Mechanisms of Sediment Movement from Topset To

Home , Progradation

MECHANISMS OF SEDIMENT MOVEMENT FROM TOPSET TO TOESET IN THE CENOZOIC CLINOFORMING NW AUSTRALIAN MARGIN, AND IMPLICATIONS FOR RESERVOIR DEVELOPMENT, A SEISMIC GEOMORPHOLOGY STUDY

by Alden Griffin

A thesis submitted to the Faculty and the Board of Trustees of the Colorado School of Mines in partial fulfillment of the requirements for the degree of Master of Science of (Geology).

Golden, Colorado

Date ______

Signed: ______Alden A. Griffin

Signed: ______Dr. Lesli J. Wood Thesis Advisor

Golden, Colorado

Date ______

Signed: ______Dr. Wendy Bohrson Professor and Department Head Department of Geology and Geological Engineering

ABSTRACT

The boundary between continental or shelf-delta scale foresets and toesets is a place of dynamic flow transition, but few studies have tried to image or examine this transition zone in seismic data. It is physically modelled as an area of high erosion where supercritical flows cascade down the clinoforming foresets and transitions to subcritical flows, eroding and depositing sediments as back-stepping cyclic steps (Ono and Plink-Bjorkland, in press; Kostic and Parker 2007). More recent modern systems work in lakes (Gardner et al., 1998), marine fjords (Mosher and Thomson, 2002) and in Cenozoic outcrops (West et al., 2019) have documented these cyclic steps and their relationship to steep shelf-margin and steep deltaic- margin slopes, and downslope current interaction. This study initially utilized a data set of extremely large shelf clinoforms offshore Guyana. The current project will use a 3D data volume of well-imaged foreset-to-toeset clinoform transition zones in the Cenozoic deposits offshore

NW Australia to examine the seismic expression of these processes. These clinoforms are deposited in ~ 6 packages that exhibit different degrees of progradation/aggradation architectures, seismic geomorphologic features and overall clinoform geometries. We will compare and contrast these different spatial and temporal clinoforming systems and examine the influence of controls on the foreset to toeset transition zone. The overall evolution of the clinoform packages, specific clinoform morphologies (sediment waves, slope gullies, mass transport deposits), and well log information will provide valuable insight into the ancient processes affecting the system and the sediment budget. This study has the following goals: 1) to utilize an integrated set of geologic and geophysical data to identify processes driving topset-to- foreset-to-toeset basin sedimentation and their resultant deposits, 2) to map the geomorphology and deposits of the complex zone of flow transitions in the foreset-to-toeset along this margin, 3)

iii

to understand any genetic links between clinoform progradation/ aggradation architectures, seismic geomorphologic features and the overall clinoform geometries, and 4) to speculate on the implications of these processes and deposits for reservoir and seal potential in such ancient margins around the world.

TABLE OF CONTENTS ABSTRACT ...... iii

LIST OF FIGURES ...... viii

LIST OF TABLES ...... ix

CHAPTER 1 INTRODUCTION ...... 1

CHAPTER 2 GEOLOGICAL OVERVIEW ...... 6

2.1 Regional Structure and Tectonism ...... 6

2.2 Regional Stratigraphy ...... 8

CHAPTER 3 DATASET AND PROPOSED RESEARCH METHODS ...... 12

3.1 Seismic Stratigraphic Framework ...... 13

3.2 Clinoform Surfaces ...... 14

3.3 Seismic geomorphology of the transition zone ...... 16

CHAPTER 4 PREVIOUS WORKS ...... 20

4.1 Clinoform architectures ...... 20

4.2 Flow transitions ...... 21

4.3 Modeling and Outcrop Work ...... 22

CHAPTER 5 OVERALL EVOLUTION OF SHELF MARGIN ...... 26

5.1 Structural influence ...... 27

5.2 Package Descriptions ...... 28

5.2.1 Package 1 ...... 29

5.2.2 Package 2 ...... 31

5.2.3 Package 3 ...... 32

5.2.4 Package 4 ...... 33

5.2.5 Package 5 ...... 34

5.2.6 Package 6 ...... 36

5.3 Progradation/aggradation (P/A) rates and effect on the sediment budget ...... 36

5.4 Log analysis ...... 40

5.5 Synopsis of Morphologies ...... 44

CHAPTER 6 SLOPE GULLIES ...... 47

6.1 Introduction ...... 47

6.2 Previous Works ...... 49

6.3 Observations ...... 51

6.3.1 Surface 2 (S2 gullies) ...... 51

6.3.2 Surface 4 (S4 gullies) ...... 57

6.3.3 Comparison of S2 and S4 gullies ...... 59

6.4 Discussion ...... 61

CHAPTER 7 SEDIMENT WAVES ...... 65

7.1 Introduction ...... 65

7.1.1 Ignative-current vs. continuous-current generated sediment waves ...... 66

7.2 Previous works ...... 69

7.3 Observations ...... 71

7.3.1 Surface 4 ...... 72

7.3.2 A-type waves ...... 72

7.3.3 B-type waves ...... 76

7.4 Discussion and Conclusions ...... 77

7.4.1 A-type waves ...... 77

7.4.2 B-type waves ...... 78

CHAPTER 8 MASS TRANSPORT COMPLEXES ...... 82

8.1 Introduction ...... 82

8.2 Previous Works ...... 84

8.3 Observations ...... 86

8.4 Discussions and Conclusions ...... 90

CHAPTER 9 CONCLUSIONS ...... 98

REFERENCES ...... 104

APPENDIX A ...... 111

vii

LIST OF FIGURES

Figure 1. 1 Typical clinoforming system ...... 2

Figure 1. 2 Seismic survey location map...... 4

Figure 1. 3 Clinoform packages in the study area ...... 5

Figure 2. 1 The Northern Carnarvon Basin outline ...... 7

Figure 2. 2 The present-day structural elements in the Dampier sub-basin ...... 8

Figure 2. 3 Cenozoic tectono-stratigraphic chart of the Northern Carnarvon Basin ...... 9

Figure 3. 1 Schematic model showing the morphometrics recorded for MTCs within this study 17

Figure 4. 1 Flow dynamics and architectures for turbidity currents...... 23

Figure 5. 1 Muderong structure map ...... 27

Figure 5. 2 Evolution of the margin through time...... 29

Figure 5. 3 Clinoform package sub-units ...... 30

Figure 5. 4 Isopach maps of each clinoform package ...... 31

Figure 5. 5 Surface 5 structure map ...... 39

Figure 5. 6 Well log analysis ...... 41

Figure 6. 1 S2 slope gullies ...... 54

Figure 6. 2 Area 1 and 2 S2 slope gullies ...... 56

Figure 6. 3 S4 slope gullies ...... 58

Figure 7. 1 A-type and B-type sediment waves ...... 75

Figure 8. 3 P5MTC 1 ...... 91

Figure 8. 6 Channels features in Package 5 ...... 96

Figure 8. 7 Package 6 slope channels ...... 97

Figure 8. 8 Schematic model of geomorphologic features and processes ...... 99

viii

LIST OF TABLES

Table 3. 1 Wells used in this study ...... 13

Table 5. 1 Calculated progradation and aggradation rates for clinoform packages ...... 39

Table 6.1 Morphometrics of slope gullies recorded in literature ...... 52

Table 6. 2 Range of morphometric parameters for the Surface 2 slope gully system vs. Surface 4 slope gully system ...... 62

Table 7. 1 Morphometrics of A-type and B-type sediment waves ...... 79

Table 7. 2 Morphometrics of sediment waves recorded in literature ...... 80

Table 8. 1 Morphometrics of MTCs in this study ...... 86

CHAPTER 1

INTRODUCTION

Clinoforms have been widely studied due to their importance in the evolution of continental margins worldwide. They occur globally in all margins throughout geologic time as basin-ward dipping horizons, forming in both carbonate and siliciclastic settings, and manifest at a variety of temporal and spatial scales (Cathro et al., 2003). In a proximal to distal orientation, clinoforms are architecturally composed of a topset, foreset, and toeset portion (Patruno et al.,

2015; Figure 1.1). Their geometries are a reflection of the nature of wave, tide and fluvial energies, sediment type and volumes, and the influence of extrinsic variables such as currents, relative or eustatic sea level changes, or regional tectonics at the time of their deposition (Cathro et al., 2003). As sediment laden fluvial waters enter the standing waters of a lake or ocean, they lose competency and are forced to drop their sediments, forming clinothems. Clinothems are clinoform-bounded sets of strata which gently prograde seaward into deeper water (Patruno et al., 2015). As these clinothems grow and steepen, they release sediment laden gravity flows down the c as submarine gravity flows. Such flows have been studied in both numerical and physical models (Vellinga et al., 2017; Hughes Clarke, 2016) and outcrop studies (Kostic et.,

2019, West et al., 2019) which show them transitioning from supercritical to subcritical velocities, a process that contributes to bypass of sediment, clinoform toeset erosion, and/or deposition of sediment on the foreset/ toe of slope. These toe of slope deposits can contain large percentages of sand and can act as prolific hydrocarbon reservoirs or may serve as prolific sand sediment sources for more basinward gravity flows. Such sand accumulations are documented occurring in the modern seafloor off the Squamish Delta, in Howe Sound Basin, Eastern Canada.

Subaerial delta Subaqueous delta clinoform clinoform

Topset

Foreset

Toeset

Figure 1. 1 Cross sectional profile showing features of a typical clinoforming system. The clinoforms in this study are subaqueous and the rollover point remains submerged during clinoform progradation (modified from Patruno et al., 2015).

There, super-critical to sub-critical flow transitions result in the deposition of “… thick sand beds with climbing dune stratification, which record high-energy turbulent waning flows under hydraulic-jump conditions during flow expansion at the mouth of the incised valley channel”

(Kostic et al., 2019). In addition, recent discoveries in offshore Guyana and Suriname may be reservoired in similar deposits at the toe of very steep clinoforming Cenozoic margins (Wood et al., 2019; Campbell, 2005).

The prolific nature of very large clinoform margins in basins around the world emphasizes the need to better understand how sediments are distributed within these potential reservoir units. Although several studies have been done examining these processes at the outcrop scale, and these processes have been documented in small scale physical models, there have been few studies that image, interpret and provide a template for identifying the deposits of

cyclic steps at a seismic scale. Likewise, seismic scale data that documents the morphology of clinoforming margins under the influence of shallow and deep water currents is very sparse.

Goals of this study are to document the seismic geomorphology of the topset-foreset-toeset system in several clinothems, prograding northwestward across the northwest Australian margin.

Observations will be used to identify the morphologic signature of margin bypass versus margin deposition. Observations will be further used to assess the relationship between clinoform character (architecture, size, lithology, dip) and the clinoform geomorphologic nature. Finally, the topset, foreset and toeset geomorphologies will be explored to investigate both shelfal and deepwater processes influencing the movement of sediments over the clinoform surface.

Although terminology can be convoluted, for the purposes of this research, we will refer to a clinoform as a single reflector that shows a general sigmoidal geometry with a basinward dip. A clinoform package is a unit of reflectors that is bounded at the top by an individual clinoform and composed internally of a variable number of reflectors that appear to have a clinoforming geometry.

This study incorporates 3D seismic data and well logs and report data from the Northern

Carnarvon Basin (NCB) of Australia, specifically the Dampier sub-basin, on the northwest shelf of Australia (Figure 1.2). Seismic data are provided by Woodside Energy and logs and supplemental data are available through the National Offshore Petroleum Information

Management System (NOPIMS). The 3D volume has well-imaged foreset-to-toeset clinoform transition zones in the Cenozoic interval (Figure 1.3). The research approach and results are not specific to this area alone and are intended to answer overarching questions regarding the nature of this transition zone that are broadly applicable to such settings around the world. This study has the following goals: 1) to utilize an integrated set of geologic and geophysical data to

Figure 1. 2 Location map and outline of Fortuna 2015 3D survey in offshore northwest Australia. Well locations, including 171 within the survey extents, as well as petroleum titles (outlined in blue) are labeled. Figure taken from Geoscience Australia, 2014.

identify processes driving topset-to-foreset-to-toeset basin sedimentation and their resultant deposits, 2) to map the geomorphology and deposits of the complex zone of flow transitions in the foreset-to-toeset along this margin, 3) to understand any genetic links between clinoform progradation/ aggradation architectures, seismic geomorphologic features and the overall clinoform geometries, and 4) to speculate on the implications of these processes and deposits for reservoir and seal potential in such ancient margins around the world.

In the NCB, the Cenozoic interval is not a working petroleum system unlike the underlying prolific Mesozoic petroleum system because of the lack of sufficient overburden, lack of source and lack of seal, but this study can be used as an analog for other basins, including but

Figure 1. 3 Six large clinothems are found constructing the study area margin during middle Eocene through Miocene time, bounded by Horizons 1-7 (marked in different colors, in ascending order from oldest to youngest on the seismic line above). The overall thickness of the clinoform package ranges from 430 to over 1100 m across the survey with the thickest strata occurring in the clinoform foresets.

not limited to the Cretaceous and Cenozoic of the Guyana and Suriname margin (Wood, 2019;

Yang and Escalona, 2011) and the Cretaceous margins plays of the western and southern Gulf of

Mexico (Janson et al., 2011), as well as Aptian-Cenomanian discoveries on the clinoforming margin of Arctic Alaska (Houseknecht, 2019).

CHAPTER 2

GEOLOGICAL OVERVIEW

2.1 Regional Structure and Tectonism The Northern Carnarvon Basin (NCB) is a rift-dominated basin that formed during late

Paleozoic extension of the northwest shelf of Australia. The NCB shelf covers an area of approximately 650,000 km2 (Hocking et al., 1987). The basin is comprised of four northeast trending sub-basins which include, the Exmouth, Dampier, Barrow, and Beagle sub-basins, from south to north respectively, and it is bordered by the Curvier, Gascoyne and Argo Abyssal plains

(Geoscience Australia, 2014; Chongzhi et al., 2013, Figure 2.1). The NCB has undergone several phases of rifting and, more recently inversion, creating a somewhat complex structural architecture. The three main structural elements include the inboard Peedamullah and Lambert

Shelf, the four sub-basins (depocenters), and the outboard Exmouth Plateau and Rankin Platform

(Geoscience Australia, 2014; Figure 2.1).The structural history of the basin is largely controlled by the gradual breakup of Gondwanaland which began in the early Jurassic and extended through

Valaginian time (Barber, 1994). The basin has evolved from a pre-rift, syn-rift, and post-rift sag basin to the present-day passive margin. Rifting began at the end of the Callovian and the extension created a series of NE-SW trending faults that compartmentalized the four sub-basins and controlled deposition into the sub-basins (Hocking et al., 1987; Barber 1994). The continental breakup triggered sea floor spreading and the formation of the Argo Abyssal Plain

(Hocking, 1987). Late Jurassic rifting was caused by the separation of the Australian and Indian plates and formed the Gascoyne and Cuvier Abyssal Plains (King, 2008; Figure 2.1). By the

early Cretaceous, rifting had terminated, and the basin became a tectonically quiescent environment.

Deposition shifted from siliciclastics to carbonates due to change in circulation patterns after the breakup of Gondwana, which ended 123.5 Mya (Hocking, 1987; Barber, 1994). During the post-rift sag phase, subsidence was dominant and the study area transitioned from a shelf to deep marine environment (Chongzhi et al., 2013, Hocking et al., 1987). The basin was quiet until recent Miocene inversion and reactivation of faults due to the collision of the Australian and

Greater India plates (Chongzhi et al., 2013, Hocking et al., 1987).

De Grey Nose

Figure 2. 1 The Northern Carnarvon Basin, outlined in blue, is composed of four sub-basins, the Exmouth, Barrow, Dampier and Beagle. The seismic survey, outlined in black, covers the northern Dampier sub-basin and part of the Rankin Platform. The red and green figures indicate known oil and gas fields respectively. Figure modified from Chongzhi et al., 2013.

The Dampier sub-basin has several structural components that control deposition into the basin. The present day structures include, from southeast to northwest, the Eliassen Orion

Terrace, anticlinal Rosemary Trend, Lewis Trough, anticlinal Madeline Trend, Kendrew Terrace, and Rankin Platform (Barber, 1994; Figure 2.2). All of the structural components were formed during the Middle and Late Jurassic rifting events which covered the entire NCB (Barber, 1994).

Figure 2. 2 The present-day structural elements in the Dampier sub-basin. There are two anticlinal trends and two linear synclinal depressions that were formed during the Jurassic rifting events. The seismic survey used in this study covers the Madeleine Trend, Kendrew Trough, and Rankin Platform. Figure taken from Barber (1994).

2.2 Regional Stratigraphy The Dampier sub-basin is mainly comprised of Triassic, Jurassic and lower Cretaceous sediments that can reach thicknesses up to 10 km (Baillie et al., 1994 and Geoscience Australia,

2014). The working petroleum system is comprised mainly of the Triassic and Jurassic syn-rift sediments and overlain by a regional seal deposited in the early Cretaceous. The interval of interest in this study spans from the middle Eocene Walcott formation to the Miocene Trealla

Limestone formation. The main stratigraphic units present in the Dampier sub-basin are outlined below and include the Paleocene-Holocene deposits. Early Triassic-Cretaceous stratigraphy can be found in Appendix A. The Cenozoic stratigraphic column is labeled in Figure 2.3.

Figure 2. 3 Cenozoic tectono-stratigraphic chart of the Northern Carnarvon Basin. The stratigraphic interval of interest in this study lies within the post-rift phase of tectonism and is outlined in the red box, including the Upper Walcott formation, Mandu Limestone and Trealla Limestone, covering ~40 Mya. Although this is a mixed system, carbonate sedimentation is dominant in the interval of interest. Figure modified from Paumard et al., 2018.

Lambert Formation (Early Paleocene): Slightly silty claystone with an upper silty sandstone member. It is only deposited in the Dampier sub-basin with no time equivalent deposit (Hocking,

1987). Deposition occurred during a marine transgression in a shallow marine environment

(Hocking, 1987).

