier-2 relief of Armitage et al. (2009); and (ii) upper surface relief and accommodation related to ‘megaclasts’ that protrude above the MTD upper

81 surface; this type of relief can be up to several hundred metres high and is equivalent to Tier-3 relief of

Armitage et al. (2009) (see also Deptuck et al., 2003; Pickering and Corregidor, 2005; Moscardelli et al.,

2006; Moscardelli and Wood, 2008; Jackson and Johnson, 2009). Due to the limited areal extent of the outcrops, field-based studies have principally stressed the magnitude (height) of relief associated with the upper surfaces of MTDs. 3D seismic analysis compliments these studies and indicate that this relief can extend for several kilometres along-strike, and can result in the formation of a series of sub-circular, almost completely-enclosed transient minibasins, which controlled gross thickness patterns in overlying strata (i.e. relief at the top of SU2, which controls thickness variations in SU3). We speculate that MTD minibasins that are related to protruding megaclasts and buckle folds are likely to be randomly distributed down and along slopes, have variable amplitudes of relief, and that the orientation of these depocentres are likely to be highly variable and not show a predictable orientation with respect to the slope (Fig. 3.15).

Fig. 3.15: Chart illustrating the styles, dimensions and orientations (relative to slope) of different styles of tectonic and deposition relief and accommodation on continental margins, based on our observations from offshore Angola.

The lack of well data precludes calibration of the detailed depositional response to the development of intra-slope depositional relief. If post-MTD sedimentation patterns are dominated by turbidity current-fed depositional systems, it is likely this complicated relief would control sediment dispersal and stratal architecture of sandstone-dominated channels and lobes (e.g. Pickering and Corregidor, 2005; Armitage

82 et al., 2009; Jackson and Johnson, 2009; van der Merwe et al. 2009). Furthermore, due to a lack of well data and, more specifically, biostratigraphic data, the timescales over which tectonic and depositional relief influenced upper slope deposition cannot be constrained. We speculate, however, that tectonic relief, which is likely to be greater than depositional relief on many continental margins, may influence depositional patterns over a significantly longer duration. There will, however, be specific times and/or location on the slope where depositional relief is more important that tectonic relief in controlling depositional patterns and, therefore, can influence the local (<5 km length-scale) distribution and stratal architecture of upper slope reservoirs. A key observation from this study is that, irrespective of when and where they occur on continental margins, tectonic and depositional relief both provide a mechanism for trapping reservoir-prone clastic sediments in an upper slope setting, in an area which is otherwise typically associated with sediment bypass to the lower slope and basin floor.

3.7 Conclusion

The key conclusions from the 3D reflection seismic analysis of the mid-Pliocene-Holocene succession in the Lower Congo Basin, offshore Angola are:

1. A series of NW-SE-striking, thin-skinned, salt-detached normal faults are developed on the upper

slope and are responsible for the formation of a series of depocentres separated by an intra-slope

horst.

2. Three main seismic packages are recognised and these are split into six seismic units based on

seismic facies transitions, and the geometry and regional extent of their bounding surfaces. The

seismic facies within this stratigraphic succession document deposition on a submarine slope.

3. The initial stages of deposition occurred when the slope was above grade and dominated by

degradational or erosional processes; deposition at this time was therefore characterised by the

emplacement of a series of MTDs that were up to 250 m thick and at least 100 km 2 in areal

extent. The downslope translation, and hence thickness and distribution of these MTDs was

controlled by the intra-slope horst, which formed a prominent at-seabed structure at this time.

4. Deposition of MTDs was associated with the entrainment of substantial volumes of substrate and

the formation of significant (150 m) depositional relief along their upper surfaces; this relief

controlled large-magnitude and abrupt changes in thickness in the overlying strata.

5. The latter stages of deposition occurred when the slope began to achieve ‘grade’ and depositional

processes dominated. Deposition at this time was from either turbidity currents and/or suspension

fallout; deposition during the latter stages of margin development was not controlled by the intra-

slope horst, which had been ‘healed’ by this time.

6. Our study highlights the role that tectonically- and depositionally-controlled bathymetric changes

can have on accommodation development on the upper slope, and the manner in which this

controls subsequent depositional patterns.

7. Tectonic and depositional accommodation provide a mechanism for capturing and trapping

reservoir-grade sediments in an upper slope setting that conventional models would predict are

areas of coarse-grained sediment bypass to the lower slope and basin-floor.

Chapter Four

3D seismic analysis of pockmark distribution and development within mid-Pliocene to modern

strata, offshore Angola

4.1 Abstract

The mid-Pliocene-to-Holocene stratigraphic succession in the upper- to mid-slope offshore

Angola is locally characterised by occurrences of pockmarks distributed across the stratigraphy. Three- dimensional seismic data allow the investigation of the possible sources of the expelled fluid which generated the pockmarks and the geologic controls that constrain their distribution.

The study area is dissected by NW-SE-trending normal faults. In the upper- to mid-slope area, these faults form a graben, which passes downslope into a salt-influenced horst on the upper-slope that divides the study area into two depocentres. Six seismic units (from base to top: SU1, SU2, SU3, SU4,

SU5, and SU6) are identified within the study interval and surface seismic attributes are extracted along their top surfaces. Pockmarks occur on the top surfaces of the uppermost four SUs and are suggested to result from migration of fluids from rapidly buried and poorly compacted sediment gravity flows in unit

SU2, and from deep-seated Miocene source rocks and leaked hydrocarbon reservoirs.

The surficial distribution of the pockmarks shows two areal patterns, curvilinear strings and disorganised complexes. These patterns are attributed to two geological controls on subsurface fluid flow.

The first control is structural, represented by large normal faults which deformed the entire Miocene-to-

Holocene stratigraphy and appear to be sealing at SU3 and SU4 levels, but non-sealing at SU5 and SU6 levels. The second control is stratigraphic, and consists of porous-permeable Mass Transport Deposits

(MTDs) in unit SU2, and carrier beds in units SU5 and SU6. The faults and the lateral terminations of

MTDs are linear features, and genetically related pockmarks are developed as strings of pockmarks above them, while the carrier beds are associated with disorganised complexes of pockmarks. The overall pockmark distribution exhibits greater morphological diversity at upper stratigraphic levels.

This work shows that analysis of 3D seismic data is an effective tool in evaluating the spatial distribution and geological controls on development of pockmarks in marine environments. The resulting insights are useful in predicting subsurface fluid migration, which has direct impacts in the exploration and production of hydrocarbons, and in the efficient prediction of shallow drilling hazards.

4.2 Introduction

Pockmarks are crater-like features created by fluids that erupt and stream through sediments on the seabed (Hovland & Judd 1988; Hovland et al., 2002, Gay et al., 2003; Gay et al., 2006; Cartwright et al., 2007; Moss and Cartwright, 2010) and are indicators of fluid flow (King & MacLean, 1970; Hovland et al., 1987; Hovland and Judd, 1988; Hovland et al., 2002). They generally occur in fine-grained sediments as cone-shaped, circular or elliptical depressions of varying diameter and depth (Hovland et al.,

2002; Gay et al., 2003; Gay et al., 2006; Moss and Cartwright, 2010; Sun et al., 20011; Moss et al., 2012) and may be found in many geological environments that may be contiguous, including petroleum- generating regions (Josenhans et al. 1978; Newton et al. 1980; Hovland et al. 1984; Ellis and McGuinness

1986; Rise et al. 1999), regions underlain by crystalline bedrock (Floden and Soderberg 1988;

Papatheodorou et al. 1993), estuaries and coastal areas (Scanlon and Knebel 1989; Hill et al. 1992; Taylor

1992; Kelley et al. 1994; García-García et al. 1999), areas of groundwater flow (Bussmann and Suess

1994) and lakes with or without hydrothermal activity (Pickrill 1993). Understanding the controls on development of pockmarks may have significant impact in the exploration and production of petroleum resources, and the identification and management of their associated potential drilling risks.

4.2.1 Sources of pockmark-forming fluids

Subsurface sources of fluid that were expelled to form pockmarks fall into two broad categories, related to (1) rapid sedimentation and (2) generation of fluids (e.g. hydrocarbons) during burial diagenesis.

4.2.1.1 Rapid sedimentation in clastic depositional systems

Clastic depositional systems in shallow- and deep-water environments are prone to rapid sedimentation of low permeability layers in which associated pore fluid cannot escape (e.g Berhmann et al., 2006) because the expulsion of pore fluid cannot keep pace with the high rate of sedimentation. The consequent build-up of pressure by the entrapped fluid as a result of rapid increase of the overburden pressure (e.g. Bjorlykke, 1993, Bolton & Maltman 1998; Berhmann et al., 2006) can lead to the generation of excess pore fluid pressures which may be sufficient to cause escape from the sand and silt bodies through the muddy cover, thereby creating pockmarks on the seafloor (Cole et al. 2000). Although rapid sedimentation takes place in both shallow- and deep-water environments, the phenomenon is common to deepwater settings where rapid, episodic, localised deposition of sands within an overall

“background” of laterally continuous mud deposition is prevalent (Shepard, 1951; Nelson 1982).

Pockmarks have been found to occur in ‘undisturbed’ sediments on shelf slopes in areas adjacent to and often upslope of slides (e.g. Bugge et al., 1987; Evans et al., 1996; Bouriak et al., 2000; Hovland et al., 2002; Gravdal et al., 2003; Mazzini et al., 2005; Mienert et al., 2010). Examples of such occurrence include the Storegga Slide, offshore Norway; the Gulf of Cadiz, offshore Spain, where numerous slides have occurred; adjacent to the Humbolt Slide, offshore California, USA; and on the large ‘collapse feature’ on the Blake Outer Ridge, offshore North Carolina, USA (Dillon et al. 1998; Baraza et al. 1999;

Gardner et al. 1999; Yun et al. 1999; Bouriak et al. 2000; Dillon and Max 2000). Pockmarks also occur above shallow buried turbiditic channels with overlying hemipelagic sediments playing the role of an impermeable seal (e.g. in the Lower Congo Basin; Gay et al., 2003, Gay et al., 2006). In the sedimentary sequences of the Naust Formation which is characterised by hemipelagites and contourites mainly deposited during interglacials, offshore mid-Norway, extremely rapid sediment loading occurred in proglacial depocenters to create overpressures in the interglacial sediments (Hjelstuen et al., 2005; Berg et al., 2005; Bryn et al., 2005; Forsberg & Locat, 2005; Solheim et al., 2005) which in turn developed fluid escape features on the seafloor around the Storegga slide (Bugge et al., 1987; Evans et al., 1996; Bouriak et al., 2000; Gravdal et al., 2003; Mienert et al., 2010).

