Timing and Mechanisms Controlling Evaporite Diapirism On

TIMING AND MECHANISMS CONTROLLING EVAPORITE DIAPIRISM ON

ELLEF RINGNES ISLAND, CANADIAN ARCTIC ARCHIPELAGO

Jennifer Anne Macauley

A thesis submitted in conformity with the requirements

for the degree of Master of Applied Science

Graduate Department of Geology

University of Toronto

Ellef Ringnes Island, Canadian Arctic Archipelago

M.A.Sc. thesis by Jennifer Anne Macauley

Department of Geology, University of Toronto, 2009

Abstract

This thesis investigates the timing and mechanisms involved in the formation of evaporite piercement structures on Ellef Ringnes Island, Canadian Arctic Archipelago. The study includes the interpretation of industry seismic reflection and borehole data to characterize the geometry of the domes, 1D backstripping of wells to investigate the role of tectonic influences on diapirism, and analogue modelling to better understand the mechanisms that drive diapirs with dense anhydrite caps. I propose that basement structures played a significant role in the formation of evaporite domes by triggering and directing salt movement. The domes developed during the Mesozoic by passive growth driven by the differential loading of salt on adjacent fault blocks, which led to their present day asymmetric geometries. Diapir growth rates in the Mesozoic were closely linked to the rate of sedimentation, which was greatly influenced by the amount of accommodation space provided by tectonic subsidence of the basin.

ii Acknowledgements

I first and foremost would like to thank my supervisor Alexander Cruden for the support and guidance he has provided over the last two years. I am very grateful for the many amazing opportunities this project has given me during my studies at the University of Toronto.

I would like to thank Tom Brent, Keith Dewing and Chris Harrison from the Geological Survey of Canada (Calgary) for their helpful insight on the project. Tom Brent, in particular, spent several weeks with me gathering and processing data that was crucial for my project. I enjoyed the many lunches we spent during my visits to Calgary.

I would like to thank Randell Stephenson for being a great host in Amsterdam. He was encouraging throughout this process and his support and expertise on the Sverdrup Basin is greatly appreciated. I would also like to thank him for the bike he lent me during my stay in Amsterdam.

I acknowledge the members of my supervisory committee, Andrew Miall and Bernd Milkereit, who have been highly supportive throughout my project.

I thank Christoph Schrank for being an awesome friend and mentor when Sandy was unavailable. He would always provide help when asked, above and beyond my expectations. I enjoyed all the random nights we spent out in Toronto.

Becky Ghent helped me with countless posters and math problems. I most appreciated her company, however, on many de-stressing tea breaks.

I thank David Boutelier for reading over multiple drafts of my thesis and always listening whenever I needed to sort out my thoughts.

Last but not least, I would like to thank all the friends and family that distracted me from my studies when I needed it most.

iii

This research was financed with an OGSST (Ontario Graduate Scholarship in Science and Technology) awarded to J. Macauley, and NSERC (Natural Sciences and Engineering Research Council Discovery) and Geological Survey of Canada Unconventional Oil and Gas Research and Development grants awarded to A. Cruden. Access to seismic and borehole data was provided by ConocoPhillips, Imperial Oil Resources and PetroCanada.

iv Table of Contents

Abstract ii Acknowledgments iii Chapter 1: Introduction 1 1.1 Background on the development of the Sverdrup Basin 1 1.2 Salt tectonics 3 1.2.1 Salt properties 3 1.2.2 Mechanics of salt flow 4 1.2.2.1 Differential loading 7 1.2.2.2 Initiation and growth of diapirs by regional tectonics 8

1.2.3 Allochthonous evaporite structures in the Sverdrup Basin 10 1.2.4 Timing and mechanisms of halokinesis in the Sverdrup Basin 11

1.3 Study Area: Ellef Ringnes Island 14 1.3.1 Stratigraphy 14 1.3.2 Mafic intrusions 16 1.3.3 Structural geology 17 1.3.4 Evaporite piercement structures 17

1.4 Research Objectives 18 1.4.1 Chapter Two: Structural and geometrical characterization of evaporite piercement structures on Ellef Ringnes Island 19 1.4.2 Chapter Three: Basin subsidence modelling 19 1.4.3 Chapter Four: Analogue modelling of diapirism 19

Chapter 2: Structural and geometrical characterization of evaporite piercement structures on Ellef Ringnes Island 20 2.1 Introduction 20 2.2 Methods 21 2.2.1 Data 21 2.2.2 Seismic processing 22 2.2.3 Analysis of borehole data 23

v 2.2.4 Interpretation of formation tops 24 2.2.5 Isopach maps 25 2.3 Results 29 2.3.1 Dome asymmetry 29 2.3.2 Salt canopies at Hoodoo Dome 30 2.3.3 Deformation and unconformities within overburden adjacent to piercement structures 32 2.3.4 Regional variations in formation thicknesses from isopach maps 34

2.4 Discussion 36 2.4.1 Basement fault-initiated and controlled diapirism 36 2.4.2 The timing of diapirism in the Sverdrup Basin 41

2.5 Conclusion 43

Chapter 3: Basin subsidence modelling 45 3.1 Introduction 45 3.2 Methods 48 3.2.1 Backstripping 48 3.2.2 Data 49 3.2.3 Input parameters for backstripping 50

3.3 Results 52 3.3.1 Regional and local changes in tectonic subsidence rates 52 3.3.2 Contour plots of 1-D subsidence curves through time 56 3.3.3 Limitations due to data quality and resolution 56

3.4 Discussion 59 3.4.1 Relation between tectonic subsidence trends and diapir growth 59 3.4.2 Axel Heiberg subsidence 65

3.5 Conclusion 65

vi Chatper 4: Analogue modelling of diapirism 67 4.1 Introduction 67 4.1.1 The role of anhydrite in diapirism and other tectonic settings 67 4.1.2 Anhydrite piercement structures in the Sverdrup Basin 68 4.1.3 Previous analogue and numerical modeling of diapirism involving anhydrite 71

4.2 Methods 73 4.2.1 Model set-up 73 4.2.2 Materials and scaling 75

4.3 Results 79 4.3.1 Limitations due to boundary conditions 79 4.3.2 Problems localizing sheared diapir margins 83

4.4 Discussion 85 4.4.1 The influence of material properties on the development of sheared dome margins 85 4.4.2 Localization of salt flow and the effect of boundary conditions 87 4.4.3 The affect of buoyancy on diapirism within the Sverdrup Basin 89 4.4.4 Insight on the development of Hoodoo Dome canopies 91

4.5 Conclusion 93

Chapter 5: Summary and final remarks 95 5.1 Summary of mechanics and timing of salt growth on Ellef Ringnes Island 95 5.1.1 Initiation of Diapirism 95 5.1.2 Diapirism in the Triassic 96 5.1.3 Diapirism in the Jurassic 97 5.1.4 Diapirism in the Cretaceous and Early Tertiary 98

5.2 Implications for oil and gas exploration 99 5.2.1 Petroleum generation in the Sverdrup Basin 99 5.2.2 Potential for future oil and gas exploration 101

5.3 Future Research 104

vii

References 105 Appendices 114

viii List of Tables

Chapter 2: Structural and geometrical characterization of evaporite piercement structures on Ellef Ringnes Island Table 2-1. Parameters and functions used for time-depth conversion 28

Chapter 3: Basin subsidence modelling Table 3-1. Parameters used for exponential depth laws in backstripping 49

Chapter 4: Analogue modeling of diapirism Table 4-1. Scaled parameters and ratios for model and nature 76 Table 4-2. Summary of physical experiments 80

List of Plates

Plate 1. Regional seismic profiles Plate 2. G(F) series seismic profiles Plate 3. I(F) series seismic profiles (I)

Plate 4. I(F) series seismic profiles (II)

ix List of Figures

Chapter 1: Introduction Fig. 1-1. Regional extent of the Otto Fiord evaporite facies 2 Fig. 1-2. Physical properties of rock salt in comparison to other lithologies 3 Fig. 1-3. 3D schematic diagram of various shapes and sizes of salt structures 5 Fig. 1-4. Diagram illustrating conditions required for active piercement 6 Fig. 1-5. Diagram illustrating conditions of hydraulic head-gradient analysis for salt tectonics 7 Fig. 1-6. Three stages of diapirism and their characteristic structures 9 Fig. 1-7. Cross-sectional profiles of passive diapirism at various aggradation rates 10 Fig. 1-8. Map and cross-section of the wall-and-basin structures on Axel Heiberg 13 Fig. 1-9. Map of Ellef Ringnes Island 15 Fig. 1-10 Stratigraphy of Ellef Ringnes Island 16

Chapter 2: Structural and geometrical characterization of evaporite piercement structures on Ellef Ringnes Island Fig. 2-1. Index map of seismic reflection and borehole data for Ellef Ringnes Island 22 Fig. 2-2. Synthetic well log for Hoodoo Dome H-37 24 Fig. 2-3. Instantaneous velocity with depth signature of the Deer Bay Formation 26 Fig. 2-4. Line drawing of a cross-section through Dumbells Dome 30 Fig. 2-5. Spatial extent of interpreted canopies and a shear zone at Hoodoo Dome 31 Fig. 2-6. Cretaceous unconformities within sediments adjacent to domes 33 Fig. 2-7. Isopach maps of Mesozoic formations 35 Fig. 2-8. Examples of basement initiated diapirism 37 Fig. 2-9. Schematic of a drape monocline 37 Fig. 2-10. Schematic of the development of asymmetric domes 39 Fig. 2-11. Sedimentary thicknesses and crustal signatures derived from Bouguer gravity data 40 Fig. 2-12. Schematic drag profile adjacent to a passive diapir 42

x Chapter 3: Basin subsidence modelling Fig. 3-1. Tectonic subsidence curves for various locations across Sverdrup Basin 47 Fig. 3-2. Locations of wells used in the backstripping study 51 Fig. 3-3. Tectonic subsidence curves backstripped from 200 Ma 53 Fig. 3-4. Tectonic subsidence for select wells dating back to 245 Ma 54 Fig. 3-5. Comparison of local and regional tectonic subsidence signatures 55 Fig. 3-6. Contour plots of tectonic subsidence 57 Fig. 3-7. Errors for tectonic and basin subsidence at Sutherland O-23 58 Fig. 3-8. Sketch showing the subsidence signature of basin uplift and erosion 59 Fig. 3-9. Apatite fission track ages for uplift in the Sverdrup Basin 61 Fig. 3-10. Interpreted tectonic phases and periods of diapirism 62

Chapter 4: Analogue modeling of diapirism Fig. 4-1. Buoyancy force for different diapiric heights of anhydrite and salt 68 Fig. 4-2. Gamma ray and sonic profiles of Hoodoo L-41 well iii Fig. 4-3. Photos of the anhydrite cap and margins at Dumbells Dome 70 Fig. 4-4. Cross-section of a centrifuge model of salt diapirism 71 Fig. 4-5. Profile and line drawing of 2D analogue experiments of salt diapirism 72 Fig. 4-6. Davison et al.’s 2D injection set-up 74 Fig. 4-7. Diagrams of the experimental apparatus 75 Fig. 4-8. Material properties of the visco-elastic plastic analogue anhydrite layer 78 Fig. 4-9. Time progression of Experiment 6 82 Fig. 4-10. The development of cracks within the anhydrite layer 83 Fig. 4-11. Photo of asymmetric dome growth in Experiment 8 84 Fig. 4-12. Development of "canopies" in Experiment 8 85 Fig. 4-13. Sketches on the development of canopies at Hoodoo Dome 92

Chapter 5: Conclusion Fig. 5-1. Schematic diagrams on the development of evaporite piercement structures 97 Fig. 5-2. Location and stratigraphic distribution of known hydrocarbon pools 100 Fig. 5-3. Examples of canopies capping dipping sedimentary beds 103

xi List of Appendices

Appendix 1. Seismic reflection surveys 114 Appendix 2. Instantaneous velocity-depth curves for Mesozoic formations 116 Appendix 3. Time-depth conversion MATLAB code 118 Appendix 4. Isopach maps derived from depth conversion of interpreted seismic horizons 121 Appendix 5. Individual well input files used in 1D backstripping 126

xii 1 Introduction

1.1 Background on the development of the Sverdrup Basin

Ellef Ringnes Island is located within the Sverdrup Basin, a pericratonic trough that formed along a northeast-southwest oriented continental rift zone. The steep-sided basin contains up to 13 km of Carboniferous to Tertiary strata at the basin axis (Balkwill, 1978). During initial rifting, non-marine conglomerates and sandstones of Serpukhovian to Bashkirian age were deposited within horst and graben basement structures along the basin flanks (Davies & Nassichuk, 1991). Within the axial region of the basin, subsidence exceeded deposition as marked by the gradual progression from non-marine to deep water deposits (Davies & Nassichuk, 1991). The Carboniferous Otto Fiord Formation, mostly composed of interbedded limestone and anhydrite, is the initial submarine transgressive fill of the developing rift basin (Davies & Nassichuk, 1975). Halite horizons of the Otto Fiord Formation were restricted to the deepest portions of the basin in the Axel Heiberg and Barrow sub-basins (Fig. 1-1). The sub-basins are thought to have been separated by a carbonate platform (Meneley et al., 1975; Davies & Nassichuk, 1975; Balkwill, 1978; Embry, 1991); however the existence of this platform has never been confirmed by deep drilling. Sedimentation was briefly interrupted and Upper Carboniferous to Upper Permian rocks were folded and faulted during the Melvillian Disturbance (Thorsteinsson & Tozer, 1970). The Late Paleozoic event is primarily recorded in the southern margin of the basin (Thorsteinsson, 1974; Balkwill, 1978) and is attributed to the formation of pull-apart basins associated with syn-extensional transcurrent faults (discussion in Davies & Nassichuk, 1991). The basin evolved rapidly in the early Mesozoic during the post-rifting thermal stage of basin subsidence (Stephenson et al., 1987) with continual burial of the basin axis by thick deposits of siltstone and shale. During the middle Mesozoic the basin slowly accumulated terrigenous clastic sediments, which overflowed the former basin margins in the late Mesozoic (Balkwill, 1978). Continuous subsidence and sedimentation was interrupted in the Cretaceous with uplift that produced numerous widespread unconformities along with normal faulting and magmatism (Embry, 1991). The Hauterivian to early Turonian Sverdrup Basin Magmatic Province consists mainly of intrusive sills and dykes, flood basalts and pyroclastic flows

1 restricted to the northeastern portions of the basin (Embry & Osadetz, 1988; Villeneuve & Williamson, 2006). The magmatic activity is believed to have originated from a mantle plume or hot spot located north of Ellesmere Island that coincided with the development of the oceanic Canada (also referred to as Amerasian) Basin (Embry & Osadetz, 1988). In the final stage of development, the basin was tectonically modified and inverted during the early Tertiary Eurekan orogeny due to the collision between Greenland and Ellesmere Island (Balkwill, 1978; Miall, 1984). The Eurekan orogeny was responsible for the uplift and erosion of the Sverdrup rim and the development of a fold-and-thrust belt on Ellesmere and eastern Axel Heiberg Islands (Fig. 1-1). Regional uplifts or arches are also significant features of the orogeny and are attributed to crustal-scale folding (Stephenson et al., 1990).

Figure 1-1. Regional extent of the Otto Fiord evaporite facies in the Sverdrup Basin. Diapirs are highlighted in black with the majority situated in central Axel Heiberg (modified from Davies & Nassichuk, 1975).

2 1.2 Salt tectonics

1.2.1 Salt properties

Once effective porosity is lost from burial to depths of 100 to 200 m, salt is relatively -3 incompressible with an average density of 2200 kg ּm , making it less dense then most carbonates and compacted siliciclastic rocks. A density inversion typically occurs from de- watering of siliciclastic sediments during burial at approximately 1200 m (Fig. 1-2a; Warren, 2006). Burial of at least 1600 m (more typically 3000 m) is required before the average density of the entire package of overlying sediments exceeds that of salt (Hudec & Jackson, 2007). Salt is also an effective seal to fluid migration and has a high thermal conductivity (Fig. 1-2b).

Figure 1-2. Physical properties of rock salt in comparison to other lithologies (Warren, 2006). a) Density variation with burial. b) Thermal conductivity. c) Viscosity.

The rheology of salt is strongly influenced by temperature and the presence of water. Dry salt deforms by dislocation creep while diffusion creep can activate with as little as 0.05 wt. % intercrystalline brine (Urai et al., 1986). Triaxial deformation experiments conducted by Urai

3 et al. (1986) confirmed that samples with ~0.05 wt. % brine (dry) and entirely saturated samples (wet) deformed by a power law creep with a power exponent n >5 at strain rates faster than 10-7 s-1. A power law creep rheology implies that the strain rate of salt is non- linearly proportional to the differential stress as governed by the following equation:

n & = Aσε (1) whereε& is the strain rate, σ is the differential stress and A is a material constant dependant on temperature and grain size. At slower strain rates, however, the same samples deformed with a power law exponent of n= 2 (dry) and n=1 (wet). The samples subjected to the lower strain rates formed similar recrystallization textures as is observed in nature through the enhancement of solution-precipitation processes. Dynamic recrystallization was strongly enhanced with increasing amounts of brine, effectively weakening the salt. As most natural salt contains 0.1-1 wt. % intercrystalline brine, salt can be approximated as having a Newtonian viscous behaviour (n ~ 1) under strain rates that are typical of geological settings (Weijermars et al., 1993). Weakening due to increasing moisture content dramatically lowers the effective viscosity of rock salt from 1019 Pa ּs for dry salt to 1013 Pa ּs for saturated salt (Warren, 2006). Due to the higher moisture content provided by meteoric water, salt is typically less viscous and flows at faster rates at the surface.

1.2.2 Mechanics of salt flow

Due to its low density and negligible yield strength, salt often accumulates most of the total strain in a tectonic system and therefore strongly influences the style of deformation in many basins. This allows salt to behave as a lubricant or detachment surface, effectively allowing the decoupling of overlying sediment from the basement. Salt structures evolve from concordant, low amplitude structures such as salt anticlines, rollers and pillows to discordant high amplitude intrusions such as diapirs, stocks and salt walls (Fig. 1-3; Jackson and Talbot, 1986). Piercement and extrusion of salt at the surface can lead to the formation of salt glaciers, sheets, tongues or canopies.

Figure 1-3. 3D schematic diagram illustrating different shapes and sizes of salt structures. Increasing structural maturity and size of the salt structures are shown for (a) linear structures and (b) point sources. (Figure taken from Hudec & Jackson, 2007, simplified from Jackson & Talbot, 1991).

The rate of salt flow is highly dependant on many factors, such as basement control, the thickness and geometry of the source layer and the rate and total amount of regional extension, compression, sedimentation and erosion. Deformation of rock salt occurs over a large range of strain rates, from 10-8 to 10-16 s-1 (Jackson & Talbot, 1991; Jackson & Talbot, 1986). Average strain rates of 10-14 to 10-15 s-1 are estimated for gravity driven diapirism, while diapiric growth assisted by tectonic compression can occur at rates that are a several magnitudes faster (Jackson & Talbot, 1986). Average vertical growth rates of diapirs range from 0.01 mm/yr to a maximum of 0.5 mm/yr during initial stages of growth and in diapirs subjected to tectonic compression (discussion in Koyi, 2001; Jackson & Talbot, 1991). The development of diapirs is often not steady-state, however, and many diapirs experience multiple surges in growth (Jackson & Talbot, 1986).

Salt diapirism was originally proposed by Arrhenius (1912) to be driven by the buoyant rise of salt into a uniform overburden, similar to the Rayleigh-Taylor instability of two viscous materials with an inverted density contrast. Buoyancy forces alone, however, are not able to initiate or drive salt diapirism as the strength of brittle overburden often resists flow (e.g.

5 Jackson & Talbot, 1986; Daudré & Cloetingh, 1994). The relative density in most tectonic settings is therefore of less importance than the relative strength of the salt and the overburden. Active piercement is often simplified using the concept of the hydraulic head of a fluid pressurized by the weight of overburden. Active piercement occurs wherever the pressure head of the fluid exceeds the strength of the roof rock, which is dependant on the thickness of the overburden. The minimum thickness that can maintain sufficient strength to hold back buoyant and pressurized salt is defined by the piercement threshold (Fig. 1-4; Vendeville & Jackson, 1992a & b).

Figure 1-4. Schematic diagram illustrating the conditions required for active piercement with variations in thickness from (a) salt ridges or (b) topographic troughs. The diagram illustrates how the ridge or trough needs to pass the piercement threshold for salt to be able to actively pierce the overburden. (c) Active piercement is more easily facilitated by extending the overburden so that it is thinned and weakened by normal faulting (Vendeville & Jackson, 1992a).

Piercement and mobilization is more easily facilitated by a combination of lateral (tectonic) and gravity forces (Jackson & Talbot, 1986; Vendeville & Jackson, 1992a; Daudré & Cloetingh, 1994). Due to the high thermal conductivity and thermal expansivity of salt, diapirism can also be enhanced through thermal convection when subjected to high geothermal gradients and elevated near-surface temperatures (Talbot, 1978).

6 1.2.2.1 Differential loading Differential (also referred to as gravitational) loading involves lateral variations in the thickness, density or strength of the overburden. This mechanism is most effective during early stages of burial when the density contrast between unconsolidated and compacted sediments is highest (Jackson & Talbot, 1986). Lateral variations are often associated with a sedimentary facies change or prograding sequences such as deltas, alluvial fans, coral reefs and volcanic deposits. Differential loading from prograding sediments, for example, leads to the evacuation of salt to areas of lower pressure (Fig. 1-5a). Lateral variations from prograding sediments often occur on continental slopes, which has the added component of an elevation head gradient in combination with a pressure head gradient exerted on the salt layer (combination of Fig. 1-5a & b). In this case, the salt acts a detachment surface that allows a thick package of overlying sediments to translate down slope through the process of gravitational sliding. This process is characteristic of salt tectonics along passive margins such as the Gulf of Mexico.

Figure 1-5. Examples of hydraulic head- gradient analysis for salt tectonics. (a) An example of the principle driving force for differential loading from the presence of a pressure head gradient, but no elevation head gradient. (b) An example of how an elevation head gradient produces gravitational sliding. (c) If neither a pressure nor elevation head gradient exists, the salt remains at rest despite the basement boundary conditions (Hudec & Jackson, 2007).

7 1.2.2.2 Initiation and growth of diapirs by regional tectonics In most halokinetic systems, basin extension and compression are required to initiate and promote salt flow. Regional extension has shown to be the most effective mechanism for the initiation of diapirism in basins such as the North Sea, Nordkapp and Dnieper-Donets Basins.