Dockrell Formation (Early-Late Paleocene): Deep-water marl and claystone with abundant planktonic and benthonic foraminifera. Deposition occurred in an overall sea-level rise

(Hocking, 1987). This unit was deposited in an outer shelf to slope environment (Apthorpe,

1988).

Wilcox Formation (Late Paleocene-Middle Eocene): Slightly silty, pyritic and glauconitic claystone with foraminifera present throughout the formation (Hocking, 1987). The Wilcox was deposited in an outer shelf to deep marine environment (Hocking, 1987). The formation is separated into two units, a lower marl unit and an upper silty claystone to siltstone unit. The distinction between the upper and lower units is due to a transition from warm to cooler oceanic temperatures and increased continental run-off (Apthorpe, 1988).

Walcott Formation (Middle Eocene): Argillaceous calcilutite with diverse foraminifera overlain by a cherty-limestone sequence. The Walcott formation marks the transition from a clastic to a carbonate dominated system (Apthorpe, 1988). Deposition occurred in an outer shelf to slope environment (Hocking, 1987).

Mandu Limestone (Early-Middle Miocene): Prograding wedge of unsorted sediment consisting of quartz silt, basal marls and claystones that coarsen upwards into calcilutites and calcarenites

(Apthorpe, 1988; Smith, 2014). This represents a prograding carbonate shelf with open shelf conditions and water depths of several hundred meters (Hocking, 1987; Apthorpe, 1988, Cathro et al., 2003).

Trealla Limestone (Middle Miocene): Bioclastic packstone to grainstone with finely banded deposits interpreted as algal mats. The limestone was deposited in a shallow marine environment

(Hocking, 1987). The thinly bedded dolomites and sandstones are interpreted as reef and reworked sediments (Tortopoglu, 2015).

Bare Formation (Middle-Late Miocene): Interbedded carbonate-siliciclastic sediments (Smith,

2014). Deposits contain medium grained quartz sandstones with interbedded calcarenite

(Hocking, 1987). The presence of both carbonates and siliciclastics indicates a nearshore to beach environment (Hocking, 1987).

Delambre Formation (Late Miocene-Holocene): Fine-grained calcilutite sequence with abundant biogenic material deposited during a rapid transgression (Hocking, 1987).

CHAPTER 3

DATASET AND PROPOSED RESEARCH METHODS

The Northern Carnarvon Basin has a vast extent of public data due to petroleum exploration beginning in the 1950’s (Geoscience Australia, 2014). The Fortuna 2015 3D seismic survey collected by Woodside Energy and their partners, public well log data, and industry well completion reports were used in this study. The wells used and digital logs available are listed in

Table 3.1. The majority of the interpretation was based on the Fortuna 3D seismic survey. The survey was completed in 2014 and was processed with the help of insights from the vintage

Demeter 3D survey with the goal to improve imaging beneath the Cenozoic succession for regional prospectivity (Fortuna 3D Seismic Interpretation Report, 2016). The processing work was completed by WesternGeco. The seismic survey area is ~4000 km2 and covers the northern portion of the Dampier sub-basin and edge of the Rankin Platform. Structurally, it covers areas of the Madeline trend, Kendrew trough, and Rankin platform (Figure 2.2). The 3D survey is a full stack depth migrated (TVD) dataset. In order to convert time data to a depth volume,

Woodside did a well tie in Petrel using the North Rankin-4 well. Sonic and density logs were used to create a synthetic log to link logs and samples data to the seismic volume (Fortuna 3D

Seismic Interpretation Report, 2016). The survey inline range is 1515-9187, crossline range 557-

25525, and the data reaches down 3500 m in depth. CMP (common mid-point) spacing is 6.25 m. For our study, the seismic data has been cropped in Petrel and includes early Jurassic-present day sedimentation. There are 171 wells in this survey area; 52 wells are used in this project, of which 8 have digital logs. Because the focus of data collection is for petroleum exploration, many wells are typically not logged through the shallower, Cenozoic interval because it is not part of the known working petroleum system. Wells with data in the younger portions of the

stratigraphic section were used in order to constrain the ages of the seismic horizons and provide information on the section’s lithology. Biostratigraphic data and stratigraphic tops were taken from the well completion reports where available.

3.1 Seismic Stratigraphic Framework Key mapped seismic horizons were selected based on clinoform geometries, seismic amplitude signatures and the continuity of reflectors. Strong, continuous reflectors mark velocity changes within the sedimentary record and provide a framework for the evolution of the prograding clinoform packages. Seven key horizons were picked on inlines and crosslines to

accurately separate the interval into six seismic packages. The reflectors that bound clinoform

packages are both downlap surfaces and condensed sections. The packages are labeled 1-6 from oldest to youngest (Figure 1.3). Although the entire seismic interval of analysis is believed to represent ~40 Mya, the sequence framework provided by this study will better parse the age of

Table 3. 1 Wells used in this study

each major clinoform package to provide a more detailed understanding of rates of sediment accommodation and intervals of missing time. Cathro et al. (2003) and Smith et al. (2014) outline a detailed sequence stratigraphic framework of the Cenozoic sedimentation in the

Dampier sub-basin. This study did not replicate this sequence stratigraphic framework, but rather chose key horizons that could be accurately mapped through the survey. Seismic data mapping and analysis was done in Petrel.

3.2 Clinoform Surfaces The key horizons were picked on tightly spaced inlines and crosslines and surfaces were generated in Petrel. Horizons were picked using the seeded autotracking tool and manually adjusted when necessary. The resulting clinoform surfaces reveal a variety of slope gradients and overall geometries. Average non-decompacted foreset slopes were measured from several different crosslines spaced evenly throughout the survey. This approach is not optimal but justified due to 1. The young nature of the sections and lack of compaction in the mid-Cenozoic section, and 2. Our approach to compare and contrast the morphology of horizons within a very narrow thickness interval of the section. We understand that care must be taken in applying these results to other basins and older settings. Distance and non-decompacted thickness data were collected using the Petrel “ruler tool” and age data from well reports were utilized to calculate rates of progradation and aggradation for each clinoforming package. The provided ages from well reports are often very broad (ie. Middle Miocene), resulting in larger ranges of error on rates of progradation and aggradation.

Three geomorphologic features were identified in the data; slope gullies, sediment waves and submarine mass failures, for additional analysis. Slope gullies were measured using a modified methodology of Field et al., (1999). The spacing, width, depth and length of slope gullies were measured in seismic cross-sections and on surface attribute maps. Gully spacing is

the distance from the center of the deepest part of one gully (thalweg) to the center of the deepest part of the adjacent gully. Gully width is the straight line distance that intersects the top of the gully sides. Gully depth is the distance along a vertical line drawn from the bottommost portion of the gully up to a straight line connecting the sides of the gully. The gully length is the plan view runout distance of the gullies as mapped on surface attribute extractions. Sinuosity of gullies within the study area is 1 or very close to 1, which means the gullies are roughly straight.

Line drawings of gullies in plan view on attribute maps extracted from Surfaces 2 and 4 were done to better document the morphology and evolution of gullies in the foreset and toeset of clinoforms. Two types of sediment waves (A-type and B-type) were interpreted on attribute extractions from Surface 4 based on their spatial extent, their shelf-parallel and shelf-oblique orientations in plan view and their undulating cross-sectional morphology. Line drawings of the sediment waves were done by hand in plan view. Sediment wave height, lateral extent, length, and orientation were measured using the approach of Prieto et al., (2016). In addition, wavelengths were measured between individual deposits. Sediment wave height is the vertical interval over which the bedforms exist in a dip-oriented cross section. Sediment wave length is the distance across the wave in plan view with a distinct start and end. Sediment wave orientation was recorded as shelf-parallel, shelf-oblique or shelf-concentric. Spatial angles between wave orientation and the shelf break were also measured. Sediment wave wavelength is the horizontal distance between the peak of one wave to the peak of the adjacent wave on a section oriented perpendicular to the strike of the adjacent waves. The sediment waves do not have a large amount of relief, so identifying the peak in cross section can be challenging.

Sediment wave shape was noted if resolvable as either symmetric, downslope asymmetric, or upslope asymmetric. Submarine mass failures were identified above Surface 1 and below

Surface 6. The total length, width, deposit thickness, slope gradient, area and volume of the mass transport complex were all measured using a modified methodology of Clare et al. (2018) and

Moscardelli and Wood (2008) (). Total mass failure length is the plan view or cross sectional straight line distance from the failure head scarp to the downslope limit of the continuous deposit. The mass failure width is the largest straight line distance across the complex, orthogonal to the total length, in plan view. The mass failure deposit thickness is the distance from the topmost extent of the chaotic facies to the basal surface in a dip oriented cross section.

Slope gradient was measured on an adjacent un-failed slope interpreted to represent the slope of the pre-failure surface. The area is simply length x width and the volume is length x width x thickness. The nature of the basal surface (ie., deeply erosional, shallowly erosional, non- erosional) and nature of the failure material (deformed, un-deformed, chaotic) were also recorded. Root mean squared (RMS) amplitude and variance slices were used to identify the extent of the chaotic material and failure edges. The estimated area was measured and overall geometry, geomorphology and nature of the failure material was recorded.

3.3 Seismic geomorphology of the transition zone Seismic geomorphology utilizes seismic data in order to reconstruct the morphology of seascapes and landscapes through time (Wood, 2007). When seismic geomorphology is combined with log and core data, the spatial and temporal variability of reservoirs as well as the sequence stratigraphic framework can be better understood (Wood, 2007). The nature of active processes, and the resultant geomorphology of the foreset to toeset regions of a submarine margin are the major focus of this study. Based upon physical and numerical modeling, field analysis and more recently remotely-sensed observations, it is believed that this physiography is a construct of flows moving downslope across these regions. Several variables are hypothesized, through numerical and physical modeling, to influence the nature of these flows. It is suggested

Figure 3. 1 Schematic model showing the morphometrics recorded for MTCs within the study area. The total length, width, deposit thickness, slope gradient, area and volume of the mass transport complex were all measured using a modified methodology of Clare et al. (2018) and Moscardelli and Wood (2008). Figure based on Clare et al. (2018).

by the authors of these studies that, under the right conditions, (West et al., 2019; Kostic et al.,

2019; Vellinga et al., 2017; Catigny et al., 2010; Postma and Cartigny, 2014) flows transition from supercritical flow to subcritical flow as the flows reach the flattening in slope and transition from the foreset to toeset regions. The location of this transition is poorly understood. One of the goals of this study is to identify any morphologic changes in the foreset to toeset deposits that may represent this zone of transition, and to examine variables in the nature of clinothems that may influence where this important transition happens. In order to identify the areas of deposition or erosion on the foreset, seismic amplitude extractions were done in Petrel. Seismic attributes were extracted using the framework of interpreted key horizons outlined above. These attributes help reconfigure the depositional environment and morphology of the system at the time of deposition, and they reveal details on structure, stratigraphy, and deposition that are otherwise hidden in the data (Koson et al., 2014; Sarhan, 2017). Attributes extracted from the volume, utilizing the key mapped horizons to guide extraction, provide a method for mapping the geomorphology of the larger depositional systems that characterize the different seismic packages. Root mean squared (RMS) amplitude and coherency/variance were used to map the morphology of the foresets. RMS is a windowed amplitude extraction that computes a gross amplitude for the desired window regardless of positive/ negative amplitude sign to measure the reflectivity of the horizon or interval (Smith, 2014; Brown, 2004). RMS amplitudes will highlight sand bodies with high porosities and the sharp stratigraphic contrasts that characterize unconformities as high RMS amplitude values (Koson et al., 2014; Sarhan, 2017). Variance, a similar attribute to semblance, measures the similarity of waveforms and can identify faults, unconformities, and channel edges (Koson et al., 2014). Attributes were projected onto the key

mapped surfaces to highlight slope gullies and their morphology, the nature and extent of sediment wave fields and the nature and extent of mass transport deposits.

CHAPTER 4

PREVIOUS WORKS

4.1 Clinoform architectures Clinoforms have been widely studied due to their importance in the evolution of continental margins worldwide. They are ubiquitous in their presence throughout geologic time, exhibit a wide range of spatial scales (1-104 m in height) and are worldwide in extent (Patruno et al., 2015). Patruno et al., (2015) examined a dataset composed of ancient and modern clinoforms at delta, shelf-prism and continental-margin scales to identify diagnostic features for each type of architecture. Patruno et al’s., (2015) goal was to be able to identify these clinoforms in relation to their grain size, in seismic or stratigraphic data. His quantitative characterization approach outlined several different morphologic and chronostratigraphic parameters for each type of clinoform. Morphologic features used by Patruno et al. (2015) in their characterization were clinoform height, dip, and cross section morphology. For example, a conclusion drawn by these authors from clinoform height and dip determined that larger-scale clinoforms are deposited in deeper waters, over longer periods of time, and on steeper slopes (Patruno et al., 2015). Rates of progradation and clinoform trajectory, which are time correlative, were calculated by the authors from clinoform architectures as well as age dates from well reports. This study creates an important tie between recent and ancient systems and highlighted the applicability of seismic geomorphology for investigating paleo-environmental conditions. Cathro et al., (2003) looked at clinoforms on 3D seismic data in the Northern Carnarvon Basin. They used clinoform architectures as well as amplitude extractions in order to interpret the evolution of the shelf margin from Oligocene-to-Miocene time. They determined that clinoform progradation occurs at paleowater depths greater than 100m and is not greatly affected by changes in sea level.

Thousands of studies on clinoform geometries and architectures have been conducted (for a review see Patruno et al., 2015), illustrating the utility of these architectures in understanding the nature and processes active along continental margins. However, the study of supercritical to subcritical flow transition along margins is a relatively new development, and is an area ripe for additional work. This study will be the first known to this author to try to characterize such processes through a seismic geomorphologic approach.

4.2 Flow transitions Although this study is highlighting the supercritical to subcritical flow transition specifically, there are also high to low density and laminar to turbulent flow transitions that occur in these settings (Waltham, 2004). These flow dynamics are driven by sediment types and volumes, Reynolds number, and Froude number (Waltham, 2004). The distinction between supercritical and subcritical flow is important because they produce very different deposits

(Postma and Cartigny, 2014). Supercritical flow occurs when “flow is faster than the speed of a wave propagating across the flow surface,” (Waltham, 2004). The transition between supercritical and subcritical flows creates a sharp decrease in flow velocity and increase in flow thickness, common at a sharp decrease in slope, which produces the hydraulic jump. The deceleration causes sediment deposition from the flow, changing the density and viscosity of the flow (Waltham, 2004). Waltham (2004) investigated the three flow transformations in the San

Dimas Reservoir gravity current and the linkage between them in order to discern the nature of their deposits. He concluded that the laminar to turbulent transition and high to low density transition are connected and can only be achieved by fluid entrainment. The Froude number is assumed to be 1 for hydraulic jumps, but Waltham (2004) found that the Froude number is only

1 when the flow is uniform with depth, and hydraulic jumps can occur at Froude numbers less than 1 in depth-variable flows. A Froude number is the ratio of flow velocity over the wave

propagation velocity (Postma and Cartigny, 2004). Postma and Cartigny (2014) connected flow transitions and their resultant deposits in an outcrop and core study. They showed that at the bed scale, depositional processes and flow dynamics in turbidity currents exhibit four distinctly different sedimentary facies. Supercritical bedforms are far beyond the scope of the bed scale and are not included in turbidite facies models (Postma and Cartigny, 2014). However, supercritical bedforms have been identified on modern deepwater slopes as antidunes, chutes and pools and cyclic steps (Postma and Cartigny, 2014; Hughes Clarke et al., 2012; Ono and Pink-

Bjorklund, 2018; West et al., 2019; Kostic et al., 2019). The four facies distinctions are a supercritical base layer (I), subcritical base layer (II), supercritical flow (III) and subcritical flow

(IV) and correspond to different sedimentary structures and depositional environments (Figure

4.1; Postma and Cartigny, 2014). Case I yields the deposition of long wavelength cyclic steps, dune-shaped bedforms with upstream migration, typical of upper-fan channels. Case II facies are typically found on unconfined lower slopes in lobes with limited discharge and the lack of a hydraulic jump. These deposits are tabular and stratified in nature and can extend for tens of kilometers. Case III, non-stratified supercritical flow deposits are long wavelength, a few kilometers, cyclic steps that occur in levee and lobe systems. Case IV, non-stratified subcritical flow occurs on basal plains and yields planar turbidite deposits (Postma and Cartigny, 2014).

Postma and Cartigny (2014) conclude that turbidite sequences should be not be grouped into an idealized Bouma sequence, but rather sorted by flow dynamics. They emphasize the importance of processes over the well-established turbidite sequence.