4.2.1.2 Fluid expulsion from diagenesis, deep-seated source rocks and leaked reservoirs

Pore water generated during diagenetic transformation of clays, particularly smectite to illite, constitutes an important source of overpressured fluid in the subsurface (e.g. Bruce, 1984; Osborne and

Swarbrick, 1997), a process governed by the following reaction:

K-feldspar + Smectite => Illite + Quartz + Water

There is a broad correlation between the onset of overpressuring and the smectite to illite transformation (Bruce, 1984), as a result of the progressive generation of water within the pore space of the host sediments. Although significant water is released in the transformation, overpressure will not result unless an overall volume expansion occurs during the reaction (Osborne and Swarbrick, 1997) and the coupling of the smectite-illite reaction and overpressure could be due to reduction in sediment permeability and change in compressibility induced by the transformation. Osborne and Swarbrick (1997) estimate that smectite to illite reactions commonly result in volume increases of 0.1-4.1% and volume decreases of 0.7-8.4%.

The primary expulsion of petroleum out of the source rocks is for the most part driven by the increased pressure caused by the phase transition from kerogen to petroleum (Bjorlykke, 1993). The migration of subsurface-generated hydrocarbons to the near-surface is well documented (Horvitz, 1939;

Rosaire, 1940; Kartsev et al, 1959; Horvitz, 1981; Jones and Drozd, 1983; Horvitz, 1985; Klusman, 1993;

Schumacher and Abrams, 1996). The expulsion of hydrocarbons from mature source rocks and their migration, storage, and remigration (leakage) through the overburden toward the surface provide important evidence of a working petroleum system (Heggland, 1997; O’Brien et al., 2005; Cartwright et al., 2007). During migration, some of the fluids may be trapped or temporarily stored in reservoirs, whereas others imprint the stratigraphic succession with pipe features and pockmarks during their path to the surface (Magoon and Dow, 1994; Cartwright et al., 2007, Andresen et al., 2011). Analysis of 3D seismic reflection data has proven that pockmarks and their associated pipe-like conduits are formed by the rapid migration of gas-charged pore fluids from deeper reservoir formations (Heggland, 1997;

Heggland, 1998; Berndt et al., 2003; Bünz et al., 2005; Ligtenberg, 2005).

4.2.2 Fluid migration pathways for pockmark development and distribution

Subsurface migration conduits convey mobilised fluids from their point of expulsion through the cover sedimentary successions to the sea floor or land surface. The development of pockmarks as a result of fluid movement through structural (e.g. faults, salt structures) and depositional features that focus fluid flow (e.g. bedrock interfaces, clastic dykes and sills) have been highlighted by Gay et al., (2006, 2007),

Hübscher and Borowski (2006), and Pilcher and Argent (2007).

4.2.2.1 Structural migration pathways

The presence of structures, especially faults and salt and shale diapirs, that developed before or at the same time as fluid movement and adjacent to potential sources of entrapped fluid can focus fluid migration (e.g. Eichubl et al., 2000; Bünz et al., 2005; Andresen and Huuse, 2011; Andresen et al., 2011).

Eichubl et al. (2000) also documented an example of a slump head scarp where increased seepage was observed to follow fault and anticlinal trends, suggesting a structural control on mass wasting caused by upward fluid migration. Fault systems can be sealing if they juxtapose permeable reservoir rocks against impermeable non-reservoir rocks (Allan, 1989; Freeman et al., 1990; Yielding et al., 1997) or if the faulting process has generated a membrane seal by cataclasis (Antonellini and Aydin, 1994), cementation by diagenesis, framework grain-clay mixing (Watts, 1987; Knipe, 1992), or clay smearing (Bouvier et al.,

1989; Sibson, 1994). Active faults and fractures provide high-permeability conduits for fluid flow during deformation (e.g., Sibson, 1994; Barton et al., 1995), and juxtaposition or membrane seals may be breached if the fault is reactivated subsequent to hydrocarbons charging the trap (Jones and Hillis, 2003).

Faults and fractures that are subjected to a state of stress conducive to reactivation are prone to act as conduits for fluid flow (Barton et al., 1995). Dilatant fracturing and brecciation in faults (Roehl, 1981) can also enhance the permeability of the faulted rock (e.g. Norton and Knapp, 1977; Howard and Nolen-

Hoeksema, 1990).

4.2.2.2 Depositional migration conduits

Interbedded sandstones act as conduits for the flow of fluids being expelled from compacting shales when excess hydraulic head is initiated as a result of substantial differences in pore pressure which

89 exist between the two lithofacies (Bredehoeft et al., 1988). The continuous nature of sediment compaction means that flow occurs between the sandstones and shales throughout the burial process. Where laterally extensive inclined aquifers exist, deep highly overpressured regions are connected to shallower, less overpressured regions by a permeable pathway (Yardley and Swarbrick, 2000). In deepwater successions, strong lithological and hydraulic contrasts exist between sandstone-rich intervals (e.g. submarine fans) and the mudstone-dominated sediments that enclose them, such that turbiditic sandstones may act as fluiod-flow conduits associated with pockmarks (e.g. Gay et al., 2003; Hübscher and Borowski, 2006;

Andresen and Huuse, 2011, Andresen et al., 2011). Fluid seepage, and consequent development of pockmarks, may also occur preferentially along slump scarps and in submarine canyons (Orange and

Breen, 1992; Eichubl et al., 2000).

4.2.3 Study aims

This study focuses on the Mid-Pliocene-to-Holocene slope succession in Block 4, offshore

Angola. The aims of this study are to: (i) document and characterise the morphology and distribution of pockmarks within the stratigraphic succession; (ii) deduce the source of the fluids that created the pockmarks; and (iii) assess the controls responsible for formation of the pockmarks.

Fig. 4.1: (a) Map of southwest African margin showing the location of the Lower Congo Basin in relation to main bounding hinge zones. (b) Map of tectonic framework of the Lower Congo Basin indicating in red box the location of Block 4 in relation to other Blocks. Modified from Anderson et al., (2000).

4.3 Geological setting

The Lower Congo Basin, which lies between 7 °30’S and 7 °45’S (Fig. 4.1) evolved as a result of early Cretaceous rifting, due to basement-involved upper crustal extension which coincided with the opening of the South Atlantic Ocean (Brice et al., 1982; Karner et al., 1997; Marton et al., 2000). The post-rift stratigraphy is subdivided into two major sequences separated by an Oligocene unconformity

(Seranne et al., 1992; Seranne, 1999). The Late Cretaceous to Early Oligocene sequence, which represented a time of stable green-house climate, is characterised by aggradational siliciclastic and carbonate ramp deposits (Seranne, 1999) that contain source rocks which produced thermogenic oils and gases (Burwood et al., 1990). These source facies include the Pre-Salt Barremian Bucomazi Formation source, Cenomanian-Turonian marine clastic Iabe Formation sourced petroleum systems (Cole et al.,

2000), and the marine-sourced (mainly post-salt) oils from the Kwanza (Schiefelbein et al., 1999). These source rocks offshore Angola probably entered the hydrocarbon maturation and generation window

91 during the Aptian-Turonian which is equivalent maturation timing for other regionally contiguous basins including the Namibe, Walvis, Luderitz, and Orange basins (e.g. Bray et al., 1998). Sedimentation during the early Oligocene to Recent sequence is typified by oscillations between dry and wet climatic settings with high amplitude (tens of metres to >100 m) and high frequency (100 kyr – 1 Myr) sea-level changes that typify ice-house conditions (Bartek et al., 1991; Miller et al., 2005; Zachos et al., 2001). This interval is characterized by progradation of a terrigenous wedge accompanied by deep incision and erosion of the

Congo submarine canyon (Seranne, 1999). This progradation and associated delivery of terrigenous clastic sediment to the Lower Congo Basin developed the large turbidite fan systems which represent the hydrocarbon habitat of a large proportion of the hydrocarbons discovered in the Lower Congo Basin

(Brice et al., 1982; Uchupi, 1992; Droz et al., 1996; Sikkema and Wojcik, 2000; Broucke et al., 2004).

During the Miocene, a period of continental uplift caused renewed thin-skinned extension and an influx of coarse clastic sediment (Fig. 4.2) (Brice et al., 1982; Duval et al., 1992; Lavier et al.,2001) and sea-level changes controlled the intensity and style of sediment delivery to deepwater turbidite fans (Broucke et al.,

2004).

Fig. 4.2: Regional geoseismic section across the Angolan margin. Modified from Lentini et al. (2010). Major structures and depositional sequences are highlighted. COB=continent–ocean boundary. Note the occurrence of a thick, highly-mobile salt layer (Loeme Formation) across the entire width of the margin. The study area lies within the updip extensional domain (see red box). The location of transect is shown in Fig. 4.1b.

Turbidites in the stratigraphic succession within Block 4 occur in three major depositional sequences: the 17.5–15.5 Ma Burdigalian- Early Miocene, the 8.2–6.3 Ma Tortonian-Late Miocene, and the 6.3–5.5 Ma Messinian-Late Miocene (Anderson et al., 2000). The Early Pliocene to Present is

92 characterised by deposition of sparse, fine-grained muddy sediment and hemipelagic sedimentation

(Jansen et al., 1984a; Pufahl et al., 1998; Uenzelmann-Neben, 1998). The western part of Block 4, offshore Angola, is characterised by a series of normal fault-bounded graben and half-graben (Fig. 4.3).