Extensional salt tectonics The initiation and development of diapirs are generally described by a three stage process involving reactive, active and passive stages of growth. Vendeville & Jackson (1992a) highlighted the role of thin-skinned regional extension in initiating and promoting diapir growth. Regional extension forms normal faults in the thick brittle overburden, leading to the most effective avenue for piercement by weakening and thinning of the overlying sedimentary units. The process of reactive diapirism was defined by Vendeville & Jackson (1992a&b) as the process of early stages of piercement where denser half-grabens and grabens sink into the pressurized salt source layer (Fig. 1-6a). During this stage, the rotation of fault blocks allows for thicker accumulations of sediments adjacent to the developing diapir. As the diapir matures and the roof rock is significantly thinned, the build-up of pressure within the salt layer and buoyancy effects leads to the forceful upheaval and shouldering aside of roof rock during active diapirism (Fig. 1-6b; Nelson, 1991; Schultz-Ela et al., 1993). During this stage, the rate of salt growth is higher than the average rate of subsidence and sedimentation, resulting in the upward drag of adjoining strata and often extrusion of salt at the surface. Once the pressure is relieved, the diapir further matures during the passive stage (Fig. 1-6c), previously referred to as downbuilding (Barton, 1933). At this stage, salt has pierced the surface and growth is driven by the response of the overburden sinking into the source layer. The shape of the diapir during this stage of growth is strongly governed by the balance between the sediment aggradation rate and net diapiric rise (i.e. rise including effects from dissolution and erosion of salt) (Fig. 1-7; Jackson et al, 1994; Koyi, 1998).

Figure 1-6. Three stages of diapiric growth and their characteristic structures. P, V, B represent stresses from pressure, salt viscosity and overburden brittle strength respectively (Jackson et al., 1994).

Compressional salt tectonics Laterally shortened salt structures are found in inverted rift basins, convergent plate boundaries or at the down-dip edge of passive margins. In the absence of pre-existing diapirs, thin salt often behaves as a décollement for thrust sheets, such as in the Zagros Mountains, Angola and in the Pyrenees. Regional convergence often modifies the shape of pre-existing diapirs, however does not lead to the initiation of new salt diapirs. In fact, shortening thickens and strengthens the overlying sediments, reducing the ability to initiate new diapirs (Hudec & Jackson, 2007).

Figure 1-7. Cross-sectional shapes of passive diapirs controlled by the relative rates of net diapir rise and sediment aggradation. (a) Where diapir rise rate exceeds aggradation rate, diapirs widen and adjacent sediments are dragged further upward. (b) Where diapir rise rate matches aggradation rate, diapir margins are near vertical. (c) Where diapir rise rate is less than aggradation rate, diapirs narrow upward and are onlapped/buried by sediments (Hudec & Jackson, 2007).

Shortening of pre-existing diapirs often leads to the nucleation of salt cored folds and thrust sheets which then propagate laterally. Also, pre-existing diapirs often link together to form anticlines or salt walls that are oblique to the convergence direction. If diapirs are buried by a sudden pulse of rapid sedimentation, regional compression can also aid in the triggering of active diapiric growth.

1.2.3 Allochthonous evaporite structures in the Sverdrup Basin

Approximately 100 piercement structures occur in the axial region of the Sverdrup Basin, the majority of which are located on Axel Heiberg and the Ringnes Islands (Fig. 1-1; Thorsteinsson, 1974). Large, concentric diapirs are most common in the western region of the Sverdrup Basin on Ellef Ringnes Island, Sabine Peninsula (Melville Island) and western Axel Heiberg Island. Except for a few cases where salt is exposed (e.g. Hugon & Schwerdtner, 1982), the domes at surface are composed of gypsified anhydrite caps cored by halite (Heywood, 1955, 1957; Kranck, 1961; Gould & DeMille, 1964; Hoen, 1964; Schwerdtner & Clarke, 1967). From their first discovery, the anhydrite caps were recognized to be of primary sedimentary origin and were most likely mobilized along with the underlying salt (Heywood, 1955, 1957). The diapiric evaporite units were first correlated by

10 Thorsteinsson & Tozer (1957) with the Carboniferous Otto Fiord Formation (Nassichuk & Davies, 1980). The anhydrite caps are estimated to range from 200 to 800 m in thickness from measured sections and gravity studies (Panarctic Oils Ltd., 1972; Schwerdtner & Clarke, 1967; Spector & Hornal, 1970).

In the central and eastern regions of Axel Heiberg Island, the style of salt structures differs from those of the western region of the Sverdrup Basin with the occurrence of linear to crooked walls separated by broad synclinal basins that formed from the piercement of anhydrite and salt into the crests of narrow anticlines (Thorsteinsson, 1974; van Berkel et al., 1984). This regional structural style is known as the “wall-and-basin structure” (WABS) and contains the densest cluster of diapirs in the basin (van Berkel et al., 1984; Jackson & Harrison, 2006). Evaporites are also present within regional scale thrust faults such as the Stolz Fault Zone and North Mokka Fault located on eastern Axel Heiberg Island (van Berkel et al., 1983).

1.2.4 Timing and mechanisms of halokinesis in the Sverdrup Basin

The development of the diapirs and salt structures on Axel Heiberg Island was originally believed to be related to the Tertiary Eurekan orogeny (Fortier et al., 1963; Thorsteinsson & Tozer, 1960), however subsequent field studies suggest that the orogenic event merely overprinted pre-existing features, deforming them into the tight anticlinal folds of the WABS region, and that diapirism was occurring by at least the Cretaceous. Evidence supporting diapirism in the Cretaceous includes thinning of the Hassel Formation towards the Isachsen Dome (Gould & DeMille, 1964) and the intrusion of Late Cretaceous dykes and sills after many domes were emplaced (Hoen, 1964; Schwerdtner & Clarke, 1967). Allochthonous evaporites in the WABS region are overlain concordantly by the Cretaceous Isachsen, Christopher and Hassel formations, suggesting that there was a long period of diapir growth that pierced the Upper Paleozoic and Lower Mesozoic units prior to cessation and re-burial of the structures (van Berkel et al., 1984; Schwerdtner & Osadetz, 1983; Jackson & Harrison, 2006). Stott (1969) also used considerable stratigraphic evidence on Ellef Ringnes Island to suggest that certain diapirs were rising as early as the Late Jurassic. Modification of the Late Triassic Heiberg Group near large domal diapirs provides the earliest constraints on diapiric

11 growth (Schwerdtner & Osadetz, 1983; van Berkel, 1989; Stephenson et al., 1992), however, Stephenson et al. (1992) have proposed that the domes rose slowly before, more than 50 Myr prior to deposition of the Heiberg Group. Schwerdtner & Osadetz (1983) suggest that diapirism started as early as the uppermost Carboniferous and lowermost Permian when the anhydrite layer existed as an “anhydrite mush” with a lower density and effective viscosity.

A common mechanism proposed for the initiation and localization of diapirism includes differential loading on top of faulted basement blocks (Schwerdtner & Osadetz, 1983; Stephenson et al., 1992). Jackson & Harrison (2006) also mention the possibility of rift- related extension as a cause of diapir initiation. Gould & DeMille (1964) and Balkwill (1978) suggest that during most of the Mesozoic, passive diapir growth occurred due to the lithostatic load of clastic overburden. Balkwill (1978) went on to suggest that salt was originally deposited on the basin shelf and subsequently squeezed (coincident with gravitational sliding) toward the basin axis by thick prograding deposits originating from the southwest and from the eastern Ellesmere region. Compressional intraplate stresses in the basin centre resulting from flexural bending of underlying strata induced by sediment loading and far-field tectonic events may have accelerated diapirism in the Mesozoic (Stephenson et al., 1992). Passive diapirism ceased by the Late Cretaceous as evidenced by the lack of deformation in cross-cutting mafic sills and dykes. Locally on central Axel Heiberg Island, extrusion onto the surface is proposed to have occurred in relation to a widespread unconformity during the Hauterivian (130 Ma; Jackson & Harrison, 2006). The extrusion of multiple diapirs coalesced into a broad canopy that was subsequently buried by onlapping Isachsen and younger formations. Salt flow was reactivated and exhumed in the Tertiary by the Eurekan orogeny, producing second generation salt structures (WABS) originating from the allochthonous salt layer (Fig. 1-8). The geometry and wavelength of the WABS region reflect the occurrence of a salt source layer at shallower stratigraphic levels than the autochthonous salt deposit (Jackson & Harrison, 2006). Compression related to the Eurekan orogeny also lead to other features such as buckle folds and lubricated thrust faults due to the decoupling of strata along less competent evaporite layers.

Figure 1-8. Generalized map and cross-section of the wall-and-basin structures (WABS) on central Axel Heiberg Island (Jackson & Harrison, 2006).

13 1.3 Study area: Ellef Ringnes Island

The evaporite structures and stratigraphy surrounding the domes on Ellef Ringnes Island were first studied by Heywood (1955, 1957), however the first comprehensive mapping of the entire island was completed by Stott (1969) (Fig. 1-9). The island extends from the north boundary of the Sverdrup Basin to the basin axis in the southern region. The northern Isachsen Peninsula is composed of Tertiary deposits that lie unconformably on uplifted and eroded Late Jurassic to Early Cretaceous units. Conformable units that overlie the Late Jurassic- Early Cretaceous rocks are exposed further south with the occurrence of the youngest conformable unit (Late Cretaceous- early Tertiary) confined to the axes of regional synclines. Stratigraphy and prominent structural features are summarized below.

1.3.1 Stratigraphy

The stratigraphy of central Ellef Ringnes Island is summarized in Figure 1-10. At depth, Ellef Ringnes Island is underlain by a sequence of Late Paleozoic to Early Mesozoic carbonate and deep marine / prodelta siltstones and shales (Davies & Nassichuk, 1991; Embry, 1991). Marine deposits are conformably overlain by the sandstone-dominant Heiberg Group. The Late Triassic to Early Jurassic Heiberg Group represents deposition of delta front siltstones grading to coarser siltstone and sandstone of delta plain and shallow marine shelf origin (Embry, 1991). In the western regions of Ellef Ringnes Island, the Heiberg Group is split into five formations representing the intercalation of prodelta / marine shale deposits and thin beach to marine shelf sandstones. The Heiberg Group is overlain by shales and siltstones of the Jameson Bay, McConnell Island, Ringnes and Deer Bay Formations, which are punctuated by thin marine shelf deposits of the Awingak Formation in the western regions of Ellef Ringnes, and by the Sandy Point Formation. At the basin margins, the Early Cretaceous Isachsen Formation lies unconformably above the Deer Bay Formation, which is suggested to mark the onset of the Amerasian rift event (Embry & Osadetz, 1988). The Isachsen Formation is dominated by medium to coarse-grained delta plain and fluvial channel deposits. Thick conformable Cretaceous formations (Christopher, Hassel and Kanguk) alternate between siltstone and sandstone dominant units of marine shelf and deltaic origin.

Figure 1-9. Map of Ellef Ringnes Island with labelled evaporite piercement structures. The bedrock geology is the unpublished compilation work of C. Harrison. The maps used in the compilation include Stott (1969) for Ellef Ringnes Island and Balkwill & Roy (1978) for King Christian Island.

The youngest unit related to the development of the Sverdrup Basin is the Expedition Formation, belonging to the Late Cretaceous to Tertiary Eureka Sound Group. The Eureka Sound Group is highly variable in lithology and thickness, suggesting that deposition took place within a series of isolated basins separated by intrabasin upwarps related to the Eurekan orogen (Miall, 1991). Ricketts & Stephenson (1994), however, argue that the Expedition Formation was deposited within a contiguous basin and that syn-tectonic sediments were not restricted between intrabasin upwarps until the Middle to Late Eocene. On Ellef Ringnes Island, the Expedition Formation is composed of non-marine sandstones and unconsolidated

15 sands (Stott, 1969). In the northern region of the island, crossbedded sands and gravels of the mid-Miocene to early Pliocene Beaufort Formation lie unconformably on the uplifted and eroded Deer Bay Formation (Stott, 1969). The Beaufort Formation is most likely fluvial in origin having been fed by northwest flowing rivers that terminate in the Arctic Ocean (Miall, 1991).

Figure 1-10. Summary of the Late Paleozoic and Mesozoic stratigraphy of Ellef Ringnes Island (modified after Dewing & Embry, 2007; Patchett et al., 2004; Embry, 1991; Embry & Osadetz, 1988).

1.3.2 Mafic intrusions

At surface, mafic intrusions include dykes and sills that are mostly concentrated within the Deer Bay Formation. They crop out as circular structures composed of diabase and gabbro and are interpreted as ring dykes by Stott (1969). However, they could also represent the erosional pattern of saucer-shaped sills. Sills and dykes are also noted locally within the Isachsen and Hassel Formations. The age of the intrusions was estimated by Larochelle &

16 Black (1963) to be between 102 and 110 Ma, which correlates with the average ages of the Sverdrup Basin Magmatic Province (Villeneuve & Williamson, 2006). All of the exposed evaporite domes on Ellef Ringnes Island contain large blocks of mafic intrusions. The blocks have chilled margins and are relatively unaltered, indicating that intrusion likely occurred after the emplacement of the diapirs. The blocks have not been dated, but are suspected to be of Cretaceous age (Stott, 1969).

1.3.3 Structural geology

Major structures on Ellef Ringnes include large, broad north to northwest trending synclines and anticlines with the youngest Mesozoic units preserved in the axes of the synclines. The greatest concentration of faults occurs within the Jurassic Deer Bay Formation on Reindeer Peninsula and the southern region of Isachsen Peninsula. The faults occur as conjugate pairs oriented northwest-southeast and northeast-southwest and coincide with the occurrence of large mafic saucer shaped sills (previously interpreted by Stott, 1969 as ring dykes). Tangential and radial faults are also associated with the evaporite piercement structures, particularly Hoodoo Dome where the cover has not been fully pierced.

1.3.4 Evaporite piercement structures

The island contains several ovate anhydrite piercement structures up to 15 km in width. Dumbells, Contour and Isachsen Domes are all exposed at surface, while Hoodoo Dome is buried under a thin cover of the Isachsen Formation. Helicopter, Haakon and Malloch Domes are only partially exposed on the east and west coast of Ellef Ringnes Island. The domes are aligned within regional anticlines and are in unconformable contact with the Isachsen Formation at surface. In the case of Dumbells Dome, the Isachsen Formation dips steeply away from the dome with orientations of up to 80 degrees (Stott, 1969). The anhydrite caps are relatively unaltered in the center of the dome with highly sheared margins (Heywood, 1955; Van Leeuwen, 2005). In addition to inclusions of mafic blocks, the domes also commonly contain large blocks of limestone. The limestone blocks may represent original beds intercalated within the anhydrite or blocks from the overlying Hare Fiord Formation that were brought up with the diapirs (Stott, 1969). A relatively large outcrop of limestone occurs on the western margin of Contour Dome. The age and the relation of the contact between the limestone and the evaporite dome are not well constrained.

17 1.4 Research objectives

Salt structures on Axel Heiberg Island have been investigated extensively (e.g. Kranck, 1961; Hoen, 1964; Gould & Mille, 1964, Schwerdtner & Clark, 1967; van Berkel et al., 1983; van Berkel et al., 1984; Jackson & Harrison, 2006). However, little work has been carried out on Ellef Ringnes Island since the original mapping by Stott (1969). The piercement structures on Ellef Ringnes Island are of interest as they are not strongly overprinted by deformation from the Tertiary Eurekan orogenic event. These structures therefore reflect diapirism related to earlier stages of basin development without the effects of late-stage regional compression. Due to the basin geometry and distribution of autochthonous salt in the basin axis, the source layer is buried to depths of up to 8-13 km (Balkwill, 1978). The initiation and rate of vertical migration of the diapirs are therefore poorly constrained due to the lack of high quality seismic and borehole data at these depths. The situation is further complicated by the presence of dense anhydrite caps, which diminishes (and possibly reverses during the onset of diapirism) the effect of buoyancy-driven diapirism.

The purpose of this research is to further constrain the timing and mechanisms involved in the formation of evaporite structures on Ellef Ringnes Island. Particular questions that are addressed include: • The timing and mechanisms required to initiate diapirism • Identification and timing of discrete phases of growth • The relation of regional tectonic subsidence to discrete phases of growth • The effect of dense anhydrite caps in the development of the piercement structures

The following questions are addressed by (1) constraining the subsurface geometry of the domes (2) analyzing tectonic subsidence signatures proximal to the domes (3) 2D analogue modeling on the development of piercement structures. These methods are briefly discussed below.

18 1.4.1 Chapter Two: Structural and geometric characterization of evaporite piercement structures on Ellef Ringnes Island

This chapter attempts to characterize the geometry of the piercement structures on Ellef Ringnes Island through the interpretation of historical seismic reflection and borehole data. The subsurface geometry and structures in the adjacent strata aid in the understanding of the kinematics and dynamics that govern their initiation and movement. The style of structures, unconformities and thinning of interpreted units toward the domes also constrain the timing of discrete periods of salt growth during the Late Triassic to Cretaceous. Isopach maps of the central and southern regions of Ellef Ringnes Island were also produced to visualize periods of sediment thinning and thickening from a regional perspective.

1.4.2 Chapter Three: Basin subsidence modelling

Backstripping of borehole data was completed to produce 1-D tectonic subsidence curves. The curves were produced in order to discern the subsidence signature related to salt movement adjacent to the domes in comparison to the average regional tectonic subsidence signature of Ellef Ringnes Island. The average trends of the tectonic subsidence curves also highlight periods of change in subsidence rates and can be directly related to different stages of diapir growth interpreted in seismic reflection profiles.

1.4.3 Chapter Four: Analogue modelling of diapirism

The final stage of the project involves scaled physical 2D models with forced injection of fluids into a down-building overburden of sand as analogues for syn-sedimentary diapiric growth. The experiments focus in particular on the emplacement of competent and relatively undeformed blocks of dense anhydrite on top of buoyant salt structures.

19 2 Structural and geometrical characterization of evaporite piercement structures on Ellef Ringnes Island

2.1 Introduction

The improvement in the quality of 2D and 3D seismic techniques has significantly advanced our understanding of salt structures. 3D seismic reflection data in particular allow interpreters to map structures to resolutions of a few tens of meters over thousands of square kilometers (Cartwright & Huuse, 2005). With recent increasing oil prices and decreasing costs for seismic acquisition and processing, many divergent margin basins are now covered extensively by 3D seismic surveys (Fetter, 2009).

Due to the isolated conditions, harsh environment and lack of infrastructure, oil and gas exploration within the Sverdrup Basin is costly and underdeveloped. A boom in oil and gas exploration occurred in the 1970’s and 1980’s with the discovery of 19 major petroleum fields (Chen et al., 2000), however the reserves have for the most part sat stagnant as the cost to recover them is still uneconomical. Due to the lack of recent research and exploration in the area, the timing and mechanisms controlling the initiation and mobilization of evaporite diapirs are still poorly understood. Seismic reflection surveys completed in the Sverdrup Basin are restricted to 2D single to multifold data with varying quality of resolution. Seismic interpretation of salt structures was recently conducted on the salt-based fold belt on Melville Island (Harrison, 1995); otherwise no known seismic interpretation of salt structures has been published on any other of the Canadian Arctic islands.

To properly understand the mechanisms controlling the formation of the evaporite piercement structures on Ellef Ringnes Island, it is important to constrain their 3D structure. In the following chapter I present a detailed analysis of archived industry 2D seismic reflection and borehole data from Ellef Ringnes Island with focus on the Dumbells, Contour and Hoodoo Domes.

20 2.2 Methods

2.2.1 Data

Legacy industry 2D seismic reflection SEG-Y data and borehole data were made available by the Geological Survey of Canada (Calgary). Seismic lines were chosen to encompass the central and southern regions of the island with detailed coverage of the Hoodoo, Dumbells and Contour Domes (Fig. 2-1). The seismic data includes surveys conducted by several oil companies from 1969 to 1973 (list of seismic surveys included in Appendix 1). The data consists of unfiltered, unmigrated, single to multi-fold data. The 100 and 1000 series seismic lines were acquired by Panarctic Oils Ltd. in 1969 and 1973 respectively. The 100 survey series consists of single fold data with shot points spaced at ~1200 m, providing a trace spacing of 25 m. This series was the first seismic shot on the island and consists of long reconnaissance lines that span across Ellef Ringnes Island. The 1000 survey series consists of 4-fold data with split spread shot geometries and a trace spacing of 33.5 m. The 1000 series is distributed throughout central and southern Ellef Ringnes Island, but are most concentrated south of Hoodoo Dome. The I(F) survey series was conducted by Imperial Oil Ltd. in 1971 and includes the majority of seismic lines surrounding Hoodoo Dome. The seismic reflection data are 3-fold for lines shot with end-on spread geometries for dip lines, and 6-fold for lines shot with split spread geometries for perimeter lines. Both types provided a trace spacing of 24 meters. The G(F) series was completed by Gulf Resources in 1972 and provides the most detailed coverage of Dumbells and Contour Dome. The data is 6-fold with split spread shot geometries and a trace spacing of 33.5 meters.

Gravity data was also collected along the majority of the seismic lines, however not all records (such as the I(F) series) were available for this study. Borehole data collected includes drill logs, geophysical logs (sonic, gamma ray, resistivity, etc.), check shot surveys and interpreted formation tops. The boreholes are located onshore and offshore with the penetration of the oldest and most complete Mesozoic sections occurring offshore of the southern region of Ellef Ringnes Island.

Figure 2-1. Index map of seismic reflection data and borehole locations. The data compiled for Chapter 2 are highlighted in green (for seismics) and red (for boreholes). The seismic profiles included in Plates 1-4 are labeled on the map and map inset of Hoodoo Dome.

2.2.2 Seismic processing

Many seismic reflection profiles were previously recovered from non-standard SEG-Y formats and contained no information within the trace headers. With the cooperation of the Geological Survey of Canada (Calgary), many of the lines located on or just offshore of Ellef Ringnes Island were quality controlled, including entry of the shot points and x/y coordinate locations within the trace headers. The lines were also analyzed for bulk shifts, dead traces 22 and mislabeling of traces and shot points. Many of the Gulf Resources (G(F) series) and Imperial Oil (I(F) series) lines approaching Dumbells, Contour and Hoodoo Dome were filtered and migrated at the Geological Survey of Canada (Calgary). High and low frequency noise below 7 Hz and above 55 Hz was removed by band pass filters. To remove diffracted arrivals from point sources (e.g. edges of mafic sills) and to restore dipping reflectors to their true vertical position, Kirchoff migrations were applied post-stack using root mean square velocities (VRMS) taken from original processing information on paper copies of the data. The

VRMS values were originally picked by individual companies from pre-stack shot gathers.