4.3 Modeling and Outcrop Work The processes moving sediment along clinoforms has attracted more interest over the last decade and has been explored at a small scale in both outcrop and modeling. The transition zone between foreset and toeset, which is the focus of this study, has been previously studied but it is

not fully understood from the micro to meso-scale. Supercritical flow deposits are poorly understood and rarely recorded in outcrop because of their poorly constrained sedimentary structures and lack of appropriate nomenclature. There is also confusion between supercritical flow deposits with other similarly featured deposits. This is not to say that they don’t exist, there have been supercritical flow deposits identified in alluvial to deepwater outcrops (Ono and Pink-

Bjorklund, 2018). However, outcrop examples are sparse due to the poor preservation potential of these high energy deposits (Vellinga et al., 2017). Vellinga et al., (2017) used a depth-resolved numerical model to investigate the mechanics and deposits of a cyclic step system, the named deposit that is believed to result from supercritical to subcritical flow transitions. Based on the

Flow direction

Figure 4. 1 Flow dynamics and architectures for high (Case 1 & II) and low (Case III & IV) concentration turbidity currents. Figure taken from Postma and Cartigny (2014). model results, the depositional signature of a cyclic step includes amalgamated concave up erosional surfaces with low angle foresets and backsets (Vellinga et al., 2017). Their findings aligned with outcrop work done on cyclic steps except for the scale of deposits. Other numerical modeling suggests that many supercritical flow deposits are transitional bedforms between antidunes and cyclic steps (Kostic et al., 2019). They have been physically modelled as an area of high erosion where supercritical flows cascade down the clinoforming foresets and transitions to subcritical flows, eroding and depositing sediments as back-stepping cyclic steps (Ono and

Plink-Bjorkland, 2018; Kostic and Parker 2007). Cyclic steps are identified in outcrop based on

their long wavelength and low amplitude geometry at different scales, low angle foreset and backsets, concave up scour and fill and convex up geometries, and internal erosion and bounding surfaces (Ono and Pink-Bjorklund, 2018). Although the numerical modeling constitutes a much smaller scale manifestation of these processes, the cyclic step observations are very similar in location, geometry and character with outcrop results. There are problems with identifying these features in core due to their long wavelengths and low amplitudes (Ono and Pink-Bjorklund,

2017). Both the numerical modeling and outcrop study above identify these deposits as cyclic steps. Cyclic steps form in higher ranges of supercritical flow and are upstream migrating, asymmetric bedforms with steeply dipping backsets and scour surfaces (West et al., 2019; Kostic et al., 2019; Catigny et al., 2011). Cyclic steps are formed as the supercritical flow velocities rapidly decrease and produce a hydraulic jump at the base of the foreset as the flow becomes subcritical (Postma and Cartigny et al., 2014). However, West et al., (2019) refers to the supercritical flow deposits he observed in slope turbidites of the Fish Creek-Vallecito Basin

(Southern California) as antidunes. An antidune forms in lower ranges of supercritical flow and has broadly symmetrical, wavelike bedforms with low angle geometries that migrate upstream

(West et al., 2019). The use of antidune or cyclic step here is a nomenclature discrepancy that can be resolved with further study, and these two deposits may be formed in a continuum (see discussion below). Kostic et al., (2019) encompasses the widest range of scales in discussing these processes and deposits, analyzing numerical experiments, seismic data, and ground penetrating radar to map the architectures at a large scale. Upper-flow regime bedforms were studied in the Squamish Delta, British Columbia, Mazzarra River Delta, Italy and the ice- marginal lacustrine deltas, Germany and used as an analogy to deep-water systems (Kostic et al.,

2019). The Squamish delta experiences frequent turbidity currents, supercritical flow events,

observable over a 12 day period. As the Squamish flow originates, antidunes are first to form, but as it persists the antidunes are reworked and cyclic steps dominate (Kostic et al., 2019). If the flow persists for long enough, all antidunes will be destroyed. Kostic et al., (2019) uses several different resources to look at the upper flow regime deposits and has identified both small and large scale features that are important in gaining a better understanding of these processes. A study in the Mazzarra delta reviewed by Kostic, showed confined upper flow regime deposits within gullies. These deposits are linked to fluvial systems and can be telling of sea level fluctuations (Kostic et al., 2019). The ice marginal lacustrine delta study observed bedforms emplaced by tractional, surge-type and sustained supercritical currents. Each flow formed cyclic steps and antidunes, however there were differences in scale and grain size distribution. Cyclic step deposits are distinguished by lenticular scours with infill of backset cross stratification which are overlain by antidunes, low angle cross stratified sheet-like beds (Kostic et al., 2019).

Kostic et al., (2019) also explained the idea of sediment partitioning in relation to clinoform progradation. Gravel and sand accumulated in the foreset and the suspended load was deposited in the thinner toeset. There have been advances made in investigating the flow transition between supercritical and subcritical flow, but there are obvious inconsistencies with nomenclature that causes confusion and misunderstanding. Clarity will grow as more study is done.

CHAPTER 5

OVERALL EVOLUTION OF SHELF MARGIN

The six prograding clinoform packages identified in this study have been constrained by two different horizons; Horizon 1 and 7. The total thickness of these six prograding packages range across the study area from 430 to 1100 m (Figure 1.3). The basal horizon, identified as

Horizon 1 in Figure 1.3, is a high amplitude, continuous reflector overlain by a low amplitude chaotic to semi-chaotic seismic facies. Horizon 1 predates clinoform progradation and developed during the sag phase of the Mesozoic rift development, which corresponds to the underlying

Kendrew Trough trend. Horizon 1 transitions from a steeply dipping reflector in the SE of the study area to a shallow dipping, near horizontal reflector towards the NW. The relative age of

Horizon 1 is middle-late Eocene (~48-34 Mya) and corresponds to the base of the Upper Walcott

Formation, resolution does not allow for a more precise age. The top horizon of the clinoform packages, identified as Horizon 7, is a high amplitude reflector that quickly loses both amplitude and continuity to the SE (landward). It separates the underlying high to moderate continuous reflectors of the prograding clinoform sequence from the chaotic to semi-chaotic seismic facies interpreted as an overlying mixed siliciclastic/ carbonate debris flow. The relative age of Horizon

7 is middle Miocene and corresponds to the Trealla Limestone. It is important to note that reflectors within the study interval become discontinuous and display lower amplitudes towards the shelf (SE). This change is caused by a loss of reflection signal due to the presence of overlying late Miocene-Pliocene mixed carbonate/ siliciclastic shoreface deposit known as the

Bare Formation (Smith et al., 2014). Well log information was integrated with the seismic interpretation allowing for the development of a chronostratigraphic and lithostratigraphic framework. The six clinoform packages encompass ~40 Mya of geologic time and contain three

stratigraphic units; the Walcott Formation, the Mandu Limestone and the Trealla Limestone.

During this time, the margin was dominated by the deposition of calcilutites, calcarenites and marls. Clinoform progradation and aggradation on this margin is not a direct function of sea level, as they have been found to outbuild in water depths of hundreds of meter, but rather sediment supply and accommodation (Cathro et al., 2003).

5.1 Structural influence The Mesozoic history of the Dampier sub-basin was dominated by multiple rifting events which resulted in large scale faulting during the Jurassic. These faults controlled sedimentation and accommodation from the Jurassic well into middle Eocene deposition. The Muderong

Formation, an early Cretaceous shale, was mapped in order to understand the influence of the underlying structure on Paleogene sedimentation. Lows on the Muderong structure map (Figure

Figure 5. 1 The Cretaceous Muderong structure map highlights the effects of large scale Jurassic faulting. The inset dip oriented cross section, marked A-A’ on the structure map, shows the underlying faults effect on overlying Eocene sedimentation and highlights the sag phase, concave up bed geometry in Horizon 1 that overlies the Kendrew Trough

5.1) represent the downthrown side of the Jurassic-age faults. In the majority of these faulted areas, the lows correspond to the concave up nature of the middle Eocene reflector. In areas where this does not hold true, the underlying deposits have been thick enough to fill the accommodation or the throw on the faults was less drastic and did not create the same magnitude of sag. The Jurassic rifting influences deposition as far up section as the middle Eocene Walcott

Formation (basal horizon). This creates more accommodation in Package 1 in comparison to the other stratal packages.

5.2 Package Descriptions Along the study margin, a transition from a clastic-to carbonate-dominated depositional system occurred in the middle Eocene. The northward drift of the plate upon which the northwest shelf rides, from ~36-40⁰ S to its current location at ~20⁰ S (Veevers et al., 1991), is responsible for the shift to warmer climate, and carbonate sedimentation in the Cenozoic. The basal horizon of the clinoform package (Horizon 1) is of middle Eocene age and is time equivalent to the

Walcott Formation which denotes this shift. The horizon marks the transition from horizontal reflectors (beds) to clinoforming (progradational) reflectors. Although the six seismic packages display different seismic character, geomorphologic features, progradation/ aggradation rates and clinoform architectures (Figure 5.2), the six seismic packages are predominantly composed of calcilutite and calcarenite interbedded with fine-grained carbonate marls, according to well reports. Formation distinctions were assigned to packages, however because this is a prograding system, thickness variations and preservation of full stratigraphic units are dependent on location within the clinoforming package (ie., topset, foreset, toeset).

Figure 5. 2 A-A’ shows the evolution of the margin from P1-P6 through ~40 Mya from the middle Eocene to the middle-late Miocene. The black squares are the clinofrom rollover points and the white arrows indicate periods of progradation and aggradation. The dotted white line is the end of sag phase infill and beginning of clinoform formation in P1. P1-P4 are overwhelmingly progradational however there are periods of aggradation present in each. The transition into P5 and P6 marks the onset of more aggradation and upbuilding of the system, which can be seen in the thicker topsets in comparison to P1-P4. The average progradation/aggradation rates for each package are in Table 5.1. Progradation/ aggradation rates were calculated using well report data and distance measurements taken directly in Petrel. Average rates (m/ Mya) were calculated based on five inlines spaced evenly across the survey to account for spatial and temporal variations.

5.2.1 Package 1 Package 1 is bounded at the base by Horizon 1 and at the top by Horizon 2 (see Figure

5.3). Horizon 1 is a high amplitude, continuous reflector that transitions from a steeply dipping reflector to a more horizontal reflector to the NW with a general concave-up shape. Horizon 2 is a relatively continuous reflector with variable amplitudes and shows sigmoidal clinoforming shape sloping at gradients that range from 2.5⁰ to 3.5⁰. Slopes generally increasing to the NE.

According to well reports, the age of this package is middle Eocene-Oligocene and represents the upper Walcott Formation and lower Mandu Limestone. The Walcott Formation is time equivalent to the shift from a siliciclastic dominated to a carbonate dominated margin. The package thickness ranges from 50 to 450 m with the thickest portion of the unit spanning from the prograding slope wedge into the topset (Figure 5.1). The clinoform topset unit in this interval

Figure 5. 3 Dip oriented cross section displays the sub-units present within Packages 2-5. Sub- units were identified solely based on seismic character and architecture. Sub-unit divisions aided in bedform identification and description of internal fabric as well as progradation/ aggradation distinction. Although sediment waves were studied in detail on Surface 4, the undulating patterns seen in P2B and P4B should be explored further as potential wave fields. Full descriptions of sub-units are provided in Package descriptions.

are steeply dipping at the SE extent of the survey and are more gently dipping into the clinoform rollover point. The base of the package has planar reflectors with reflector strength varying across the survey. Associated with this package are four areas of highly chaotic seismic facies.

These facies extend updip to end sharply against a spatially extensive discontinuity. The detailed morphology of these deposits are discussed in the mass transport complexes chapter, and they are herein interpreted as submarine mass failures. The package foreset is a series of basinward- dipping, prograding reflectors rapidly thinning towards the toeset. Package 1 shows both temporal and spatial variation in seismic character across the survey. On the inlines, the package transitions from low-moderate amplitude, continuous reflectors with variable thickness on the shelf to a thinner, higher amplitude package approaching the slope. The deposits resulting from submarine failures, here termed mass transport deposits, can be tracked in 3D space and provide more information on their size and character. In the package topset, a ~170 m reflector package onlaps to the SW onto Horizon 1. In the package foreset, Horizon 2 is very undulatory and

Figure 5. 4 Isopach maps of each package labeled 1-6, depoceners are indicated with a dashed white line. 1. The increased accommodation due to underlying faulting allows for thick sedimentation in the topset-foreset transition zone and rapidly thins into the bottom of the foreset and toeset. 2. Drastic thinning in the topset and toeset with most deposition in the foreset, indicating the onset of progradation. 3. Most deposition still occurring in the foreset, but slower progradation rates were observed in comparison to Package 2. 4. Deposition begins to increase in the topset as aggradation increases. 5. More aggradation drives deposition in the topset and foreset at it builds up and out. 6. Aggradation rates are highest and the thickest interval is in the vertically stacked channelized foreset.

incises into the underlying surface, creating slope gullies. In the package toeset, the reflectors are more chaotic and thin onto a structural high in the middle of the survey.

5.2.2 Package 2 Package 2 is bounded at the base by seismic Horizon 2 (previously described) and at the top by Horizon 3 (see Figure 5.3). Horizon 3 is the uppermost reflector of Package 2, and shows excellent continuity in the foreset but poor continuity in the topset and toeset. However, although continuity is poor, amplitudes in the topset and toeset are moderate to high reflection strength.

The horizon topset is gently dipping at the SE extent of the data and transitions to a horizontal reflector in the NW direction for a length of ~6.7 km before the shelf break. The slope gradients range from ~2.9⁰ to 4.1⁰ with no definitive trend. According to well reports, the age of this

package is middle-upper Oligocene and represents the lower Mandu Formation. The package thickness ranges from 50-300 m with the thickest portion in the prograding foreset. The clinoform package displays very drastic thinning in the topset and foreset (Figure 5.4). The package topset is composed of two thick reflectors with moderate amplitude and continuity. The package foreset is purely progradational and composed of interbedded, high-low amplitude reflectors with moderate-good continuity. The entire foreset package can be divided into a lower sub-unit and an upper sub-unit, both of which are progradational (Figure 5.3). The lower sub- unit, P2A, has a strong reflector at the base with low amplitude reflectors downlapping onto it and is both progradational and aggradational. The upper sub-unit, P2B, has thicker, higher amplitude reflectors with wavy bases and higher slope gradients with a falling trajectory implying progradation. The undulating nature of the reflectors in P2B could indicate the presence of sediment waves, however, they were not explicitly studied in this research. P2B is more aggradational than the lower sub-unit. The package toeset transitions from thin, very low amplitude, discontinuous reflectors in the SW to thicker, moderate amplitude, discontinuous- continuous reflectors in the NE. On inlines, the package topset is a thin package of low-moderate amplitude, continuous, parallel reflectors. The package foreset is the thickest portion and aside from the undulating basal surface of the package (Horizon 2), the reflectors are moderate-high amplitude, continuous, parallel reflectors. The package toeset is a thin, low amplitude package.

5.2.3 Package 3 Package 3 is bounded at the base by seismic Horizon 3 (previously described) and at the top by Horizon 4 (see Figure 5.3). Horizon 4 is a high amplitude and continuous reflector in the foreset and toeset, and it transitions to a lower amplitude, discontinuous reflector in the topset.

Horizon 4 has a sigmoidal clinoform architecture. The horizon topset is gently dipping for ~24 km and relatively flat for ~9 km before is encounters the rollover point. The slope gradients of

Horizon 4 range from ~3.5⁰ to 5.2⁰, increasing to the NW. According to well reports, the age of this package is upper Oligocene and represents the Mandu Formation. The package thickness ranges from 25-250 m with the thickest portion in the prograding foreset (Figure 5.4). The largest interval of progradation is 4 km in length. The clinoform package has a thin topset and toeset. The package topset has discontinuous, chaotic, low amplitude reflectors that transition into thick, parallel, high amplitude reflectors towards the foreset. The package foreset is interbedded low-high amplitude reflectors with good continuity and can be divided into and upper and lower sub-unit. The lower sub-unit, P2A, has high-moderate amplitude interbedded thin reflectors losing continuity downdip. The upper sub-unit, P2B, of the foreset has moderate amplitude thick reflectors interbedded with lower amplitude, discontinuous reflectors with undulated bases (Figure 5.3). The package toeset is composed of two high amplitude, continuous reflectors. On inlines, the package topset has low amplitude, sub-parallel reflectors with laterally variable continuity. The package foreset is composed of 2-3 thick moderate amplitude, continuous, parallel-wavy reflectors. The package toeset thins and displays high amplitude, continuous reflectors.

5.2.4 Package 4 Package 4 is bounded at the base by seismic Horizon 4 (previously described) and at the top by Horizon 5 (see Figure 5.3). Horizon 5 is a moderate amplitude, continuous reflector, and separates a low amplitude seismic facies below from a high amplitude, chaotic seismic facies above. The horizon topset has a gentle long wavelength undulating geometry moving SE-NW which is more prominent than in the underlying packages. The horizon topset is flat for ~16 km before the rollover point. The clinoform reflector transitions from a low amplitude, discontinuous reflector in the topset to a moderate amplitude, continuous reflector in the foreset and toeset. The slope gradients range from 5.7⁰ to 6.5⁰ with no definitive trend. According to well reports, the

age of this package is upper Oligocene-early Miocene and represents the upper Mandu

Formation. The entire package thickness ranges from 30-300 m with the thickest portion in the prograding foresets (Figure 5.4). The entire package foreset can be divided into a lower sub-unit and an upper sub-unit. The lower sub-unit, P4A, is high amplitude, continuous parallel reflectors.

The upper sub-unit, P4B, is interbedded low-moderate amplitude, discontinuous prograding reflectors. Each reflector in P4B appear to have an internal basinward stepping fabric, similar to

P2B (Figure 5.3). There does appear to be overall basinward thinning of Package 4, but most of the thinning appears to occur in the toeset rather than the foreset. The package toeset thins and has 2-3 high amplitude reflectors. On inlines, aside from package thickness changes moving down the slope, the reflectors are moderate-high amplitude, continuous, consistent thickness reflectors. There are small lenticular channel bodies above Horizon 4 in the foreset, which are interpreted as slope gullies.

5.2.5 Package 5 Package 5 is bounded at the base by seismic Horizon 5 (previously described) and at the top by Horizon 6 (See Figure 5.3). Horizon 6 is a high amplitude, continuous reflector that is faulted at the shelf break, which makes it challenging to jump correlate into the shelf. In addition, the horizon is incised by overlying processes creating a more discontinuous horizon as we move landward. The horizon topset is a moderate amplitude, continuous-discontinuous, and relatively horizontal for ~13 km before reaching the clinoform rollover point. The horizon foreset is represented by a steeply dipping, faulted, high amplitude reflector. The horizon toeset is a high amplitude, continuous reflector but the basinward most portions of the toeset extend beyond the data’s spatial limits. The slope gradients of the individual clinoforming reflectors range from ~5.6⁰ to 7.7⁰, and are generally increasing in dip to the NE. According to well reports, the age of Package 5 is early-middle Miocene and represents the Mandu Formation and Trealla

Limestone. The thickness ranges from 50-400 m with the thickest portion in the prograding foreset, and the average thickness is ~150 m (Figure 5.4). The thickness from the topset to foreset remains fairly consistent and shows a great increase in topset thickness in comparison to

Packages 1-4. This package is very complex and contains channels, mass failures, and faulting.