The upper- to mid-slope of the study area is deformed into a fault-bound graben located above a salt weld between two raft blocks of the late Cretaceous Pinda Formation. The fault-bound graben switches downslope into a fault-bound horst directly overlying a large raft block. The horst subdivides the upper slope into two main depocentres, which are denoted here as the eastern and western depocentres (Fig.

4.3).

Fig. 4.3: (a) Structural map illustrating the setting of the study area. The major and minor normal faults developed during the Mid-Pliocene are marked. Note the subdivision of the lower slope area into two depocentres separated by a horst domain. The axis of the eastern depocentre during the Mid-Pliocene is indicated on the map. (b) Oblique-dip oriented geo-seismic profile, across the lower slope, showing the analysed seismic units (SUs) and structures in the study area. Location of the geoseismic section is indicated in Figure 4.4a. The series of faults denoted by ‘S’ are synthetic faults while those denoted ‘A’ are antithetic faults. Modified from Olafiranye et al. (2013).

Middle Pliocene-to-Holocene strata of the clastic-dominated Malembo Formation are represented in the study area by ca. 620 m high clinoforms that dip towards the SE (Fig. 4.4). This unit was deposited during thin-skinned extension and comprises a succession of sandstone-dominated, submarine channel-fill deposits, mudstone-dominated MTDs and hemipelagic mudstones (Jansen et al. 1984b; Uenzelmann-

Neben 1998; Olafiranye et al., 2013). Riverine hyperpycnal plumes also delivered fine-grained sediment to the Lower Congo Basin (Cooper 1999) during the Middle Pliocene-to-Holocene.

Fig. 4.4: Strike oriented vertical profiles of the analysed seismic units (SUs) and structures in the study area. (a) Uninterpreted seismic and (b) interpreted geoseismic sections across the shelf, lower and upper slope. Location of the seismic section is indicated in Figure 4.3a. Modified from Olafiranye et al. (2013).

4.4 Dataset and methodology

The dataset used for this study is a single full stack volume of 3D seismic which cover an area of

760 km 2 within 500 – 700 m water depth. The quality of the data is good-to-excellent particularly in the uppermost 1200 msecs that represent the interval of the study. The seismic survey was sampled at intervals of 4 msecs TWT (milliseconds two-way travel time) with inline and crossline bin spacing of

12.5 m. The seismic volume is time-migrated but not depth-migrated. Therefore, where necessary, true depth and thickness are estimated using an interval acoustic velocity of 2000 m/s that is applicable to deepwater channel systems in offshore Angola (Bakke et al., 2008). This interval velocity and dominant data frequency of 45 Hz result in a vertical resolution of approximately 11 m with an error margin of +/- 1 m. Based on the acoustic impedance (reflectivity) of the seafloor identified in the seismic volume, the polarity of the data is SEG normal whereby a downward increase and downward decrease in acoustic impedance are denoted by blue and red reflections respectively (Figs. 4.4-4.6).

Fig. 4.5: Description of the six mapped seismic units (SUs) and their constituent seismic facies. The bold, dashed orange line above “SU2a” sub-unit in the chart indicates the transition from sheared deposits that contain preserved kinematic indicators to turbidity current deposits. Key surfaces are A = Basal Shear Surface (BSS), B = top SU1, C = top SU2, D = top SU3, E = top SU4, F = top SU5, G = top SU6. Modified from Olafiranye et al. (2013).

This study uses the seismic facies and seismic stratigraphic schemes of Olafiranye et al. (2013), the former based on the well-calibrated seismic facies scheme of Prather et al. (1998) for similar deposits at equivalent burial depths in the Gulf of Mexico. Olafiranye et al. (2013) subdivided the studied interval into six seismic units bounded by seven mappable seismic reflections (SU1-6; Fig. 4.5). Thickness or isochore maps (in two-way time) are generated for all of the seismic units to analyse their thickness variation. The seismic expression, geometry and distribution of these units, the characteristics of their bounding surfaces, and the interpreted depositional processes of the units are documented in detail in

Olafiranye et al. (2013) and summarised in Figure 4.6. The absence of well data and robust quantitative seismic parameters to calibrate seismic reflectivity in the study dataset preclude the possibility of directly interpreting lithology.

Fig. 4.6: Illustration of stratal architecture in the analyzed seismic units (SU1-SU6). Note the rugose nature of the “SU2b” sub- unit, and the wedge-shaped geometry and basal rotational plane of the “SU2c” sub-unit. SU2 is the most internally deformed of the seismic units in the succession. Modified from Olafiranye et al. (2013).

In addition, this study uses two seismic attributes,‘variance’ and ‘chaos’, to detect disruption of reflectors around pockmarks and other features. The variance attribute, also referred to as ‘edge’

(Bahorich and Farmer, 1995; Luo et al., 1996) or ‘coherence’ (Chopra and Marfurt, 2007; Suarez et al.,

2008), is calculated across the top surface of each seismic unit SU1 to SU6 (Fig. 4.5) and variance maps generated. The variance attribute is a quantitative measure of the coefficients of lateral amplitude changes

(Bahorich and Farmer, 1995; Chopra and Marfurt, (2007) which is calculated through normalized cross- correlation between adjacent traces in the same survey, thereby eliminating the variability of source- wavelet amplitude and phase, to quantify wave-form continuity (Bahorich and Farmer, 1995). Variance attribute calculation is sensitive to seismic textures and morphology (Chopra and Marfurt, 2008) and therefore it is amenable to discriminating structural heterogeneities. The chaos attribute is a measure of the lack of organisation in seismic reflections, whereby regions with low reflection consistency correspond to zones with chaotic signal patterns (Randen and Sønneland, 2005). The chaos attribute is assessed within a seismic unit by the use of iso-proportional slices that are generated with reference to the top and base surfaces of the unit. The chaos attribute can identify spatially confined and discontinuous seismic facies (e.g. Back et al., 2011). In the study dataset, the chaos attribute is applied to seismic unit

SU2 (Figs. 4.4, 4.5 and 4.6) in order to map MTDs that have been previously identified in cross-sections through this unit (Olafiranye et al., 2013).

Pockmark density was assessed by manual counting their occurrences on the variance maps of the top surfaces of the seismic units. Pockmark density at each surface has been calculated using a grid of 2 x

2 km squares overlain on the study area. The morphological parameters of the pockmarks (diameter, ellipticity, height) have also been measured at each surface (see Appendix for details). The average diameter of all pockmarks is also deduced for the different top surfaces or pockmark fields. Diameter and ellipticity have been measured for all pockmarks, but height only for those that can be resolved in cross- section. Measurements of pockmark diameter and ellipticity were taken from the edge of their lower variance rim, in order to account for the full diameter of the pockmarks. These morphological observations allow the individual pockmarks to be characterised and classified using a modified version of the scheme of Hovland et al. (2002) (Fig. 4.7). Modification of the existing classification scheme was

98 necessary because Hovland et al. (2002) intermixed pockmark sizes with their morphological characteristics in classifying pockmarks. The spatial distributions of pockmarks that have resolvable

Fig. 4.7: Classifications of (a) pockmark type based on their morphological and reflector characteristics, and (b) pockmark spatial distributions. Modified from Hovland et al. (2002). heights are characterised using the classification scheme of Hovland et al. (2002). Both “strings” of aligned pockmarks (denoted S1, S2, S3, and S4 in the following sections) and “complexes” of pockmarks

(denoted C1, C2, and C3 in the following sections) are recognised.

4.5 Results

4.5.1 Density and distribution of pockmarks

No pockmarks are developed at the top surfaces of the SU1 and SU2 units (Fig. 4.8a, b). Instead, high values of variance are noted only along fault traces, which are described later. Pockmarks are identified in the four upper mapped surfaces (top surfaces of SU3, SU4, SU5, and SU6) (Fig. 4.8c, d, e, f), where they are marked by high values of variance (cf. Fig. 4.7b).

Fig. 4.8: Variance maps of the top surfaces of (a-f) seismic units SU1-SU6. The red polygons enclose the area of pockmark occurrences on the top surfaces of SU3, SU4, SU5, and SU6. Note that some areas of anomalously high variance at the tops of SU1, SU2, and SU6 do not occur as circular features, and therefore do not represent pockmark development.

The density and pattern of distribution of pockmarks are observed to be similar on the top-SU3 and top-

SU4 surfaces (Fig. 4.9a, b). Their density and distribution on the top-SU5 and top-SU6 surfaces also

follow a similar pattern (Fig. 4.9c, d). Hence, these pairs of surfaces are described and interpreted

together below.

100

Fig. 4.9: Maps showing the spatial distribution and density of pockmarks at the top surfaces of (a-d) seismic units SU3-SU6. Maps illustrate the relationships between potential conduits for fluid flow and the occurrence of pockmarks. Note in particular the high density of pockmarks around faults and the downdip termination of MTDs in underlying unit SU2.

4.5.1.1 Top SU3 and Top SU4

Description: All the pockmarks on the top-SU3 and top-SU4 surfaces occur in the western depocentre where the maximum density of pockmarks ranges between 6 and 9 pockmark/km 2 (Fig. 4.9a, b). The average density of pockmarks on each surface is 1 pockmark/km 2 and the densest concentrations of

101 pockmarks are located above the trace of the down-dip termination of MTDs in the underlying SU2 unit

(Fig. 4.10a, b).

Fig. 4.10: Detailed variance maps illustrating the distribution of pockmarks developed at the (a) top-SU3 and (b) top-SU4 surfaces. S1 denotes a string of pockmarks. (c) Uninterpreted seismic and (d) interpreted geoseismic section (located in Fig. 4.10a, b) along a string of pockmarks.

Some anomalous features that superficially resemble pockmarks are noted within the eastern depocentre on variance maps of the top-SU3 and top-SU4 surfaces (Fig. 4.8c, d), but these features lack a circular form and resolvable height, which precludes their classification as pockmarks using the criteria applied in this study (Fig. 4.7a). Pockmarks on the top-SU3 and top-SU4 surfaces occur as a curvilinear

“string” on the variance maps of these surfaces (Fig. 4.10a, b).