Imperial Oil seismic records surrounding Hoodoo Dome did not include VRMS picks from shot gathers, and therefore interval velocity models (Vint) were produced based on interpreted horizon picks. The average velocities of each formation were estimated from the Hoodoo

Dome H-37 sonic logs. The Vint were then converted to VRMS and various versions of migration were completed for each seismic line with varying reductions of the initial VRMS values (from 100 % to 90 %) to remove the affects of over migration at depth. Several regional lines, such as 127 pt 1-5 and transect 1097-I(F)66-1060, were also migrated.

2.2.3 Analysis of borehole data

Synthetic well logs were produced for the following wells: Hoodoo Dome H-37, Hoodoo L- 41, Hoodoo N-52, Elve M-40, Dumbells E-49, Helicopter J-12 and Cape Allison L-50. An example of the Hoodoo Dome H-37 synthetic well log is included in Figure 2-2. Formation tops interpreted by Dewing & Embry (2007) were evaluated alongside gamma ray and sonic records. Cretaceous units generally increase in velocity with depth due to compaction, while the Deer Bay and Ringnes Formation are close to fully compacted at average velocities of m ּs-1. The most noticeable signature is the jump to higher velocities (4000-4500 3000-3500 m ּs-1) at the transition from the Jameson Bay Formation to the Heiberg Group. Depth intervals were converted to time using velocity/time pairs from check shot surveys to allow for comparison and overlay of well logs on intersecting seismic reflection profiles. Synthetic seismic signatures were produced for each well by frequency matching of neighbouring seismic traces.

Figure 2-2. Synthetic well log for Hoodoo Dome H-37 with formation tops from Dewing & Embry (2007).

2.2.4 Interpretation of formation tops

Formation horizons within seismic profiles were interpreted based on the picked tops from wells within the central and southern regions of the island using the Kingdom Software Suite program (Seismic Micro-Technology Inc.). The majority of wells located within these regions were drilled to the top of the main oil and gas hosting reservoir rock, the porous sandstone-dominant King Christian Formation and correlative Heiberg Group. Control on formation tops therefore dates as far back as the Late Triassic-Early Jurassic. Lines that do not intersect with wells were tied in with neighboring and cross-cutting seismic reflection surveys. The Christopher and younger formation horizons are not intersected by any wells and were therefore approximated from surface outcrops. Digital elevation models were also used to tie the top of the Isachsen Formation to the surface when contacts from previous mapping (Stott, 1969) did not agree with the location of interpreted horizons.

24 Limitations in the interpretation of horizons are mostly due to the low quality and resolution of the data. The resolution of the data most likely smears detailed structures, such as faults and unconformities, especially in the top second where stronger first arrivals are not easily filtered. This is most problematic in imaging faults related to salt movement near the surface, such as at Hoodoo Dome. Permafrost layers are also problematic in the top 500 milliseconds of the data as seismic waves are often amplified due to the increase in velocity from frozen water-filled pore spaces of permeable units, such as the sandstone dominant Hassel Formation. The above issues make it difficult to accurately tie formations to the surface, in particular formations with no stratigraphic control from wells. Mafic sills are also common, especially within Jurassic formations northeast of Dumbells and Contour Domes. The highly reflective sills reach thicknesses of up to 140 m, making it difficult to identify underlying horizons. Interpretation of evaporite bodies is complicated around Hoodoo Dome, where salt sills and canopies may be present. The signals of sedimentary units that underlie such structures are often diminished as most of the energy is reflected at the salt boundary. Steeply dipping dome margins and adjacent upwarped sediments are also difficult to image if the angle of incidence with respect to the boundary is greater than the critical angle of refraction. For some seismic profiles, such as along seismic profile 123, gravity was recorded in conjunction with seismic acquisition, which aids in the interpretation of the boundary between the salt bodies and adjacent sedimentary units. The Bouguer anomaly data, however, are of low spatial resolution. Without additional information on the regional Bouguer gravity field it is also difficult to define residual anomalies associated with the presence of salt or deeper crustal signatures.

2.2.5 Isopach maps

Isopach maps of each formation were produced to identify regional thinning/thickening trends across central and southern Ellef Ringnes Island. The isopach maps are based on interpreted formation horizons from seismic reflection profiles (distribution of lines included in Figure 2-1). For consistent ties between all of the seismic surveys, only interpreted horizons of unmigrated data were used. To convert interpreted tops from two-way-travel time to depth, average interval velocities for each formation are required. Available velocity data

25 collected in boreholes include check shot surveys and sonic logs. Sonic logs provide the most detailed velocity profiles for each formation with high enough resolution to identify increasing/decreasing trends in velocities due to facies changes and compaction. Instantaneous velocity versus depth curves were produced for each unit by extracting the desired portion of sonic logs for a particular formation. For each individual formation, a composite sonic log was created from multiple wells in an attempt to capture velocity profiles over the largest range of depths. To characterize the general trend in velocity with depth, curves were fit to the composite sonic plots (e.g. Fig. 2-3). The composite sonic logs for each formation are included in Appendix 2.

Figure 2-3. Instantaneous velocity with depth signature of the Deer Bay Formation derived from a compilation of sonic logs from multiple wells. The best fit power law curve parameters are included in Table 2-1. Velocity-depth profiles for other Mesozoic formations are included in Appendix 2.

Sonic logs record the instantaneous velocity within a formation; however, interval velocities are required for depth conversion of formation tops in two-way-travel time. An expression for interval velocity is obtained by integration of instantaneous velocity. For a linear approximation of the increase in instantaneous velocity vz with depth z,

oz += kzvv (1)

26 where vo is the surface velocity and k is the velocity gradient. This expression for instantaneous velocity is integrated to obtain the following interval velocity: Δz 1 V = = ekzv Δtk −+ )1)(( (2) int Δ Δtkt to

where Δt is the one-way-travel time of the layer, Δz is the unit thickness and zt is the depth to the top of the layer (Hillis et al., 1995). For multilayer time-depth conversion, the velocity travel path through overlying layers needs to be considered, therefore a layer cake depth conversion is required to approximate formation depth:

v Δtk Δtk z o )1( +−= eze (3) n k n−1

where subscript n denotes the nth layer. Linear functions, however, may over-estimate velocities at greater depths as velocities often do not increase continuously with depth indefinitely (Al-Chalabi, 1997). The rate of velocity increase slows with depth and often becomes constant once units are fully compacted. Power law functions are therefore chosen for older units as they often fit the data better than linear functions and the velocity trend decreases and levels off with respect to the linear function (1):

/1 m z = Azv (4) where A is the pre-exponent constant and m is the exponent that characterizes the decay of the velocity gradient. Integration of (4) provides the following relationship of depth with time for individual layers (Al-Chalabi, 1997):

q /1 q ( nn −1 Δ−= tqAzz ) m −1 (5) q = m A MATLAB code was written to implement the time-depth conversion of each formation (Appendix 3) with the appropriate formation curve parameters listed in Table 2-1. Contour plots were produced using an “inverse distance to power” gridding method that closely retains original values of data points. The interpolation method applies weighted averages, which is appropriate for this study as data points are not uniformly distributed within the

27 interpolated region. The resulting contours were clipped to the contact boundaries of each formation and a Gaussian low-pass filter was applied to smooth any extreme outliers.

Table 2-1. Parameters and functions used for time-depth conversion of individual formations and estimated fit errors (coefficient of determination, R2).

-1 2 Linear Vo (m ּs ) k R Expedition Fm 2500 0 - Kanguk Fm 1305 1.405 0.54 Hassel Fm 2140 0.537 0.22 Christopher Fm 1937 0.706 0.54 Isachsen Fm 2545 0.526 0.14 Power Law A (m ּs-1) m R2 Deer Bay Fm 1677 9.952 0.28 Ringnes Fm 1686 11.54 0.13 McConnell Island Fm 884 5.301 0.35 Jameson Bay Fm 1355 7.672 0.31

Inaccuracies in the contour plots can be divided into errors related to interpretation of horizon picks (i.e. time at which a horizon occurs in a seismic section), and errors related to the approximation of instantaneous velocity from sonic logs. Errors related to the interpretation of horizon picks have been briefly mentioned in previous sections and are related to the resolution of the seismic data (i.e. fold and spacing of geophones) and occurrence of highly reflective bodies that mask reflections from sedimentary horizons, such as salt canopies and mafic sills. The wide spread of data values in sonic logs lead to low coefficients of determination, R2 (column 4 in Table 2-1), however, this method of curve fitting is used in an attempt to approximate an average velocity profile with depth and therefore the R2 values are less relevant for this study. Increases in velocities from permafrost zones and intrusive sills northeast of Dumbells and Contour Dome could lead to under estimated thicknesses. The range in velocity data with depth is also quite sparse for the youngest formations (Hassel,

28 Kanguk and Expedition Formations), and therefore the trend may only be accurate for limited depth ranges. The Expedition Formation was assumed to have a constant velocity of 2500 m ּs-1 as no velocity data was made available. The inaccuracies of this assumption apply only to a limited area, within the Christopher Syncline and east of Hoodoo Dome.

2.3 Results

The interpreted seismic profiles discussed in the text below are included within Plates 1 to 4. Larger versions of the isopach maps included in this chapter are found in Appendix 4.

2.3.1 Dome asymmetry

Interpreted E-W oriented seismic profiles across Dumbells and Hoodoo Dome reveal asymmetric geometries within onlapping sedimentary units. This asymmetry is best portrayed in seismic profile 123 and in the composite profile of G(F)9B and G(F)11B (Fig. 2-4) for Hoodoo Dome and Dumbells Dome respectively. Both profiles share similar characteristics with generally thicker sedimentary packages occurring at deep structural levels on the eastern side of the domes. For Dumbells Dome, this thickness variation is also accompanied by the occurrence of a shallow rim syncline located approximately 5 km from the edge of the piercement structure. The structural offset on the eastern side of Dumbells Dome is accommodated by dramatic thinning of upturned beds toward the dome margin. At Hoodoo Dome, the noticeable structural offset of formations on either side of the dome is accommodated by a near vertical fault or shear zone where the coherency of the sedimentary reflectors are lost on the seismic section. The eastern fault block is offset downward and can be traced through multiple E-W oriented seismic sections, giving an average fault trace orientation trending N-NW to S-SE (Fig. 2-5), parallel to the axial trace of the Hoodoo Dome anticline. The effects of this zone can be seen at surface with the apparent “narrowing” of the Hassel and Kanguk Formations. The shear zone is also detected as a perturbation of the Bouguer gravity profile along seismic profile 123 (see Plate 1). The shear zone terminates at the interpreted evaporite dome margin in profiles I(F)85 and 123. The shear zone is not apparent north of seismic profile I(F)85 and may be cut off by the NE-SW trending fault mapped by Stott (1969) on the northeast side of the dome.

Figure 2-4. Line drawing of seismic profiles G(F)9B and G(F)11B illustrating the asymmetry of Dumbells Dome. Seismic profiles are migrated with a vertical exaggeration of 1.5. Mafic sills are highlighted in dark green and the basic outline of evaporites is highlighted in yellow.

In addition to thicker units on the eastern side of Hoodoo Dome, a consistent pattern of “diverging” sedimentary horizons along the north-northeast margin is visible within seismic profiles I(F)77, I(F)79+A, I(F)80, I(F)81A, I(F)84 and 122. The beds diverge approaching the dome margins, with units dipping downward beneath the interpreted top of the Heiberg Group. With migration, this feature collapses into broad open synclines that appear to extend beneath interpreted salt canopies (e.g. profiles I(F)80 and I(F)84; see Section 2.3.2 for more details on Hoodoo Dome canopies).

The 3D geometry of Hoodoo Dome is complicated by faulting, which is not resolved well in 2D seismic sections. An asymmetry is also apparent along strike of the Hoodoo Dome anticline (profile 122). Units along this NW-SE section demonstrate similar attributes with greater thicknesses occurring at deeper structural levels along the north side of the dome. Formations also thin more toward the north margin of the dome in relation to the south margin.

2.3.2 Salt canopies at Hoodoo Dome

Due to the lack of seismic reflection surveys acquired over top of the domes, the geometry of the salt bodies underlying Dumbells and Contour Dome are not constrained. Little evidence is present in cross-secting G(F) series seismic lines to suggest that a salt body extends beneath the adjacent sedimentary sequences. Seismic line 127 pt. 3 crosses the link between Dumbells

30 and Contour Dome, however salt does not appear to pierce interpreted Triassic to Cretaceous horizons.

The partial sedimentary cover of Hoodoo Dome allows for better coverage of seismic reflection data on the top of the dome. Strong reflectors on seismic profiles I(F)75, I(F)77, I(F)80 clearly mark the transition from low velocity overlying sediments to contrasting higher velocities of anhydrite and halite at the top of the dome. The salt margins, however, are typically difficult to identify, which led to the notion of irregular evaporite dome boundaries. Hoodoo Dome is suspected to have multiple salt canopies branching off the main diapir stock to produce complex “Christmas tree” geometries. The tips of the canopies are identified in migrated profiles by diverging units above and below intruded salt bodies. Velocity pull-ups from the lateral variation of clastic sediment and evaporite velocities also aid in interpreting canopy boundaries. An example of an interpreted salt canopy is recorded in seismic line I(F)71, where the Deer Bay Formation flexes upward while the Heiberg Group appears to warp below the edge of a body where the strength of reflective surfaces is diminished.

Figure 2-5. Spatial extent of the interpreted canopies at subsurface and the trace of the shear zone along the eastern margin of Hoodoo Dome.

The number and extent of canopies surrounding Hoodoo Dome are complex and highly variable between neighboring seismic lines. On the southwest margin of the dome, an interpreted canopy extends beneath the Heiberg Group, constraining the age of the oldest

31 interpreted canopy to at least the Late Triassic. Along seismic profile I(F)67 and I(F)68, the Heiberg and younger formations gently lap onto this interpreted canopy. Further north, the number and complexity of canopies increase. Canopies occurring within Jurassic formations are broad, with many possible higher frequency canopy “offshoots” (e.g. profile I(F)75). The offshoots are difficult to identify and the interpretations are open to debate, however the first order canopy signature is still highly convincing. Profile I(F)75 is a good example of possible canopy offshoots that overlie the boundary of the Ringnes-Deer Bay Formations, and Jameson Bay-McConnell Island Formations. I(F)85 may also contain many higher frequency canopy offshoots, however the strength of sedimentary reflectors makes it difficult to be confident in this interpretation. Canopies that occur within the Heiberg Group are difficult to trace further north in seismic profiles due to masking from younger canopies; however Triassic canopies, such as the one found further south along profile I(F)67, are also suspected to occur along the north and northeastern margins of Hoodoo Dome (e.g. profile I(F)75). Thinner and less developed canopies may exist on top of the Deer Bay Formation on the eastern margin of Hoodoo Dome as marked by the downward flexure of the Deer Bay reflector (profile I(F)85). This interpreted canopy is the youngest identified along the margin of Hoodoo Dome.

2.3.3 Deformation and unconformities within overburden adjacent to piercement structures

On the dome margins, formations are thinned and dragged upward during syn-halokinetic sediment deposition. Thinning is most evident during the Late Jurassic and Early Cretaceous within the Deer Bay and Isachsen Formations. Early to Middle Jurassic formations are also dragged upward, however the imaging quality of these formations degrades with depth in the seismic sections and in areas where canopies overlie these formations. As previously mentioned, thinning along the eastern margin of Dumbells Dome is quite dramatic where the vertical offset of formations is expected to have occurred due to salt withdrawal in the source layer. At Hoodoo Dome, thinning of units is most apparent within formations that onlap salt canopies. In the Hoodoo N-52 well, the Jameson Bay and McConnell Island Formations pinch out due to thinning on top of a canopy on the northeast margin of Hoodoo Dome. This

32 well is used to constrain the tops of thinning formations on the margin of the dome in seismic profiles I(F)81A and I(F)82.

Brittle faulting is rare on the margins of the domes where deformation has occurred predominantly by ductile mechanisms (e.g. profiles G(F)2, I(F)67). Minor faults exist on the west margin of Dumbells Dome (profile 127 pt.3), however the seismic profiles that suggest the presence of faults are single fold data and some of the offset reflectors may be an artifact of the low imaging quality and imprecise migration. Faults within sediments overlying Hoodoo Dome are difficult to make out in the uppermost second of the I(F) series surveys as the first arrivals mask shallow reflections. Faults inevitably exist, however, as indicated from previous regional mapping of Hoodoo Dome (Stott, 1969). In addition to the northwest- southeast trending shear zone identified on the east margin of Hoodoo Dome, an extensive fault is visible in seismic profile I(F)82 from the offset of Cretaceous formations. The location of this fault in the seismic profile matches that of a fault trace at surface interpreted by Stott (1969).

Figure 2-6. Cretaceous unconformities within the Isachsen Formation (black arrow) along (a) the east margin of Dumbells Dome in profile G(F)11B and (b) northeast margin of Hoodoo Dome in profile I(F)75. The seismic profiles are migrated with a vertical exaggeration of ~1.5.

Unconformities are rarely visible within the seismic profiles. The majority of units thin toward the domes without any apparent erosional surfaces. The most evident unconformities occur within the Isachsen Formation, cutting off dragged and thinned Deer Bay deposits (e.g. I(F)75) and capping eroded Isachsen rim syncline deposits (G(F)11B; Fig. 2-6). The canopy

33 offshoots visible within some seismic profiles also suggest that there are unconformities capping dragged sediments on the dome margins and they are suspected to be related to uplift and erosion during diapirism. For example, the Heiberg and Jameson Bay Formations may truncate underlying, steeper dipping beds within profiles I (F)75, I(F)77, I(F)79+A, and I(F)85.

2.3.4 Regional variations in formation thicknesses from isopach maps

Isopach maps of Mesozoic formations are presented in Figure 2-7. Seismically approximated formation thicknesses are similar (+- 50 m) to observations from boreholes (Fig. 2-7). The largest deviation from known thicknesses occurs in the northeastern region where seismically determined thicknesses are under-estimated for the Ringnes Formation by approximately 100-150 m. This error is most likely due to the occurrence of thick mafic sills that led to difficulties in interpretations and under-estimation of velocities. The isopachs show a progressive thickening of units from the south at Meteorologist Peninsula to the northeastern margin of Dumbells and Contour Domes. Another common signature is the localized thickening of units, such as the Christopher Formation, within the Christopher Syncline west of the offshore Malloch Dome. More detailed observations include the thinning of units toward the domes. This is most apparent for the Deer Bay, Isachsen, Chrisopher and Hassel Formations surrounding Hoodoo Dome where isopach contours are tightly spaced. The contours do not always follow the shape of the formation outcrop patterns (e.g. Deer Bay and Isachsen Formations) due to lack of control points on the southern margin of Hoodoo Dome (see Fig. 2-1 for distribution of seismic surveys). The contours suggest that thinned units at the margins of Hoodoo Dome thicken more rapidly on the eastern side, in agreement with the asymmetric profile of Hoodoo Dome discussed in Section 2.3.1. Thinning of units toward the Dumbells and Contour Domes is not as obvious and may be due to the low resolution of the reflections in the seismic profiles close to the dome margins. Localized thickening (e.g. Isachsen, Deer Bay and Ringnes Formations) within the concave curvature of the two domes matches observations of rim syncline deposits adjacent to Dumbells Dome (profile G(F)11B).

Figure 2-7. Isopach maps derived from the depth conversion of seismic horizon picks. The colour scale on the right margin of the diagram represents the thickness of each formation. Formation thicknesses from boreholes are included on individual diagrams. X/Y coordinates (in meters) are generated with a Lambert Conic Conformal map projection. Larger versions of these plots are included in Appendix 4.

35 The resolution of picked horizons decreases for the Ringnes, McConnell Island and Jameson Bay Formations, especially northeast of Dumbells and Contour Dome where strong reflections from intrusive sills mask those of the Jurassic formations. In general these units do not show a large thickness variation compared to Late Jurassic to Early Cretaceous formations, however a gradational N-S pattern of thickening still exists. Though diverging sedimentation is mostly recorded within formations underlying the Heiberg Group, thickening of Jurassic formations on the north side of Hoodoo Dome is consistent with the location of surveys that record this feature.

2.4 Discussion

2.4.1 Basement fault-initiated and controlled diapirism

The spatial distribution of Hoodoo, Dumbells and Contour Dome may be closely linked to major basement faults as supported by the near linear outcrop patterns and asymmetric subsurface geometries of the domes. Thick-skinned deformation (i.e. basement involved deformation) during regional extension is a suggested trigger mechanism for diapirism in other salt basins such as Dniepr-Donets, North Sea, Danish and Nordkapp Basins (Stovba & Stephenson, 2003; Bishop et al., 1996; Koyi & Petersen, 1993; Koyi et al., 1993b). Numerical and analogue models have also confirmed the importance of basement geometry and structures induced by regional extension on the development of salt diapirs (Gaullier et al., 1993; Koyi & Petersen, 1993; Koyi et al., 1993a; Koyi, 1991; Nalpas & Brun, 1993). The concept of thick-skinned deformation is in contrast to the widely accepted mechanism of thin-skinned deformation where the stretching of overburden leads to faulting and sinking of half grabens into pressurized salt (Vendeville & Jackson, 1992a). Jackson & Vendeville (1994) argue that the effect of basement faults on salt diapirism is indirect unless salt is thin or depleted and that diapirism is more commonly triggered by thin-skinned extension. However, recurring observations of diapir asymmetry, lack of extensive overburden faulting and the spatial correlation of salt diapirs to basement faults suggests that thick-skinned extension is also an important and common mechanism in triggering diapirism in many basins (e.g. Fig. 2-8; Stovba & Stephenson, 2003). Though salt often obscures underlying basement faults in seismic data, basement faults have been identified in seismic reflection profiles from the Dniepr-

36 Donets, Danish and North Sea Basins (Stovba & Stephenson, 2003; Koyi & Peterson, 1993; Nalpas & Brun, 1993).

Figure 2-8. Examples of basement initiated diapirism and resulting asymmetry for diapirs in the (a) Pripyat Trough, Belarus and (b) Dniepr-Donets Basin, Ukraine (Stovba & Stephenson, 2003).