The package topset is represented in seismic by moderate-high amplitude, discontinuous- continuous, wavy reflectors. The bottom of the topset is represented in seismic by moderate amplitude, parallel reflectors that are truncated by the overlying scour surface. The top of the topset has a large scour surface that is infilled with sediments represented in seismic by moderate amplitude, continuous, parallel reflectors. As the topset flattens and crosses the fault, it is characterized by high-to moderate-amplitude reflectors downlapping onto a chaotic basal failure surface. The entire package foreset can be divided into lower and upper sub-units. The lower sub-unit, P5A, is composed of moderate amplitude, continuous reflectors that show a prograding geometry. The upper sub-unit, P5B, is composed of low-to-moderate amplitude, discontinuous reflectors, and the interval contains isolated areas of high amplitude continuous reflectors, interpreted to be rafted blocks (Figure 5.3). The basinward-most portions of the toeset are not present due to the limited extent of the survey. On inlines, the package topset is composed of seismic reflectors showing high-to moderate-amplitude comprising geomorphologic elements that include stacked channels, small-scale faulting, mass failures and elongate lenticular channel forms. The mass failure complexes (Cardona et al., 2020; Moscardelli et al., 2006) are composed of high amplitude, continuous reflectors interbedded with low-to-moderate amplitude, discontinuous reflectors. The edge of the mass failure is very sharp and can be tracked with ease on plan view attribute maps.

5.2.6 Package 6 Package 6 is bounded at the base by Horizon 6 (previously described) and at the top by

Horizon 7 (See Figure 5.3). Horizon 7 is a high amplitude, continuous reflector where present, but incision by the overlying deposition makes continuous mapping difficult. The horizon topset is gently dipping for ~ 14 km before reaching the clinoform rollover point. The horizon foreset is a high amplitude, continuous reflector but is incomplete basinward as it progrades beyond the extents of the data. The slope gradients range from ~5.7⁰ to 6.6⁰, but are not truly representative because the full slope cannot be imaged in the eastern half of the dataset. Inconsistencies in the gradient calculations also exist due to the highly channelized and discontinuous nature of the uppermost clinoform foreset slope. According to well reports, the age of this package is middle

Miocene and represents the Trealla Limestone. The package thickness ranges from 50-400 m with the thickest portion in the prograding foreset. Similar to Package 5, we also see a thickness increase in the topset. The average package thickness is ~150 m (Figure 5.4). The package topset is characterized by moderate amplitude, continuous, parallel reflectors. The package foreset is both aggradational and progradational and contains very high amplitude, highly channelized reflectors. The package toeset is not present due to the limited extent of the survey. On inlines, the package topset is composed of low-to-high amplitude, discontinuous reflectors that are affected by the overlying incising surface and transitions to moderate amplitude, continuous, parallel reflectors downslope. The package foreset is composed of spectacular moderate-to-high amplitude, vertically stacked, short wavelength channels that are continuous across the entire length of the survey.

5.3 Progradation/aggradation (P/A) rates and effect on the sediment budget Estimated progradation and aggradation rates were calculated using relative age dates from well reports and measurements made by the author using Petrel software. The rates vary

spatially across the survey. The driving mechanisms behind the rates of progradation/aggradation are directly linked to the sediment budget. The partitioning of the sediment budget between the shelf and deepwater has important implications for the presence or absence of deepwater basinal sands which can serve as hydrocarbon reservoirs (Steel and Olsen, 2002). The clinoform architecture, geometry, and trajectory can be used to determine the partitioning of the sediment supply between the shelf and basin. There are several different controls on the evolution of the margin; sediment supply, shelf processes, relative sea level, tectonism and climate change (Steel and Olsen, 2002). The clinoform trajectory is defined by the change in shelf/ slope break or rollover point over time (Steel and Olsen, 2002). Although the term shelf-slope break is used, it does not imply a continental shelf margin setting; clinoforms are present in many different environments. Figure 5.2 shows the evolution of the study area margin, over ~40 Mya and a total progradation of ~17 km and a total aggradation of ~1 km. Progradation and aggradation rates were calculated on several different crosslines because of the spatial change in slopes across the survey and the average rate is displayed in Table 5.1. Package 1 has two distinct packages; a stage of sag phase infill and then the initiation of clinoform geometries, which has a progradation rate of 397 m/ Mya and an aggradation rate of 86 m/ Mya. Package 2 is mainly progradational and has an average progradation rate of 1326 m/ Mya and a short interval of aggradation accumulating at a rate of 26m/ Mya. Package 3 has an average progradation rate of 480 m/ Mya, and progrades at a slower rate than Package 2 due to a decrease in sediment supply. Following progradation, Package 3 sees a period of aggradation at a rate of 53 m/ Mya. This slower progradation rate is illustrated architecturally as the Package 3 toeset does not thicken, as in

Package 2, indicating limited bypass in the foreset and thus limited basinward deposition.

Package 4 shows an early phase of aggradation, accumulating sediment at a rate of 74 m/Mya,

before becoming progradational with an average rate of 920 m/ Mya. There is a shift in the system between Packages 1-4 and Packages 5 and 6. Packages 5 and 6 show rising trajectories, thicker topsets and higher amplitudes with more channelization, whereas Packages 1-4 show falling trajectories, thinner topsets and lower amplitudes with little channelization. Packages 5 and 6 have average progradation rates of 200 m/ Mya and 290 m/Mya and average aggradation rates of 273 m/Mya and 146 m/ Mya respectively. These packages show the lowest progradation rates and highest aggradation rates in the system which is indicative of their rising trajectories

(Table 5.1). We have outlined a couple working hypotheses to explain the shift in the system.

The increased accommodation landward in the topset could be related to the previously discussed structural low in Horizons 5 and 6 overlying the deeper Kendrew Trough allowing for an increase in accommodation (Figure 5.5). However, a combination of other factors such as sea level, tectonics, sediment supply, climate change and subsidence could be responsible for the increase in accommodation and onset of aggradation. The overall increase in accommodation and large scour surface within Package 5 would indicate that there is a landward variable that is turning on in the system. The bright amplitudes are related to the introduction of clastics into the system. The balance of the increased accommodation and increased sediment supply in the system allows for aggradation within Packages 5 and 6.

Based on non-decompacted water depth averages and eustatic sea level cures, the system shows an overall shallowing upward with a deepening cycle in the early middle Miocene. Based on the study done by Cathro et al. (2003), we know that clinoform rollover points were deeply submerged during margin evolution and we assume shelf water depths to be ~20->100 m.

Figure 5. 5 Surface 5 structure map shows the consistent increase in slope gradient moving towards the shelf break and grades into a low that extends across the shelf for up to ~12 km in some areas. The structural low creates more accommodation and results in thicker deposition in the topset of Package 5.

Table 5. 1 Calculated progradation and aggradation rates for clinoform packages

5.4 Log analysis The six clinoform packages comprise portions of three different stratigraphic units in the

Dampier sub-basin, the Walcott Formation, Mandu Limestone and the Trealla Limestone (Figure

6), from oldest to youngest. Well reports from wells that penetrate these clinoforms provide information on stratigraphic tops in the wells and their ages, and lithology of the penetrated clinoform intervals. Based on well reports, the entire interval of interest is predominately composed of calcilutites and calcarenites interbedded with marls that were deposited from the middle Eocene-to-middle Miocene. Digital well logs were available for eight of the wells and these logs were correlated to the seismically defined clinoform packages to produce type logs of clinoforms and their constituent packages (Figure 5.6). No core data were available in this study, but lithologies for the intervals of interest were interpreted from log response and well reports.

A well report from the Lady Nora 2 well provided formation tops and their ages, however tops were not provided for the seven other wells. Gamma ray (GR), sonic and resistivity logs were used to improve definition of formation boundaries in the wells, and subsequent interpreted clinoform lithologies. Seismic Horizon 1 coincides with a drastic overlying clean GR, increase in sonic velocity and decrease in density. According to the Lady Nora 2 well report, the base of the upper Walcott is marked by a transition from low sonic velocity below to higher sonic velocity within the formation. The upper Walcott is a calcilutite with traces of calcareous claystone at the top of the upper Walcott. The upper Walcott has a blocky GR signature that sits immediately overlying Horizon 1. The boundary between the underlying Walcott and the overlying Mandu limestone is marked in logs and reported in well reports as a decrease in sonic velocity as well as an increase in Gamma Ray moving from the base of the upper Walcott into the Mandu. The GR log response of the overlying Mandu limestone is “ratty” at the base due to higher interbedded clays but cleans upward. The contact between the Mandu limestone and the overlying Trealla

Figure 5. 6 A) The well log correlation is organized from topset to toeset and the spatial orientation of the wells can be seen on the base map. The black horizons indicate stratigraphic formations and the colored horizons correspond to H2, H3 and H4. The Upper Walcott, Mandu and Trealla bases are labeled. B) The Eaglehawk 1 type log penetrates all of the packages and shows the formation tops overlain on the seismic data. The Upper Walcott has a characteristic blocky GR signature and overall decrease in sonic velocity. The upsection transition into the Mandu Limestone is marked by an increase in GR and sonic velocity. The Trealla Limestone does not have a clear transition point, but typically shows a slight decrease in GR. C) P2 is most completely logged with penetrations in the clinoform topset, foreset and toeset and shows the comprehensive progradation of the system. P2 is a mainly progradational package and is composed of the Mandu calcarenite. The topset show a blocky GR signature near the base, with an upward increase in GR response as thin marls begin to interbed with calcarenite beds up section. The foreset is thicker and shows a similar log response to the topset intervals but with slightly less obvious increase in gamma up section. The toeset shows a low GR response with consistent, thin, high gamma interbeds.

formation was not obvious in well log signatures. It appears to be associated with a very subtle decrease in GR at the base of the Trealla but shows no strong sonic or resistivity response.

Interestingly enough, the transition on seismic is marked by a change to higher amplitude reflectors, as seen in clinoform package 5. The Trealla formation is a clastic/ carbonate deposited following a sea level rise in the Middle Miocene (Hocking et al., 1987).

The eight wells with digital logs through the Cenozoic interval penetrate different portions (ie., topset, foreset, toeset) of the clinoform packages providing lithologic information and log character for each portion. Although every clinoform package is penetrated by wells,

Package 2 is the most completely logged with penetrations in the topset, foreset and toeset in multiple locations. Package 2, being dominantly progradational and composed of Mandu limestone has a caclcilutite and calcarenite composition containing, by definition, 50% or more transported, clay-sized and sand-size carbonate grains respectively. These calcilutites and calcarenites are interbedded with thin marls. Because of the thoroughness of data over Package 2 we will use these data to do a log comparison between the topset, foreset and toeset portions of these clinoforms (Figure 5.6).

Portions of clinoforms can often have very different lithology and log character. The cross section in Figure 5.6 provides the basis for observations on the character of different portions of the calcilutite-calcarenite rich Clinoform Package 2. The P2 topset intervals show a blocky GR signature near the base, with an increase upward in gamma response as thin marls begin to interbed with calcarenite beds up section. The increase in gamma response is accompanied by a decrease in resistivity and a decrease in sonic velocity. Foreset intervals show log response similar to those of topset intervals but with slightly less obvious increase in gamma up section. Lithology of these clinoform portions vary laterally. The Lady Nora 2 well shows a

very blocky GR signature with far less interbedding of thin marl beds than what is seen in the

North Rankin 2 and Eaglehawk 1 wells. Compared to the foreset portion of the clinoform, the toeset portions show more rapid thinning in the interval. Gamma log signatures are blocky and show consistent, thin, high gamma interbeds. In some areas, the P2 toeset is the uppermost upper

Walcott and transitions into the Mandu limestone.

The lithologies and log signatures observed and described above display lateral and temporal variations in grain-size distribution and interval thickness which have important implications on reservoir characterization. As discussed above, the foreset is the thickest interval and shows lateral variations in gamma response (ie., blocky, smooth GR vs. rattier GR), which are due to the absence or presence of interbedded marls. For example, it appears that the Lady

Nora 2 has a smooth, consistent GR response which could indicate a dominantly calcarenite interval vs. the North Rankin 2 well, which shows an overall increase in GR and the consistent presence of marl interbeds. The presence of marl interbeds within the calcarenite-rich interval will lower the net-gross value. However, marls also have the ability to act as seals between the calcarenite beds due to their lower porosity which can be an important factor in the determination of flow units and hydrocarbon production and recovery. The thickness of the foreset interval and lateral variability could also be a function of autocyclicity within the system, which is a local response to changes in energy that results in changes in sedimentology, occurring during the overall progradation of the margin (Cecil, 2003). The topset and toeset show drastic thinning in comparison to the foreset intervals. The blocky GR response and presence of few marl interbeds up-section in the topset vs. the consistent presence of interbeds and higher GR in the toeset appears to be a function of grain size and proximity to the source.

The muddy intervals and fine-grained calcilutites appear to increase basinward into the toeset

whereas the more coarse-grained deposits are concentrated in the proximal topset. The thinning of the intervals is a function of the accommodation in the system and progradational nature of the package. The changes in thickness and lithology across the clinoform package occur over distances of a couple kms or less and are vital to accurate reservoir modeling.

5.5 Synopsis of Morphologies Deep-water depositional systems have become better understood in recent years due to the increased amount of deepwater 3D datasets and the potential for petroleum exploration in deepwater deposits (Posamentier and Kolla, 2003). 3D seismic data allow detailed imaging of deepwater depositional systems and seismic geomorphology can be utilized to identify specific depositional elements that are the building blocks of deepwater systems (Posamentier and Kolla,

2003). The seismic geomorphology of shelf to slope to basin transitions are highly complex due to the influence of gravity transitioning flows from subcritical to supercritical to subcritical, and the generation of gravity induced morphologies in the form of slumps, slide, debris flows and turbidites. In addition, the encroachment of sediment depositional systems into areas influenced by subaqueous currents create process complexities that result in development of shelf, slope and basin sediment waves (Wynn and Stow, 2002), and can create cross-currents that influence the migration of deepwater lobes and channel systems. The data within the study area enable identification of three primary architectures; slope gullies, sediment wave fields and submarine mass failures, all of which are discussed below in further detail. Larger features which appear to be submarine channels were identified on this margin but were not the focus of this study.

Submarine slope gullies are straight, regularly spaced, small, low-relief channel features found on continental slopes and in steeply dipping areas of the seafloor (Field et al., 1999;

Shumaker et al., 2017). Their role in slope sedimentation is poorly understood but has received more attention in the last decade (Field et al., 1999). Slope gullies serve as either areas of

submarine erosion and bypass from the slope to the deep basin or as areas of aggradation and deposition (Lonergan et al., 2013). Slope gullies were identified on Horizons 2 and 4 within the study area and on the associated interpolated surfaces. The gullies on each surface have different characteristics, interpreted to reflect difference in sediment supply and flow velocities during development (Shumaker et al., 2017).

Deepwater sediment waves are undulating sediment features with a rhythmic nature that occur in several different submarine environments, including submarine levee-fan systems, floors of fjords, and other deep water settings (Kostic et al., 2019; Lee et al., 2002). Sediment waves have been reported to migrate upslope or downslope and can be oriented in various directions depending on current directions, flow velocity, and source direction of sediment feeding the deposits (Howe, 1996). They are typically fine-grained, and a product of suspension fall out, but coarse-grained waves have been reported in various settings (Wynn and Stow, 2002;

Habgood, 2003). Sediment waves can be formed by both Froude-supercritical and subcritical flow and have typically been associated with turbidity currents and bottom currents (Howe,

1996; Wynn and Stow, 2002). Sediment waves can be further classified as antidunes and cyclic steps, however, the complexity and variability of sediment waves within this study does not always allow for a specific classification. Sediment waves have been modeled and studied at several scales, but the processes that form them are still poorly understood. Sediment waves were identified on an RMS amplitude map extracted on Surface 4 of the seismic dataset (Refer to

Figure 6.3).

Mass transport complexes (MTCs) and mass transport deposits (MTDs) are two scales of submarine mass failures (see Cardona et al, 2020 for discussion) that occur along continental margins or in areas of steep slopes prone to gravitational failures. MTC/MTD can be composed

of a variety of different types of deposits, including slides, slumps and debris flows, (Moscardelli and Wood, 2015). MTCs, seismic scale features, are recognizable in seismic data along many deep water margins and can constitute up to 70% of the entire deep water stratigraphic sequence in some basins (Moscardelli et al., 2006). MTCs were identified as part of Package 6 in the study area. We will discuss these architectures in more detail in the next chapters.

CHAPTER 6

SLOPE GULLIES

6.1 Introduction Continental slopes are dominated by gravity driven transport processes which include, slides, slumps, debris flows and turbidity currents. These flows can form a variety of erosional features such as gullies, channels, canyons and valleys (Galloway, 1998). These downslope processes are also responsible for deposition of chaotic slope deposits with broad sedimentary textures, channel/ gully levees, sediment waves and basin floor fans (Field et al., 1999).

However, because these deepwater processes are rarely studied in real time, their role in downslope sedimentation is poorly understood. Submarine canyons have long since been the focus point of continental slope morphology studies, but the importance of slope gullies has more recently been acknowledged due to the widespread extent of the gullying process and the recognition of hydrocarbon reservoir potential in these slope gullied settings (Pratson et al.,

2007; Surpless et al., 2009). In comparison to submarine channels and canyons, gullies are smaller scale in terms of width, length and depth (Field et al., 1999; Lonergan et al., 2013).

Aside from their low morphometric relief, gullies are also typically formed and infilled with fine-grained sediments which makes them hard to distinguish in outcrop (Hubbard et al., 2010).

Because of their small scale, gullies have only been detected recently and mapped in detail due to advances in imaging techniques. Such imaging has allowed for the identification of hundreds of submarine slope gullies developed along the NW Australian margin within our study area.