Interpretation: The comparable density and spatial distribution of the pockmarks in these two deepest surfaces suggest that they were created by the same fluid expulsion pathways and, potentially, by the same events. The dense concentration of pockmarks in a curvilinear ”string” that is positioned directly above the trace of the down-dip termination of MTDs in underlying unit SU2 (Fig. 4.10a, b) establishes a developmental connection between the pockmarks and the underlying MTDs.

4.5.1.2 Top SU5 and Top SU6

Description: The majority of pockmarks on the top-SU5 and top-SU6 surfaces occur within the western depocentre with some minor occurrences upslope of the eastern depocentre (Figs. 4.8e- f, 4.9c- d). The

102 density of pockmarks on the top-SU5 surface ranges between 1 and 3 pockmarks/km 2 while the top-SU6 surface has an average density of 2 pockmarks/km 2. The densest occurrence of pockmarks on both surfaces is concentrated around fault traces within the western depocentre (Figs. 4.9c-d, 4.11a, 4.12a).

Fig. 4.11: Detailed variance map illustrating the distribution of pockmarks at the (a) top-SU5 surface. S1, S2, S3, and S4 denote strings of pockmarks, while C1, C2, and C3 denote complexes of pockmarks. (b) Uninterpreted seismic and (c) interpreted geoseismic section (located in Fig. 4.11a) showing fault systems in relation to pockmarks.

Some of the pockmarks on the top-SU5 and top-SU6 surfaces occur in four distinctive curvilinear

“strings” denoted S1, S2, S3, and S4, while others are scattered as “complexes” without any preferred orientation, denoted C1, C2, C3 (Figs. 4.11a, 4.12a).

Fig. 4.12: Detailed variance map illustrating the distribution of pockmarks at the (a) top-SU6 surface. S1, S2, S3, and S4 denote

103 strings of pockmarks, while C1, C2, and C3 denote complexes of pockmarks. (b) Uninterpreted seismic and (c) interpreted geoseismic section (located in Fig. 4.12a) showing fault systems in relation to pockmarks.

The C1 “complex” is developed in the eastern depocentre, while the C2 “complex” is located in the southwestern corner of the pockmark field (Figs. 4.11a, 4.12a). The C3 “complex” occurs between the S3 and S4 “strings” (Figs. 4.11a, 4.12a). Although the pockmarks on the two surfaces share similar spatial configurations, pockmarks are more prevalent on the top-SU6 surface than on the top-SU5 surface

(compare Figs. 4.11a and 4.12a). The scattered high variance features in the upslope part of the top-SU6 surface (Fig. 4.8f) are not interpreted as pockmarks, rather they are considered to represent facies changes due to paralic sub-environments on the upper slope.

Interpretation: The S1 “string” mimics the curvilinear position of the down-dip termination of MTDs in unit SU2 (Figs. 4.11a, 4.12a) and in intervening units SU3 and SU4 (Fig. 4.10a, b), suggesting that a similar geologic control created this string of pockmarks. The curvilinear trends of pockmark “strings”

S2, S3, and S4 mimic the position of fault traces, implying that the fault planes acted as migration pathways for subsurface fluid flow. The distributions of the C1, C2, and C3 pockmark “complexes” are not readily matched with underlying geological features. Permeable and porous carrier beds in the SU5 and SU6 units may have caused lateral movement away from fault planes that is reflected in the geometry of these complexes. The increased density of pockmarks on the top-SU6 surface compared to the top-SU5 surface can be attributed to decreased fault throw, which resulted in a weaker sealing capacity of the faults at top-SU6 level and which allowed enhanced fluid flow to the seafloor (cf. Hubscher and

Borowski, 2006; Gay et al., 2007, their Figs. 6 and 7).

4.5.2 Morphology of pockmarks

The morphology of pockmarks imaged at the top-SU3 and top-SU4 surfaces is similar, as is pockmark morphology at the top-SU5 and top-SU6 surfaces.

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4.5.2.1 Top SU3 and Top SU4

Description: The majority of pockmarks on the top-SU3 and top-SU4 surfaces occur as circular and elliptical features with anomalously high variance in the centre and a rim of relatively low variance (Fig.

4.10a, b), and are characterised as eyed pockmarks (cf. Fig. 4.7a). A few pockmarks occur without a high variance centre (Fig. 4.10a, b), and are characterised as normal pockmarks (cf. Fig. 4.7a). The average diameter of the pockmarks on the top-SU3 surface is between 110 and 280 m, while pockmarks on the top-SU4 surface have average diameters between 90 and 240 m (Figs. 4.13a-b, Appendix).

Fig. 4.13: Graphs of long axis versus short axis (cf. Fig. 4.7a) for pockmarks developed at the top surfaces of seismic units (a-d) SU3, SU4, SU5, and SU6. The points plot in close proximity to lines indicating circular plan-view geometries on all four surfaces.

Pockmarks on both surfaces have average ellipticity ratios of 1.2, but pockmarks developed on the top-

SU4 surfaces are closer to being circular than those on the top-SU3 surface (i.e. the former plot closer to the line indicating perfect circularity in Fig. 4.13a, b). Pockmarks developed on the two surfaces have comparable average heights of 16 m (Fig. 4.14a, b). Pockmark height increases as average diameter increases (Fig. 4.14a, b). Though some anomalous features that resemble pockmarks are noted within the eastern depocentre especially on the variance maps of Top SU3 and Top SU4 (Figs. 8c and d), a closer

105 dissection of these features shows that they lack any form of circularity or identifiable height which precludes them to be classified as pockmarks.

Fig. 4.14: Graphs of pockmark heights versus average diameter for the pockmarks developed at the top surfaces of seismic units (a-d) SU3, SU4, SU5, and SU6. Pockmark height increases as the average diameter of pockmarks increases.

Interpretation: The comparable diameter, height, and ellipticity of the pockmarks on the top-SU3 and top-

SU4 surfaces suggest that the pockmarks formed under similar conditions on both surfaces. The high variance in the centre of the eyed pockmarks can be related to remnant coarser material from the process of winnowing or from biological and authigenic precipitation of carbonates (Hovland et al., 2002). The seismic stratigraphic context of the pockmarks imply that the formative processes operated on the lower continental slope (Figs. 4.4, 4.10c-d). Pockmarks on both surfaces have average ellipticity ratios of 1.2, but pockmarks developed on the top-SU4 surfaces are closer to being circular than those on the top-SU3 surface (i.e. the former plot closer to the line indicating perfect circularity in Fig. 4.13a, b). Pockmarks developed on the two surfaces have comparable average heights of 16 m (Fig. 4.14a, b). Pockmark height increases as average diameter increases (Fig. 4.14a, b). Though some anomalous features that resemble pockmarks are noted within the eastern depocentre especially on the variance maps of Top SU3 and Top

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SU4 (Figs. 8c and d), a closer dissection of these features shows that they lack any form of circularity or identifiable height which precludes them to be classified as pockmarks.

4.5.2.2 Top SU5 and Top SU6

Description: On the top-SU5 surface, the average diameter of the pockmarks is between 60 and 210 m

(Fig. 4.13c, Appendix). The measured heights of all the pockmarks are between 5 and 30 m (Fig. 4.14c,

Appendix). The pockmarks on the top-SU6 surface have average diameters between 60 and 280 m and an average height of 12 m (Figs. 4.13c, 4.14c, Appendix). Both eyed and normal pockmarks occur on the two surfaces (Figs. 4.11a, 4.12a). The pockmarks exhibit circular and elliptical shapes with average ellipticity ratios of 1.3 and 1.5 respectively for the top-SU5 and top-SU6 surfaces. The majority of pockmarks have near circular geometries in plan view (i.e. they plot close to the line indicating perfect circularity in Fig. 4.13c, d). Pockmark height shows a moderate-to-strong positive corrrelation

(correlation coefficient, R 2 = 0.43-0.99) with average pockmark diameter (Fig. 4.14c, d). The scattered high variance features in the upslope part of Top SU6 (Fig. 4.8f) are not identified as pockmarks, rather they are considered to be increased facies changes of a paralic sub-environment.

Interpretation: The large average diameters of pockmarks imaged on the top-SU5 and top-SU6 surfaces are interpreted as the surface expression of cone-shaped pockmarks in cross-section that have propagated through the underlying units (Figs. 4.11b-c, 4.12b-c).

4.5.3 Thickness of Seismic Units

Isochore maps were generated for each seismic unit (Fig. 4.5) in order to identify spatial trends in thickness, and to compare these thickness trends to the development and distribution of pockmarks across the study area. Isochore maps are converted from time to depth using a velocity of 2000 m/s.

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Fig. 4.15: Depth-converted isochore maps illustrating the thickness of seismic units (a-f) SU1 to SU6. The white dotted polygons delimit areas of pockmark development on the top surfaces of units SU3, SU4, SU5, and SU6 (Fig. 4.8c-f). Note that seismic units are thickest within the eastern depocentre, whereas pockmarks occur within the western depocentre (Fig. 4.9).

4.5.3.1 Seismic Unit SU1

The maximum thickness of unit SU1 is about 95 m, in the eastern depocentre of the lower slope

(Fig. 4.15a). Elsewhere the unit is uniformly thin (<10 m) (Fig. 4.15a). The unit downlaps and onlaps onto the Middle-Pliocene basal surface (Fig. 4.6). Pockmarks are not developed on the top-SU1 surface

(Fig. 4.8a). The apparent lack of pockmarks on the top surface of the unit may be attributed to the absence of sufficiently high pore fluid pressure to cause catastrophic fluid expulsion through the overlying hemipelagic deposits, even in the thickest part of the unit.

4.5.3.2 Seismic Unit SU2

Unit SU2 exhibits a maximum thickness of 300 m within the eastern depocentre, and thins to c.

90 m within the western depocentre around the trace of the down-dip termination of an MTD on the lower slope (Fig. 4.15b). On the mid-slope, updip of the MTD pinchout, the thickness of the unit is fairly

108 uniform (c. 150 m), but it decreases to c. 70 m in the northwestern corner of the study area. Despite the large thickness of MTDs within unit SU2 in the western depocentre (Fig. 4.15b), there are no recognizable pockmarks on the top surface of the unit (Fig. 4.8b).