Effective weakening of the overburden within thick-skinned deformation settings occurs by the collapse of salt and overburden over basement faults (i.e. drape monoclines; Jackson & Vendeville, 1994; Gaullier et al., 1993), leading to localized extension and fracturing within the overburden (Fig. 2-9). Displacement along basement faults and the thickness ratio of the salt and overburden controls localization of overburden deformation and salt flow (Nalpas & Brun, 1993; Koyi et al. 1993a). One or a combination of these two factors most likely governed the size and shape of the diapirs on Ellef Ringnes Island.

Figure 2-9. Schematic of a drape monocline, illustrating the localized weakening of overburden in response to the collapse of the overburden over displaced basement fault grabens (Jackson & Vendeville, 1994).

37 The rotation of basement blocks perturbs an unstable balance between salt and overburden, causing lateral variations in the thinning of both layers (Stovba & Stephenson, 2003). Movement along basement faults would also displace the salt and overlying sediments on downthrown fault blocks by extensional collapse, creating accommodation space for thicker packages of sediments to accumulate. The differential loading of the eastern fault blocks with respect to the western fault blocks at Ellef Ringnes Island would have caused further subsidence into the underlying evaporite layer, which would further perpetuate loading of the fault block through a positive feedback mechanism (Fig. 2-10). Due to the positive feedback mechanism that drives increased subsidence along the eastern margin of the domes, the basement fault is not required to slip continually during the history of dome growth. Differential loading would result in the expulsion of salt on top of the eastern fault block, effectively driving the growth of the neighboring salt diapir. Such a mechanism would result in the following asymmetries observed in E-W oriented seismic reflection profiles across the domes on Ellef Ringnes Island: • Thickening of sedimentary units on the downthrown fault blocks (eastern margins) • Formations occurring at deeper structural levels on the eastern margins • Development of more pronounced rim synclines and dramatic thinning of sediments at Dumbells Dome due to salt withdrawal beneath the eastern margin • Diverging patterns of sedimentation northeast of Hoodoo Dome due to rotation from the gravitational collapse of the eastern fault block • Formation of near vertical shear zones adjacent to Hoodoo Dome due to the gravitational collapse of the eastern fault block • More frequent occurrence of salt canopies within sediments overlying the eastern fault block at Hoodoo Dome

Predominant ductile deformation of clastic sediments and the lack of brittle structures along the margins of the domes also support the notion of basement-controlled extension in contrast to thinning and fracturing of the overburden via thin-skinned extension as proposed by Jackson & Vendeville (1994).

Figure 2-10. (a) Schematic of a dome illustrating the positive feedback mechanism that drives excess sediment loading and salt withdrawal on an eastern fault graben. The feedback mechanism leads to rotation of sedimentary beds and structures within overburden related to gravitational collapse. (b) Progressive stages of diapir initiation and growth controlled by basement structures (Koyi et al., 1993a).

The basement structures underlying Ellef Ringnes Island are thought to be northwest- southeast oriented normal faults (perpendicular to the basin axis) dipping to the east toward the Axel Heiberg depocenter (Figure 2-11a). The linear arrangement of diapirs in the Sverdrup Basin has been previously suggested to be evidence for control by normal faults in the basement during the Mesozoic (van Berkel et al., 1983; Stephenson et al., 1992; Schwerdtner & Osadetz, 1983). Though deformation patterns in the overburden suggest the existence of underlying basement faults and the orientation of the proposed faults correlate well with the occurrence of the Axel Heiberg depocenter, no major basement structures have been resolved in past gravity and seismic refraction studies conducted on Ellef Ringnes Island (Forsyth et al., 1979; Sobczak & Overton, 1984).

Figure 2-11. Figure adapted from Oakey & Stephenson (2008) highlighting the linear gravity low on southern Ellef Ringnes Island (blue arrow). (a) Mesozoic sedimentary thicknesses show a depocenter over Axel Heiberg Island. (b) Crustal Bouguer anomaly map in which the gravitational effects of the sediments is removed. Location of seismic refraction lines are in black.. The axes of linear gravity highs and lows are highlighted in black and white dashed lines respectively. Arches and gravity highs within the Sverdrup Basin include the Cornwall Arch (CA), West Axel Heiberg High (WAH), Princess Margaret Arch (PMA). The Eurekan Frontal Thrust (EFT) is also included in the figures. (c) Depth-to Moho from the inversion of the ‘crustal Bouguer’ anomaly. EU marks the location of a receiver function station. (d) Crustal thickness.

More recent gravity modeling of the Canadian Arctic Innuitian region and Greenland (Oakey & Stephenson, 2008) highlights several N-S trending linear features that were originally interpreted to be due to crustal-scale folding that perturbed the Moho during the Tertiary Eurekan Orogeny (Stephenson et al., 1990). One of these features intersects Meteorologist Peninsula in southern Ellef Ringnes Island, just east of Hoodoo Dome (Fig. 2-11). The linear feature is a gravity low that is associated with a higher than average depth-to-Moho and slightly elevated crustal thickness. Such a regional gravity anomaly could also be related to basement structures that existed prior to the Eurekan Orogeny. The lack of significant deformation and thrusting within basin sediments on Ellef Ringnes Island suggests that effects of compression of the Eurekan Orogeny is

40 limited to the eastern regions of the Sverdrup Basin and that crustal-scale folding beneath Ellef Ringnes Island is highly unlikely. The wavelength of lithospheric folds, however, is on the order of 100 km (Oakey & Stephenson, 2008) and may therefore not be easily detected within basin sediments at smaller scales.

2.4.2 The timing of diapirism in the Sverdrup Basin

Due to the lack of data sampling at depth, units older than the Heiberg Group were not identified in seismic reflection profiles and borehole data, therefore the initial stages of diapirism are not constrained. The diverging pattern of sedimentation toward the north- northeast margin of Hoodoo Dome (e.g. I(F)79+A) indicates that salt withdrawal beneath the overlying sedimentary formations began as early as the Triassic. This observation is consistent with other authors who observe modification of Triassic formations near large domal diapirs (e.g. Schwerdtner & Osadetz, 1983; van Berkel et al. 1989; Stephenson et al., 1992; Harrison, 1995). The extrusion of salt within or below the Heiberg Group either suggests a decrease in sedimentation rate during the Triassic, or that the upward migration of evaporites was occurring at a faster rate than regional sedimentation due to collapse of the margins from salt withdrawal (i.e. peripheral sinks).

The formation of multiple canopies at Hoodoo Dome during the Jurassic suggests that the diapirs at that stage were at or close to the surface. Multiple canopy offshoots located along the north-northeast margins of Hoodoo Dome indicate that salt diapirism was most likely episodic or occurred in response to episodic sedimentation. Where canopies are absent, such as at Dumbells and Contour Dome, formations thin toward the domes in response to the dragging of the adjacent sediments by the growing diapirs. The ductile deformation of the marginal sediments also suggests that diapirism was actively occurring at or close to the surface when sedimentary units were not fully lithified. The narrow pseudo-plastic behaviour of drag zones is a deformation style indicative of passive diapirism or downbuilding, where the growth of a diapir keeps pace with sedimentation (Fig. 2-12; Alsop et al., 2000).

Figure 2-12. Schematic drag profile adjacent to a salt diapir illustrating the common deformation mechanisms within incompetent, homogenous overburden typical of passive diapirism (Alsop et al., 2000).

The gravitational collapse of the eastern regions of Hoodoo, Dumbells and Contour Domes continued during the Jurassic with the development of rim synclines at Dumbells Dome and further offset along the shear zone located on the east margin of Hoodoo Dome. At Hoodoo Dome, the excess subsidence in the northeast in comparison to the south led to the accumulation of thicker packages of Jurassic sediments (Fig. 2-7) and may have allowed further migration of the Jurassic canopies.

The continual onlap of the Hoodoo Dome canopies by the Late Jurassic-Early Cretaceous formations (Deer Bay and Isachsen Formation) marks a period of passive diapir growth where aggradation rate (vertical accumulation of sediments) exceeded net diapiric rise (Koyi, 1998). Continual periods of onlap were punctuated by regional unconformities during the deposition of the Isachsen Formation, and extrusion of salt on the east side of Hoodoo Dome. The domes were eventually buried by the upper members of the Isachsen Formation and the Christopher Formation.

During burial and up until the deposition of the Kanguk Formation, thicker deposits continued to accumulate on the east side of Hoodoo Dome as accommodation space was continually provided by the offset of the shear zone along the east margin. The Hassel Formation also appears to thin toward the dome (e.g. I(F)73; Gould & DeMille, 1964),

42 suggesting that diapirism was still occurring during the Late Cretaceous, most likely by continued downbuilding from adjacent sediments. Diapirism may have continued into the Early Tertiary with gentle uplift and thinning of the Kanguk and Expedition Formations; however, active piercement of overlying Cretaceous sediments most likely occurred during the Tertiary Eurekan Orogeny. Exposure of anhydrite caps at the surface is most likely associated with uplift and erosion during this period with a possible added component of active diapirism driven by regional compression.

2.5 Conclusion

While little is known about Carboniferous to Triassic formations at depth, the asymmetry of the evaporite structures on Ellef Ringnes Island suggests that their initiation may have been localized on top of pre-existing basement faults. More observations supporting thick-skinned (i.e. basement controlled) over thin-skinned extension as a trigger for salt flow includes the lack of brittle faulting within the basin sedimentary deposits across Ellef Ringnes Island. Basement faults would have localized deformation within the overburden, creating weaknesses for migrating viscous salt to exploit. Basement faults would also have established differential loads on the two opposing fault blocks as a driving mechanism for diapir growth. Differential loading produced rim synclines, shear zones, structural offset and thickening of the overburden on the eastern sides of Dumbells, Contour and Hoodoo Dome.

Diverging sedimentation and the formation of canopies within or beneath the Heiberg Group constrains a minimum Late Triassic age for the initiation of diapirism at Ellef Ringnes Island. From the Late Triassic to the Early Cretaceous the domes continued to develop under passive diapirism (i.e. downbuilding). This is supported by the plastic deformation and thinning of the Jurassic and Cretaceous units adjacent to the dome, and by the formation of numerous canopies at Hoodoo Dome due to episodic extrusion of salt at the surface. During the Late Jurassic to Early Cretaceous the domes were slowly onlapped and buried by the Deer Bay, Isachsen and Christopher Formations. Thicker deposits continued to accumulate on the eastern margin of Hoodoo Dome, suggesting that evacuation of salt beneath overlying sediments continued until the Late Cretaceous.

43 Active piercement of Cretaceous units most likely occurred from regional compression in the Tertiary produced by the Eurekan orogenic event.

The largest improvement that could be made to this study is the acquisition of higher resolution 3D seismic reflection data that would further elucidate the geometry and structure of the domes. Improved acquisition techniques and processing of higher quality seismic data may also aid in imaging deeper structures, such as faults within the basement. Due to time constraints, gravity modelling was not completed, however the margins of the domes may be easier to identify if residual gravity anomalies were to be extracted from the regional Bouguer gravity field. Further evaluation of formation thicknesses should include correcting sonic logs to check shot surveys and porosity studies (to determine relative compaction and uplift of formations). Further modifications should also include corrections for higher velocities close to the surface due to permafrost zones. A more extensive look at the interpolation methods and relative weighting would also aid in producing more refined contour plots.

44 3 Basin subsidence modelling

3.1 Introduction

The first attempt at basin subsidence analysis of the Sverdrup Basin was conducted by Sweeney (1977) and included regional 1D subsidence versus time curves, based on measured stratigraphic sections and unpublished well logs with age constraints from dated fossil horizons. Sweeney (1977) described semi-quantitatively the degree of subsidence that is accounted for by isostatic compensation and correlated changes in subsidence patterns with the timing of tectonic events, such as the initial rifting of the Canada Basin. Stephenson et al. (1987) produced regional and local isostasy models and estimated the stretching factor of the lithosphere required to isostatically accommodate the syn-rift strata. The models were thermo-mechanical and investigated the effects of thermal in-plane forces and bending moments on basin subsidence and lithosphere thinning. Further 1-D backstripping and tectonic subsidence modeling of various locations across the basin was completed by Stephenson et al. (1994; Fig. 3-1). The stratigraphic thicknesses used for backstripping were also taken from unclassified exploration wells and measured sections. Detailed stratigraphic sections of Cenomanian and younger strata measured in the eastern region of the Sverdrup Basin were merged with data from the rest of the basin to approximate the thickness of eroded formations. Estimations of eroded thicknesses were also taken from compaction and vitrinite-reflectance studies. Gentzis & Goodarzi (1993) calculated subsidence curves for North Sabine H-49 and Roche Point J-43 wells, located on Melville Island. Their study, however, focused more on the relationship of the burial history with oil maturation and did not specifically discuss tectonic subsidence.

Six distinct periods of tectonic subsidence were identified by Stephenson et al. (1994). The first stage was the initial rifting phase (311-269 Ma) followed by more subdued post-rifting thermal subsidence (1 and 2 in Fig. 3-1). Subsidence due to thermal relaxation was interrupted in the Late Jurassic by a second stage of rifting lasting between the Oxfordian and Albian (157-97 Ma; 3 in Fig. 3-1). This stage of rifting was most pronounced in the eastern region of the basin and is associated with Cretaceous igneous activity (Embry & Osadetz, 1988) and is thought to be a signature related to rifting of the Canada Basin. Passive thermal

45 subsidence continued from the Cenomanian to the Maastrichtian (4 in Fig. 3-1). The last stage of subsidence was compressional resulting from seafloor spreading in the Labrador Sea and Baffin Bay that led to the impingment of Greenland against Ellesmere Island. The rates of subsidence during this period (65-43 Ma; 5 in Fig. 3-1b) were equal to or greater than the Permo-Carboniferous subsidence rates due to initial rifting. The build-up of compressional stresses produced by the Eurekan Orogeny eventually led to large scale crustal failure in the Middle Eocene (6 in Fig. 3-1b), producing several basin highs (Stephenson et al., 1990; Oakey & Stephenson, 2008) that restricted sedimentation within intramontane sedimentary basins (Miall, 1991; Ricketts & Stephenson, 1994). A more detailed study on the relative timing of subsidence and sedimentation of the Expedition and younger formations during the Tertiary was completed by Ricketts & Stephenson (1994) and is consistent with cooling rates obtained from apatite fisson-track thermochronology and vitrinite reflectance studies by Arne et al. (2002). A more extensive apatite fission-track and (U-Th-Sm)/He dating study also highlighted periods of uplift of the southern Sverdrup Basin rift flanks in response to subsidence from (1) the initial rifting in the Late Carboniferous to Early Permian, (2) a period of increased subsidence in the Early Triassic linked to higher sediment supply within the basin, and (3) local responses to Late Cretaceous rifting within Baffin Bay (Grist & Zentilli, 2005).

More detailed 1-D backstripping of exploration wells was completed for this study in an attempt to identify localized changes in subsidence patterns next to evaporite domes. Deviations from regional subsidence patterns can potentially identify modifications to the thicknesses of sedimentary units by diapirism and may be used to identify specific periods of active evaporite dome growth or cessation. Subsidence analysis consisted of the backstripping of 19 wells located on and just offshore of Ellef Ringnes Island and neighboring King Christian Island.

Figure 3-1. Observed (solid) and modeled (dashed) tectonic subsidence curves for nine locations (AH, northern Axel Heiberg Island; RP, Raanes Peninsula; PP, eastern Prince Patrick Island; CI, Cornwall Island; ER, Ellef Ringnes Island; BM, Blue Mountains, Ellesmere Island; MK, Mackenzie King Island; SP, Sabine Peninsula, Melville Island; BC, basin centre). (a) Model tectonic phases in the Paleozoic and Mesozoic: 1, extension; 2, thermal subsidence; 3, renewed extension; 4, thermal subsidence. Dashed lines highlight age limits from geological observations. (b) Composite tectonic subsidence curves with Cenomanian and younger strata extracted from nearby exposures. The curves show additional tectonic phases of: 5, onset of horizontal compression and 6, culmination of the Eurekan Orogeny and demise of Sverdrup Basin. (c) Basin and tectonic subsidence curves for Mackenzie King Island with the water-depth profile used in backstripping. This diagram outlines the two possible periods (3a and 3b) identified by Stephenson et al. (1994) for the initial rifting of the Canada Basin (Stephenson et al., 1994).

47 3.2 Methods

3.2.1 Backstripping

The subsidence history of the Sverdrup Basin is recorded by its stratigraphy and can be separated into subsidence related to isostatic effects from the loading of sediments/water and tectonic events through a method known as backstripping. Quantitative backstripping methods were developed in the 1970’s as improved paleontological databases became available (van Hinte, 1978; Steckler & Watts, 1978; Bond & Kominz, 1984). Backstripping methods often assume that the vertical sediment load is compensated locally by Airy isostasy. If the lithosphere is assumed to support the sediment load by regional flexure, the separation of tectonic and sediment load contributed subsidence is more complex and requires the input of the flexural rigidity of the lithosphere and the wavelength of the sediment load (Allen & Allen, 1990). Subsidence curves are generally produced by applying three corrections to present day compacted stratigraphic thicknesses: (1) decompaction, (2) paleobathymetry and (3) absolute sea level fluctuations or eustasy (Allen & Allen, 1990). To correct for compaction, porosity curves for various lithologies are required. Change in porosity with depth is often modeled as -cz an exponential relationship φ=φoe where porosity φ is dependant on depth z, surface porosity φo and the coefficient rate c, which is dependant on lithology. Paleobathymetry is often adjusted using data from micropaleontological studies, however it can also be estimated from sedimentary facies and specific geochemical signatures. Eustatic corrections are often the least precise due to the lack of reliable global sea level curves.

For this study, backstripping was carried out with the Petroprob code developed by Van Wees (cf. Van Wees et al., 1996; Van Wees & Beekman, 2000). The code utilizes a multi-1D probabilistic lithosphere tectonic model that is capable of incorporating heat flow parameters and basin inversion. The forward modeling approach is based on the pure-shear lithosphere thinning model of McKenzie (1978), however the code has the added capability of incorporating different finite amounts of thinning of the crust and mantle lithosphere (e.g. Royden & Keen, 1980) defined by stretching factors δ and β respectively. In the code, Airy isostasy is adopted for basin loading with an assumed mantle density of 3400 kg ּm-3. In

48 Petroprob, porosity curves for backstripping are calculated using a double exponential porosity depth law (Van Wees et al., 2009; after Bond & Kominz, 1984):

z − zscale ϕ z)( z<=zscalechange = ϕ0e (1)

zscalechange ⎛ − ⎞ z ⎜ zscale ⎟ − e zscale2 ϕ z)( z>zscalechange = ⎜ϕ0 zscalechange ⎟e (2) − ⎜ zscale2 ⎟ ⎝ e ⎠

where φ is porosity, z is depth in meters, φo is surface porosity, zscale and zscale2 are the coefficient rates and zscalechange is the depth at which the porosity-depth trend changes. The values of each parameter are function of the lithology (Table 3-1). The porosity of each unit is therefore determined from the percentages of the lithologies defined by the user.

Table 3-1. Parameters used for double exponential porosity depth law for various lithologies within Petroprob code (taken from Van Wees et al., 2009).

Lithology φo (%) zscale (m) zscale error (m) zscalechange (m) zscale2 (m) Sandstone 30 2022 667 1000 3356 Shale 61.5 930 16 500 2116 Siltstone 42.5 2038 621 500 2871 Limestone 49 1238 377 500 1993 Salt 6 7000 0 20000 5000 Coal 6 7000 0 20000 5000 Anhydrite 6 7000 0 20000 5000 Carbonate 49 1238 377 500 1993

3.2.2 Data

Ellef Ringnes Island and the surrounding area (including Thor Island and King Christian Island) contain over 35 wells (Fig. 3-2). Approximately 10 of the wells were drilled for permafrost studies and therefore only reach maximum depths of 500 m. The other 25 wells were drilled for exploration purposes with depths ranging from approximately 900 to 4300 m. Information gathered from each exploration well varies, but most often includes the depth

49 of formation tops, description of core logs and various geophysical logs (including sonic, gamma ray, resistivity and self potential). Drilling and logging were completed by several companies, therefore the quality of data and interpretation and nomenclature of formation tops vary between wells. A stratigraphic study conducted by Dewing & Embry (2007) led to the reinterpretation and regional correlation of wells on Ellef Ringnes Island and for the rest of the Sverdrup Basin. The majority of wells extend to just below the oil and gas-bearing reservoir rocks of the King Christian Formation and correlative Heiberg Group. Wells located just offshore of the southern region of Ellef Ringnes Island intersect with some of the oldest formations of Early Triassic age (Blind Fiord Formation). The Isachsen J-37 and Pollux G-60 wells located on Isachsen Peninsula in North Ellef Ringnes Island penetrate even older formations dating back to the Late Carboniferous, however most of the Mesozoic units were eroded during uplift of the Sverdrup Rim and the wells are therefore irrelevant for this study.

3.2.3 Input parameters for backstripping

The relevant information required for backstripping includes stratigraphic time-depth pairs (thickness and age of formations), lithology, compaction/porosity, and depositional water depth (Stephenson et al., 1994). Thicknesses of units were based on reinterpreted formation tops from exploration wells (Dewing & Embry, 2007). The ages of the formation tops were estimated from stratigraphic columns produced by Embry & Osadetz (1988), Embry (1991), Patchett et al. (2004) and Dewing & Embry (2007). Lithologies were broken down into average percentages of sand, silt, clay, carbonate and salt recorded in individual drill logs. Information on lithology was extrapolated from neighboring wells if drill logs were not included and for projected thicknesses of eroded units. Paleo-bathymetry and eustasy were excluded from backstripping. According to Stephenson et al. (1994), estimates of paleo-water depth for the Early Triassic was less than 250 m dropping to 0 m (Fig 3-1c) during the deposition of Late Jurassic to Early Cretaceous terrestrial deposits. Paleo-water depth is therefore negligible during most of the Mesozoic. Compaction/ porosity analysis was not completed for this study due to time constraints.

Figure 3-2. Locations of wells on Ellef Ringnes Island and the surrounding region. The wells used for this study are labeled and highlighted in red.