Submarine slope gullies commonly develop on upper continental slopes and in steep areas of the seafloor (Longergan et al., 2013; Shumaker et al., 2017; Chiocci and Casalbore,

2011). They are present on both delta slopes and slopes with no apparent fluvial input (Surpless

et al. 2009). Slope gullies play a significant role in slope sedimentation and play a part in a slope’s evolution over long periods of time (Field et al., 1999). Gullies are typically evenly spaced and widespread on continental slopes which sets them apart from single, larger channel morphologies (Lonergan et al., 2013). Gully formation ranges from erosional to aggradational or a continuum of the two (Shumaker et al., 2017). Erosional gullies include submarine rills and dendritic gullies. Rills are downslope gullies that converge with other gullies or canyons at low angles and are found on open slope areas. Dendritic gullies have a branching pattern and intersect canyons at high angles on steep slopes (Pratson et al., 2007). Aggradational gullies can be attributed to dispersive sedimentation of fine-grained material that results in a draping geometry of reflectors (Field et al., 1999). Gullies evolve over time and can change from an erosive feature that delivers sediment downslope to an inactive conduit draped with hemipelagic sediment (Surpless et al., 2009).

Slope gullies are present on the modern seafloor and can also be preserved in the ancient sedimentary record as can be seen within our study area. Likewise, the gullies observed in this study are well-preserved ~1800 m below the seafloor. In this study, gullies are not ephemeral features easily destroyed by transport and erosion processes, but rather important, well-preserved conduits for slope sedimentation and margin evolution.

Gullies can be separated into two morphologic categories; u-shaped and v-shaped, based on their cross-sectional geometry (Lonergan et al., 2013). V-shaped gullies are formed by mass wasting processes close to the shelf edge or at a change of dip along the continental slope

(Lonergan et al., 2013). They typically show frequent bifurcation and feed into large canyons or channel systems. U-shaped gullies are straight, poorly developed gullies that do not typically bifurcate, and don’t feed into larger systems (Lonergan et al., 2013). The gullies observed in this

study have variable architectures and occur at different points in time but could be dominantly classified as u-shaped gullies. Although gullies can be identified based on this morphological difference, gullies evolve over time and can change shape. For example, a u-shaped gully could have originally been a v-shaped gully that had increased infill in the gully axis and the response on seismic is in turn a u-shaped gully. This creates ambiguity when assigning processes strictly based on gully shape.

Although interest in submarine gullies has increased in the past decade, the processes responsible for their formation, infilling and relation to deepwater sedimentation is still poorly understood (Shumaker et al., 2017). Slope gullies are rarely observed in modern systems and due to their fine-grained nature, preservation in outcrop is not guaranteed, which makes process detection challenging. Sediment gravity flows, turbidity currents, plunging hyperpycnal flows from rivers, internal waves, nepheloid transport and submarine slides have been hypothesized as processes responsible for gully initiation (Galloway, 1998; Shumaker et al., 2017; Lonergan et al., 2013; Chiocci and Casalborne, 2011; Field et al., 1999). This is a rich topic of study. Aside from the lack of understanding on the processes driving gully formation; gullies are not well defined in literature. The gully terminology encompasses a wide range of scales and morphologic features with no formal definition. The lack of a well-defined scale or formation process for gullies creates ambiguities when differentiating them from other channelized features.

6.2 Previous Works High resolution seismic data and developments in high resolution multibeam seafloor imaging (Hughes Clarke et al., 1996) have allowed for a detailed study of the continental slope as well as the role of gullies in sediment transport (Surpless et al., 2009). Several studies have been done on the aggradational nature of slope gullies present on the northern California margin

(Spinelli and Field, 2001; Field et al., 1999; Burger et al., 2003). Researchers identified an initial

phase of erosion due to gravity flows during a lowstand followed by largely aggradational gully growth attributed to nepheloid deposition during higher sea levels. The late Miocene conglomeratic gully fills at Gaviota Beach, California were analyzed in outcrop for insights into the processes responsible for gully fill (Surpless et al., 2009). Due to the large grain size and erosional nature of these gullies, Surpless (2009) proposed oversteepening of the slope to have initiated slope failure, and gully formation by subsequent sediment gravity flows. Chiocci and

Casalbore (2011), investigated the slope gullies on the upper continental slopes offshore the

Tiber and Voltruno Rivers in the Tyrrhenian Sea and the Simento River in the Ionian Sea, and suggested that the gullies were formed during a period of highstand and low sediment supply.

They considered hyperpycnal flows as the mechanism of these depositional slope gullies, independent of sea level fluctuations. Over 1100 individual slope gullies on the Antarctic continental margin were quantitatively analyzed and researchers proposed slope character, large- scale spatial characteristics, ice-sheet history and sediment yield as the main driving factors behind differences in gully morphology (Gales et al., 2013). They identified five different gully types that range from simple single channel to complex branching systems (Gales et al., 2013).

The gully initiation was attributed to sediment-laden glacial meltwaters that triggered turbidity currents (Gales et al., 2013). Modern and subsurface u-shaped aggradational gullies on the passive margin of Gabon, West Africa were deposited in response to spatially-variable deposition rather than erosion (Lonergan et al., 2013). The authors noted that continuous flows can have changes in velocity and sediment concentration which causes an irregular pattern of erosion and deposition. Interestingly enough, these authors noted sediment waves cross cutting erosional and aggradational gullies similar to those documented in the study by this researcher.

In Gabon, the authors attributed the formation of sediment waves to sheet-like turbidity currents.

Shumaker et al. (2017) observed a similar interaction between submarine slope gullies and sediment waves on the continental slope in the Taranaki Basin, New Zealand. Following the logic of Lonergan et al. (2013), Shumaker et al. (2017) also attributed gully initiation and the presence of sediment waves to sheet-like turbidity currents. The large expanse of slope gullies worldwide and the different processes that form them give valuable insight into continental shelf and slope systems. Such variability also highlights the variations in processes in gullies of similar scale. The measurements taken on slope gullies in the studies mentioned above are organized in Table 6.1, and will be used to place our observations into a more global context.

6.3 Observations Slope gullies were interpreted on Surfaces 2 and 4 within the clinoform packages of the study area (Figures 15 and 17). Gullies were mapped using both strike and dip oriented cross section profiles in order to 1) denote gully initiation points on the slope, and 2) analyze full gully morphology. Gully shape was noted and spacing, width, depth, length and slope gradients were measured. These data are meant to provide insights into the spatial and temporal variability of gullies, and help link their development to the processes that formed them.

6.3.1 Surface 2 (S2 gullies) The S2 slope gullies are well developed across the entire length of the seismic survey in a region that constitutes the slope, ranging from the shelf roll over point to the base of the clinoform toesets. 157 gullies initiate in the upper topset of the clinoform, and converge to form

112 through going gullies that reach the basin floor. Gullies are straight to slightly sinuous, parallel each other, are evenly spaced and trend northwest, perpendicular to the isobaths. Based upon non-decompacted clinoform heights, the gullies extend from ~250 msl at the top of the foreset to as deep as ~400 msl in the distal toeset. Gully spacing ranges from ~150-1000 m and

Table 6.1 Morphometrics of slope gullies recorded in literature

gullying appears more concentrated to the west. Gully width ranges from 100-400 m and run out lengths range from ~1-7 km but they typically average 2-3 km. Gully depths range from 5-50 m.

Based on a dip angle map generated in Petrel, the gullies occur on slopes of 2-5.7⁰. The slopes in the foreset average 4-5⁰. Although slope measurements are made on non-decompacted clinoforms, the shallow depth of burial and data from wells suggest limited compaction. Gullies appear to initiate on the slope in the middle-to lower-foreset region, well below the shelf break rollover point (Figure 6.1C) where they have very low relief. Interestingly, although we do not know what it means, the initiation points form a sinuous planform trail along the slope, irrespective of the planar planform of the shelf break roll over points. (Figure 6.1A). The basinward extent of the gullies varies but the gully depth remains consistent and spacing greatly increases downslope. The increase in spacing is evidence of the concentration of flow from multiple gullies into single primary pathways.

Gullies are typically u-shaped with gently sloping to steep gully margins, as shown in

Figure 6.1B. In areas where the gully margins are steeper there is more aggradation and relief in the inter-gully area. Gully margins are asymmetric in some areas near the point of gully initiation and have a seismic wavelet doublet where gullies have yet to bifurcate. Where amplitudes are inconsistent, the gully axis and gully margin are both higher amplitude than the inter-gully areas.

The gullies maintain high amplitudes along their entire basinward lengths. Gullies were studied in great detail across the extent of the survey, and lateral variations were encountered. General trends within the southwestern and northeastern parts of the survey (Areas 1 and 2 on Figure 6.2) will be discussed here in detail. In the southwestern half of the survey, gullies are narrow with steep margins that can persist for a maximum vertical extent of ~70 m. It is important to note that in some areas, there is gully formation beneath Horizon 2, which increases their vertical extent.

Figure 6. 1 A) Surface 2 structure map with the shelf, shelf break and basin floor annotated. The shelf break is relatively straight whereas the gullies form a sinuous trail in planform. a) S2 gullies are evenly spaced, straight to slightly sinuous, parallel one another and trend NW. B) A- A’ reiterates the even spacing and highlights the draping overlying reflector. Horizon 2 (H2) is the blue gullied reflector and is indicated with the black arrow. b) The enlarged image shows the degree of basal erosion as well as areas of infill. C) The gully initiation point is well beyond the shelf break and occurs in the middle-to-lower foreset region.

Gullies initiate in the middle-to-lower foreset and can extend into far reaches of the toeset, however, most terminate in the upper toeset region. Incision into the basal surface is variable. Gullies can fully erode the underlying reflectors up to ~ 60 m below Horizon 2 or gullies may only partially incise causing minor to no discontinuities in the underlying reflector.

Reflectors overlying gullied intervals are typically aggradational and drape gullied intervals with minor changes in thickness between the gully axes and intergully areas. The deepest incision and area of most infill in the gully axes occurs in the middle foreset. The downslope evolution of the gullies across the study area, but typically, erosion and incision into the underlying reflectors decrease in the lower foreset, and gullies become very wide (up to ~400 m) basinward and are overlain by a very aggradational reflector. In the furthest reaches of the toeset, intergully areas are barren and the all deposition is in the widening gully axis as flows become unconfined and begin depositing in toe of slope fans. Gullies still exist over vertical intervals of 30-60 m basinward. Gullies in the northeastern half (Area 2) of the study show slight changes in their morphology, where gullies have more gentle margins and only persist for a maximum vertical extent of ~50 m. Gully incision does not have as much of an impact into underlying reflector continuity and in this area, gullies only incise into the reflector directly beneath Horizon 2. In this area, gullies are mainly aggradational overlain by draping reflector with uniform thickness.

Gullies in Area 2 typically do not extend as far basinward as those in Area 1, but gullies that persist into the toeset widen and shallow, accumulating up to 20 m of fill within the gully axes and showing no intergully deposition. The gullies in this area also overlie MTC 1 in Package 1

(discussed below). The overall nature of the Surface 2 gullies is aggradational, indicated by the consistent draping geometry of the overlying reflector. The gullies do not occur in complexes but are rather distributed over the length of the survey in between a maximum vertical window of

~70 m of sedimentation. The distinction between Area 1 and 2 slope gullies shows the lateral variation present in the system. Gullies in Area 1 show deeper basal incisions and longer runout lengths overlain by differential drape fill thicknesses and gullies in Area 2 are low relief and show aggradation, followed by a drape of very consistent thickness downslope.

Figure 6. 2 A) Surface 2 structure map with Area 1 and 2 indicated with white boxes. S2 gullies show slightly different geomorphologies between Areas 1 and 2, however both gully types are still considered aggradational. B) Cross Section A-A’ shows the deeper incision into underlying reflectors as well as periods of infill above draping reflectors as opposed to C) Cross section B- B’ which shows the more aggradational nature of the gullies with less incision into the underlying reflector moving to the NE.

6.3.2 Surface 4 (S4 gullies) Approximately 30 S4 slope gullies were identified on Surface 4 over an area of ~30 km2.

The gullies are short, straight, parallel one another and trend northwest. Unlike the gullies on

Surface 2, these gullies do not show upslope bifurcation. Gully spacing ranges from ~220-2000 m, and they are more abundant in western portions of the study area (Figure 6.3). Gully widths range from 100-600 m and they have run out lengths from 800-2000 m. Gully depth ranges from

7-30 m, with the average depth of ~14 m. Based on a dip angle map generated in Petrel, the gullies occur on slopes of 2.9-7.8⁰, but the gully initiation points are in steeper sloping areas, typically averaging ~6.5-7⁰. The Surface 4 gullies are confined to the foreset region; they initiate in the upper foreset and extend into the middle-lower foreset. They appear to be positioned between A-type sediment waves upslope and B-type sediment waves downslope. Strictly based on the planview image, the gullies appear to intersect the downslope extent of the oblique A- waves and disrupt or overprint B-type wave development in the middle foreset region.

Hypotheses on the relationships between gullies and sediment waves will be expanded in the discussion. Basinward, the gullies decrease their height and increase their width, creating low morphologic relief.

Gully morphologies initially appeared to be overprinted by the A-type sediment waves upslope based on RMS amplitude map interpretations, but after point tracing the thalwegs of several gullies, it is evident that the gully shape in cross section is the true gully morphology and the gully reflections are influenced by the overlying sediment wave deposits. In shelf parallel cross section, the gullies are asymmetric with gully margin gradients changing frequently. They typically show infill in the gully axes and erosion or thinning in the intergully areas. Moving downslope, the gully relief is very minimal, approaching the seismic resolution limits, and makes the identification of gully fill challenging. Although there is evidence of minimal gully infill, the

Figure 6. 3 A) Surface 4 RMS amplitude map with slope gullies outlined in black. The presence of the A and B-type sediment waves makes the planview identification of gullies more challenging. B) A-A’ shows the u-and v-shaped gullies with variable margin gradients and obvious gully infill. The S4 gullies persist for ~30 m. Downslope, they become very low relief and document small amounts of incision into the underlying reflector. C) B-B’ shows the spatial interaction between the S4 gullies and the B-type sediment waves downslope. The S4 gullies initiate in the upper foreset and extend into the middle-lower foreset.

vertical interval over which these gullies persist is only ~30 m. Horizon 4 gullies show consistently high amplitudes throughout their extent. The underlying reflector has moderate-low amplitudes and is very discontinuous by nature which makes it more difficult to determine whether or not Horizon 4 is incising into it and causing it to break even further apart. In some areas, it is clear that the gullies do not incise into the underlying reflectors. There is evidence for gully fill near the shelf and the overlying reflector displays a thickening in the gully axes and thinning on the margins with an overall flat top. This may be suggestive of sandier fill in the proximal reaches of the gullies with fill decreasing downslope. Gully widths increase drastically, and gully depths increase slightly downslope. It is hard to determine the basinward extent of the gullies because the sediment waves have reworked original deposition.

6.3.3 Comparison of S2 and S4 gullies The morphometrics of gullies on Surface 2 and gullies on Surface 4 are fairly similar, with only small-scale differences. The S2 gullies lie stratigraphically at the base of the progradational, Clinoform Package 2, and they interact with the complex topography formed by the emplacement of several MTCs at the top of Clinoform Package 1. The S4 gullies lie stratigraphically at the base of Clinoform Package 4, which shows more aggradation in comparison to P2. In plan-view, the morphologies of the Surface 2 and the Surface 4 gully systems are very different. Surface 2 is a complex system of bifurcating, closely spaced, long gullies that occur across the entire width of the survey in a ~115 km wide swath that extends from downslope of the shelf break rollover into the toeset regions. In strike-oriented cross sections, the S2 gullies are closely spaced, symmetrical, gullies with a draping overlying reflector mimicking Horizon 2 topography. In contrast, the S4 gully system has short, straight, asymmetric and symmetric gullies that only exist over a limited area in the western portion of the survey. Dip-oriented cross sections across the Surface 4 gully system displays concave up, low

relief features similar to those seen in Area 2 of Surface 2. However, the S4 gullies are less aggradational, initiate further upslope, have a shorter runout length, and are less developed spatially. The morphologic differences, differences in the slope initiation points and differences in the spatial extent of the gullies suggest different conditions of development. The infill of the gullies is very interesting. When the overlying reflector has a draping geometry, as seen in

Surface 2 gullies, there is an additional very minor, but detectable amount of infill above that reflector before returning to horizontal reflectors (Figure 6.1B). Although the draping reflector thickness appears uniform, there is a ~1-5 m increase in the gully axis depth, resulting in a total axial sediment accumulation of ~20 m. This draping, aggradational gully morphology in the S2 gullies indicates that fine-grained sedimentation dominated, and the gullies were passively filling with sediment across the entire slope. It is possible that prior to being filled with a slope drape, the gullies bypassed sediment across the slope and that sediment drapes were deposited during a quiescent, post-bypass time when the slopes were dominated by fine-grained sedimentation. In contrast to the gullies on Surface 2, Surface 4 gullies show little to no deposition in intergully areas and all sedimentation occurs in the gully axes. However, the infill in the gully axes are ~10 m. The shallow nature of these gullies may imply a shorter time of development; however, it could also imply fewer or less tractional flows moving through the gullies. Alternatively, the substrate may be less prone to erosion, or the slope may encourage more suspensional and less tractional (erosive) processes. The erosive basal nature of the induvudial S2 and S4 gully systems varies, but it is interesting to compare and contrast them as different populations. The S2 gullies have steep margins and are deeper than the S4 gullies, which are typically asymmetric with lower relief. Both gully populations incise and disrupt the continuity of their underlying reflectors. However, the extent of S2 gully incision is greater (~70 m) than S4 gully incision

(~30 m), incising and truncating reflectors much deeper than the S4 gully systems. Exceptions to such erosion of underlying reflectors do exist in both populations. Differences in depth of gully incision could be due to lower frequency or lower magnitude of flows or flows with more suspension than tractive character, resulting in less downcutting energy occurring on Surface 4.