4.5.3.3 Seismic Unit SU3

The greatest thickness (up to 130 m) of unit SU3 occurs in three localised areas in the eastern depocentre (Fig. 4.15c). The unit thins to c. 30 m in the western depocentre. Upslope of the western and eastern depocentres on the lower slope, the thickness of unit SU3 is c. 60 m on the middle and upper slope, but it thins rapidly to <10 m on the shelf to the northeast (Fig. 4.15c). Thickness variations in the eastern depocentre are attributed to infilling of rugose topography at the top-SU2 surface (Fig. 4.8b). The three local thickness maxima abut the steep margins of MTDs in unit SU2, and record infilling and

“healing” of topographic breaks on the seafloor during deposition of unit SU3. Pockmarks on the top surface of unit SU3 are restricted to the western depocentre (Figs. 4.8c), where the unit is relatively thin

(white polygon in Fig. 4.15c). Thus, pockmark occurrence does not appear to be directly related to the thickness of unit SU3, which forms the immediate overburden of MTDs in unit SU2. It is possible that fluid released by dewatering of the thick sediment pile in the eastern depocentre developed sufficient momentum to force lateral fluid flow through sediments draping the intervening horst and into the western depocentre thus contributing to fluid flow and pockmark development here.

4.5.3.4 Seismic Unit SU4

Unit SU4 exhibits a local maximum thickness of 90 m within the eastern depocentre, and is relatively thin (c. 15 m) within the western depocentre. The unit gradually thins to <10 m on the upper slope towards the north. Pockmarks on the top surface of unit SU4 (Fig. 4.8d) occur in the western depocentre, where the unit is relatively thin (white polygon in Fig 4.16d), such that the thickness of unit

SU4 does not directly control the distribution of pockmarks. As for unit SU3, lateral migration of expelled fluid from the thicker sediment pile in the eastern depocentre may have occurred.

4.5.3.5 Seismic Unit SU5

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Unit SU5 reaches a thickness of c. 200 m in the eastern depocentre and c. 70 m in the western depocentre on the lower slope (Fig 4.16e). The unit thins to <10 m on the upper slope. Pockmarks on the top surface of unit SU5 are confined to the western depocentre (Fig. 4.8e), where unit SU5 is only c. 70 m thick (white polygon in Fig 4.16e). As in the underlying units, pockmark distribution does not correspond to unit thickness.

4.5.3.6 Seismic Unit SU6

SU6 is thickest within the eastern depocentre (c. 130 m) and reaches about 60 m within the western depocentre of the lower slope and on the middle slope (Fig. 4.15f). The unit is thinner (<10 m) on the upper slope and shelf. Pockmarks at the top surface of SU6 occur within the western depocentre (Fig.

4.8f), where unit SU6 is moderately thick (white polygon in Fig. 4.15f). Again, the thickest part of the unit occurs in the eastern depocentre, and does not coincide with the area of pockmark development in the western depocentre.

4.5.4 Fault distribution

Several normal faults with synthetic (downthrown to the SW) and antithetic (downthrown to the

NE) senses of displacement deform the Mid-Pliocene-to-Holocene succession of the study area (Figs. 4.3,

4.4). Synthetic faults have a general NW-SE strike and dip towards the SW (Fig. 4.3). Seventeen such faults are identified in the studied interval (labelled S1 to S17 in Fig. 4.3). The S1 fault is about 40 km in length and cuts across the entire study area to form the inboard limit of the eastern depocentre in the upper slope and shelf. It developed a maximum throw of c. 90 m on the base Mid-Pliocene surface (i.e. base of unit S1) around the Graben (Fig. 4.3a). However, the throw of the fault increases downwards through the underlying succession, and it reaches over 300 m at the base-Late Miocene surface (Fig.

4.3b). The fault terminates in the Early-Mid Miocene succession and is welded on the salt decollement that tops the pre-salt sequence. It has two genetically-related splays (S2, maximum throw 80 m, and S3, maximum throw 85m) that terminate within the Late Miocene succession (Fig. 4.3). The S4 and S5 faults each have maximum throws of about 40 m on the base Mid-Pliocene surface. S4 reaches to the top of the pre-salt sequence while S5 terminates within the Early-Mid Pliocene succession. The two faults, about 10

110 km each in length, mark the outboard limit of the horst in the lower slope (Fig. 4.3a). The S6 to S14 faults have an average length of about 20 km and average maximum throw of c. 15 m and are located upslope of the eastern depocentre. S15 to S17 are located upslope of the western depocentre and have average throw and length of 15 m and 5 km respectively.

Antithetic faults trend parallel to the synthetic faults (NW-SE) and dip towards the NE (Fig. 4.3).

Ten antithetic faults are identified in the study area (labelled A1 to A10 in Fig. 4.3) and are. The biggest of the faults is A1 which is approximately 30 km in length and developed a maximum throw of 55 m around the Graben at the base Mid-Pliocene surface (i.e. base of unit S1) (Fig. 4.3a). This fault marks the eastern extremity of the horst between the western and eastern depocentres on the lower slope and the outboard limit of the Graben on the upper slope. The A2 fault terminates laterally on A1 within the eastern depocentre and is of length c. 10 km and maximum throw of 30 m at the base Mid-Pliocene surface. The A3and A4 faults have lengths of 15 and 5 km respectively, and occur in the upper slope updip of the eastern depocentre. Their average maximum throw is 30 and 10 m, respectively. The A5 fault is approximately 5 km in length and occurs, probably as a conjugate of S2, in the lower slope within the western depocentre. It has a maximum throw of 35 m at its centre and dies vertically within the Late

Miocene succession (Fig. 4.3a). The A6 to A10 faults are located on the mid-slope updip of the western depocentre. Their length ranges between 2 and 10 m, and their maximum throw of range between 7 and

15 m.

Although both the synthetic and antithetic normal faults are prevalent in the area upslope of the eastern and western depocentres (Figs. 4.3a, 4.4b), pockmarks are not present here. In contrast, pockmarks occur downslope in the western depocentre where fewer faults are developed (Figs. 4.3, 4.9).

One explanation for this discrepancy between the distribution of faults and pockmarks is that the large throws of the S1, S2, S4, S5 and A5 faults in the western depocentre enabled them to be effective in transmitting fluid, because the deformation that develops large fault throws creates vertical migration

111 pathways from the deep fluid source into shallower permeable beds . The upslope faults on the other hand, have far smaller throws within the studied succession.

4.5.5 MTD distribution in unit SU2

MTDs in the Mid-Pliocene to Holocene succession of the study area are confined to seismic units

SU1 and SU2 (Fig. 4.5). Based on their seismic facies character, they are interpreted to comprise a range of en masse deposits which include slides, slumps, and debris flows with embedded mega-clasts that were emplaced in the lower slope within a “background” of hemipelagites (Olafiranye et al., 2013). The spatial and temporal distribution of MTDs in unit SU2 is assessed through three maps of the “chaos” attribute taken from iso-proportional slices through the lower, middle, and upper parts of the unit (Fig. 4.16).

Fig. 4.16: Maps of (a-c) the chaos seismic attribute, and (d-f) interpreted deposits from iso-proportional slices taken at three levels within seismic unit SU2: (a, d) lower slice, taken at 25% thickness from the base of the unit; (b, e) middle slice, taken at 50% thickness from the base of the unit; and (c, f) lower slice, taken at 75% thickness from the base of the unit.

In the lower part of the unit (corresponding approximately to sub-unit SU2a, Fig. 4.5), c. 80% of the

MTDs are distributed within the eastern depocentre with c. 20% occurring on the horst and in the western

112 depocentre (Fig. 4.16a, d). In the middle part of the unit (corresponding approximately to sub-unit SU2b,

Fig. 4.5), MTDs dominantly occur within the eastern depocentre and on the horst (Fig. 4.16b, e). In addition, the northeastern corner of the upper slope, updip of the eastern depocentre, is characterised by a chaotic seismic facies in the lower and middle parts of unit SU2 (Fig. 4.16a, b). Because it is typified by continuous reflections (Figs. 4.17, 4.18), this chaotic seismic facies is not interpreted as MTDs; rather it is interpreted as a progradational wedge, possibly comprising turbidites, on the upper slope. In the upper part of unit SU2 (corresponding approximately to sub-unit SU2c, Fig. 4.5), MTDs occur mainly on the lower and mid slope, within and updip of the eastern depocentre (Fig. 4.16c, f). Pockmarks are developed predominantly within the western depocentre (Fig. 4.8c-f), where there is only a minor occurrence of

MTDs in unit SU2 (Fig. 4.16a). Nonetheless, fluid expelled by dewatering of the MTDs may have attained sufficient hydraulic head to flow vertically from the down-dip pinchout of the MTDs and to develop pockmarks in the overlying seismic units.

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Fig. 4.17: (a) Uninterpreted seismic and (b) interpreted geoseismic sections of seismic facies in the area upslope of the eastern depocentre (Fig. 4.16), aligned along depositional dip. (a) Depositional strike seismic section. Note that seismic reflections within unit SU2 are continuous.

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Fig. 4.18: (a) Uninterpreted seismic and (b) interpreted geoseismic sections of seismic facies in the area upslope of the eastern depocentre (Fig. 4.16), aligned along depositional strike. Note that seismic reflections within unit SU2 are continuous.

4.6 Discussion: Sources of fluid and conduits for the development of pockmarks

Based on the analysis of the available 3D seismic data two possible sources of fluid for the pockmarks in the Mid-Pliocene to Holocene succession are identified. These two fluid types have an implicit association with different migration conduits, which are controlled by structural and depositional features.

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4.6.1 Water expelled from MTDs and associated sediment gravity flows

MTDs and associated sediment gravity flows in unit SU2 are a potential source of the fluids that may have lead to the development of pockmarks at shallower stratigraphic levels. MTDs and sediment gravity flows contain a large volume of water when first deposited, prior to compaction.

The estimated volume of the sediment which has been removed to cause the well developed pockmarks (i.e. pockmarks having measurable heights) is approximately 0.22 km 3 (see appendix).