Backstripping was carried out on a subset of 19 wells located on central and southern Ellef Ringnes Island and King Christian Island (highlighted in red in Fig. 3-2). The wells were chosen based on their intersection with the oldest unit. Wells that did not penetrate deeper than the Heiberg Group or King Christian Formation were excluded. The erosion level at which the wells were drilled varies throughout the data subset. None of the wells intersect formations younger than Cenomanian age (Hassel, Kanguk and Expedition Formations); therefore estimates of pre-erosion thicknesses were extrapolated from isopach maps (Chapter 2). Eroded portions of formations older than Cenomanian age were also estimated from

51 isopach maps and/or extracted from adjacent wells where no erosion of that layer had occurred. Several of the wells, such as Helicopter J-12, contain numerous mafic sills up to 140 m thick. These sills are assumed to have intruded at a later stage of basin development and therefore did not contribute to the overall thickness at that specific time period of sedimentation. The effects of these sills were therefore removed and the thicknesses of units adjusted to consider only the sedimentary influence on basin loading. The input files for the Petroprob backstripping code are included in Appendix 5.

3.3 Results

Due to restriction of most well data to the Late Triassic, results from backstripping of various wells dating back to 200 Ma are compared to examine general trends of subsidence. A few wells with formations dating as far back as 245 Ma are also briefly discussed. The total subsidence for a distinct period of basin evolution can be examined quantitatively, however true tectonic subsidence rates and amounts are dependant on the entire basin loading history and therefore the general trends and changes in tectonic rates for this study can only be examined qualitatively.

3.3.1 Regional and local changes in tectonic subsidence rates

Regionally, the majority of the wells share similar trends in subsidence dating from 200 Ma to 95 Ma (Fig. 3-3). Relative rates of tectonic subsidence can generally be described in four stages. (1) The Rhaetian to Barremian (200 to 128 Ma) is considered to be a period of moderate tectonic subsidence with slight oscillations that vary between individual subsidence curves. For example, several curves located further south (e.g. Char G-07 and King Christian N-06) show a brief period of faster subsidence until 185 Ma followed by slightly slower rates until ~150 Ma where the rates increase again until 128 Ma. These slight variations decrease further north toward the Sverdrup Rim. (2) Subsidence patterns change in the Aptian during deposition of the Isachsen Formation from 128 Ma to 113 Ma. Many, but not all wells, record a brief period of rapid tectonic subsidence for 2 Myr that abruptly decreases until 113 Ma. This period of quiescence is most pronounced in wells located further north in Noice Peninsula (e.g. Dome Bay P-36, Mocklin Point D-23, Noice D-41 and Louise O-25) and further east toward the Dumbells and Contour Domes (Helicopter J-12 and Dumbells E-49).

Figure 3-3. Tectonic subsidence curves backstripped from 200 Ma to 95 Ma. The graphs are split into wells located in (a) the north (Noice Peninsula and Dumbells/ Contour Dome region) (b) central (encompassing the region surrounding Hoodoo Dome) and (c) south (Meteorologist Peninsula, King Christian Island and offshore). See Figure 3-2 for locations of wells. Wallis K-62 (central and south) and Kristoffer Bay B-06 (north and central) are included in two plots for comparison of different regions. Dotted lines indicate estimated eroded formations.

(3) The third stage in the Albian (113 to 103 Ma) records the highest rates of tectonic subsidence within the time period examined for this study. (4) The last stage (103 to 95 Ma) illustrates a decrease in tectonic subsidence rate, with almost no subsidence during the deposition of the Hassel Formation.

King Christian N-06, Kristoffer Bay B-06 and Sutherland O-23 wells demonstrate that prior to 200 Ma, tectonic subsidence was much higher (Fig. 3-4) at rates that are comparable to the rapid subsidence recorded during the deposition of the Christopher Formation (113 to 103 Ma). This rapid tectonic subsidence may have only lasted until ~225 Ma, however, as shown by the King Christian N-06 curve that dates back even further to 245 Ma (Fig. 3-4b).

Figure 3-4. Tectonic subsidence for select wells dating back to (a) 228 Ma and (b) 245 Ma. Dotted lines indicate estimated eroded formations

54 We had hoped to identify local subsidence signatures related to salt movement, however the distance of wells to the domes proved to be problematic. Hoodoo N-52 is the only well where a local signature that deviates from regional subsidence trends could be identified with confidence. This deviation from the regional trends provides the only constraint on the timing of diapir growth from the subsidence curves. Hoodoo N-52 records a period of little to no tectonic subsidence from 185 Ma to 128 Ma during which average regional tectonic subsidence rates remain relatively steady (see comparison to Hoodoo Dome H-37 in Figure 3-5a).

Figure 3-5. (a) Tectonic subsidence of Hoodoo N-52 and Hoodoo Dome H-37. Hoodoo N-52 records a local tectonic signature that deviates from regional trends, represented by Hoodoo Dome H-37. (b) The difference in tectonic subsidence between Hoodoo N-52 and Hoodoo Dome H-37. (c) The change in the rate of local tectonic subsidence signature deviations (Hoodoo N-52) from the regional tectonic signature (Hoodoo Dome H-37). A negative deviation implies that subsidence at Hoodoo N-52 is occurring at a slower rate (due to diapirism) and a positive deviation implies that subsidence is occurring at a faster rate than regional tectonic subsidence trends.

55 Deviations from the regional curves are most pronounced from 185 to 161 Ma, followed by a short period (161 to 153 Ma) where the tectonic subsidence rate roughly matches regional rates (Fig. 3-5c). After 138 Ma, negative deviations in tectonic subsidence rates decrease and the difference in tectonic subsidence between Hoodoo Dome H-37 (regional signature) and Hoodoo N-52 (local signature) also decreases (Fig 3-5b & c).

3.3.2 Contour plots of 1-D subsidence curves through time

The similarity in the trends of tectonic subsidence curves suggest that tectonic influences were regional and not related to salt movement (with the exception of Hoodoo N-52), however the relative amounts of tectonic subsidence varies across Ellef Ringnes Island and King Christian Island. Figure 3-6 plots the regional trend in tectonic subsidence at discrete time periods between 200 Ma and 95 Ma. It should again be emphasized that these tectonic subsidence values are not quantitative as they do not include any effects of basin loading prior to 200 Ma and show only the relative relationships of tectonic subsidence between the individual wells. The contour plots were created using a natural neighbor interpolation method that retains values close to the original data set. Natural neighbor interpolation is similar to triangulation (linear interpolation), however interpolation between data points is weighted based on the size of Thiessen polygons. Hoodoo N-52 is not included as it is the only well that is strongly influenced by salt movement adjacent to Hoodoo Dome. The gradual trend toward higher subsidence rates in the east-central region of Ellef Ringnes Island is consistent through out the Mesozoic. By 75 Ma, the highest tectonic and basement subsidence are concentrated around Helicopter J-12, east of Dumbells and Contour Domes. Subsidence decreases toward the south offshore of both King Christian and Ellef Ringnes Island.

3.3.3 Limitations due to data quality and resolution

The limited depth of wells prevents quantitative analysis of rates and amounts of tectonic subsidence. In addition, limitations of the data also impose restrictions for further basin subsidence modeling. For example, the study was originally intended to include forward tectonic modeling to estimate stretching factors for the crust and sub-crustal lithosphere (i.e. similar to the study conducted by Stephenson et al., 1987); however it is difficult to estimate

Figure 3-6. Contour plots of tectonic subsidence at various periods in basin development. Plots are produced from wells backstripped from 200 Ma to 95 Ma. X/Y coordinates (in meters) are generated with a Lambert Conic Conformal map projection. and incorporate the stretching values inherited prior to the Late Triassic/Early Jurassic and were therefore not computed. Errors related to constrained formation ages are due to the average approximation of ages taken from published stratigraphic columns (Embry & Osadetz, 1988; Embry, 1991; Patchett et al., 2004; and Dewing & Embry, 2007). Though the stratigraphic columns are based on ages obtained from paleontological studies, these studies are basin-wide and therefore ages are summarized on a regional basis. No known paleontological studies have been completed on drill cores from Ellef Ringnes Island, therefore approximations taken from stratigraphic columns may vary slightly from true ages (~ ± 2 Ma; Figure 3-7). Rates of subsidence within each formation are also averaged due to age constraints, therefore rapid changes in basin subsidence restricted to even shorter periods of time may not be resolved. 57

Figure 3-7. Example of error ranges for formation ages and porosity-depth uncertainties for Sutherland O-23 tectonic and basin subsidence curves. Errors are discussed in text.

Errors due to the original estimates of formation thicknesses are related to unidentified unconformities within wells. Unconformities and the amount of erosion within the stratigraphic column are difficult to identify and approximate as contacts are often conformable. Sandy Point Formation, for example, may represent strandplain deposits bound by an unconformity that formed from forced regression during the Toarcian-Aalenian. Unconformities are also reported during the Hauvertian, Aptian and Albian-Cenomanian, however the extent of these unconformities are often restricted to the basin margins in most basin-wide stratigraphic columns (e.g. Embry, 1991). Errors that may arise from possible unconformities would also affect the estimated ages of formation tops. Under estimation of formation thicknesses during periods of basin uplift and erosion will lead to errors in subsidence and uplift rates for specific time periods, however the total tectonic subsidence that occurred would average out to similar values (see Figure 3-8 for explanation). The extrapolation of formation thicknesses from neighboring wells and isopach maps where units were eroded is the least precise. The depositional extent of Late Cretaceous formations, for example, may not have covered certain wells and thinning toward the margins of the basin may also have occurred.

Figure 3-8. Sketch illustrating how basin uplift (associated with erosion) and subsidence equate to periods of little to no tectonic subsidence if estimates of eroded thicknesses are not included in backstripping input files.

In addition to the limitations related to data inputs, errors are also produced within the backstripping process implemented in Petroprob. Uncertainties from porosity-depth curves are incorporated within Petroprob (zscale error in Table 3-1), allowing for sensitivity and calibration analysis on these parameters (van Wees et al., 2009). Data calibration and sensitivity analysis is carried out using Monte Carlo Markov Chain techniques. The effect of errors from porosity-depth sampling on tectonic and basin subsidence curves is demonstrated in Figure 3-7 for Sutherland O-23 well.

3.4 Discussion

3.4.1 Relation between tectonic subsidence trends and diapir growth

Despite limitations in the quantitative analysis of tectonic subsidence and modeling of lithospheric thinning, the primary focus of the study is to examine first order subsidence signatures and relate them to observations from seismic reflection profiles regarding diapir growth. The general trends in tectonic subsidence rates are therefore sufficient to identify

59 possible tectonic triggers that have affected and changed patterns of diapir growth. Diapir growth in relation to tectonism is subdivided into four stages below.

Diapirism in the Triassic The lack of stratigraphic resolution during the Early Triassic makes it difficult to comment on possible links between tectonic subsidence and salt growth during this period. For the small number of wells that do intersect strata of Triassic age, however, there is a consistent signature of rapid basin and tectonic subsidence rates that date from at least 228 Ma to 202 Ma and may continue as far back as 250 Ma (Fig. 3-4). Sweeney (1977) records multiple regions with higher than average basin subsidence rates from 225 Ma to 200 Ma, however tectonic subsidence curves were not produced. The study completed by Stephenson et al. (1994) shows that some regions, such as Ellef Ringnes Island, Cornwall Island and Axel Heiberg Island, record a short period of relatively rapid tectonic subsidence (245 to 240 Ma) followed by basin uplift. This period of rapid tectonic subsidence does not continue any later than 235 Ma, which is inconsistent with the ages obtained from our results. The change in subsidence rate at 202 Ma may be related to a switch from rift-related subsidence to a period of thermal relaxation subsidence. However, Stephenson et al. (1994) suggest that the initial rifting phase only lasted until the mid-Permian (311 to 269 Ma). Sweeney (1977) suggests that the high subsidence rate in the Late Triassic reflected a period of basin extension; however Sweeney’s results were also limited to the Middle Triassic due to lack of data at depth in the basin axial region. Cooling ages from apatite fission-track studies completed in eastern Ellesmere Island and western Greenland record a period of cooling interpreted to indicate increased subsidence and sediment supply within the Sverdrup Basin beginning in the Middle Triassic (Fig. 3-9; Grist & Zentilli, 2005). This event is separated from the initial rifting event and is inferred to reflect epeirogenic exhumational cooling of the distal cratonic sediment source areas during periods of high sediment supply from well-developed river systems. Grist & Zentilli’s (2005) findings correlate well with timing of high rates of Middle to Late Triassic tectonic subsidence identified in this study. Balkwill (1978) also supports the notion of a second rifting phase in the Early Mesozoic with the accumulation of a thick load of sediment in the axial trough and suggests that this period may mark the onset of evaporite migration and diapirism.

Figure 3-9. Apatite fission track ages versus mean track length for samples taken from eastern Ellesmere Island, Devon Island and western Greenland (black circles) and for the Sverdrup Basin (white circles). Group 3 records a period of cooling from increased subsidence and sediment supply associated with epeirogenic exhumational cooling of the distal cratonic sediment source areas from well-developed river systems (Grist & Zentilli, 2005).

This period of rapid tectonic subsidence precedes with the formation of a canopy occurring within the Heiberg Group as interpreted in multiple seismic sections surrounding Hoodoo Dome (e.g. profiles I(F)67 and I(F)68; Chapter 2). The formation of a canopy suggests that diapir growth rate was greater than the rate of sedimentation during the Late Triassic-Early Jurassic. Prior to the formation of the canopy, rapid accumulation of sediments would be associated with a faster rate of tectonic subsidence in the Late Triassic, which would load and pressurize the underlying Otto Fiord evaporites. As sedimentation slowed during the Early Jurassic (beginning ~200 Ma), the rate of evaporite displacement may have exceeded the rate of sedimentation, causing salt to extrude at the surface until a pressure equilibrium was reached with the surrounding sediments.

Figure 3-10. Interpreted tectonic phases and periods of diapirism superimposed on the Sutherland O-23 tectonic subsidence curve.

Diapirism during the Jurassic to Early Cretaceous

The Jurassic and Early Cretaceous are marked by a period with a relatively steady rate of subsidence, which is interpreted by Stephenson et al. (1994) as a post-Paleozoic-rift thermal subsidence phase. In this study, the Jurassic to early Cretaceous is interpreted as a period of passive diapirism where diapir growth and tectonic subsidence (and hence sedimentation) rates are similar. This is supported by subsidence trends from Hoodoo N-52 well, where diapir growth is suggested to match tectonic subsidence to produce a recorded signature of tectonic quiescence for the majority of the Jurassic and Early Cretaceous (Fig. 3-5). During this period, evaporites were most likely unpressurized and at surface in equilibrium with the adjacent sediments. Slow aggradation of sediments on top of the extruded salt at Hoodoo Dome would have terminated the lateral flow of salt at the surface and slowly buried the allochthonous layer to produce salt canopies. Multiple subsequent episodes of salt extrusion and burial during this period may have been triggered by slight variations in tectonic

62 subsidence rates. Passive diapirism during this period is supported by characteristic patterns of thinned and ductile deformed drag folds adjacent to the domes within seismic profiles (Chapter 2; Alsop et al., 2000).

Diapirism and onset of Canada Basin rifting in the Early Cretaceous

The increase in subsidence rate in the Albian (113-103 Ma) is recognizable in most curves produced by Stephenson et al. (1994) for the Sverdrup Basin. Stephenson et al. attributed this signature to a period of possible renewed extension (3b in Fig. 3-1). The event is one of two possible periods identified by Stephenson et al. (1994) that could indicate the initial rifting of the Canada Basin, the other beginning at 157 Ma (3a in Fig. 3-1). Only major subsidence signatures were identified and discussed by Stephenson et al. (1994) as they were examining the tectonic subsidence curves dating back to the initial Permo-Carboniferous rifting event. For this study, smaller variations in tectonic subsidence are more prominent due to the restriction of data to the Mesozoic. For example, the 126 to 113 Ma (Aptian) period of little to no tectonic subsidence leading up to possible renewed extension in the Albian was not previously recognized. This quiescence in tectonic subsidence occurs during the deposition of the Isachsen Formation (Walker Island Member) and may correlate with the unconformity that superimposes flat lying sediments on top of eroded rim syncline deposits east of Dumbells Dome (Chapter 2). If the Aptian event is related to this unconformity, then the event may mark a period of basin uplift and subsequent erosion. Possible uplift in the Aptian is followed by rapid tectonic subsidence during which the Christopher Formation buries the evaporite domes on Ellef Ringnes Island.

Basin uplift and erosion may have been restricted to an even shorter time span, but since the tectonic subsidence events are averaged for each formation they may not be fully resolved. In fact, uplift must be counterbalanced by subsidence during the Aptian to produce a tectonic rate close to zero in subsidence curves (Fig. 3-8). If basin uplift was restricted to a shorter time span than suggested by the subsidence curves, rapid subsidence may have initiated earlier with the deposition of the youngest Isachsen members in the late Aptian. Burial of the domes by both Upper Isachsen members and Christopher Formation correlates with seismic and borehole observations, as well as with surface outcrop patterns. On the contrary, the even 63 shorter period of rapid subsidence prior to the Aptian event recorded within some wells may have continued much longer into the Aptian prior to uplift and erosion.

Basin uplift followed by rapid subsidence are both interpreted as signatures that mark the onset of the Canada Basin seafloor spreading event (period highlighted in Fig. 3-10). Rifting of the Canada Basin north of Ellesmere Island would explain why the Aptian signature is much stronger in tectonic subsidence curves located further north toward the Sverdrup Rim (e.g. Dumbells E-49). If this interpretation is correct, then the initial seafloor spreading of the Canada Basin would date to ca. 128-126 Ma, in between the two time periods proposed by Stephenson et al. (1994). Paleomagnetic data suggests that sea-floor spreading occurred post- Valanginian (Halgedahl & Jarrard, 1987) and ceased prior to the Late Cretaceous (Harbert et al., 1990), constraining initial breakup to the Early Cretaceous. Embry & Dixon (1994) listed three unconformities (Hauterivian, Aptian and Cenomanian) that could possibly represent the breakup unconformity that marks the initial onset of Canada Basin seafloor spreading. Embry & Dixon (1994) interpreted the Hauterivian as the main rift onset or breakup unconformity based on the significant increase of normal fault activity following the unconformity. They also argue that the Hauterivian breakup unconformity (138 to 135 Ma) was also supported by associated volcanism in the Sverdrup Basin. More recent 40Ar-39Ar dating of mafic magmatism in the Sverdrup Basin by Villeneuve & Williamson (2006) suggest that the major pulse of igneous activity peaked later at 129-127 Ma. Hubbard et al. (1987) also determined an age of 128 Ma for the Canada Basin breakup unconformity observed in north Alaska and northwest Canada. The aforementioned peak ages of igneous activity and the unconformity recognized by Hubbard et al. (1987) correlate well with our suggested ages of Canada Basin rift initiation.

Diapirism in the Late Cretaceous

The resolution in the data decreases significantly in the Late Cretaceous as eroded unit thicknesses are approximated from isopach maps presented in Chapter 2. The Albian- Cenomanian unconformity between the Hassel and Christopher Formations (Embry, 1991) correlates with a short period of no tectonic subsidence dating from 103 to 95 Ma. As suggested by Embry & Dixon (1991), the unconformity may represent cessation of seafloor

64 spreading in the Canada Basin. The subsidence data does not have the resolution to suggest that basin-wide tectonic events significantly influenced diapir growth in the Late Cretaceous. As mentioned previously in Chapter 2, diapirism most likely continued during the Late Cretaceous to Early Tertiary with gentle doming and thinning of overlying strata that switched to active piercement of overlying Cretaceous sediments during the Eurekan Orogeny.

3.4.2 Axel Heiberg subsidence

Though the stratigraphic resolution is not high enough to identify subsidence signatures related to salt movement or discrete features such as stratigraphic offset along faults, the first order pattern of subsidence appears to have a convincing regional trend. The trend in tectonic subsidence is consistent with the isopach maps presented in Chapter 2, with the greatest accumulation of sediments on Ellef Ringnes Island located east of the Dumbells and Contour Domes toward the Mesozoic Axel Heiberg depocenter. Possible errors in subsidence related to unidentified unconformities are of less concern as the unconformities are most likely uniform and regional and therefore affect discrete periods of absolute subsidence rates but not the overall relative trend in subsidence between wells.

3.5 Conclusion

Tectonic subsidence curves determined for Ellef Ringnes Island and the surrounding region highlight many possible tectonic events that have affected the rate of diapir growth. A switch to slower tectonic subsidence rate in the Late Triassic may have triggered salt to extrude at the surface of Hoodoo Dome as the rate of sedimentation decreased. From the Jurassic to Early Cretaceous, salt growth matched subsidence rates during a passive stage of diapirism. The mid-Cretaceous is marked by sudden basin uplift followed by rapid subsidence, which correlates with an unconformity within the Isachsen Formation and subsequent burial of the salt domes by the youngest Isachsen members and Christopher Formation. The sudden basin uplift followed by rapid increased subsidence rates are interpreted to be a signature related to the onset of Canada Basin seafloor spreading north of the Sverdrup Basin. The decline in tectonic subsidence during the Cenomanian coincides with the deposition of the Hassel Formation. Reactivation of salt growth and piercement of overlying Cretaceous units is most

65 likely related to the Eurekan Orogeny; however such an analysis is beyond the resolution of subsidence data used in this study. Subsidence during the Mesozoic has a systematic regional trend with higher rates of tectonic subsidence occurring in the eastern region of Ellef Ringnes Island with slower rates occurring south of the island. These results are consistent with sedimentation directed toward the Axel Heiberg depocenter, which records the thickest accumulation of Mesozoic strata of up to 7 km within the eastern Sverdrup Basin (Oakey & Stephenson, 2008).

Further improvements that can be made in the future include increasing the resolution of the subsidence dataset by increasing the number of 1D sampling locations. Synthetic wells could be created to sample formation thicknesses from isopach maps closer to the domes to pick up tectonic subsidence signatures related to salt growth. The 1D subsidence trends could increase temporal resolution and put further constraints on periods of diapiric growth. Thicknesses of units could be extracted from isopach maps presented in Chapter 2, however it is difficult to interpret thinning units proximal to the piercement structures without control from wells. Formation thicknesses can also vary significantly for steeply dipping units depending on the precision and spacing of Vrms values used for migration. The isopach maps presented in Chapter 2 can also be implemented into Petroprob directly to produce higher resolution contoured maps of tectonic subsidence across Ellef Ringnes Island. Vitrinite reflectance studies should also be incorporated for more accurate estimates of compaction and erosion.