An alternative hypothesis might be that the reflector underlying the S4 gully system is simply not as erodible as the reflectors underlying the S2 gully systems, or that the S4 system evolved over a shorter period of time not allowing for deep incision. S2 gullies appear to be more aggradational and persist longer through time based on the draping infill character and basal incision. The S4 gullies occur on slopes of 4-5⁰, whereas S2 gullies occur on slopes of ~2.5⁰. The steeper slope gradient may also explain the short, straight nature of the S4 gullies. The increased frequency of upslope bifurcation in gullies on Surface 2 suggest flows coming from multiple sources updip and feeding the middle slope. Likewise, such ornamentation suggests a longer time period of development, enabling a more mature development of slope gullies.

6.4 Discussion The slope gully analysis done in the northwest shelf of Australia’s Dampier Sub-basin reveals ranges of morphometrics (Table 6.2) that fall within the bounds of those measured for other gully systems in modern and ancient basins worldwide (see previous discussion). Strictly based on scale alone, the gullies most closely compare to those documented in the Taranaki

Basin by Shumaker et al. (2017). However, the wide ranges (or high variance) of sizes in gullies speaks to the difficulty in classifying gullies on the basis of morphometrics alone. Our study does reveal differences in gully cross-sectional geometry that might prove useful in determining differences in depositional processes along continental margins.

Surface 2 gullies show three very important characteristics. These include their lateral extent across the entire study area, their updip bifurcated nature and the draping overlying

Table 6. 2 Range of morphometric parameters for the Surface 2 slope gully system vs. Surface 4 slope gully system

reflector. Their lateral extent suggests that the slope is being fed by a shelf process active across the entire study area. The regular spacing and expansive spatial extent of the Surface 2 gullies indicate a line source as opposed to a point source, feeding these slopes. Moreover, this margin did not have a fluvial input that would have served as the point source during this time that could have provided basinward sedimentation across the survey extent (Hocking et al., 1987). The updip bifurcating nature of these gullies supports the suggestion that the source of sediments feeding to the upper slope was more of a line source than a point source and supports the convention that sediments are being transferred through these gullies for a long period of time, allowing them to build secondary gullies off of the primary gully trunk streams. They were active for a fair period of time, but were dominated by bypass in nature, as indicated by the lack of detectable infill of parallel reflectors, like those identified in gullies of the Taranaki Basin by

Shumaker et al. (2017). The undulating nature of the seismic reflector that forms the overlying drape suggests that the gullies were not filled when the system was abandoned, allowing for a thick pelagic drape to develop overlying the S2 gully system. Although there is gully infill via the draping sediment, it is a slope-wide period of fine-grained sedimentation as opposed to localized gully infill.

In contrast to the S2 gullies, gullies on Surface 4 do not cluster and are more broadly spaced. Spacing is very irregular and they occur over a limited area in the northwestern region of the study area. This limited spatial occurrence might suggest a fairly localized source of sediment transfer from shelf to slope. Gullies on this surface are on average shallower but wider than those developed on Surface 2. Lower amounts of vertical incision could be due to more tractional flows developing over lower slopes (Wynn et al., 2000), however this difference could also be a function of a more erosive-resistant older slope lithology limiting the vertical erosion of gullies on Surface 4. The short-lived nature of these gullies would imply an erosive flow that was quickly infilled in the gully axis following incision and development and a return horizontally bedded sedimentation shortly thereafter.

We conclude that submarine slope gullies are an integral part in downslope sedimentation. Although the S2 and S4 gullies have different defining characteristics and exist over different spatial and temporal extents, they have similar morphometrics overall. The gully formation is attributed to different processes and shows evidence of tractional and suspensional flows filled by both passive, aggradational drapes and active gully infill. Based on the observations made, gullies are not limited to one depositional environment; they can form during different stages of clinoform progradation and aggradation and interact with other processes such as sediment waves and mass failures. These findings agree with the conclusions made by

Shumaker et al. (2017). Although it was not the focus of this project, slope gullies can also serve as prolific hydrocarbon reservoirs under certain conditions (Surpless et al., 2009). Gullies can act as a compartmentalized reservoir if they are infilled with coarse-grained material with evidence of tractional flow in the system (Surpless et al., 2009). Based on the infill in the S2 and S4 gullies, the fine-grained drape present in the S2 gullies would not serve as an ideal reservoir,

however the heterogeneity of the S4 gully active infill and presence of coarser-grained material could lend itself to a potential hydrocarbon reservoir. However, the small scale of these features would require further petrophyscial and core investigation. Up-dip hydrocarbon migration though gullies can also pose a large risk to exploration, and the presence of an up-dip stratigraphic trap or seal is necessary (Surpless et al., 2009). The upper foreset initiation point of the S4 gullies poses a large risk of hydrocarbon losses updip and the integrity of a viable stratigraphic trap would need to be confirmed to prevent migration through the top of the system.

In contrast, the middle- lower foreset initation point of the S2 gullies could mitigate the risk for updip hydrocarbon migration.

CHAPTER 7

SEDIMENT WAVES

7.1 Introduction Sediment waves have been identified in studies for several decades in a variety of subaqueous settings worldwide. These include submarine levee-fan systems, canyon floors, floors of fjords, flanks of volcanic arcs and other deep water and slope environments (Cartigny et al., 2011; Lee et al., 2002). However, there are large variations in their scale, morphology and grain size (Cartigny et al., 2011). Sediment waves have been defined as transverse bedforms with crests aligned perpendicular to the current flow direction that typically migrate upslope

(Lonergan et al., 2013). However, this is not always true and sediment waves can migrate both upslope and downslope and have crests aligned in various directions (Howe, 1996; Xu et al.,

2008). Upslope migration is indicative of the influence of supercritical flow processes which are not an environment specific process (Symons et al., 2016). Sediment waves are common at or close to the seafloor and can also be preserved in the ancient rock record, yet there is a lack of organization on the formation and classification of them.

The initiation mechanism for the deposition of sediment waves is poorly understood, but typically involves continuous currents flowing at or close to the seafloor and something on the seafloor that impedes or changes the nature of that flow causing it to drop sediment (Cartigny et al., 2011). Sediment waves, originally solely attributed to the flow of alongslope bottom currents, are recognized as also being deposited by downslope flowing currents (Wynn and Stow, 2002).

Such downslope flowing currents are believed to have a more turbiditic origin. Because turbidites have been proven hydrocarbon reservoirs, data on turbidite-originating flows are more abundant, and many of the recent studies focus on turbidite-flow origin sediment waves (Wynn

and Stow, 2002). Finally, other studies identify a complex continuum of continuous and non- continuous current processes, combined with other gravity processes such as slumping and creep to form “wave-like undulations” (Wynn and Stow, 2002; Xu et al., 2008). Although there have been positive strides in sediment wave classification and identification, the case-to-case variability indicates that there is still ambiguity with nomenclature.

7.1.1 Ignative-current vs. continuous-current generated sediment waves Although there is not a clear classification scheme, sediment waves are commonly classified based on process and grain size. The majority of studies have attributed the formation of sediment waves to either ignative currents or continuous currents, which occur along most continental margins (Mulder et al., 2008). Ignative currents are episodic, short-lived events that include turbidity currents (Boggs, 1995; Prieto, 2016). Turbidity currents are the process responsible for the deposition of turbidites. They are rapid mass movements consisting of turbid

(sediment laden) water (Heezen and Ewing, 1952). Sediment from a turbidity current is deposited as a turbidite when there is a density difference between the turbid flow and the ambient water it enters. Turbidity currents are typically ignative surge events only lasting up to a few days, but more stabilized turbidity flows (non-ignative flows) can occur, lasting up to a few weeks (Mulder et al., 2008). These are typically hyperpycnal turbidity flows formed during floods (Mulder et al., 2008). In contrast to turbidity currents, continuous deep ocean currents are long-lived events that include along-slope currents that produce and rework sediment into sediment wave deposits (Martorelli et al., 2016). Deposits from all of these flow types may be classified into sub-types based upon grain size (fine or coarse). Deposits resulting from ignative flows may have similar morphologies and scale to those deposits resulting from continuous currents, however differences in flow energy, competency and duration are defining characteristics (Mulder et al., 2008).

Sediment waves formed by turbid gravity flows have been identified in many studies and are grouped based on their grain size (Wynn and Stow, 2002; Cartigny et al., 2011, Lee et al.,

2002; Lonergan et al., 2013; Shumaker et al., 2017). These currents flow downslope and deposit approximately flow-perpendicular bedforms that are either fine-grained or coarse-grained. Fine- grained ignative flow sediment waves are typically mud-and silt-dominated deposits and are typically found in areas of unconfined turbid flow, such as, in association with channel levees.

Ignative flow sediment wave morphology is related to the slope gradient and sediment supply

(Normark et al., 1991; Wynn and Stow, 2002; Cartigny et al., 2011; Wynn et al., 2000). Wave crest alignment is normally slope parallel and the waves are composed of interbedded turbidites and pelagic sediment (Wynn and Stow, 2002). These finer-grained, ignative flow-generated sediment waves can be up to 7 km long and 80 m high (Wynn and Stow, 2002). Coarse-grained ignative-flow sediment waves are sand and gravel dominated and are found in proximal channels or channel lobe systems. These types of sediment waves migrate upstream and exhibit very straight crests (Wynn and Stow, 2002; Cartigny et al., 2011). Because they are laid down in a confined setting, the flow velocity controls the wave geometry; slower flows will generate sediment waves with shorter wavelengths (Wynn and Stow, 2002). Coarse-grained waves are smaller scale than fine grained waves; averaging ~ 1 km long and 10 m high, and they persist for only short periods of time (Wynn and Stow, 2002).

A second family of sediment waves are those deposited by persistent/continuous current occurring at or very near the sea floor. These sediment wave types are also classified based on grain size (Wynn and Stow, 2002; Habgood et al., 2003; Mosher and Thompson, 2002). Fine- grained continuous current-deposited sediment waves are dominantly composed of mud and/or fine sand and are typically associated with sediment drifts, or with deep basins not affected by

ignative gravity flow currents (Wynn and Stow, 2002). They have straight crests that are normally oriented oblique to the slope, they migrate upslope and can extend over several thousand square km (Wynn and Stow, 2002). Continuous current deposited, fine-grained sediment waves can be up to 10 km long and 150 m high (Wynn and Stow, 2002). In contrast coarse-grained, continuous current sediment waves are typically composed of sand and can exhibit a wide range of orientations. These deposits can morphologically be straight-crested features or can occur in the form of a barchan-shaped sediment waves. However, the latter have not been frequently identified (Habgood et al., 2003). They are typically ~ 200 m in length and a few meters high (Wynn and Stow, 2002) and they tend to occur in unconfined areas where persistent bottom currents are present (Symons et al., 2016).

Sediment waves provide compelling insight into the type of deepwater currents and slope processes present at the time of or following clinoform progradation. Sediment waves have been interpreted as antidunes and cyclic steps in numerical modeling and outcrop studies (Cartigny et al., 2011; West et al., 2019; Vellinga et al., 2017; Kostic, 2019; Postma and Cartigny, 2014). For example, West et al., (2019) proposed the term antidunes as a possible interpretation for upstream-migrating sediment waves in a turbidite system, further breaking down the umbrella term, “sediment waves”. In this study, the term sediment wave will include both cyclic steps and antidunes, however there is not enough supporting evidence to make the distinction between the two within this study. Unfortunately, seismic-scale evidence is scarce to differentiate between the two processes on the basis of seismically discernable features. Ignative gravity current- deposited waves have a similar morphology to continuous current-deposited sediment waves.

The key may lie in their occurrence relative to gravity flow sourcing slope elements (clustered in association with slope channels), their spatial distribution on the slopes (paralleling slope or

perpendicular to slope) and their overall depositional setting (bottom current waves have been proven to occur on slope gradients as low as ~0.025⁰, whereas turbidity current waves require slopes > 0.1⁰ (McCave, 2017)).

7.2 Previous works Because deep water flows are seldom directly observed in nature, many studies use numerical and physical modeling as well as seismic surveys and core to investigate the dynamic nature of sediment waves. Sediment waves have been identified in datasets since the 1950’s, however, they were initially classified as ripples (Ewing et al., 1968) and hills (Johnson and

Schneider, 1969). As imaging techniques improved, sediment waves were increasingly identified in different studies and the main formation processes were attributed to ignative gravity currents and continuous bottom currents. However, there were still people who believed that sediment waves were generated by slumping (Bouma and Treadwell, 1975). Advancements in remote imaging (seismic and side-scan) has allowed for more detailed studies of sediment waves. Wynn et al. (2000) investigated two different wave fields in the western Canary Islands, La Palma and

El Julan. Both wave fields were interpreted as deposited by ignative gravity currents, however the La Palma wave field was deposited in an unconfined setting and the El Julan wave field was deposited in a confined, channelized setting (Wynn et al., 2000). Wynn et al. observed that flow confinement produces smaller-scale sediment waves (Wynn et al., 2000). In the Gulf of Cadiz,

NE Atlantic, sidescan sonar, 3.5 kHz profiles, piston cores and bottom photographs were used by

Hagbood et al. (2003) to document a large occurrence of sediment waves that showed a variety of scales and grain sizes. They concluded that the sandy-muddy sediment waves, as well as sand sheets and channels were created by both downslope and alongslope continuous bottom currents.

This study shows the dynamic currents that can occur on steeply dipping regions of the

continental margin, exhibiting continuous current flows parallel as well as perpendicular to slopes (Hagbood et al., 2003).

Many studies have attempted to create a classification scheme for sediment waves. Wynn and Stow (2002) characterized 14 deep water sediment wave examples based on both process and grain size. They created criteria for differentiating between the two processes and grain sizes based on depositional environment, wave morphology, wave sediments and migration (Wynn and Stow, 2002). The details of this study are mentioned in the above section. This analysis allows for the interpretation of depositional environment based on wave morphology; however, this classification scheme is based on a small dataset and only applies to deepwater waves. In an attempt to create a more statistically valid classification scheme, Symons et al. (2016) did a robust analysis of 82 sediment waves in variable water depths and environments and identified small sediment waves, large sediment waves and scours. The large gap in scale between the sediment waves is attributed to confinement (ie. Smaller waves form in more confined areas) and is independent of process (Symons et al., 2016). Although they were able to draw more conclusions from their larger dataset, their work largely agreed with the classification provided by Wynn and Stow (2002) who used a much smaller data set.

The above studies distinguish sediment waves based on morphometrics and attribute them to either ignative or continuous currents. Prieto (2016) did a quantitative analysis of sediment waves from 31 sediment wave fields identified in peer-reviewed literature. The author divided the waves into contour-conforming and non-contour-conforming waves, which had not been discussed in previous works (Prieto, 2016). The contour-conforming waves were associated with contourite drifts and typically occurred in the lower slope to basin floor, whereas the non- contour-conforming waves are typically associated with leveed channels and can also be formed

by unconfined turbidity currents on the continental slope and rise (Prieto, 2016). The contour- conforming sediment wave lengths and heights were larger; however, the range of wave lengths and wave heights were very similar (Prieto, 2016).

Sediment waves, deposited as cyclic steps and antidunes through processes of supercritical flows may be very different internally from waves deposited by continuous bottom currents during subcritical flow. Although these differences are not visible at a seismic scale, such differences have been seen in physical models and in outcrop studies. Cartigny et al. (2011) did a comparative study between simple ignative gravity current deposited sediment waves and those deposited as supercritical flow influenced cyclic steps using a numerical model to simulate cyclic step deposition. The geometry and grain size of the simulated deposits were compared to datasets from both fine and coarse-grained sediment waves. They concluded that previously identified upslope migrating sediment waves contain certain internal structures and geometries caused by supercritical flow processes forming cyclic steps (Cartigny et al., 2011). Therefore, although difficult to discern seismically, differentiation of these two types of deposits is important for understanding their reservoir character.

7.3 Observations Features identified as sediment waves were observed on Horizon 4 within the prograding clinoform packages. Reflectors with undulating character similar to the sediment waves on

Horizon 4 were identified in P2B, P3B and P4B, however they were not closely studied and could be a result of another process (Figure 5.3). Sediment wavelength, height, wave-crest length and wave-crest orientation and shape were measured. Morphometrics, as well as spatial distribution, occurrence relative to other slope morphologies and setting of deposition were used to interpret the processes responsible for their deposition.

7.3.1 Surface 4 There are two morphologically different sediment waves identified on Surface 4. The two different wave forms have very distinct geometries that can be easily interpreted on plan view attribute maps. Two wave types, A-type and B-type have been identified (Figure 7.1). A-type waves are sediment waves present directly beneath the shelf break with crests oriented oblique to the shelf break, and B-type waves present in the toe of slope and show crests oriented parallel to the shelf. There is evidence of interaction between the two wave types, as illustrated by sharp directional changes when they encounter each other. Water depths were calculated from clinoform heights using non-decompacted seismic data and were believed to be ~250-400 m in the foreset-upper toeset region. Sediment wave types A and B are described in more detail below.

7.3.2 A-type waves The sediment waves are well developed over an area of ~60 km2. Approximately 75 shelf-edge oblique sediment waves are recognizable on Surface 4 confined to a band approximately 65 km wide located in the upper slope directly basinward of the shelf break. They are elongate bedforms that are evenly spaced, averaging ~150 m between crests. Although they occur over the length of the survey, they are most well-developed in the middle third of the study area (Figure 7.1). Depositionally, they are developed below the rollover point in the upper portion of the margin foresets (Figure 7.1). The waves appear to interact with slope gullies in the

SW portion of the survey, as slope gully formation appears to predate sediment wave development. The relationship between the two geomorphologic features is looked at further in the Discussion section. In northeastern portions of the study area, A-type sediment waves terminate into B-type sediment waves (discussed below), but the temporal relationship between the A and B waves is not clear. The A-type sediment waves have wavelengths of ~ 150-500 m

and wave heights of ~ 12-20 m. The wave crests are straight with lengths of ~ 2 km and they are oriented ~ 30⁰ from North. Moving to the NE, the crest orientation is ~45⁰ and the point of convergence with the B-type waves is more gradual. The slope gradient at the top of the foreset is ~ 5-7⁰. The A-type waves exist exclusively in the upper-to-middle foreset regions and appear to be truncated by the more downslope-located B-type waves. However, a few of the A-type waves do extend into the lower foreset before assuming a shelf-parallel orientation.