However, the total volume of the removed sediments should be in excess of the above volume given that there are many other pockmarks with heights that are poorly imaged. On the other hand, the maximum volume of the fluid expelled from the MTDs in SU2 is estimated to be between 7 and 10 km 3 (see Fig.

4.16 for areal distribution of MTDs and Scherer (1987) for porosity of Pliocene sandstones). The volume of removed sediments that led to the development of the pockmarks is thus between 0.2 and 0.3% of the expelled fluid from the MTDs. The significant disparity in these volumes where the potential volume of the ejected fluid from the MTDs outweighs the volume of the removed sediments suggests that the fluid expelled from the MTDs is a plausible source of fluid for the development of the mid Pliocene-Holocene pockmarks.

Pockmark “string” S1 is present at the top surfaces of units SU3, SU4, SU5 and SU6 (Figs. 4.10a- b, 4.11a, and 4.12a), and its position coincides with the downdip termination of MTDs in unit SU2 (Fig.

4.8). Although the maximum thickness of unit SU2 (c. 300 m), which contains abundant MTDs, is located within the eastern depocentre (Fig. 4.15b), the thickness of the unit where it directly underlies pockmark

“string” S1 (c. 90 m) was potentially sufficient to generate the pore fluid necessary for the creation of the pockmarks through lateral migration from the eastern depocentre. Although the timing of fluid expulsion is difficult to resolved, the occurrence of pockmark “string” S1 in the same vertical position at the top of units SU3 through SU4, SU5, and SU6 (Figs. 4.10a-b, 4.11a, 4.12a) is a strong indication that fluid expulsion from the MTDs in unit SU2 occurred as a singular event which post-dated the deposition of the entire mid Pliocene-Holocene interval. Alternatively, multiple fluid expulsion events may have followed the same pathway during deposition of units SU3 to SU6. This interpretation implies that overpressured

116 water flowed laterally through the porous and permeable MTDs until their down-dip termination was reached, and the water was expelled vertically. Similar processes have been documented at the margins of channelised turbiditic sandbodies in the Lower Congo Basin (Gay et al., 2003) and elsewhere (Sun et al.,

2011).

Although the identified MTDs is thickest within the eastern depocentre (Fig. 4.16), the dynamic depositional nature of the lower slope around this area may have created a situation where the regime is not capable of preserving the continuous or pulsating fluid flushing events required to develop pockmarks. Such destructional slope systems are characterized by local to sub-regional shelf margin and upper slope retreat with consequent development of a significant bounding surface (composite slide plane, canyon-cut surface, or regional mass-wasting surface (Galloway, 2002).

4.6.2 Hydrocarbons from deep-seated source rocks and breached reservoirs

The presence of normal faults of large throw that deformed the studied Mid-Pliocene-to-Holocene succession and deeper petroliferous Miocene strata (Fig. 4.3) suggest that the pockmarks may have formed by the primary expulsion and secondary migration of hydrocarbons from Miocene source rocks.

This inference is supported by the work of Andresen et al. (2011), which documents hydrocarbon plumbing systems that combine Miocene source rocks to the development of pockmarks and associated features such as gas chimneys and high reflectivity zones in offshore Angola. Pockmarks are rarely imaged along faults on the top surfaces of Pliocene units SU1 to SU4 (Figs 4.8a-d, 4.9a-b), possibly because many faults have large throws at these levels and may be sealing. Although many reports (e.g.

Walgenwitz et al., 1990; Cole et al., 2000; Cartwright et al., 2007) ascribe the timing of hydrocarbon migration offshore Angola to the Miocene and Andresen et al. (2011) attributed fluid escape to palaeo- pockmarks of Pliocene-Pleistocene time. Due to the lack of pockmarks in Pliocene units SU1 to SU4, this current analysis points to an indication of a Pliocene hydrocarbon migration events in the Lower Congo

Basin of Angola.

Pockmarks are abundant along faults imaged at the top surfaces of units SU5 and SU6, where faults have smaller throws (“strings” S2 to S4 in Figs. 4.11, 4.12). The identification of surface

117 hydrocarbons in the Lower Congo Basin (e.g. Gay et al., 2006; Gay et al., 2007; Olu-Le Roy et al., 2007) suggests that fluids generated by deeply buried source rocks are reaching the surface at the present day and may be capable of generating pockmarks. Although there is no evidence for remigration of hydrocarbons from breached reservoirs to form pockmarks in the study dataset, this process is reported in the Lower Congo Basin by Andresen et al., (2011).

The close proximity of faults of large throw to the MTDs in unit 2 implies a strong possibility of mixing of the two sources of fluid identified above. This study therefore concurs with Gay et al. (2003) that the expected nature of fluids escaping from seafloor pockmarks should be a mixture of interstitial water (expelled from the MTDs in unit 2), shallow biogenic gases (produced by bacterial degradation of organic matter) and thermogenic gases or oil from deeply buried source rocks and/or breached reservoirs.

The pockmark “complexes” C1, C2, and C3 resolved at the top surfaces of units SU5 and SU6 can be attributed to the lateral movement of expelled fluid along thin carrier beds, which are below seismic resolution, within the SU5 and SU6 units. Smaller sized pockmarks, with diameters less than 50 m, may develop in this way via expulsion of fluid from sediment piles that are only 5 m thick (Josenhans et al.,

1978; Nelson et al., 1979; Rogers et al., 2006, their Fig. 9) if a sufficiently large volume of fluid is available and sufficient overburden is deposited to generate overpressure and trigger pore fluid expulsion.

Apart from the fact that the eastern depocentre (Fig. 4.19) in the lower slope is essentially a dynamic depositional regime which may have inhibited the development of pockmarks (see section

4.5.3.2) during the mid Pliocene-Holocene, the absence of pockmarks in this area may also be an indication that deep-seated hydrocarbon source rocks is the sole supplier, through the large throw normal faults, of the fluid which created and concentrated pockmarks within the western depocentre.

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Fig. 4.19: Model for the development of pockmarks in the mid-Pliocene to Holocene succession in Block 4, Lower Congo Basin, offshore Angola. A sequence of three events is interpreted: (a) MTDs and sediment gravity flows in unit SU2 are subjected to overburden pressure which causes the development of hydraulic head around their downdip termination; (b) a first set of pockmarks is developed in units SU3 and SU4 around the trace of the downdip termination of the MTDs in unit SU2, by the vertical expulsion of overpressured water. During this time, the development of pockmarks in SU5 and SU6 is contemporaneous with SU3 and SU4. Alternatively, the development of pockmarks above the downdip termination of the MTDs in unit SU2 is sustained during deposition of units SU5 and SU6, and additional pockmarks are developed by the movement of hydrocarbons along faults that intersect deeply buried Miocene source rocks and/or breached reservoirs, and by movement of these fluids along thin carrier beds in units SU5 and SU6.

4.7 Conclusions

1. The mid Pliocene-Holocene succession in Block 4, offshore Angola was analysed for the

development and spatial distribution of pockmarks with the aid of a good quality 3D seismic data.

Two key attributes -variance and chaos- allowed for surficial characterization of pockmarks and

internal sedimentology of the potential sources of pockmark-forming fluids. The development of

pockmarks in relation to the distributions of MTDs and faults was also investigated.

2. Normal and eyed pockmarks are identified on the four interpreted pockmark-bearing surfaces

with similar morphologies of circular and elliptical shapes of near-perfect circularity. Pockmark

density is high around fault tips and increases from the deeper surfaces to the shallower

stratigraphic levels with maximum density 1 pockmark/km 2 and 2 pockmark/km 2 respectively.

Pockmarks are distributed preferentially as curvilinear strings along structural constraints and as

random complexes as a consequence of carrier-bed fluid fluxing.

3. Whereas the identified MTDs are predominantly located within the eastern depocentre in the

lower slope of the study area, the vast majority of pockmark occurrence is within the western

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depocentre where they are located along the fault traces and the downdip limit of the MTDs. In

addition, the location of pockmarks within the western depocentre is in contrast to the

development of the thicker part of the units within the eastern depocentre across the mid

Pliocene-Holocene succession.

4. The potential sources of fluids that created the pockmarks are the MTDs and their associated

pelagic and hemipelagic carrier beds. In addition, the deep-seated Miocene source rocks in the

Lower Congo Basin are thought to have contributed expelled thermogenic fluids for the

development of the pockmarks. In both fluid sources, overpressure is identified as one plausible

mechanism of fluid expulsion while hydrocarbon remigration from breached reservoirs is also

suggested as a mode of fluid migration. The array of normal synthetic and antithetic faults that

deformed the mid Pliocene-Holocene strata acted as the conduits for the migration of the expelled

fluids.

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Chapter Five

Conclusions and future work

5.1 Conclusions

The major conclusions from the investigation of the shallow interval (1500 msecs) of high quality

760 km 2 3D reflection seismic data of the mid-Pliocene-Holocene succession in the Lower Congo Basin, offshore Angola are summarized below.

5.1.1 Seismic facies, stratigraphic architecture and bathymetric controls on upper slope depositional

patterns

Two major depocentres separated by an intraslope horst are formed due to a series of northwest- southeast striking, thin-skinned, salt-detached normal faults which are developed on the upper slope of the study area in Block 4, within the Lower Congo Basin, offshore Angola. In this area, three main seismic packages are recognised and are sub-divided into six seismic units (SUs) based on seismic facies transitions, and the geometry and regional extent of their bounding surfaces. These seismic facies record depositional patterns within the stratigraphic succession of the submarine slope which is imaged by the seismic dataset in this area. The initial stages of deposition occurred when the slope was above grade and dominated by degradational or erosional processes where deposition at this stage was characterized by the emplacement of a series of MTDs that were up to 250 m thick and at least 100 km 2 in areal extent. The intra-slope horst, which formed a prominent at-seabed structure at this time largely controlled the downslope translation, and hence thickness and distribution of these MTDs across the two depocentres.