4 Analogue modelling of diapirism

4.1 Introduction

4.1.1 The role of anhydrite in diapirism and other tectonic settings

Anhydrite, CaSO4, often forms from secondary or diagenetic dehydration of primary gypsum at low temperatures of 50-60 oC (Warren, 2006). Dehydration of gypsum to anhydrite results in a 25 % increase in density (~ 2900 kg ּm-3) and a release of large volumes of water. In geological settings, anhydrite is known to deform in a ductile manner, even at relatively low temperatures (Muller et al., 1981). During deformation, anhydrite often acts as one of the weakest layers within sedimentary sequences and accommodates the highest strains (Heidelbach et al., 2001). Décollement surfaces beneath thin-skinned fold-and-thrust belts are often localized in anhydrite horizons, such as the basal detachment surface in the Jura Mountains, Switzerland (e.g. Dell’Angelo & Olgaard, 1995), the Anatalya Thrust System in Southern Turkey (Marcoux et al., 1987) and along thrusts in Umbria-Marche Apennines, Central Italy (Hildyard et al., 2009). In the Sverdrup Basin, the ductile and strain-localizing nature of anhydrite is apparent in anhydrite-lubricated faults along the Eureka Sound fold- and-thrust belt, such as the Stolz Thrust and North Mokka Fault (Thorsteinsson, 1974; van Berkel et al., 1983). Anhydrite layers also commonly produce various types of fold patterns in diapir stocks (Talbot & Jackson, 1987; discussion in Zulauf et al., 2009).

Buoyancy is argued to provide a fundamental driving force for salt tectonics, however the presence of dense anhydrite caps reduce the effect of buoyancy driven diapirism (Fig 4-1; Stephenson et al., 1992). This influence is most significant at the initial stages of diapirism, where the buoyancy force is possibly reversed, producing a counteracting force to initial perturbations that would otherwise develop into diapirs. The existence of anhydrite caps alone suggests that other mechanisms are involved, possibly triggered by tectonism.

This chapter presents results from 2D scaled analogue models that investigate the role of anhydrite caps on the development of evaporite piercement structures in Ellef Ringnes Island. In particular, we focus on the mechanics of uplift of competent and relatively undeformed

blocks of dense anhydrite on top of buoyant salt structures. The models involve forced injection of fluids into a downbuilding overburden of sand as an analogue for syn- sedimentary diapiric growth.

Figure 4-1. Buoyancy force for different diapiric heights of anhydrite and salt. The anhydrite-salt curve represents a salt diapir with a 2 km thick anhydrite hood. All three curves are calculated with a 10 km thick package of overlying sediments. Density of salt: 2200 :kg ּm-3, anhydrite: 2900 kg ּm-3, overburden .(kg ּm-3 (Stephenson et al., 1992 2500

4.1.2 Anhydrite piercement structures in the Sverdrup Basin

Anhydrite caprocks often form as a dissolution residue atop diapirs (e.g. Gulf of Mexico, Texas; Warren, 2006), however the anhydrite caps within the Sverdrup Basin were recognized early on to be of primary sedimentary origin and were most likely mobilized along with the underlying salt (Heywood, 1955, 57). The diapiric anhydrite and salt units are correlated with the evaporites of the Carboniferous Otto Fiord Formation (Thorsteinsson & Tozer, 1957), which are interpreted to have formed in a hypersaline subaqueous environment (Davies & Nassichuk, 1991). On Ellesmere Island, the Otto Fiord Formation has a minimum thickness of 400 m and is dominated by thick units of anhydrite and limestone (Davies & Nassichuk, 1975). Incomplete sections closer to the eastern margin of the evaporite belt at Wood Glacier are 460 m thick (Davies & Nassichuk, 1975). In the basin axial region, interbedded anhydrite and limestone are thought to overlie halite sequences in the restricted Barrow and Axel Heiberg sub-basins (Davies & Nassichuk, 1975; Davies & Nassichuk, 1991). In the Sverdrup Basin, bedded anhydrite contains a nodular structure, consisting of irregular masses of anhydrite separated by thin partings of carbonates, clay and organic detritus (Nassichuk & Davies, 1980; van Berkel et al., 1986). The nodules are excellent strain indicators and have been used in finite strain analysis of diapirs on Axel Heiberg Island (van

Berkel et al., 1986; Schwerdtner et al., 1988) and in anhydrite lubricated thrusts (Schwerdtner et al., 1988).

The thickness of the anhydrite hood on the northeastern margin of Hoodoo Dome was confirmed to be approximately 280 m from the Hoodoo L-41 well (Fig. 4-2; Panarctic Oils Ltd., 1972).

Figure 4-2. Gamma Ray and instantaneous velocity (derived from sonic) profiles with depth and two-way-travel time for Hoodoo L-41 well. The anhydrite cap (thickness ~280 m) is highlighted in orange. The green lines represent tops of intrusive sills within the salt.

Estimates of up to 800 m of anhydrite-dominant layers were given by Schwerdtner & Clark (1967) from deeply incised valley exposures on Mokka and South Fiord Domes, western Axel Heiberg Island. Gravity studies conducted over Cape Colquhoun (Sabine Peninsula), Isachsen (Ellef Ringnes Island) and South Fiord (Axel Heiberg Island) Domes gave a range of 200-550 m of anhydrite with the thickness of the caps increasing towards the eastern region of the basin (Spector & Hornal, 1970).

Deformation styles within the anhydrite caps vary regionally. In the western region of the Sverdrup Basin, the domes are generally concentric and composed of thick anhydrite- dominant hoods bordered by a rim or sheath of steeply dipping bedded anhydrite on the order of half a kilometer wide (Fig. 4-3; Gould & DeMille, 1964). For piercement structures on Ellef Ringnes Island, the steeply dipping rims of anhydrite significantly decrease in dip toward the center of the domes (Heywood, 1955; Van Leeuwen, 2005). Primary sedimentary layers of anhydrite and limestone are thicker in the core and are relatively undeformed to gently folded (Van Leeuwen, 2005; Heywood, 1955). The intensity of folding increases toward the dome margins (Heywood, 1955). The center of the domes are faulted and deeply incised by streams that chaotically break up blocks of relatively undeformed to gently folded bedded evaporites (Heywood, 1955; Gould & DeMille, 1964).

Figure 4-3. Photos of Dumbells Dome, illustrating the (a) steeply inclined and sheared anhydrite- dominant margins and (b) gently folded anhydrite-dominant cap with visible primary bedding (courtesy of A.R. Cruden).

In the eastern region of the basin, the internal structures of the anhydrite caps are much more deformed due to late-stage compression during the Eurekan Orogeny. Schwerdtner & Clarke (1967) conducted strain analyzes of buckle folds and radially oriented boudins in the Mokka Fiord and South Fiord Domes and suggested that the ductile strain patterns resulted from thickening of the anhydrite-dominant caps. Another study conducted by Schwerdtner et al. (1988) also suggests that extreme buckle folding occurred within the southern Muskox Ridge Dome due to E-W compression of the Eurekan Orogenic event. At the Stolz Diapir, upright 70

to strongly inclined folds occur within the anhydrite hood at 1 to 100 m wavelengths (Schwerdtner & Kranendonk, 1984). Radial fractures, faults and joints are also common in anhydrite hoods and cross-cut ductile fabrics.

4.1.3 Previous analogue and numerical modeling of diapirism involving anhydrite

Analogue experiments have been carried out on the interaction of anhydrite beds within and on top of salt diapirs. Schwerdtner & Osadetz (1983) addressed the emplacement of anhydrite caps in the western Sverdrup Basin using silicone putty with similar magnitudes of viscosities as analogues for salt, anhydrite and overlying sediments. The experiments were run in a centrifuge with scaled densities and thicknesses. In the model the anhydrite analogue thinned on top of the dome and was eventually pierced by the salt analogue, contrary to what is observed in the Sverdrup Basin. After piercement, the anhydrite analogue collected within synclines between diapirs (Fig. 4-4).

Figure 4-4. Cross-section through the final stage of a centrifuge model of salt diapirism involving ductile analogues for anhydrite, salt and clastic overburden (Schwerdtner & Osadetz, 1983).

Koyi (2001) carried out downbuilding experiments with a sand layer embedded within silicone as an analogue for anhydrite interlayered within salt. At the onset of diapiric rise, the layer was entrained within the diapir. As the model progressed, however, the layer broke apart and as the rate in diapiric rise decreased, sections sank within the diapir (Fig. 4-5). These results are also dissimilar to observations made in Sverdrup Basin where the cap remains competent and therefore does not sink as discrete blocks of anhydrite within the

diapir. These observations also contradict the dominance of ductile anhydrite structures within salt stocks worldwide (Zulauf et al., 2009).

Figure 4-5. (a) Profile and (b) line drawing of Koyi’s (2001) experiment of 2D diapirism with entrained dense sand layers that were used to model brittle anhydrite layers.

Further evaluation of anhydrite entrainment in salt domes was conducted numerically (Chemia et al., 2008; Chemia & Koyi, 2008; Chemia et al., 2009). Chemia et al. (2008) and Chemia & Koyi (2008) examined the influence of sedimentation rate, salt viscosity, perturbation width and stratigraphic position of the anhydrite layer within the salt source on entrainment rates. Their models, however, eventually led to boudinage and descent of dense entrained blocks of anhydrite as the source layer became depleted and the rate of diapir ascent decreased. The numerical studies demonstrated that the entrainment of the dense anhydrite blocks is possible if the rate of diapiric rise is high enough to overcome the rate of anhydrite descent (governed by density and viscosity difference between salt and anhydrite and the size/geometry of the diapir). A maximum height of anhydrite entrainment produced by the models scaled to 3.5 km, much lower than the expected stock heights of the diapirs in the Sverdrup Basin. In addition, they concluded that entrainment of embedded anhydrite was most effective in cases where the anhydrite layer occurs within the middle of the salt source layer. This contradicts the observation of uplifted anhydrite hoods on top of salt diapirs in the Sverdrup Basin.

The analogue experiments above most likely did not uplift competent and undeformed anhydrite blocks due to the use of viscous and brittle end-member analogues for anhydrite. Brittle end-member analogues of anhydrite (Koyi, 2001) did not demonstrate the competence observed in anhydrite hoods located in the Sverdrup Basin, suggesting that cohesion of the sand employed was too low. Schwerdtner & Osadetz (1983) used a viscous analogue due to the fact that many of the domes and salt structures on Axel Heiberg Island show ductile deformation within the anhydrite layer. In the case of Ellef Ringnes Island, the sheared anhydrite margins suggest that anhydrite behaves in a ductile manner, however the lack of ductile fabrics within the anhydrite caps indicate that the deformation in anhydrite is localized within regions of high strain, suggesting that a ductile plastic analogue may be more appropriate. Numerical models conducted by Chemia et al. (2008), Chemia & Koyi (2008), and Chemia et al. (2009) assigned a power law rheology where the effective viscosity is non-linearly dependant on the differential stress with a power exponent of 2. Power law rheology was chosen due to limitations of implementing brittle rheologies in the numerical code and anhydrite was therefore modeled with an effective viscosity ranging from 1019 to Pa ּs. Implementation of a power law viscous rheology may be an improvement over 1021 the use of purely Newtonian materials as it is able to better localize at regions of high stresses, however even better strain localization may occur with the use of a plastic rheology.

4.2 Methods

4.2.1 Model set-up

Due to the lack of imaging and data from the Middle Triassic and earlier, little is still known about the first stages of diapir development. From the Late Triassic to the Cretaceous the diapirs were already well established and grew during a passive stage of diapirism (Chapter 2 & 3). The experiments in this study therefore do not focus on the dynamics involved in the growth of the diapirs, but instead the interaction between the dense anhydrite layer and the developing diapir. To monitor the growth of a single diapir, a 2D set-up, similar to setups developed by Davison et al. (1993) and Schultz-Ela et al. (1993), was used where a salt analogue is forced through an opening to control the general diapir dimensions and rate of growth. Davison et al. (1993) were interested in the patterns of fault splays produced due to a

rising diapir (Fig. 4-6), and therefore chose a rigid indentor forced into an overburden of sand to model salt dome growth. The goal for this study was to produce a similar result in which a more competent pre-movement layer, which in our experiments would represent the anhydrite, could be captured as an undeformed hood. Modifications to Davison et al.’s setup included the use of a Newtonian salt analogue similar to the set-up of Schultz-Ela et al. (1993), with an appropriate average sedimentation rate for Ellef Ringnes Island. By utilizing silicone as a salt analogue instead of a rigid block indenter, we hoped to test that despite the density contrast with the anhydrite, the anhydrite layer would remain a competent hood and not sink within the salt diapir.

Figure 4-6. Davison et al.’s (1993) set-up showing the development of a central horst on top of a rigid indenter, simulating salt dome growth and overburden deformation.

Experiments were carried out in a Plexiglas tank measuring 48 × 25 × 12.5 cm, with two rigid horizontal plates separating a bottom chamber that houses the fluid salt analogue from the top compartment (Fig. 4-7). The two plates are separated by a 7 cm wide gap that predefines the flow field and hence the average geometry of the diapir. The size of the gap is scaled to estimates of the average diameter of Ellef Ringnes Island piercement structures at surface. The forced injection of the fluid salt analogue is kinematically controlled by a single piston at an average rate of 2 cm ּhr-1. A 7 mm thick layer is placed on top of the base plate to represent a pre-halokinetic layer of anhydrite. Sedimentation is implemented during the experiment by sifting granular material at a constant rate of ~5 mm every 15 to 20 minutes (see scaling in Table 4-1). The granular materials are scraped after each 5 mm interval of sedimentation to produce a flat surface, which creates variations in thickness along the dome margins.

Figure 4-7. (A) Photo of the experimental apparatus. (B) Dimensions of the apparatus. (C) Sketch showing the various material layers. Silicone is forced by a piston through a 7 cm wide slot at a rate of 2 cm ּhr-1. Sedimentation is applied at a similar rate to simulate passive diapiric growth (where sedimentation and salt growth are approximately equal).

The rate of forced injection of the fluid salt analogue is chosen to roughly match the sedimentation rate and to appropriately model a situation where dome growth is dominated by passive diapirism close to the surface, as is suspected for the growth of the Ellef Ringnes Island piercement structures during the Mesozoic. The experiments are monitored from the side by a time lapse camera to capture the cross-sectional progression of the diapir.

4.2.2 Materials and scaling

The geometric dimensions and material properties of the analogue models are scaled to nature (Table 4-1). Scaling ratios are calculated by dividing the model parameters by the nature parameters. For example, the length scale ratio L= lm/lp, where lm is the model length and lp is the prototype length (nature). Material properties, density and viscosity, are scaled similarly to give scaling factors P and M respectively. The duration and time scale ratio of the model are dependant on scaled ratios of the chosen materials and model dimensions and can be calculated using the relationship T=M/PLG where in this particular case, G=1 (scale

ratio of the gravitational constant). Stress and strain rate scale ratios can also be determined from relationships LGP and LGP/M respectively (Table 4-1).

Table 4-1. Scaled parameters for model and nature with corresponding scale ratios. Subscripts h, a, s denote halite, anhydrite, and sediments respectively.

Units Model Nature Ratio

-3 2 -5 Thickness, ha m 7.0 ×10 5.00 ×10 1.30 ×10 - Kg ּm ּs 2 3 Density ρh 1 9.75 ×10 2.20 ×10 0.44 - Kg ּm ּs 3 3 Density, ρa 1 1.28 ×10 2.90 ×10 0.44 - Kg ּm ּs 3 3 Density, ρs 1 1.06 ×10 2.40 ×10 0.44

18 -17 Viscosity, ηh Pa ּs 56 2.75 ×10 2.04 ×10 Strain rate, ε s-1 2.59 ×10-3 1.00 ×10-14 2.59 ×1011 8 -5 Yield stress, σa Pa 80 1.50 ×10 5.28 ×10 Velocity, v m ּs-1 4.63 ×10-6 1.25 ×10-12 3.70 ×106 Time, t s 1.0 3.55 ×1012 3.55 ×10-12 Gravitational constant, g m ּs-2 9.81 9.81 1.00

The dome diameter and unit thicknesses are scaled so that 1 cm in the model is approximately equal to 1 km in nature. An average thickness of 500 m was assumed for the anhydrite layer. This value is higher than the estimates of 250-300 m given for the thickness of the anhydrite hoods for Isachsen and Hoodoo Domes respectively (Spector & Hornal, 1970; Panarctic Oils Ltd., 1972); however, other domes unaffected by shortening from the Eureka Sound fold belt may contain up to 500 m thick anhydrite hoods (such as the Cape Colquhoun Dome on Sabine Peninsula; Spector & Hornal, 1970). An approximation of 500 m for the initial thickness of the anhydrite layer is therefore appropriate for this study. An average sedimentation rate of 0.05 mm ּyr-1 was assumed for Ellef Ringnes Island. This rate was also estimated for the vertical displacement of salt to model passive diapirism, which is .(within the range of rates estimated for salt flow in diapirs (0.01-0.5 mm ּyr-1; Koyi, 2001

Salt analogue Salt has an average density of 2200 kg ּm-3 and is often modeled assuming a Newtonian ,.viscosity in the range of 1017 to 1019 Pa ּs (e.g. Bahroudi & Koyi, 2003; Gaullier et al 1993; Koyi et al., 1993; Nalpas & Brun, 1993; Sani et al, 2007; Sans, 2003; Talbot, 1992; Vendeville & Jackson, 1992 a&b; Weijermars et al., 1993). Strain rates and stresses for Newtonian materials are linearly dependant and are generally used for modeling materials deforming by diffusion creep (e.g. Ranalli & Fischer, 1984). As was previously discussed in Chapter 1, diffusion creep in natural rock salt is activated with extremely low amounts of intercrystalline brine, which will dominate at the low strain rates characteristic of geological deformation (Weijermars et al., 1993). We also adopt a Newtonian viscous behavior for the upwelling salt with a viscosity within the range mentioned above (see Table 4-1). A silicone polymer or polydimethylsiloxane (PDMS) silicone oil (“pure silicone 60,000 cSt fluid” from Clearco Products Co.) was used in this study as an analogue for salt. The silicone is a -transparent Newtonian viscous fluid with a viscosity of 56 Pa ּs and a density of 975 kg ּm 3 (Boutelier et al., 2008). The low viscosity silicone was preferred to the commonly used high viscosity PDMS (viscosity = 2.5×104 Pa ּs; Weijermars, 1986) in order to obtain a reasonable experiment duration as defined by the scaling.

Overburden analogue A mixture of quartz sand (#505 Silica 50M from Bell and McKenzie Ltd.) and ceramic microspheres (Z-light ceramic microspheres from 3M Ltd.) was used as an analogue for sedimentary overburden. Z-light microspheres were added to quartz sand to decrease the density to an appropriately scaled value to model an overburden with a natural density of .(kg ּm-3 (Table 4-1 2400 Granular materials are used to model rocks that deform as a Mohr-Coloumb material with brittle failure (Byerlee, 1978). Granular materials commonly used for analogue modeling, however, do not maintain constant frictional properties and often exhibit strain hardening followed by strain softening before stable strength is reached (Lohrmann et al., 2003). Properties of granular materials are also dependant on handling techniques; poured sand often has a lower bulk density and a lower peak friction angle and cohesion than sifted sand (Krantz, 1991; Lohrmann et al., 2003). The friction angle and cohesion of the two granular

materials were derived from ring shear tests conducted at the GeoForschungsZentrum, Potsdam, Germany. The quartz sand has a cohesion of 3 Pa and a friction angle of 36° at peak strength, which decreases to 23° at stable strength. Z-Light microspheres have a friction angle of 26° at peak strength and cohesion of 82 Pa. The friction angle of Z-Light microspheres remains relatively constant, reducing to 25° at stable strength. The cohesion and frictional properties of the mixture are assumed to be within the range of the two end member materials.

Anhydrite analogue

Anhydrite is assumed to have a density of 2900 kg ּm-3 with a plastic-ductile rheology as is suggested by geological observations. To simulate a plastic rheology, a yield strength of 150 MPa was adopted at which anhydrite is expected to significantly weaken in nature (Dell’Angelo & Olgaard, 1995).

Figure 4-8. Material properties of the analogue anhydrite layer from oscillatory stress sweep tests at 20 oC and 0.1 Hz using a TA AR1000 rheometer. The material is visco-elastic below 70 Pa as the phase shift delta is approximately 45o (0o purely elastic, 90o purely viscous). At stresses above 70-90 Pa the material reaches its plastic yield strength σa, becomes viscous (delta ~90o) and weakens significantly. The spread of the data below 5 Pa is an artifact of poor measurement resolution of the rheometer at those stresses.

A mixture of mineral oil (heavy lab grade M-9300 from ACP Chemicals Inc.), Vaseline and bentonite clay was used to simulate a visco-elastic plastic layer. The clay was added to control the density at the appropriate scaled value of 1280 kg ּms-1 (Table 4-1). Materials were mixed at temperatures over 60 °C and poured into a mold in which the material was allowed to cool and harden. Material properties of the analogue anhydrite layer were determined from oscillatory stress sweep tests using a TA AR1000 controlled-stress rheometer. Once plastic materials fail, their strength is significantly reduced; therefore the rheometer’s heating plate was set past the reset temperature of the material and allowed to cool to 20 °C prior to conducting stress sweep tests. Results from the stress sweep tests show that the material behaves as a visco-elastic material until the material reaches its plastic yield strength, σa, at 70-90 Pa. Once it reaches this yield stress, the material becomes viscous and weakens significantly (Fig. 4-9).

4.3 Results

Seven experiments were completed that examine the behaviour of an overlying anhydrite layer on the development of a single diapir (Table 4-2). Multiple experiments were intended to test the affects of different sedimentation rates and thicknesses of the anhydrite layer. The chosen set-up, however, proved to have unexpected challenges that prevented full implementation of the experimental program. The following sections discuss the general behaviour and problems encountered during the experiments.

4.3.1 Limitations due to boundary conditions

The experiments were designed as a 2D set-up with free-slip side walls in order to monitor the rise of the diapir. If free-slip boundary conditions are perfectly implemented in each model, then the rise of the diapir is expected to be cylindrical. However, diapirs that formed above the base plate most often developed an asymmetric geometry due to variation in the degree of adhesion of the analogue anhydrite layer along each Plexiglas side wall. As the experiments progressed with sedimentation, the rising diapirs became more asymmetrical, producing a gap between the ascending side of the anhydrite layer and the side wall.

Table 4-2. Summary of experiments including the variable rates of sedimentation/ compression of the piston and the dimensions and rheology of the anhydrite layer used for each experiment.