Although A-type sediment waves are present over large extents of the survey, there are lateral variations in geometry which could correspond to alterations in flow velocities and sediment supply. In planview, the A-type waves are less developed and have a minor presence in the southwestern most portion of the survey. This could be due to the presence of slope gullies in this portion of the survey allowing more basinward sedimentation instead of upper slope deposition. The plan view morphology becomes more pronounced and frequent moving to the

NE and the occurrence of gullies greatly lessens. To examine the cross sectional geometry of these sediment waves, a random line was picked perpendicular to the A-type waves (Figure 7.1).

In cross section, the A-type waves are very low morphologic relief and have a wide range of geometries. Typically, they are broad, asymmetric bedforms with the highest amplitudes occurring in the downdip trough (lee side). The overlying reflector fills in the low relief bedforms and shows non-deposition or thinning on wave crests and elongate, consistent thickness infill with minor thickening downslope. Sediment waves forming Horizon 4 display low relief concave up shapes with a symmetrical to slightly downslope-oriented asymmetrical cross-sectional geometry.

Figure 7. 1 RMS amplitude map extracted on Surface 4. A) Unannotated map showing the lateral extent of sediment waves B) Enlarged annotated map shows orientation of sediment waves on the slope and the spatial relationship between A-type and B-type waves. C) B-B’ shows the cross-sectional view of the A-type sediment waves which are oriented oblique to the shelf edge. The sediment waves exist in the middle foreset and are typically broad and symmetric in shape. The arrows denote the geomorphic boundary between the two wave types. D) A-A’ shows the B-type sediment waves in cross section. They are upslope asymmetric with symmetry increasing downslope. The black line indicates the initiation of B-type sediment waves.

7.3.3 B-type waves B-type sediment waves are well developed over an area of 70 km2. Approximately 70 individual, roughly shelf-parallel sediment waves were identified on RMS amplitude maps

(Figure 7.1). The sediment waves show straight-to-sinuous crests and frequently bifurcate in planview. These B-type waves are located further downslope than the A-type waves, and are found as far basinward as the lower foreset regions. The S4 slope gullies are terminated or overprinted by the B-type waves. The A-type waves also appear to be truncated by the B-type waves. The B-type waves are not well developed in the inter-gully areas of the survey’s southwestern portion. However, to the NE where gullies are less frequent, the B-type waves are better developed in the upper-middle foreset. The B-type waves have wavelengths ranging from

~ 100-500 m and wave heights ranging from ~ 3-20 m. The sediment waves appear to be continuous for crest lengths up to ~ 4 km and are oriented roughly shelf parallel. The slope gradient in the lower foreset-to-toeset regions ranges from ~ 1.2-5⁰ over ~ 1.5 km. B-type waves have irregular spacing and crest lengths, which creates areas of wave clustering and areas of isolated wave development. A seismic cross section oriented perpendicular to the crest of the waves shows that waves are asymmetric with a steep, short stoss side and a long lee side creating a pointed crest, and upslope asymmetry (Figure 7.1). As the slope gradient decreases toward the toe of slope, the waves become more symmetric with rounded crests and troughs. The most basinward waves distinguishable are in the upper toeset and have broad, very low relief geometries. In the northeastern fourth of the survey, the waves become more symmetric and evenly spaced with shorter wavelengths. The wave amplitudes are consistently high with a slight decrease in amplitude on the downslope limb of the wave. The top of the reflector (top of

Horizon 4) has a slightly undulating character where the sediment waves exist and the overlying reflector mimics that geometry. However, this is a very low relief contact and the overlying

reflector does not display any substantial thickness changes or erosive nature which indicates the absence of infill in the wave troughs. The sediment waves persist over a vertical interval of ~20 m.

7.4 Discussion and Conclusions 7.4.1 A-type waves The temporal relationships between the A-type and B-type waves and the slope gullies are not well understood, but hypotheses can be developed based on study observations. A-type sediment wave morphometrics most closely align those of the fine-grained continuous current morphometrics observed by Wynn and Stow (2002), however, we know that A-type waves are deposited in a more proximal depositional environment than what is proposed by Wynn and

Stow for sediment waves in their own study. Likewise, A-type sediment waves appear to resemble in size those termed “non-contour-conforming sediment waves” discussed by Prieto

(2016). Based on the orientation and morphology of the A-type waves, and their “attached” appearance to the distal shelf, we propose that down-slope directed off-shelf currents are responsible for their formation. These currents are transferring sediments from the distal shelf to the upper slope or upper clinoform foreset and remolding these sediments into fields of sediment waves. The shelf-derived currents are not necessarily bottom hugging gravity directed density flows but likely currents in the shelfal water column. Another hypothesis is that these upper foreset sediment waves are part of a continuum, with flows moving downslope to provide sediment to contour currents reworking slope sediments into shelf- parallel B-type sediment waves. In the southernmost regions of the study area, sediment waves do not develop, as all the available sediment for formation is bypassed basinward via well-developed slope channel systems. This implies that the gullies either predate or are simultaneously active with the A-type sediment wavefields.

7.4.2 B-type waves The chronostratigraphic relationship between the slope gullies and the B-type sediment waves is not well understood. There is evidence that, at least in the northeastern portions of the study area, the slope gullies could post date the B-type waves. The B-type sediment waves are not well defined on RMS amplitude maps and appear only moderately developed in the SW corner of the dataset. Coincidentally, this is a region where slope gullies are very well developed.

In this area, sediment wave troughs can be tracked on seismic cross-lines, but these sediment wave troughs flatten dramatically when they migrate over the top of any gully features. Such geometries are interpreted to indicate slope gully truncation of underlying sediment waves, causing seismic continuity disruption. However, an alternative interpretation considered is that the sediment waves postdate the slope gullies and utilize sediments to fill the gully topography, thieving sediments from the wave building process, resulting in poor wave height development in conjunction with slope gullies.

The B-type sediment waves do not occur in isolated pockets, but rather across the entire extent of the seismic survey. Although their wave morphology does vary across their survey, their occurrence throughout the entire survey points to a line source rather than a point source for their sediment. Their slope parallel, contour-conforming orientation would also indicate a flow moving roughly perpendicular to the slope. A sheet-like turbidity current could have moved downslope and deposited these sediment waves perpendicular to flow (Shumaker at al., 2017,

Lonergan et al., 2013, Lee et al., 2002, Wynn and Stow, 2002), followed by continuous current reworking of these deposits into a more sheet-like geometry.

In comparison with previous works discussed above, looking at morphologic parameters and scale alone (Table 7.1), the sediment waves observed in this study are on the lower end of the

Table 7. 1 Morphometrics of A-type and B-type sediment waves

ranges given for wavelength, but the wave height, wave crest length and orientation fall within values reported worldwide for both ignative gravity-current-generated sediment waves or continuous current-generated sediment waves. The range of sediment waves morphometrics shown in Table 7.2 highlights the poor understanding of the relationship between scale and process in these phenomena. Based on the observations and inferred modes of deposition, we have created a 3D depositional model for A and B-type waves to simplify this multi-dimensional problem (Figure 8.8).

Sediment waves have been hypothesized as potential hydrocarbon reservoirs, but no true exploration has been conducted on sediment waves known to this author (Migeon et al., 2006).

Sediment waves can constitute large, inter-connected accumulations of sand depending on the environment of deposition, which could act as a reservoir (Migeon et al., 2006). The sediment waves in this study are well-developed across ~65 km of the survey and display heights up to

~20 m and appear to be continuous for up to 4 km. The A-type sediment waves are sourced from off-shelf currents and are deposited directly beneath the shelf break. Based on their proximal location to the shelf, the A-type waves could host large accumulations of coarse-grained material and offer high reservoir potential. In contrast, the B-type sediment waves are source from unconfined turbidity currents deposited further downslope which might imply finer-grained sedimentation and could correspond to lower reservoir potential. However, the B-type sediment

Table 7. 2 Morphometrics of sediment waves recorded in literature

waves in this study show a high amplitude response which could imply otherwise. The thickness and continuity as well as the spatial extent of sediment waves observed in this study makes them a prospective reservoir target. Additional work is needed to fully understand the role of sediment waves in petroleum exploration.

CHAPTER 8

MASS TRANSPORT COMPLEXES

8.1 Introduction Subaqueous mass failures (mass transport deposits and mass transport complexes) occur on all continental margins worldwide and play an important role in the generation and destruction of seafloor topography and in deepwater sedimentation (Moscardelli and Wood,

2008; Posamentier and Martinsen, 2010). They can be comprised of a variety of grain sizes and lithologies and occur on scales from meters (sub-seismic-scale mass transport deposits) to kilometers (seismic-scale mass transport complexes) (see Cardona et al., 2020 for review). They are responsible for delivering large volumes of sediment from the shelf to the deep basin

(Cardona, 2015; Moscardelli, 2006). Mass failures are not restricted to any one environment and can occur on carbonate and siliciclastic margins or in mixed lithology margins like the one presented in this study. Mass failures are composed of any gravity-induced deposits, which includes, slides, slumps, debris flows and turbidites (Moscardelli and Wood, 2008). Mass transport deposits (MTDs) and mass transport complexes (MTCs) are historically used interchangeably in the literature, but in this study a mass transport deposit is a sub-seismic scale deposit, and mass transport complexes are an amalgamation of MTDs that create a seismic-scale deposit (Cardona et al., 2020). Independent of their seismic-scale definition, the terms are used interchangeably within this study. Instabilities in the shelf can trigger mass failures and lead to the deposition of MTCs at any time regardless of relative sea-level (Posamentier and Kolla,

2003). Slope instabilities can be attributed to a number of factors including sedimentation or structural oversteepening, tectonic movements of the seafloor, or a fall in sea level (See

Posamentier and Martinsen, 2010 for complete list). Large mass-wasting processes can also

initiate tsunamis, avalanches, and rock slides that all significantly affect onshore and offshore geotechnical infrastructure, and that can be proven to be environmental and economic hazards

(Posamentier and Martinsen, 2010; Moscardelli and Wood, 2008). Mass failure deposits have, in some settings, been known to constitute up to 70% of the entire slope and basin stratigraphy

(Moscardelli, 2006).

Interest in deepwater sedimentation for petroleum exploration, specifically the presence of mass failure deposits and their influence on hydrocarbon trapping, sourcing and production, has increased in recent years (Alves et al., 2014; Moscardelli and Wood, 2008). Mass failure deposits are typically fine-grained, low permeability and porosity rocks aligned more closely with a seal rather than a reservoir rock (Posamentier and Martinsen, 2007; Cardona, 2015), however exceptions do exist. One such exception is the Chalk Group mass failures within the

German Central Graben which were initiated by salt movement and tectonic uplift and serve as a good reservoir sourced by hydrocarbon from the Kimmeridge Clay (Arfai et al., 2016). The chalk did not have time to form a cement before being moved downslope and was able to maintain a porosity of ~50% at the time of re-deposition, providing a prolific reservoir (Arfai et al., 2016).

On seismic, mass transport complexes are very distinct due to their large size and chaotic, low amplitude to semi-transparent reflectors (Moscardelli and Wood, 2008; Alves et al., 2014).

The overall geometry of mass failure deposits is very easily identified due to their striking seismic character, however the internal complexities of the deposits are not completely resolvable at the seismic scale. Outcrop and core studies are better able to capture the internal structure of mass failure deposits. Independent of scale, similar structures and architectures are observed in all mass failures (Posamentier and Martinsen, 2010). Integrating multiple scales of

data helps to gain a more comprehensive understanding of the processes and architectures that characterize mass failures (Posamentier and Martinsen, 2010). In this study, there is extensive seismic data coverage of the MTCs, and therefore we will use the seismic to sub-seismic scale templates developed by colleagues to interpret the nature of these deposits in our study area (Bull et al., 2009; Cardona et al., 2020; Moscardelli and Wood, 2016).

8.2 Previous Works There have been many attempts made to classify mass failure deposits based on different parameters. Posamentier and Martinsen (2010), examined mass failures at both seismic and outcrop scale and noted that the broader processes and architectures can be deduced from seismic data. For purposes of our seismic-based study, we will review a classification scheme for mass failures that is based on seismic scale criteria. Moscardelli and Wood (2008) created such a classification scheme based on geomorphological and morphometric dimensions, causal mechanisms, and source area location. They identified attached and detached mass transport complexes (MTCs). Attached MTCs can be either shelf-or slope-attached and are thousands of km2 in area, whereas detached MTCs are sourced from seafloor topography that is well removed from shelf and slope sediment sources, and are orders of magnitude smaller, covering areas <10 km2 (Moscardelli and Wood, 2008; Figure 19). This classification narrows the possibilities for causal mechanisms (Moscardelli and Wood, 2015). Moscardelli and Wood (2015) compiled morphometric parameters from more than 300 published accounts of mass failures in various environments to create a statistically valid assessment of the morphometrics of attached and detached mass failures. They measured length, area, volume and thickness and tested the correlations between the different parameters. For example, there were good correlations between area and volume and a very poor correlation between thickness and volume

(Moscardelli and Wood, 2015). The authors illustrated the use of such quantitative data to predict

the nature of mass failures in very data poor areas. Clare et al. (2018) aimed to standardize subaqueous landslide morphometry on a global scale in an attempt to create an interdisciplinary language for mass failures. They identified many parameters for morphometric studies, and although all are contingent on the resolution of data and scale of the mass transport complex, they provide a standard by which to measure the mass failures in our study area (Clare et al.,

2018, refer to Figure 3.1). We will utilize the approach of Moscardelli and Wood (2015) and

Claire et al. (2018) to place the mass failures within our study area in some global context.

Figure 8. 1 Examples of attached and unattached mass failures. The failures in this study appear to be both shelf and slope attached. Figure taken from Moscardelli and Wood (2008).

8.3 Observations Ten mass transport complexes have been identified in the gross interval between seismic

Horizon 1 and Horizon 7, all exhibiting different scales, morphometric dimensions and internal geometries. For the purposes of this study, we are only looking at MTCs that are fully contained within the seismic data. Simple morphometric dimensions (total length, width, deposit thickness, area, volume and slope gradient) were collected for 6 of the MTCs and recorded in Table 8.1.

We will not describe every MTC in detail but rather look at them as a process that is moving sediments from the outer shelf and upper slope to the deeper basin.

Table 8. 1 Morphometrics of MTCs in this study

In Package 1 alone, bounded below by Surface 1 and above by Surface 2, there are four mass transport complexes that occur in various proximities to the shelf. Some are exclusively occurring in the most basinward extents of the survey. In comparison to the overlying surfaces,

Surface 1 is the most affected by underlying faulting and topography that creates complexities in this surface. There are faults that cut through the surface and in some cases truncate MTCs.

Figure 20, a variance map windowed 90 m above Surface 1, shows the spatial extent of the four failures. Mass failures in this time frame range in size as shown in Table 8.1. MTC 4 on Surface

1 was not included in Table 8.1 because the failure is not fully contained within the dataset. No

MTCs are associated with Surface 2 and only occur beneath it in Package 1.

There are three mass failures that occur within Package 5, bounded below by Surface 5 and above by Surface 6. These MTCs are very complex and have many channelized geomorphic features at their margins that have been preserved. Windowed RMS amplitude maps were used to identify the extent of the mass failures and the complexity of channelization within. Three

MTCs were identified in P5 and named the P5MTC 1, P5MTC 2 and P5MTC3. P5MTC 1 and

P5MTC 2 appear to have several failures occurring within their spatial extents. The dimensions of P5MTC 1 were based on an RMS amplitude map windowed 40 m above Surface 5, and dimensions of P5MTC 2 and P5MTC3 were based on an RMS amplitude map windowed 75 m below Surface 6 (Figures 21, 22, and 23). Although the thickness of the MTC may exceed the window range, we wanted to avoid interference between P5MTC 1 and P5MTC 2. P5MTC 1 is not a continuous easily traceable failure and has isolated sections of chaotic material. In the chaotic intervals, reflector continuity is poor and shows signs of channelization (Figure 20).

P5MTC 1 is an interval of failures that extends across ~35 km of the shelf and slope. It appears that some of the failures are a result of simultaneous flows, however, individual failures are spatially separated by regions of intact strata. P5MTC 1 appears to have occurred isolated within channels from updip to downdip. P5MTC 2 appears to be a very wide (~65 km), elongate feature with an erosive, chaotic, low-amplitude basal surface and overlying high amplitude, continuous reflectors with intervals of channelization (Figure 8.3). The outermost margins of the failure are very steep and there is an impedance contrast between the failed material and the unfailed material outside the margins.

Figure 8.2 continued

Figure 8. 2 Variance map windowed 90 m above Surface 1. MTCs 1-4 are outlined and labeled in white and cross sections of the failures are in black. A) A’-A’’ shows the complexity within MTC 1. Extensional and compression faulting, and rafted blocks can be identified in the chaotic material. B) MTC 2 is downdip from MTC 1 and could be genetically related to MTC 1, occurring as a smaller failure overtop. MTC 2 shows normal faulting and thrusting. B) B-B’ shows the wedge-like geometry of MTC 3 that can be up to ~150 m in height. The material’s internal character is very chaotic and appears to have slid downslope before slumping. C) MTC 4 is the most basinward failure and C-C’ shows extensional and compressional faulting present.