The MTDs deposition was associated with the entrainment of substantial volumes of substrate and the formation of significant depositional relief along their upper surfaces which consequently controlled large-magnitude and abrupt changes in thickness in the overlying strata. During the latter stages of deposition, upon the attainment of ‘grade’, depositional processes pre-dominated and was characterized by either turbidity currents and/or suspension fallout in which the intra-slope horst, which had been

121

‘healed’ by this time, played little or no role during the latter stages of slope margin development. This aspect of the study therefore highlights the role that tectonically- and depositionally controlled bathymetric changes can have on accommodation development on the upper slope, and the manner in which this controls subsequent depositional patterns. In addition, both tectonic and depositional accommodation provide a mechanism for capturing and trapping reservoir-grade sediments in an upper slope setting that conventionally are modelled as areas of coarse-grained sediment bypass to the lower slope and basin floor.

5.1.2 Spatial distribution and development of pockmarks and controls on their evolution

Using two key seismic attributes, variance and chaos, extracted from the 3D seismic data, surficial characterization of pockmarks and internal sedimentology of the potential sources of pockmark-forming fluids were analysed. Furthermore, the development of pockmarks in relation to the distributions of

MTDs and faults was also investigated. On the four interpreted pockmark-bearing surfaces, ‘normal and eyed pockmarks’ are identified with similar morphologies of circular and elliptical shapes of near-perfect circularity. The density of pockmarks around fault tips is high and increases from the deeper surfaces to the shallower stratigraphic levels. Along structural constraints, pockmarks are distributed preferentially as curvilinear strings, but as random complexes as a result of carrier-bed fluid fluxing. Although the identified MTDs are predominantly located within the eastern depocentre in the lower slope of the study area, the majority of pockmark occurrence is within the western depocentre where they are located along the fault traces and the downdip limit of the MTDs. Likewise, the location of pockmarks within the western depocentre is in contrast to the development of the thicker part of the units within the eastern depocentre across the mid Pliocene-Holocene succession. The MTDs and their associated pelagic and hemipelagic carrier beds as well as the deep-seated Miocene source rocks in the Lower Congo Basin are identified as the potential sources of fluids that created the pockmarks. Overpressure is identified as one plausible mechanism of fluid expulsion while hydrocarbon remigration from breached reservoirs is also

122 suggested as a mode of fluid migration in the two identified fluid sources. The conduits for the migration of the expelled fluids are the array of normal faults that deformed the mid Pliocene-Holocene successions.

5.2 Future Work

The 3D seismic analysis reported in this thesis can be further developed through the integration and analysis of well data, relevant field analogues, and geochemical data.

5.2.1 Seismic inversion and forward modelling

The availability of well data including seismic velocities within the study area should enable seismic forward modelling to generate acoustic characteristics by transforming quantitative rock-property to synthetic seismic reflection which can be linked to the seismic facies in the 3D seismic volume (e.g.

Prather et al, 1998, Armitage and Stright, 2010; Bakke et al, 2008). This modelling allows for ground- truthing of interpreted seismic facies and for a more robust lithological interpretation to be made to reduce the overall uncertainty in the interpretation of seismic facies.

5.2.2 Linking sediment gravity flows stratigraphic succession and eustatic sea-level curve

Well data including wireline logs, biostratigraphic bugs and picks are important for constraining the identified seismic units (SUs) to specific portions of the eustatic sea-level curve or a locally- developed relative sea-level curve such that the SUs are properly tied to the genetic stratigraphy of the

Lower Congo Basin, offshore Angola. This would ensure there is improvement in the interpretation of the depositional environments of identified seismic units. In relation to this, the practice of super-imposing the Exxon sequence stratigraphic model, which speculate that sea-level falls may expose the shelf and triggered fluvial degradation of the shelf edge or upper slope and emplacement of MTDs, should be applied with caution particularly given that a number of other factors (section 2.2) may initiate sediment gravity flows.

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5.2.3 Field analysis of relevant analogues

Although sediment gravity flows, especially MTDs, are better analysed, because of their large dimensions, with the aid of large seismic datasets, field outcrops are vital in understanding some slide initiating processes (e.g. Van der Merwe et al., 2009) and subseismic scale kinematic elements that are often associated with them. Closer collaboration with other research groups, for example the StratGroup at the University of Liverpool where there is enormous field and outcrop work going on on deepwater tectono-stratigraphy can provide the necessary data required to fully understand sediment gravity flows processes and products. Some relevant work has been done (e.g. Van der Merwe et al., 2009) in analyzing the large scale Mass Transport Deposits of the Permian Vischkuil Formation (Karoo Basin, South Africa) in which outcrop investigations have been carried out to add an integrated depositional and deformational model to the mechanisms generating widespread syn-sedimentary deformation in Mass Transport

Deposits.

5.2.4 Geochemical analysis of fluid samples

Whereas slope instability and pockmarks could be related to gas hydrate disassociation, we have no direct seismic evidence for the presence of an extinct or active gas hydrate zone in this area.

Furthermore, pockmarks can form due to fluid flux associated with a range of processes as they are not exclusively associated with one source or the other. To minimize the ambiguity surrounding fluid source, geochemical analysis involving fluid typing and finger-printing of fluid samples taken in the vicinity of the pockmark fields within the Lower Congo Basin is important to pinpoint the actual sources of the fluids responsible for the development of the pockmarks. Linking pockmark development with identified geologic features and deducing potential source of fluid (e.g. Gay et al, 2006; Gay et al., 2007; Andressen and Huuse, 2011) can be substantiated with the analysis of in-situ fluid samples. This can also enhance exploration strategy and the result of basin modelling to improve the result of oil and gas exploration and development.

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Acknowledgements

Recognition and appreciation is due Chris and Dave for their major input in developing the main ideas in this project. Immense thanks go to Gary who helped to streamline the scope of the work to the end.

Thanks to Canadian Natural Resources (CNR), Seismic-Micro Technology (SMT) and

Schlumberger for respectively providing the seismic data, Kingdom Suite and Petrel software used for this study.

I am especially grateful to the Petroleum Technology Development Fund (PTDF Nigeria) for their financial support.

125

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Appendices

6.1 Morphological table of Top SU3

Diameter (m) Avg. Height Ellipticity Depth (msec) S/N Long Short Diameter (msec) ratio Base Top

163

axis axis (m) S1.1 207 172 190 15 1.2 1217 1202 S1.2 244 177 211 20 1.4 1222 1202 S1.3 116 97 107 15 1.2 1220 1205 S1.4 202 162 182 16 1.2 1198 1182 S1.5 143 117 130 16 1.2 1227 1211 S1.6 129 117 123 14 1.1 1223 1209 S1.7 331 237 284 20 1.4 1230 1210 S1.8 116 110 113 13 1.1 1206 1193 S1.9 166 158 162 17 1.1 1207 1190 S1.10 153 136 145 14 1.1 1189 1175 S1.11 141 105 123 15 1.3 1186 1171 S1.12 PR PR PR PR PR PR PR S1.13 217 189 203 15 1.1 1171 1156 PR= Poorly Resolved

6.2 Morphological table of Top SU4 Diameter (m) Avg. Height Ellipticity Depth (msec) S/N Long Short Diameter (msec) ratio Base Top axis axis (m)

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S1.1 179 179 179 19 1 1183 1164 S1.2 221 195 208 20 1.1 1190 1170 S1.3 138 118 128 14 1.2 1187 1173 S1.4 187 156 172 18 1.2 1198 1180 S1.5 88 88 88 11 1.0 1189 1178 S1.6 114 90 102 10 1.3 1187 1177 S1.7 251 230 241 26 1.1 1199 1173 S1.8 139 118 129 9 1.2 1173 1164 S1.9 187 157 172 13 1.2 1176 1163 S1.10 181 133 157 14 1.4 1167 1153 S1.11 137 133 135 11 1.0 1160 1149 S1.12 PR PR PR PR PR PR PR S1.13 214 172 193 24 1.2 1155 1131 PR= Poorly Resolved

6.3 Morphological table of Top SU5 pockmarks Diameter (m) Ave. Ellipticity Depth (msec) Height S/N Long Short Diameter ratio Base Top (msec) axis axis (m)

165

S1.1 PR PR PR PR PR PR PR S1.2 PR PR PR PR PR PR PR S1.3 PR PR PR PR PR PR PR S1.4 PR PR PR PR PR PR PR S1.5 86 83 85 1.0 1045 1040 5 S1.6 PR PR PR PR PR PR PR S1.7 266 229 248 1.2 1075 1043 32 S1.8 136 121 129 1.1 1048 1038 10 S1.9 180 171 176 1.1 1069 1040 29 S1.10 170 161 166 1.1 1052 1030 22 S1.11 133 57 95 2.3 1048 1031 17 S1.12 143 108 126 1.3 1028 1016 12 S1.13 167 159 163 1.1 1034 1023 11 PR= Poorly Resolved

Diameter (m) Ave. Ellipticity Depth (msec) Height S/N Long Short Diameter ratio Base Top (msec) axis axis (m) S2.1 218 205 212 1.1 869 845 24 S2.2 111 104 108 1.1 942 930 12 S2.3 116 116 116 1.0 948 938 10 S2.4 82 71 77 1.2 983 967 16 S2.5 91 83 87 1.1 1010 999 11

Diameter (m) Ave. Ellipticity Depth (msec) Height S/N Long Short Diameter ratio Base Top (msec) axis axis (m) S3.1 PR PR PR PR PR PR PR

166

S3.2 PR PR PR PR PR PR PR S3.3 PR PR PR PR PR PR PR S3.4 PR PR PR PR PR PR PR S3.5 117 113 115 1.0 977 960 17 S3.6 118 106 112 1.1 973 962 11 S3.7 PR PR PR PR PR PR PR S3.8 120 116 118 1.0 1001 981 20 S3.9 139 92 116 1.5 1015 999 16 S3.10 118 82 100 1.4 1015 1000 15 S3.11 200 148 174 1.4 1033 1010 23 S3.12 150 79 115 1.9 1024 1010 14 S3.13 176 127 152 1.4 1053 1024 29 S3.14 170 152 161 1.1 1056 1034 22 S3.15 PR PR PR PR PR PR PR S3.16 127 124 126 1.0 1073 1059 14 S3.17 PR PR PR PR PR PR PR S3.18 PR PR PR PR PR PR PR S3.19 198 185 192 1.1 1113 1088 25 S3.20 148 123 136 1.2 1108 1092 16 S3.21 123 97 110 1.3 1108 1095 13 S3.22 97 95 96 1.0 1113 1098 15 S3.23 PR PR PR PR PR PR PR S3.24 183 103 143 1.8 1156 1144 12 PR= Poorly Resolved