Exp. Anhydrite analogue Experimental rates Summary/ notes # characteristics

1 PDMS mixed with BaSO4 Piston: 18 mm/hr I had difficulties placing anhydrite layer within the tank. The layer (Newtonian viscous Sedimentation: 20 mm/hr had no slip boundaries along the tank walls, making it difficult to analogue) 9 mm thick. monitor the behaviour of the developing dome. Instabilities within Layer placed 1 cm above the viscous layer led to “downwelling plumes” of the anhydrite base plate. analogue into the thin salt layer above the base plate.

2 Plastic analogue (see text), Piston: 21 mm/hr The diapir rose faster at the back wall of the tank. Little to no 9 mm thick. Sedimentation: 20 mm/hr dome development occurred along the front wall (recorded side). Layer placed on base plate. (5 mm every 15 min) Two small cracks formed at the crest of the dome; however the escape of silicone along the back wall prevented further propagation of the cracks toward the front wall.

3 Plastic analogue, 7 mm Piston: 21mm/hr, The experiment rose asymmetrically and a crack formed in the thick. Sedimentation: 20 mm/hr middle of the plastic layer, initiating from the highest risen side Layer placed on base plate. (5 mm every 15 min) (back wall). Little to no dome development along the front wall occurred (recorded side). Salt began escaping after 3 hours, preventing further uplift of the sediment/ anhydrite analogues.

4 Plastic analogue, 7 mm ----- Anhydrite layer broke while being placed in the tank. thick. Layer placed on base plate.

Table 4-2 cont’d.

Exp. Anhydrite analogue Experimental rates Summary/ notes # characteristics 5 Plastic analogue, 7 mm Piston: 24 mm/hr, The anhydrite analogue layer split down the middle, starting at the thick. Sedimentation: 15 mm/hr front wall and propagating to the back wall. Once the crack Layer placed on base plate. (5 mm every 20 min) developed, silicone began to flow at the surface of the model. Sedimentation rate was low enough to have a partially exposed anhydrite “hood”. This experiment was the least asymmetrical out of all the experiments.

6 Plastic analogue, 7 mm Piston: 21 mm/hr, A crack developed on the uplifted margin perpendicular to the thick. Sedimentation: 15 mm/hr front wall, close to center of the dome. Uplift of anhydrite layer Layer placed on base plate. (5 mm every 20 min) ceased at 4 hours due to escape of silicone along the front wall. Many “canopies” formed during the development of the dome (while the anhydrite analogue continued to rise).

7 Plastic analogue, 7 mm Piston: 23 mm/hr The experiment rose asymmetrically, allowing salt to migrate thick. Sedimentation: 20 mm/hr along the front wall after 2.5 hours. No cracks developed within Layer placed on base plate. (5 mm every 15 min) the crest of the dome. One fault began to propagate close to the end of the experiment at the edge of the diapir where the anhydrite layer was lifted off the base plate; however this appears to be a product of the anhydrite layer locally adhering to the base plate and the exploitation of pre-existing weaknesses in the layer.

8 Plastic analogue, 7 mm Piston: 21 cm/hr The cap rose asymmetrically in both the length and width of the thick . Sedimentation: 15 mm/hr box, with one corner adhering to the base plate. The right side of Layer placed on base plate. (5 mm every 20 min) the dome rose so that salt escaped along the margin of the dome, Dome margins cut. producing canopies that continued to thicken as the experiment progressed.

Figure 4-9. Time progression of Experiment 6 illustrates the general development of the salt diapirs in the conducted models. This particular experiment developed a tear in the anhydrite layer, allowing for silicone to migrate into the overlying sand layer, producing a variety of salt “sills” and “canopies”. The experiment also shows salt escaping along the wall due to asymmetrical uplift.

The gap eventually allows silicone to escape along the margin where maximum uplift occurs and stops further growth of the dome. The non-cylindrical geometry of the diapir is therefore a consequence of the boundary conditions rather than a significant feature of dome growth. In nature a diapir forms as a 3D structure without “free-slip sidewalls” and the salt must overcome the strength of the overlying anhydrite by breaking its integrity. Due to these boundary effects, care must be taken in the interpretation and quantification of the results.

4.3.2 Problems localizing sheared diapir margins

At initial stages of the experiments, the anhydrite layer is partially lifted up off the base plate allowing silicone to flow and distribute laterally. We do not want the anhydrite to adhere to the base plate, as the main objective for the base plate is to localize flow and not fix the anhydrite material in space. The observed lateral flow between the base plate and the bottom of the anhydrite layer is therefore valid, however it reduces the rate of vertical diapir growth. The initial perturbation in the uplift of the anhydrite layer from the localized flow through the gap in the base plate allows for differential loading of sedimentation along the length of the box. Ongoing sedimentation on the margins of the developing dome decreases the rate of lateral growth between the base plate and the anhydrite layer as the experiments progress, forcing the salt laterally toward zones of lower pressure (i.e. the developing dome; see Figure 4-9 for the general progression of the models).

Figure 4-10. Experiments failed to develop sheared margins. Instead the anhydrite layer failed at the area of maximum curvature near the center of the dome for (a) Experiment 6 and (b) Experiment 5.

Shear or fault zones were not localized along the margins of the diapirs as the strength of the anhydrite layer proved to be too strong. This prevented further development from a broad wavelength pillow stage to a more focused diapir. Instead, faults developed only at the top of the diapir where maximum curvature occurred, allowing the silicone to pierce through and prevent further progressive uplift of the anhydrite and sediments (Fig. 4-10).

In order to overcome the inability to break the integrity of the anhydrite layer along the developing dome margins, the anhydrite layer was cut along the dome margins prior to commencement of Experiment 8. In this case, we assume that other mechanisms control the shear thinning and weakening to form marginal shear zones, and examine the behaviour of the plastic hood with respect to the upwelling silicone. Boundary conditions again proved to be an issue in this experiment and prevented the symmetrical development of a dome. In this case, asymmetry is present in both the width and length of the tank, with one corner adhering to the base plate and side wall (Fig. 4-11). Despite these challenges, the competent hood did not show signs of negative buoyancy effects overcoming the development of the dome. The anhydrite hood remained a competent block that “rafted” on the top of the rising salt even after silicone began escaping along the Plexiglas side wall. For this particular experiment, asymmetrical rise allowed silicone to leak out at the margins of the uplifted anhydrite dome onto the surface of the model, which produced numerous discrete “canopies” once buried by ongoing sedimentation. Canopies that remained connected to the salt stock continued to thicken throughout the duration of the experiment (Fig. 4-12).

Figure 4-11. An example of how complex side wall and basal boundary interactions and internal rheological behaviour can result in asymmetric dome growth for Experiment 8. The anhydrite analogue layer remained fixed at one corner throughout the duration of the experiment (red circle).

Figure 4-12. Experiment 8 with pre-faulted margins shows the uplift of the anhydrite layer and extrusion of salt on the right margin. The extruded salt produces canopies once buried by ongoing sedimentation. The anhydrite is highlighted in orange, 5 mm sedimentation periods are marked in green and salt “canopies” and “glaciers” are highlighted in pink.

4.4 Discussion

4.4.1 The influence of material properties on the development of sheared dome margins

Deformation experiments conducted on natural rock samples illustrate the complexity of the rheology of anhydrite and its dependence on temperature, strain rate, confining pressure and grain size (Muller & Siemes, 1974; Muller & Briegel, 1978; Muller et al., 1981; Dell’Angelo & Olgaard, 1995; Heidelbach et al., 2001). Experiments conducted at temperatures expected in nature are inconsistent with the observations made in the field. In low temperature deformation experiments, anhydrite is shown to deform in the brittle-plastic regime by twinning, kinking and microfracturing (Muller & Siemes, 1974; Muller & Briegel, 1978; Ross et al., 1987). Embedded anhydrite layers in halite rock samples are also found to form fracture boudins under bulk constriction (Zulauf et al., 2009). The formation of fracture boudins and faults, however, is less common than curtain folds and other ductile structures frequently found in salt diapir stems (Zulauf et al., 2009; Talbot & Jackson, 1987). Brittle structures such as faults and joints are observed in exposed anhydrite caps within the Sverdrup Basin; however, the structures were superimposed on earlier ductile fabrics at final stages of emplacement (Schwerdtner & Clarke, 1967).

Muller et al. (1981) noted that fine grained anhydrite weakens significantly at temperatures of 100-200 oC (depending on grain size and strain rate). At temperatures greater than 300 oC, fine grained anhydrite samples become strongly temperature and strain-rate dependant (Muller et al., 1981). Dell’Angelo & Olgaard (1995) identified two flow regimes at high temperatures: (1) dislocation creep and twinning at high stresses with a power exponent of 5 and (2) diffusion creep at low stresses. Dell’Angelo & Olgaard (1995) were the first to suggest that diffusion creep is a dominant deformation mechanism for anhydrite at geological strain rates and described the mechanism to be similar to “superplastic flow” where materials can deform to very high strains under tensile stresses without breaking. This description of “superplastic flow” is similar to what is observed at the anhydrite dome margins, however the anhydrite caps show an opposite affect, displaying high strength that resists or opposes the formation of ductile flow structures.

In addition, the temperatures at which “superplastic flow” is observed is much higher than temperatures expected from normal geothermal gradients within basin sedimentary rocks. Due to the complexity and inconsistencies of laboratory deformation experiments, a simplified plastic material was used to match what is observed in nature; however, the anhydrite material failed to replicate the anhydrite sheared margins and competent cap. While using a visco-elastic plastic material as an analogue for anhydrite may be an improvement over past physical models, the (1) yield strength, (2) strain softening and (3) elastic properties of the Vaseline-oil-clay mixture may still not accurately reproduce the behaviour of anhydrite under geological conditions:

(1) It is difficult to extrapolate laboratory deformation of rock samples to nature as these experiments are not performed at appropriate geologic conditions and rates. Such laboratory techniques, in particular, are not capable of conducting experiments at lower strain rates more appropriate for diapirism (10-16 to 10-12 s-1; Jackson & Talbot, 1986). For modeling a ductile plastic anhydrite analogue, a yield stress must be approximated and in our case was therefore estimated from flow laws derived by such deformation experiments. This approximation may therefore overestimate the strength of anhydrite in low strain rate settings such as diapirism. Little is also known at what depth and temperature diapirism began in the Sverdrup Basin.

The strength of the anhydrite in nature may have varied over a large range during the Late Paleozoic to Cenozoic as tectonic conditions changed and diapirs matured.

(2) Other strain softening mechanisms may have further weakened the dome margins at the onset of diapirism, which are difficult to reproduce in laboratory materials. A switch from dislocation to diffusion controlled creep at high strains (e.g. facilitated by reduction in grain size from dynamic recrystallization) would cause extreme softening in natural anhydrite (Dell’Angelo & Olgaard, 1995; Heidelbach et al., 2001; Barnhoorn et al., 2005). Crystallographic preferred orientations of minerals that develop with ongoing strain also facilitate deformation by grain boundary sliding (Bruhn & Casey, 1997; Heidelbach, 2001; Barnhoorn et al., 2005; Hildyard et al., 2009). Increased water content may lead to pressure solution in areas influenced by high stresses that effectively weaken the material at the dome margins. Pressure solution textures are recognized within rock samples from the marginal regions of Dumbells Dome (Van Leeuwen, 2005).

(3) The elastic properties of the Vaseline-oil-clay mixture may have had a significant affect on the inability to localize sheared anhydrite margins. Figure 4-9 shows the broad wavelength response of the analogue anhydrite and overlying sediments to the flow of silicone through the gap in the base plate. The broad wavelength suggests that the analogue anhydrite layer is highly elastic and aided in the diffusion of stresses laterally beneath the layer. The broad convex flexure of the dome margins may have instead focused stresses within the analogue anhydrite layer at the very center of the diapir, forcing it to break due to built-up tension. The anhydrite layer in nature may have been more rigid, preventing the flexural response observed in the analogue models.

4.4.2 Localization of salt flow and the influence of boundary conditions

The experiments were unable to localize sheared anhydrite margins along the developing domes, which prevented the transition from an initial pillow stage to a fully developed diapir. Instead the initial perturbation was diffused under the anhydrite layer by lateral viscous flow. Localization of marginal shear zones likely did not occur because shear stresses at the margins were insufficient to cause failure of the anhydrite layer. This is in part related to the

rheological properties of the analogue anhydrite layer mentioned above; however localized silicone flow is highly dependant on the pressure gradients produced by the base plate geometry.

Quick tests were completed to see if dynamic topography (perturbation) could be produced without a confining top layer (anhydrite). Dynamic topography, however, was only produced at unrealistically high rates of silicone injection required to counteract the lateral spread of the silicone fluid. A measure of the amount of dynamic topography that can be produced is given by the Argand number (England & McKenzie, 1982; Buck & Sokoutis, 1994): ρgh 2 Ar = (1) u p μ where ρ is the density, g is the gravitational constant, h is the height of the layer, μ is the viscosity and up is the imposed velocity boundary condition. The dimensionless Ar predicts whether buoyancy or viscous forces will dominate and cause a topographic perturbation to form under the imposed velocity boundary conditions. If Ar is >> 1, then gravity collapse of the perturbation dominates and dynamic topography will not form. For our experiments, even with an initial layer thickness of 1 mm, the Ar number suggests that a significant perturbation will not form (Ar>35). The addition of an overlying anhydrite layer with an elastic strength would further prevent the formation of a dynamic topography. The anhydrite layer was therefore placed directly on top of the plate, and with the addition of differential loading from sedimentation, the dome was forced to localize through the gap of the base plate, despite the counteracting forces from the elasticity of the analogue anhydrite layer and normal (gravitational) stresses produced by the perturbation.

A major problem, however, is what plays the role of the base plate in nature? The wavelength of the diapir is dependant on the width of the gap, and therefore our results are highly dependant on the imposed boundary conditions. With a smaller gap, for example, the pressure gradient would be more localized and may be able to produce stresses that overcome the strength of the anhydrite layer. The development of the diapir and the structures we observe in the anhydrite layer are a result of the dynamics of diapir initiation and can not be easily simplified in a 2D injection set-up. Similar approximations are implemented in

dynamically driven models, such as in the numerical models of Chemia et al. (2008) and Chemia & Koyi (2008). Though the models were allowed to develop dynamically by the downbuilding of sedimentation, an initial perturbation width is also implemented to initiate diapirism. Multiple numerical models with various perturbation widths showed that the geometry of the dome is strongly dependant on this parameter. In the case of the Sverdrup Basin, sedimentation was most likely uniform in the basin axial region where the domes are found to occur, therefore other mechanisms are required in addition to down-building to focus the flow of salt and allow the transition from an initial perturbation to a fully developed diapir. Basement faults and grabens would localize flow and create differences in the initial depositional thickness of the evaporites and overlying clastic sediments. As suggested in Chapter 1, faults that propagate into overlying sediments would also provide weaknesses for diapirism to exploit and affect the resulting geometry of developing diapirs. If there is extensive faulting within the overburden, however, the anhydrite layer also has a higher chance of developing weaknesses for salt to pierce through. The influence of thick-skinned over thin-skinned extension, however, tends to diminish and localize deformation within overburden and may have aided in maintaining the integrity of the anhydrite layer.

4.4.3 The affect of buoyancy on diapirism within the Sverdrup Basin

Hoen (1964) argued that the strength and specific gravity of anhydrite was too great to permit diapirism, and suggested that diapirism originated while anhydrite still existed as primary gypsum that did not begin to dehydrate to anhydrite until the sedimentary overburden reached thicknesses of 5 km. He argued that diapirism ceased once mixed gypsum and anhydrite achieved comparable densities to overlying sediments. Thorsteinsson (1974), however, discredited this argument as anhydrite-gypsum equilibrium relationships do not permit the existence of gypsum at depths greater than 600 m. Warren (2006) states that the depth of transformation ranges between a few meters to a kilometer, however, a burial of at least 1600 m is required before the average density of the entire overlying column of sediments exceeds that of salt (Hudec & Jackson, 2007); therefore it is highly unlikely that a large component of diapirism was initially driven by the positive buoyancy of gypsum and salt. Schwerdtner & Osadetz (1983) suggested that alteration of gypsum at the onset of burial would have produced a “crystal mush” or mixture of anhydrite and water. Retention of water

within the layer would have significantly decreased the strength and density of the layer. After early stages of diapirism, the water would have slowly percolated through fractured overburden to form dense competent, brittle anhydrite caps that acted as passive plugs to upwelling salt (Schwerdtner, 1983; Schwerdtner & Osadetz, 1983). Several authors, however, suggest that the retention of water would be short-lived under the conditions of a rising diapir (Stephenson et al., 1992; Williams-Stroud & Paul, 1997). Also, if water created from dehydration is not able to drain freely, then beds often convert to a quicksand consistency that leads to intense deformation (Warren, 2006). The lack of intense deformation within anhydrite caps on Ellef Ringnes Island further suggests that water retention did not occur (Williams-Stroud & Paul, 1997).

Stephenson et al. (1992) argues that buoyancy driven diapirism is highly unlikely as anhydrite caps would diminish and possibly reverse the effects of buoyancy at the initial stages of diapirism; however the affect of anhydrite on diapirism is greatly dependant on the thickness of the original anhydrite layer. Buoyancy pressure curves produced by Stephenson et al. (1992; Fig. 4-1) assume a 2 km thick anhydrite hood, which is more than double the reported thicknesses observed on top of diapirs in the Sverdrup Basin. In addition, the thickest anhydrite hoods (~800 m) are all located in the eastern region of the Sverdrup Basin and are reported to have undergone extensive thickening due to compression from the Eurekan orogenic event (Schwerdtner & Osadetz, 1983). Thicknesses of undeformed anhydrite caps are more in the range of 200-500 m, which has a much less impact on diminishing the affect of buoyancy forces. Despite the challenges faced with the set-up, our experiments show that anhydrite layers scaled to present day thicknesses are capable of being uplifted if the cap remains competent, such as in Experiment 8. Though the extents of erosion/dissolution of the anhydrite caps remain unknown, we consider that buoyancy is of secondary importance, and that the ability to uplift anhydrite caps is more dependant on the rheology, in particular the strength, of the layer. Though the mechanisms that localize and weaken anhydrite margins remain unclear, as Schwerdtner (1983) suggested, the hood acts as a competent unit that may have behaved like a battering ram that penetrated the clastic overburden. Numerical experiments conducted by Chemia et al. (2008), Chemia & Koyi (2008) and by Chemia et al. (2009) suggest that the entrainment of anhydrite layers is highly

dependant on the relative flow rate and viscosity ratio of salt, however their anhydrite layers were segmented into discrete blocks that allowed the flow of salt to surround the blocks. The anhydrite caps in the Sverdrup Basin, however, act as plugs that are impermeable to salt and suggest that the material is much stronger and more resistive than the rheology implemented in the numerical models. The stratigraphic position of the anhydrite layer may also have a major impact on its behaviour. Anhydrite layers embedded within salt diapirs are susceptible to boudinage as the salt exerts a shear stress on both surfaces of the embedded layer. This shear stress is only applied to the bottom surface if the anhydrite caps the salt source layer and the anhydrite strength may therefore be able to resist shear stresses produced by diapirism.

4.4.4 Insight on the development of Hoodoo Dome canopies

Though Experiment 8 did not develop as expected due to challenges with boundary conditions, it provides insight on how canopies on Hoodoo Dome may have formed while still preserving an anhydrite cap (Fig. 4-14). Like the anhydrite units that cap the Isachsen, Dumbells and Contour Domes, the analogue anhydrite layer remained intact and continued to rise, albeit asymmetrically due to boundary effects, on top of the developing diapir. The anhydrite hood does not pierce the surface at Hoodoo Dome; however, drill logs from the Hoodoo L-41 well do confirm that at least 280 m of anhydrite is present between the salt diapir and the overlying Isachsen Formation. The relation and extent, however, of the anhydrite cap over the salt canopies identified on the margins of Hoodoo Dome (Chapter 2) is unclear. A possible scenario is that the anhydrite hood at some point existed at the surface as a competent block, rafted on top of the salt dome. Similar to observations derived from the analogue model, salt glaciers may have formed locally along the dome margin. This would imply that the integrity of the anhydrite had been locally broken along the margins. There is no evidence that canopies formed at Dumbells and Contours Dome, suggesting that the sheared anhydrite margins may have remained intact at surface. It is speculative as to why the anhydrite sheared margins of Hoodoo Dome were locally broken. The margins, for example may have broken due to extensive shearing or from fractures that developed at near surface conditions.

Due to meteoric water content, the viscosity of salt glaciers is many orders of magnitude less than salt within diapirs, allowing for rapid salt flow across the surface. Salt at surface is capable of gravity spreading down slopes as low as 3 degrees and spreads under its own weight due to its low viscosity (Jackson & Talbot, 1986). The salt glaciers were subsequently buried to form intruded salt horizons as the rate of sedimentation exceeded that of the vertical salt flow that fed the salt glaciers.

Figure 4-13. (A) Seismic reflection profile I(F)-71 where the Deer Bay (pink) and Isachsen (blue) Formations are deflected above an interpreted canopy. The profile is un-migrated with a vertical exaggeration of 1.5. (B) Experiment 8 with pre-faulted margins shows how the development of salt canopies may have formed while still preserving a rafted anhydrite cap. The canopies continued to thicken after burial. (C) Interpretive sketches on the general development of salt canopies while maintaining an uplifted anhydrite cap (after Jackson & Talbot, 1986).

As is observed in the analogue model, these salt horizons continued to thicken to form canopies as salt was continually fed into the diapir, lifting the overlying anhydrite and clastic units and deflecting sediments at the tip of canopies. Seismic profile I(F)71 provides an example of canopy thickening where the Deer Bay Formation is deflected on top of an interpreted canopy (Fig. 4-13). This additional lateral component of diapir growth may explain why Hoodoo Dome does not pierce the surface in comparison to the other domes located on Ellef Ringnes Island.

4.5 Conclusion

Analogue experiments were conducted in an attempt to understand the development of diapirs with dense anhydrite caps. It was originally hypothesized that despite the large density contrast, anhydrite hoods remained competent and were passively uplifted by rising diapirs. The uplift of the competent hood was thought to be facilitated by the formation of localized shear zones along the dome margins as is observed in nature. A plastic analogue was produced with a yield stress that scales to the strength of anhydrite approximated from rock deformation experiments. The experiments, however, were unable to localize shear zones within the anhydrite. Free-slip boundary conditions were also not properly implemented and prevented the progression of experiments past the initial pillow stage.