Figure 8.4 shows the proximal to distal evolution of the P5MTC 1 and P5MTC 2. In some areas, the two mass transport complexes are in very close proximity and appear to be one complex, but they are spatially separated by regions of intact strata. P5MTC 1 and P5MTC 2 become very channelized in the basinward direction. P5MTC 3 is isolated to the northeastern portion of the survey. P5MTC 3 is a shelf failure confined by steep margins and has a semi-continuous basal surface with evidence of scour. The medium amplitude reflectors of the MTC display an internal geometry of both compressional and extensional faulting. Channel forms can be identified in the

NE region of P5MTC3.

8.4 Discussions and Conclusions The four MTCs in Package 1 have different internal geometries and scales (Figure 20).

MTCs 1, 3, and 4 appear to be genetically independent of each other, and MTC 2 could be genetically related to MTC 1 occurring as a smaller failure over top of MTC 1. MTC 1 shows the most internal deformation; internal extensional and thrust faulting were identified (Figure 20A).

MTC 1 was overlain with the first prograding reflectors of the clinoform packages as well as the

Surface 2 slope gullies. The occurrence of MTC 1 could herald the onset of margin progradation which can change the pressure regime of the shelf edge by loading it with sediment and causing it to fail. Sediment loading at the shelf could have induced MTCs 3 and 4. MTC 2 dislocates further basinward than MTC 1 and is much smaller in size than the other MTCs on Surface 1

(Figure 20B). As noted above, MTC 2 could be genetically related to MTC 1, possibly as a block or post emplacement portion of MTC 1 that was remobilized and transported further basinward in close temporal proximity to MTC 1 emplacement. MTC 3 is located in the lower foreset and forms a wedge-shaped failure between Horizons 1 and 2. The basinward extent of MTC 3 is bounded by a structural high influenced by underlying normal faulting. MTC 3 also occurs beneath the Surface 2 slope gullies, but their relationship is unclear. Although the most landward

Figure 8. 3 A) RMS amplitude map windowed 40 m above Surface 5. The outline of the P5MTC 1 interval is marked in white and has a cookie-bite upslope architecture due to the presence of channelization. The shelf break is denoted with a dashed white line. Beyond the shelf break, channelization occurs across the extent of the failure. B) A-A’ shows the channelization, channel thalwegs indicated with a black arrow, and the higher relief channel bank margins along the upslope margin of P5MTC 1

reaches of MTC 3 show evidence of internal deformation, basinward the deposits show little deformation. MTC 4 travels the furthest basinward and has chaotic, low-to-high amplitude reflectors that are deformed due to underlying normal faulting creating post emplacement instability. This failure is isolated to the clinoform toesets and occurs on a slope of < 1⁰. The

Package 1 MTCs appear to have been influenced by underlying faulting on the shallow slope gradients and were followed by a period of clinoform progradation. Temporally, Package 1

MTCs would be considered part of the rift sag phase infill. Along with extensive underlying faulting, the presence of gas escape and sediment loading at the shelf edge likely created shelf margin instabilities inducing margin failure.

P5MTC 1 and P5MTC 2 extend over a wide region of the study margin and are morphometrically wider along the margin strike than they extend basinward (length). Both exhibit a cookie-bite ornamented, upslope boundary and channels are present in the headwall

(Figures 8.3 and 8.4). P5MTC 1 appears to be an interval of several failures, however it is difficult to differentiate the individual flows without considerable focus and mapping of these units. The failures show intervals of channelization. Channelization is present on the margins as well as in the headwalls. Figure 8.3 shows the channel thalwegs, which are typically bright amplitudes, and the channel bank margins which are low-medium amplitude higher relief features. Figure 8.4` shows the basinward evolution of P5MTC 1 and P5MTC 2, channels are present throughout and the complex vertical relationship between the two failures is well displayed. There are several meandering channels within Package 5 that occur along the margins of the failures. Figure 8.6 shows the meandering nature of the system and cross sections display the northward migration of the channel over time. The channels do not show any evidence of subaerial exposure and are interpreted to be submarine meandering channels. The absence of lobate features downslope of the channels may point to a bypassing of sediments to depocenters beyond the extent of the dataset. The absence of deposits downslope of a larger channelized feeder belt could just as easily be due to bottom currents reworking and transporting sediment to areas beyond the obvious sink (Cathro et al., 2002). Based on the heightened impedance contrast

Figure 8. 4 A) RMS amplitude map windowed 75 m below Surface 6. The outline of P5MTC 2 is marked in white. The failure also shows an upslope cookie-bite architecture due to channelization at the failure margin. Three cross sections are labeled which capture the evolution of P5MTC 2 basinward. A-A’ goes across the margin channel and displays sharp failure margins. The chaotic basal surface is overlain by high amplitude continuous reflectors. B-B’ displays the same sharp channel margins, however there is a larger interval of failure material beneath the high amplitude reflectors. It also displays the vertical relationship between P5MTC 1 and P5MTC 2, which appear to be separated by planar strata in between. Moving towards the shelf break, C-C’ shows the initiation of channelization in P5MTC 1 and P5MTC 2.

between Packages 5 and 6 and the underlying carbonate-dominated clinoform packages as well as the drop in sea level and aggradation of the system, we hypothesize that more clastics were being transported into the system at this time. P5MTC 3 appears to be the result of a shelf failure, resulting in the deposition of a large, deformed shelf-sourced slump block (Figure 8.5). Package

6 also has a large expanse of slope channels that vertically aggrade (Figure 8.7). Packages 5 and

6 are the most aggradational packages in the system. They have distinct bright amplitudes along with an abundance of channels, which is due to the addition of clastics into the system as well as an overall increase in sediment supply during this time. The increase in sediment supply could produce shelf and slope instabilities causing large-and small-scale failures to occur. Based on the overall clinoform geometry and internal morphologies within Packages 5 and 6, there appears to be a relationship between the increased presence of channels and aggradation of the margin.

The mass failures appear to be temporally ubiquitous and are not limited to a certain package or clinoform geometry and do not appear to be closely tied to progradation/aggradation, however a few assumptions can be made. The mass failures that occur in Package 1 appear to be closely tied to the underlying Jurassic faulting that in some cases cut through Horizon 1 and sediment loading at the shelf. MTCs 1 and 3 are both overlain by slope gullies, but the relationship between the failures and gullies is unclear. The four mass failures are present over the majority of Package 1 indicating their widespread occurrence before the beginning of progradation, a possible indication that they are tied to some larger phases of tectonics in the basin. The Package 5 MTCs are the most complex in this study and show the close relationship between channelization and failure; failures redirect channels to their margins. There are also channels present outside the MTCs delivering sediment downslope.

RMS amplitude A’ Max B

Min

A B’

5 km

Figure 8. 5 A) RMS amplitude map windowed 75 m below Surface 6 shows P5MTC3 isolated to the northeast corner of the dataset. B) A-A’ shows the sharp western contact and obvious amplitude change as well as the chaotic nature of the basal surface and overlying high amplitude reflectors. C) P5MTC3 appears to be a result of a shelf failure; the material appears to have failed on the right side and is sliding downslope.

Figure 8. 6 A) RMS amplitude map windowed 75 m below Surface 6. B) Cross section A-A’ is an example of one meandering channel, and it is coming out of the plane and turning back to the right (direction of flow in and out of plane marked). The downlapping to the North (yellow lines) indicates the direction of channel migration. There are many more of these examples in Package 5.

Figure 8. 7 A-A’ shows the large expanse of vertically aggrading, high amplitude slope channels present in the foreset of Package 6. Although their geometries were not studied in detail, their vertical stacking lends itself to a more aggradational margin character. The slope channels are present across the extent of the survey in the foreset of Package 6. Horizons 6 and 7 are marked (H6 and H7), however Horizon 7 is dashed because it is harder to follow within the channel complexes.

CHAPTER 9

CONCLUSIONS

Deepwater sedimentation involves complex processes that deliver sediment from the shelf- to-slope-to-basinfloor. Deepwater processes and deposits have been studied and better understood in recent years due to their role in hydrocarbon exploration, advances in imaging and availability of deepwater cores (Posamentier and Kolla, 2003). There are various processes responsible for basinward deposition and reworking of sediment. These include, sediment-driven gravity flows (MTCs and turbidites), bottom currents, and fine-grained pelagic and hemipelagic passive sedimentation during slope abandonment (Prieto, 2016; Moscardelli and Wood, 2008;

Wynn and Stow et al., 2002; Rebesco et al., 2014). The transition from shelf-to-slope displays several different sediment sources, flow types and transitions, and can be either a dynamic area of sediment deposition, or a conduit for more basinward sediment movement. Three main geomorphologic processes responsible for transferring sediment are reflected in the shelf and slope deposits of this study area. 3D seismic data interpretation allowed for the identification of three geomorphologic features that provide clues to processes active along this margin; slope gullies and associated deposits, current-deposited sediment waves and wave fields, and mass failure deposits. The processes responsible for these deposits include gravity driven currents, off- shelf currents and along-slope and upslope bottom currents. The features identified in this study can serve as downslope conduits for basinward sedimentation (slope gullies and channels) and potential reservoir and seal quality deposits. These elements are also known to provide reservoirs along many margins in the world.

Figure 8. 8 Illustration showing currents that influence sediment dispersal, sediment sources and sediment deposits across the clinoform topset, foreset and toeset. Currents at the toe of slope, both along-slope (1) and upslope (2), rework sediments to form toe of slope sediment wave fields (D). Additionally, sediment wavefields form in the upper foreset, immediately basinward of the clinoform break (E) due to shelfal currents sweeping obliquely off the distal shelf and clinoform topsets (4). Gully systems (C) focus downslope gravity currents that generate turbidity currents (3) and can result in turbidity current sediments being deposited as wave fields in the toe of slope apron (not shown). Finally, both shelf-attached (A) and slope-attached (B) mass failures contribute large amounts of sediment basinward.

This study aimed to image the transition zone between the foresets and toesets of clinoforming margins at a seismic scale with a 3D dataset from offshore NW Australia. The clinoforms were deposited in ~6 packages covering ~40 Mya of sedimentation. They display different progradation/ aggradation architectures, seismic geomorphologic features and overall clinoform geometries. We compared these different spatial and temporal clinoforming systems and examined the controls on the foreset-to-toeset transition zone. In addition, we documented a variety of process-indicative morphologies (ie., slope gullies, sediment waves and mass failures), and discussed their effect on margins and the sediment budget. These deposits and processes are summarized in the schematic model in Figure 8.8. This study has arrived at several important conclusions regarding the geomorphic features identified in this 3D seismic dataset that have implications on the basinward movement of sediment along clinoforming margins.

• Slope gully formation processes can be deduced based on spacing, height, length, spatial

extent, slope gradient and type of gully infill

• Two types of slope gullies were identified occurring along this margin. These are the S2

and S4 gully systems

• S2 gullies occur on Surface 2 and are well-developed across the entire study area in the

middle-to-lower foreset but can extend into the toeset region. They are typically straight

to slightly sinuous, paralleling one another with bifurcation in the upslope areas. They are

relatively symmetric in cross-section showing steep margins and are overlain by a

continuous draping reflector. They are well developed for vertical extents up to ~70 m.

Their spatial extent and regular spacing indicate a line-source providing sediments to the

slopes. The lack of detectable infill and upslope bifurcating pattern would indicate that

by-pass through the gullies continued for a long period of time. The draping of the

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overlying reflector suggests gullies were not infilled following incision, but rather were

overlain with a passive drape during post-incision, post-bypass, slope abandonment. The

S2 gully initiation point could have large implications on hydrocarbon migration. The

gullies do not appear to be linked to the topsets, and therefore would not be a conduit for

updip hydrocarbon migration into the topset.

• S4 slope gullies are developed on Surface 4 but are isolated in the northwestern region of

the survey. They initiate on the upper foreset and extend into the lower foreset. They are

straight, short, low relief features with ~10 m of infill in the gully axis with minimal

inter-gully development. The irregular spacing and limited coverage of Surface 4 gullies

suggests that these gullies were either formed by a localized flow. The lower amounts of

incision into the underlying reflector could indicate fewer or less tractional flows through

the system or a less erosive underlying substrate. These gullies appear to be short-lived

and were quickly infilled by subsequent flows following incision and returned to

horizontal bedding shortly thereafter. S4 gullies initiate in the upper foreset which could

pose a higher risk for updip hydrocarbon migration, and a viable seal or stratigraphic

strap would need to be confirmed if the gullies were potential reservoirs.

• Two types of sediment waves; A-type and B-type, occur on Surface 4 with each

population showing different orientations, morphometrics, and spatial extent.

• A-type sediment waves are evenly spaced, elongate waves oriented oblique to the shelf

break, occurring in ~65 km swath in the upper foreset region of the slope. They are

typically broad, symmetric or slightly asymmetric downslope, low relief bedforms. Based

on the orientation and morphology of the A-type waves, and their “attached” appearance

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to the distal shelf, we propose that down-slope directed off-shelf currents are responsible

for their formation.

• B-type sediment waves show roughly shelf-parallel, straight-to-sinuous crests. They

frequently bifurcate in planform and are confined to a ~62 km swath extending from the

middle-lower foreset to upper toeset. They are upslope asymmetric and persist over a

vertical interval of ~20 m. The spatial extent and shelf-parallel orientation of the B-type

waves suggests that a downslope sheet-like turbidity current could have deposited these

sediment waves perpendicular to flow (Shumaker at al., 2017, Lonergan et al., 2013, Lee

et al., 2002, Wynn and Stow, 2002), followed by continuous bottom current reworking of

these deposits.

• Sediment waves could be a potential reservoir target based on the thickness, continuity

and spatial extent recorded for the A-type and B-type waves within this study.

• Mass failure deposits occur along the clinoform topset, foreset and toeset intervals of

several different clinoform packages within this study. Packages 1 and 5 contain mass

failures of different scales and internal geometries. The MTCs that occur in Package 1

precede clinoform progradation and are widespread across the survey, suggesting slope

instabilities due to underlying faulting and gas escape. In addition, MTCs 3 and 4 appear

to be contemporaneous with the onset of clinoform progradation, and sediment loading at

the shelf edge likely created shelf margin instabilities inducing margin failure

• Package 5 MTCs have an abundance of channels present at the failure margins. During

this same time, an influx of clastic sediment appears to be introduced into the system. An

increase in sediment supply could produce shelf and slope instabilities causing large-and

small-scale failures to occur.

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• The dynamic transition zone between supercritical and subcritical flow that takes place in

the lower foreset-to-toeset region can be postulated based on the presence of upslope

asymmetric B-type sediment waves. Upslope asymmetry and migration has been reported

as a key indicator of supercritical bedforms in comparison to subcritical bedforms

(Cartigny et al., 2014). Cyclic steps also show in an increase in symmetry downslope on a

lower gradient, which we observe in the B-type sediment waves as the slope gradient

decreases. However, we cannot be certain of the origin of the flow or presence of a

hydraulic jump because we are not able to see internal bedforms and structures.

• The flow transition zone can be hypothesized with confidence solely based on seismic

data; however, the integration of smaller scale data could provide more concrete evidence

for flow type and increase the confidence of the location of this transition zone.

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APPENDIX A

Locker Shale (early/ Middle Triassic): Marine claystone and siltstone with minor interbedded limestone and sandstone deposited during a regional marine transgression (Hocking, 1987). The

Locker shale grades into the overlying Mungaroo formation.

Mungaroo Formation (Middle/ Late Triassic): The thick fluvio-deltaic sequence of sandstone, siltstone and claystone with minor coal deposits. The northward prograding fluvio-deltaic system covered most of the offshore NCB during a period of abundant sediment supply. The Mungaroo formation is a Type III gas-prone source rock and hosts giant gas accumulations (Geoscience

Australia, 2014).

Brigadier Formation (Late Triassic): Thinly bedded shelfal siltstone, claystone and marl deposited on the shelf during rapid subsidence due to the continental breakup (Hocking, 1987).

Deposition of thin sandstones along the Rankin Platform is known as the North Rankin

Formation (Seggie et al., 2007). The Brigadier can serve as both a gas source and reservoir

(Geoscience Australia, 2014).

Murat siltstone (Late Triassic/ Early Jurassic): Deposited in a regressive period in deep basinal areas (Hocking, 1987).

Legendre Formation (Toarcian): Regressive deltaic sandstone with an upper and lower member.

The lower member is a thick, deltaic sandstone prograding west and southwest from the De Grey

Nose (Figure 3). The upper member is a near shore sandy deposit. The Legendre can serve as a potential source rock (Geoscience Australia, 2014).

Calypso Formation (Oxfordian): Interbedded claystone and sandstone deposited during the main phase of syn-rift (Geoscience Australia, 2014).

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Eliassen Sandstone (Oxfordian): Massive sandstones interbedded with minor claystones basin floor fan deposit (Geoscience Australia, 2014).

Dingo Claystone (Oxfordian): Thick, deep marine sequence deposited during rapid tectonic subsidence.

Angel Formation (Tithonian): Turbidite sand formed by west to southwest progradation of fluvial and shallow marine sands from the De Grey Nose (Hocking, 1987). It is the main oil and gas reservoir in the Dampier sub-basin (Geoscience Australia, 2014).

Forestier Claystone (Tithonian/ Berriasian): Interbedded silty claystone and claystone with argillaceous sandstone and minor limestone and dolomite stringers. This unit unconformably overlies the Angel Formation and was deposited in a shelfal marine environment (Apache

Energy, 2004).

Birdrong Sandstone (Valanginian): Inner shelf transgressive sands with highly glauconitic

Mardie Greensand. The sands are well sorted and have up to 20 percent porosities (Barber,

1994).

Muderong Shale (Valanginian-Aptian): Basin-wide transgressive interbedded silty claystone and calcareous claystone. The Muderong shale serves as the regional seal in the NCB (Apache

Energy, 2004 and Geoscience Australia, 2014).

Windalia Radiolarite (Albian): Basin-wide transgressive radiolarian siltstone. This unit unconformably overlies the Muderong shale (Hocking et al., 1987).

Haycock Marl (Cenomanian-Coniacian): Two argillaceous calcilutite and marl units separated by a thin claystone interval. It is the more calcareous lateral equivalent to the Gearle Siltstone

(Hocking et al., 1987).

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