Diameter (m) Ave. Ellipticity Depth (msec) Height S/N Long Short Diameter ratio Base Top (msec) axis axis (m)

167

S4.1 PR PR PR PR PR PR PR S4.2 PR PR PR PR PR PR PR S4.3 PR PR PR PR PR PR PR S4.4 PR PR PR PR PR PR PR S4.5 184 131 158 1.4 1031 1020 11 S4.6 PR PR PR PR PR PR PR S4.7 PR PR PR PR PR PR PR S4.8 PR PR PR PR PR PR PR PR= Poorly Resolved

Diameter (m) Ave. Ellipticity Depth (msec) Height S/N Long Short Diameter ratio Base Top (msec)

168

axis axis (m) C1.1 PR PR PR PR PR PR PR C1.2 152 82 117 1.9 890 879 11 C1.3 110 91 101 1.2 905 895 10 C1.4 106 101 104 1.0 907 896 11 C1.5 PR PR PR PR PR PR PR C1.6 165 113 139 1.5 930 915 15 C1.7 96 35 66 2.7 935 919 16 C1.8 PR PR PR PR PR PR PR C1.9 PR PR PR PR PR PR PR C1.10 PR PR PR PR PR PR PR C1.11 PR PR PR PR PR PR PR C1.12 PR PR PR PR PR PR PR C1.13 PR PR PR PR PR PR PR C1.14 PR PR PR PR PR PR PR C1.15 PR PR PR PR PR PR PR C1.16 PR PR PR PR PR PR PR C1.17 PR PR PR PR PR PR PR C1.18 PR PR PR PR PR PR PR C1.19 128 71 100 1.8 911 898 13 C1.20 PR PR PR PR PR PR PR C1.21 117 130 124 0.9 905 895 10 C1.22 PR PR PR PR PR PR PR C1.23 115 84 100 1.4 892 880 12 C1.24 PR PR PR PR PR PR PR C1.25 PR PR PR PR PR PR PR C1.26 PR PR PR PR PR PR PR C1.27 86 70 78 1.2 931 918 13 C1.28 PR PR PR PR PR PR PR PR= Poorly Resolved

Diameter (m) Ave. Ellipticity Depth (msec) Height S/N Long Short Diameter ratio Base Top (msec)

169

axis axis (m) C2.1 131 100 116 1.3 1088 1079 9 C2.2 92 48 70 1.9 1061 1052 9 C2.3 PR PR PR PR PR PR PR C2.4 PR PR PR PR PR PR PR C2.5 128 109 119 1.2 1078 1068 10 C2.6 140 111 126 1.3 1092 1082 10 PR= Poorly Resolved

Diameter (m) Ave. Ellipticity Depth (msec) Height S/N Long Short Diameter ratio Base Top (msec) axis axis (m) C3.1 156 102 129 1.5 952 943 9 C3.2 153 122 138 1.3 958 951 7 C3.3 PR PR PR PR PR PR PR C3.4 PR PR PR PR PR PR PR C3.5 81 66 74 1.2 1000 987 13 C3.6 PR PR PR PR PR PR PR C3.7 PR PR PR PR PR PR PR C3.8 PR PR PR PR PR PR PR C3.9 130 99 115 1.3 1031 1021 10 C3.10 99 50 75 2.0 1053 1039 14 C3.11 168 141 155 1.2 1064 1044 20 C3.12 97 75 86 1.3 1043 1027 16 PR= Poorly Resolved

7.4 Morphological table of Top SU6 pockmarks

170

Diameter (m) Avg. Ellipticity Depth (msecs) Height S/N Long Short Diameter ratio Base Top (msecs) axis axis (m) S1.1 PD PD PD PD PD PD PD S1.2 PD PD PD PD PD PD PD S1.3 PD PD PD PD PD PD PD S1.4 PD PD PD PD PD PD PD S1.5 PD PD PD PD PD PD PD S1.6 PD PD PD PD PD PD PD S1.7 286 272 279 1.1 1075 1046 29 S1.8 124 89 107 1.4 1048 1039 9 S1.9 221 195 208 1.1 1069 1040 29 S1.10 200 178 189 1.1 1054 1033 21 S1.11 136 99 118 1.4 1048 1035 13 S1.12 115 91 103 1.3 1027 1018 9 S1.13 175 152 164 1.2 1033 1025 8 PR= Poorly Resolved

Diameter (m) Avg. Ellipticity Depth (msecs) Height S/N Long Short Diameter ratio Base Top (msecs) axis axis (m) S2.1 275 236 256 1.2 870 845 25 S2.2 157 113 135 1.4 943 930 13 S2.3 126 84 105 1.5 948 938 10 S2.4 189 140 165 1.4 983 967 16 S2.5 106 72 89 1.5 1011 1001 10 PR= Poorly Resolved

171

Diameter (m) Avg. Ellipticity Depth (msecs) Height S/N Long Short Diameter ratio Base Top (msecs) axis axis (m) S3.1 122 87 105 1.4 922 914 8 S3.2 106 76 91 1.4 925 923 2 S3.3 128 89 109 1.4 947 940 7 S3.4 80 56 68 1.4 952 946 6 S3.5 184 158 171 1.2 978 962 16 S3.6 150 108 129 1.4 975 963 12 S3.7 108 47 78 2.3 975 965 10 S3.8 202 172 187 1.2 1001 983 18 S3.9 175 142 159 1.2 1015 1002 13 S3.10 180 136 158 1.3 1015 1002 13 S3.11 207 167 187 1.2 1033 1011 22 S3.12 183 125 154 1.5 1024 1012 12 S3.13 214 208 211 1.0 1052 1026 26 S3.14 226 206 216 1.1 1057 1034 23 S3.15 128 117 123 1.1 1060 1047 13 S3.16 139 125 132 1.1 1072 1060 12 S3.17 115 98 107 1.2 1084 1077 7 S3.18 58 55 57 1.1 1088 1080 8 S3.19 208 197 203 1.1 1113 1088 25 S3.20 139 128 134 1.1 1108 1088 20 S3.21 154 147 151 1.0 1108 1092 16 S3.22 96 89 93 1.1 1113 1097 16 S3.23 120 97 109 1.2 1113 1099 14 S3.24 124 79 102 1.6 1118 1108 10 PR= Poorly Resolved

172

Diameter (m) Avg. Ellipticity Depth (msecs) Height S/N Long Short Diameter ratio Base Top (msecs) axis axis (m) S4.1 125 125 125 1.0 977 972 5 S4.2 87 75 81 1.2 993 987 6 S4.3 108 75 92 1.4 1004 999 5 S4.4 103 102 103 1.0 1012 1006 6 S4.5 167 122 145 1.4 1031 1019 12 S4.6 98 56 77 1.8 1048 1041 7 S4.7 121 65 93 1.9 1063 1054 9 S4.8 118 71 95 1.7 1096 1087 9 PR= Poorly Resolved

173

Diameter (m) Avg. Ellipticity Depth (msecs) Height S/N Long Short Diameter ratio Base Top (msecs) axis axis (m) C1.1 154 102 128 1.5 856 852 4 C1.2 141 130 136 1.1 890 879 11 C1.3 145 139 142 1.0 905 896 9 C1.4 116 85 101 1.4 908 897 11 C1.5 80 55 68 1.5 908 899 9 C1.6 169 146 158 1.2 931 915 16 C1.7 112 98 105 1.1 934 919 15 C1.8 173 168 171 1.0 957 941 16 C1.9 120 58 89 2.1 956 949 7 C1.10 134 82 108 1.6 958 952 6 C1.11 102 55 79 1.9 969 963 6 C1.12 129 116 123 1.1 980 963 17 C1.13 112 102 107 1.1 957 947 10 C1.14 83 60 72 1.4 947 935 12 C1.15 108 103 106 1.0 944 931 13 C1.16 146 100 123 1.5 932 922 10 C1.17 134 108 121 1.2 935 924 11 C1.18 106 84 95 1.3 925 915 10 C1.19 137 125 131 1.1 910 897 13 C1.20 126 88 107 1.4 892 879 13 C1.21 166 137 152 1.2 906 893 13 C1.22 102 94 98 1.1 876 867 9 C1.23 119 74 97 1.6 892 880 12 C1.24 116 98 107 1.2 900 892 8 C1.25 148 136 142 1.1 925 911 14 C1.26 188 164 176 1.1 930 924 6 C1.27 120 108 114 1.1 931 917 14 C1.28 112 62 87 1.8 934 924 10 PR= Poorly Resolved

174

Diameter (m) Avg. Ellipticity Depth (msecs) Height S/N Long Short Diameter ratio Base Top (msecs) axis axis (m) C2.1 122 72 97 1.7 1088 1080 8 C2.2 107 86 97 1.2 1061 1052 9 C2.3 85 43 64 2.0 1057 1051 6 C2.4 87 86 87 1.0 1067 1058 9 C2.5 107 71 89 1.5 1077 1067 10 C2.6 97 72 85 1.3 1092 1083 9 PR= Poorly Resolved

Diameter (m) Avg. Ellipticity Depth (msecs) Height S/N Long Short Diameter ratio Base Top (msecs) axis axis (m) C3.1 107 67 87 1.6 952 941 11 C3.2 116 67 92 1.7 957 951 6 C3.3 122 69 96 1.8 987 970 17 C3.4 91 78 85 1.2 990 983 7 C3.5 114 94 104 1.2 999 986 13 C3.6 138 116 127 1.2 1016 1007 9 C3.7 82 53 68 1.5 1011 1003 8 C3.8 113 69 91 1.6 1020 1017 3 C3.9 114 104 109 1.1 1031 1022 9 C3.10 142 136 139 1.0 1052 1038 14 C3.11 175 158 167 1.1 1063 1044 19 C3.12 134 122 128 1.1 1043 1027 16 PR= Poorly Resolved

175