The inability to localize marginal shear zones is most likely a product of the high elasticity of the material and the boundary conditions implemented to control the flow gradient of silicone. Though we initially intended to examine the interaction of anhydrite with the rising salt without concerning ourselves with the dynamics that initiate and drive diapirism, it appears that these initial boundary conditions have a major influence on the geometry and structures associated with the domes. Despite density contrasts, the experiments did demonstrate that competent anhydrite caps are capable of being uplifted and may act as plugs that seal off salt flow at the surface. The experiments also brought to attention a possible mechanism of canopy growth while maintaining a competent anhydrite cap at Hoodoo Dome. What is more intriguing than the density contrast of salt and anhydrite, however, are the mechanisms that control the localization of salt flow and development of sheared anhydrite margins (instead of yielding within the centre of the dome). These mechanisms are most likely controlled by the rheological properties of anhydrite (including the interaction of strength, strain softening, and elasticity) and the initial tectonic setting and pre-existing architecture of the basin sediments.

The experiments highlight practical challenges in physical modeling of the development of a single diapir where dynamics and kinematics of a system are appropriately scaled to nature. To improve the model set-up, more constraints are required on the dynamics that initiate and control dome growth in the Sverdrup Basin and the rheology of anhydrite at the onset of

diapirism. Experimentation with different materials that are properly scaled, yet can be frozen to produce cross-sections, may aid in producing 3D dynamically scaled experiments that remove the boundary condition issues mentioned above. 2D numerical models have an advantage of implementing free-slip boundary conditions and may be more flexible in testing various rheological parameters; however many codes still have difficulties implementing sedimentation.

94 5 Summary and final remarks

5.1 Summary of mechanics and timing of evaporite dome growth on Ellef Ringnes Island

The evaporite piercement structures on Ellef Ringnes Island have experienced a long period of development that dates to at least the Late Triassic and most likely initiated even earlier within the Permian/Early Triassic. Observations made in this study point to tectonic influences that led to the initiation and controlled the mechanisms and rates of diapirism on Ellef Ringnes Island. The development of the evaporite structures examined in this study is summarized below.

5.1.1 Initiation of diapirism

While little is known about the period prior to deposition of the Triassic formations, the asymmetry of the evaporite structures on Ellef Ringnes Island suggests that their initiation may have been localized on top of basement faults (i.e. thick-skinned deformation) that most likely formed due to extension during basin rifting. The basement structures are assumed to be grabens that formed at the margins of the Axel Heiberg depocenter. This correlates with the regional patterns of isopach and tectonic subsidence maps that indicate a greater accumulation of sedimentary units on the eastern margin of Ellef Ringnes Island. Diverging sedimentation approaching the domes and the formation of canopies within or beneath the Heiberg Group places a minimum age for the initiation of diapirism on Ellef Ringnes Island in the Late Triassic. Diapirism most likely began earlier in the Late Permian/ Early Triassic, perhaps triggered during a period of increased subsidence in the Early Triassic. The dominant occurrence of ductile fabrics over brittle structures in the anhydrite layer suggests that the evaporites were buried at the onset of diapirism; therefore it is unlikely that diapirism was triggered at the onset of initial basin rifting (Carboniferous to Early Permian).

The collapse of sediments and evaporites over basement fault grabens would have created local tension that was sufficiently large enough to overcome the strength and localize deformation and thinning within the overburden (Fig. 5-1a). Localized deformation then

95 created the necessary weakness for the upward migration of viscous and buoyant salt. The lack of normal faults in sediments adjacent to the developing domes (that normally form during thin-skinned extension) aided in maintaining the integrity of the anhydrite layer. The flow of salt was highly localized along basement faults and may have allowed the diapirs to form without a pillow-stage equivalent (Koyi et al., 1993a). Localized flow along basement faults and the sharp variation of sedimentary thicknesses across fault blocks may have resulted in the localization of shear stresses at the margins of the domes, causing local shearing and thinning of anhydrite.

5.1.2 Diapirism in the Triassic

The major driving mechanism of diapirism during the Mesozoic was differential loading between opposing fault blocks. Excessive sediment loading over top of down-thrown fault blocks led to the collapse of overburden into the salt, forcing the salt to withdraw and rise toward less pressurized regions (i.e., developing evaporite domes). Characteristic features that formed due to this process include rim synclines, structural offset of thickened sedimentary formations, and shear zones within sediments overlying the eastern downthrown fault blocks. Once the strength of the overburden was overcome, buoyancy may have also aided in diapir growth. The diminishing effect of buoyancy by overlying anhydrite is dependant on the thickness of the layer, which may have been thinner than previously suggested (e.g. Schwerdtner & Clarke, 1967; Stephenson et al., 1992).

A switch to a slower tectonic subsidence rate in the Late Triassic/Early Jurassic (~200 Ma) may have triggered salt to extrude at the surface of Hoodoo Dome (Fig. 5-1b). Prior to this switch in subsidence rate, diapirism was driven by loading from sedimentation which effectively pressurized the underlying salt layer. As the subsidence rate decreased, the evaporites continued to flow upward due to built-up pressure and their growth may have exceeded the rate of sedimentation. This led to salt extrusion until pressure in the evaporite source layer eventually equilibrated with the sediment load, at which time the Heiberg Group gradually lapped onto the extruded salt. The extrusion of salt at the surface suggests that the anhydrite layer was locally broken at the margins of Hoodoo Dome by this period.

96 5.1.3 Diapirism in the Jurassic

From the Late Triassic to the Early Cretaceous the domes continued to develop under passive diapirism (i.e. downbuilding), which led to thinning and ductile drag zones along the margins of the domes. The rate of diapirism during this period was not steady, leading to multiple cycles of salt extrusion and sediment onlap at the surface of Hoodoo Dome (Fig. 5-1c). Along the margins of Dumbells Dome, episodic surges in the rate of diapirism are less obvious. Switches from higher rates of net diapiric rise to higher rates of sediment aggradation were most likely controlled by tectonic triggers recorded as minor oscillations in tectonic subsidence curves. Buried canopies continued to thicken as salt was continual fed into Hoodoo Dome, creating a lateral dome growth component. During the Jurassic, the anhydrite hood at Hoodoo Dome remained competent despite broken margins and continued to act as passive plug on top of the salt stock.

Figure 5-1. Schematic diagrams illustrating the development of evaporite piercement structures (in particular Hoodoo Dome) on Ellef Ringnes Island.

97 5.1.4 Diapirism in the Cretaceous and Early Tertiary

During the Late Jurassic to Early Cretaceous the domes were slowly onlapped and buried by the Deer Bay, Isachsen and Christopher Formations due to increased rates in subsidence and sedimentation. The rate of salt withdrawal above the down-thrown fault block at Dumbells Dome decreased in the Early to Late Cretaceous as marked by the deposition of fairly uniform upper Isachsen members on top of eroded rim syncline deposits (Fig. 5-1d). This unconformity correlates with a sudden decrease in tectonic subsidence recorded in multiple wells, suggesting that this event is a regional tectonic signature and not solely a response to diapirism. The erosional event within the Isachsen Formation was followed by a period of rapid subsidence interpreted as a signature related to the onset of Canada Basin seafloor spreading north of the Sverdrup Basin. The renewed increase in subsidence correlates with the subsequent burial of the salt domes by the youngest Isachsen members and the Christopher Formation (Fig. 5-1d). The decrease of salt withdrawal is most likely a product of burial, where pressure is applied to the top of the domes thereby counteracting upward passive growth. At Hoodoo Dome, thicker deposits continued to accumulate on the eastern margin, suggesting that salt withdrawal beneath overburden continued until the Late Cretaceous. Thickening of canopies and the resulting uplift and deformation of onlapping sediments most likely facilitated dome growth during this period, while the collapse of the downthrown block was accommodated by a shear zone along the eastern margin of Hoodoo Dome.

The decline in tectonic subsidence during Cenomanian to Maastrichtian times coincides with the deposition of the Hassel, Kanguk and Expedition Formations. The crests of buried domes were most likely uplifted due to built up pressure beneath the anhydrite caps, resulting in the thinning of the Hassel and Kanguk Formations toward the dome. Active removal of overlying sediments, however, most likely occurred due to reactivation or accelerated rates of diapirism triggered by regional compression from the Eurekan Orogenic event (Fig. 5-1e). The erosion of overlying formations at this time most likely aided in increasing the contribution of buoyancy to driving mechanisms of diapirism. Reactivation and inversion of basement structures during the Tertiary could explain the initiation of lesser developed and deeper buried salt anticlines located offshore of Ellef Ringnes Island that have yet to pierce

98 Mesozoic strata (see discussion on salt anticlines in the following section). At the surface, anhydrite caps gypsified and underwent extension due to doming, forming fractures that crosscut older ductile fabrics. The strength and resistance to weathering allows the anhydrite to continue acting as a passively rising plug that has prevented the extrusion of salt at Dumbells and Contour Domes. The net lateral growth by canopy thickening may explain why the anhydrite cap at Hoodoo Dome is still partially covered by the Isachsen Formation.

5.2 Implications for oil and gas exploration

Petroleum exploration occurred in the Sverdrup Basin between 1969 and 1986 with a total of 119 wells drilled and the discovery of 19 petroleum fields, consisting of 8 oil and 25 gas pools. Estimated total proven reserves are 294 × 106 m3 oil and 500 × 109 m3 gas (Chen et al., 2000). The majority of discoveries occur in the offshore region west of Ellef Ringnes Island to northeastern Melville Island (Fig. 5-2a). Despite the extraordinary success of discoveries over a short period of time, unfavorable economic factors led to the abandonment of exploration within the Sverdrup Basin (Chen et al., 2000).

5.2.1 Petroleum generation in the Sverdrup Basin

The primary source rocks include bituminous marine shales of the Triassic Schei Point and Blaa Mountain Groups that range from immature on the basin margins to over mature in the basin axis (discussion in Chen et al., 2004). Secondary source rocks include oil shales in the Carboniferous Emma Fiord and Jurassic Jameson Bay Formations (Majorowicz et al., 2002). Petroleum migration pathways and maturation is complex, often in response to periods of basin uplift and localized heat-flow anomalies related to igneous intrusions and highly conductive salt structures (Chen et al., 2004; Dewing & Sanei, 2009). The main stage of petroleum generation occurred between the Early and Late Cretaceous as most source rocks reached or passed the maturity window at that time (Brooks et al., 1992; Gentzis & Goodarzi, 1993; Goodarzi et al., 1989, 1993).

Existing undeveloped oil and gas discoveries are located within Triassic to Lower Cretaceous strata in the southwestern Sverdrup Basin. Proven petroleum reservoirs consist of shallow marine, delta front and delta plain sandstones of the Heiberg Group, King Christian, Awingak and Isachsen Formations. 99

Figure 5-2. a) Location of known hydrates and the distribution of underlying seismic closures with conventional hydrocarbon and non-hydrocarbon accumulations (Majorowicz et al., 2002). The two largest gas fields, Hecla and Drake, are labeled. b) The stratigraphic distribution of discovered petroleum reserves. The equivalent gas reserve is indicated in dark grey and the oil is indicated in white on the right side of the diagram (Chen et al., 2000).

The majority of plays are located within porous fluvial deltaic deposits of the Heiberg Group and correlative King Christian Formation capped by argillaceous strata of the Jameson Bay Formation (Fig. 5-2b). The porosity of the Heiberg Group, however, becomes unfavorable in

100 the northeast due to deep burial and diagenesis. The remaining majority of pools are hosted within the Upper Jurassic Awingak Formation located along the southern margin of the Sverdrup Basin.

A broad low-relief flexure along the southern basin margin hosts the Hecla and Drake reservoirs, the largest natural gas field discoveries in the Sverdrup Basin. The largest oil discovery, the Cisco, is located over a broad salt cored dome that fails to pierce the reservoir rock. Common structures offshore of Ellef Ringnes Island also include salt cored anticlines and pillow structures that do not pierce reservoir rocks of the Triassic Heiberg Group and King Christian Formation. The salt cored anticlines are oriented orthogonal to the principle shortening direction of the Eurekan Orogeny. The oil and gas fields are sealed off by the thick argillaceous shales of the Jameson Bay Formation, however these structures are poorly filled due to consequence of late growth and possibly from surface seepage through extensive crestal faulting (Embry, 1982).

5.2.2 Potential for future oil and gas exploration

Predictive object-based modeling conducted for the Sverdrup Basin (Chen et al., 2004) suggests that major risks include source rock maturity and reservoir quality, and predict that conditions are more favorable in the southeastern flank of the basin for potential undiscovered oil and gas resources. The petroleum potential is reduced in the eastern region of the Sverdrup Basin due to the effects of Mesozoic magmatism and Tertiary Eurekan orogenic deformation that led to over-maturity and break down of trap integrity. Deeper erosional levels have also removed or breached the most prospective reservoir strata in many structures (Chen et al., 2000).

The major area of improvement for future reserve growth is an increase in exploration activity. Due to the low concentration of wells and seismic grid density, only large scale features were detected during the 1969-1986 exploration boom. There is still, however, a significant potential for the discovery of higher frequency and more complex structures and traps with the advancement of imaging acquisition and processing techniques, in particular the ability to map the geometric complexity of traps with higher resolution 3D seismic

101 surveys. Harrison (2001) suggests that new exploration plays could be developed in Carboniferous graben and horst settings, Permian inversion structures, Upper Paleozoic unconformities and stratigraphic pinch-outs. Potential for complex sub-salt structures and associated new hydrocarbon plays are also expected west of Axel Heiberg Island.

Passive diapirs create the potential for many small oil and gas traps within adjacent deformed sediments (Stewart, 2006). Dragged and overturned sedimentary formations allows for the migration of oil and gas up-dip along bedding planes. Petroleum resources accumulate beneath unconformities where more impermeable discordant formations cap upturned beds. Within the roof zone of a diapir, radial extension faults dipping away from the crest are common and may provide potential traps above rising salt diapirs and anticlines. Deeper diapir root structures could also include rim synclines and extensional growth faults. Salt canopies, such as the ones interpreted at Hoodoo Dome, are excellent structural traps for the accumulation of oil and gas. The impermeability of salt canopies terminating and capping dragged sedimentary beds provides an excellent seal to oil and gas seepage (e.g. Fig. 5-3). Canopies also have the potential to be much more laterally extensive than segmented fault blocks, which may result in the accumulation of petroleum in broader/larger pools. For Hoodoo Dome in particular, canopies that truncate sandstone dominant formations such as the Heiberg Formation would make excellent reservoirs. The traps would have developed in the Late Triassic to Early Jurassic, predating the main stage of petroleum generation and migration. The Heiberg Formation may also be locally truncated by impermeable shales of the Jameson Bay formation by unconformities that formed due to changes in the relative rates of diapirism and sedimentation.

The location of major fault structures within the evaporite sub-basins may also aid in predicting the occurrence of deeper salt structures, such as the salt cored anticlines west of Ellef Ringnes Island. The application of various geophysical techniques could perhaps resolve deeply buried grabens within the Sverdrup Basin basement.

102

Figure 5-3. Examples of canopies capping dipping sedimentary beds within (a) the Dniepr- Donets basin in Ukraine (Stovba & Stephenson, 2003) and (b) the Pricaspian basin (Ismail- Zadeh et al., 2004).

103 5.3 Future Research

The three most significant areas of improvement and room for expansion of this study are:

1. Gravity modeling to extract residual gravity anomalies related to the evaporite piercement structures. Residual gravity anomalies may confirm the existence and spatial extent of canopies interpreted at Hoodoo Dome and provide us with additional information on the 3D geometry of the Dumbells and Contour Domes. Residual gravity anomalies may also provide the best resolution of the depletion and thickness variations of the underlying salt source layer.

2. Basin modeling to determine the amount of lithosphere thinning that has resulted from the initiation of basin rifting in the Carboniferous to present. Such basin modeling will require the “padding” of input backstripping data with estimated thicknesses for formations dating from the Permo-Carboniferous to the Late Triassic. Forward modeling of lithosphere stretching factors should be completed in an attempt to match tectonic subsidence events recorded within backstripped curves.

3. Numerical modeling to further understand the involvement of anhydrite in diapirism in the western region of the Sverdrup Basin. The models would be dynamically and kinematically scaled to nature, and test several diapirism mechanisms (e.g. passive diapirism or downbuilding, growth under thick-skinned extension) and rheologies of anhydrite. The code must have the capability to include sedimentation/erosion and be able to handle a range of rheologies to test different end member material properties of anhydrite.

104

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Appendix 1

A table of seismic reflection surveys interpreted in this study, including the fold and spacing of geophones, is included below. The company and the year that the data was originally acquired are also included.

Seismic profile Fold of data Trace spacing (m) Company acquired Year 1056 6-fold 33.5 Panarctic Oils 1973 1057 6-fold 33.5 Panarctic Oils 1973 1058 6-fold 33.5 Panarctic Oils 1973 1060 6-fold 33.5 Panarctic Oils 1973 1071 6-fold 33.5 Panarctic Oils 1973 1074 6-fold 33.5 Panarctic Oils 1973 1076 6-fold 33.5 Panarctic Oils 1973 1077 6-fold 33.5 Panarctic Oils 1973 1078 6-fold 33.5 Panarctic Oils 1973 1091 6-fold 33.5 Panarctic Oils 1973 1097 6-fold 33.5 Panarctic Oils 1973 1099 6-fold 33.5 Panarctic Oils 1973 1100 6-fold 33.5 Panarctic Oils 1973 1215 6-fold 33.5 Panarctic Oils 1973 1217 6-fold 33.5 Panarctic Oils 1973 113 pt 1 & 2 1-fold 24 Panarctic Oils 1969 114 1-fold 24 Panarctic Oils 1969 115 1-fold 24 Panarctic Oils 1969 116 1-fold 24 Panarctic Oils 1969 117 1-fold 24 Panarctic Oils 1969 118 pt 1 & 2 1-fold 24 Panarctic Oils 1969 119 1-fold 24 Panarctic Oils 1969 122 1-fold 24 Panarctic Oils 1969 123 1-fold 24 Panarctic Oils 1969 127 pt 1-5, ext 1-fold 24 Panarctic Oils 1969 131 1-fold 24 Panarctic Oils 1969 133 1-fold 24 Panarctic Oils 1969 134 pt 1 & 2 1-fold 24 Panarctic Oils 1969 136 1-fold 24 Panarctic Oils 1969 137 1-fold 24 Panarctic Oils 1969 138 1-fold 24 Panarctic Oils 1969 139 1-fold 24 Panarctic Oils 1969 141 1-fold 24 Panarctic Oils 1969 143 1-fold 24 Panarctic Oils 1969 144 1-fold 24 Panarctic Oils 1969

114

Seismic profile Fold of data Trace spacing (m) Company aquired Year G(F)1er 6-fold 33.5 Gulf Resources 1972 G(F)2 6-fold 33.5 Gulf Resources 1972 G(F)4 6-fold 33.5 Gulf Resources 1972 G(F)5 6-fold 33.5 Gulf Resources 1972 G(F)6 6-fold 33.5 Gulf Resources 1972 G(F)7 6-fold 33.5 Gulf Resources 1972 G(F)8A 6-fold 33.5 Gulf Resources 1972 G(F)9B 6-fold 33.5 Gulf Resources 1972 G(F)10 6-fold 33.5 Gulf Resources 1972 G(F)11B 6-fold 33.5 Gulf Resources 1972 G(F)12 6-fold 33.5 Gulf Resources 1972 G(F)13 6-fold 33.5 Gulf Resources 1972 G(F)13B 6-fold 33.5 Gulf Resources 1972 G(F)14 6-fold 33.5 Gulf Resources 1972 G(F)R-4 6-fold 33.5 Gulf Resources 1972 G(F)R-2 6-fold 33.5 Gulf Resources 1972 I(F)66 3-fold 24 Imperial (Esso) 1971 I(F)67 3-fold 24 Imperial (Esso) 1971 I(F)68 3-fold 24 Imperial (Esso) 1971 I(F)70 3-fold 24 Imperial (Esso) 1971 I(F)71 6-fold 24 Imperial (Esso) 1971 I(F)72 3-fold 24 Imperial (Esso) 1971 I(F)73 3-fold 24 Imperial (Esso) 1971 I(F)74 6-fold 24 Imperial (Esso) 1971 I(F)75 3-fold 24 Imperial (Esso) 1971 I(F)76 6-fold 24 Imperial (Esso) 1971 I(F)77 3-fold 24 Imperial (Esso) 1971 I(F)78 6-fold 24 Imperial (Esso) 1971 I(F)79+A 6 and 3-fold 24 Imperial (Esso) 1971 I(F)80 3-fold 24 Imperial (Esso) 1971 I(F)81A 3-fold 24 Imperial (Esso) 1971 I(F)82 3-fold 24 Imperial (Esso) 1971 I(F)83 6-fold 24 Imperial (Esso) 1971 I(F)84 3-fold 24 Imperial (Esso) 1971 I(F)85 3-fold 24 Imperial (Esso) 1971 I(F)86 6-fold 24 Imperial (Esso) 1971

115

Appendix 2

Instantaneous velocity-depth curves for Mesozoic formations. The plots are derived from a compilation of sonic logs from multiple wells. The best fit parameters of the chosen curve functions are included in Table 2-1.

116

117

Appendix 3

Time-depth conversion MATLAB code used to calculate formation thicknesses and depths in Chapter 2. The script is developed by C. Schrank and J. Macauley.

clear;

% The following three lines specify file location and name and load it wp=['E:\USERS\JENNY\DepthConversion\MATLAB\']; % Path where to find the data files name=['AllTopsFinal.txt']; data=importdata([wp,name]); dirname=['My funky results']; mkdir(dirname);cd(dirname);

[L,W]=size(data); % L: number of lines in data, W: number of columns in data

% Fit parameters must be present in a matrix with three columns: 1) Layer number (starting from top), 2) k for linear, 1/m for power law, 3) Vo or A (pre- exponential constant), 4) code for which eqn to use 1=constant, 2=linear, 3=power law fitdata=importdata([wp,'fitdata_pow.txt']);

% This loop produces a table with the interval time for each layer; each column represents a layer in the order that is given in data counter=1; % Counter (per formation) for i=3:W; if counter==1 dt=data(:,i)/2; % Division by 2 converts to OWT; this is the interval time for the 0th layer (not present in data) else dt=(data(:,i)-data(:,i-1))/2; end; filter=find(dt<0);dt(filter)=NaN; % All negative time differences are converted to NaNs timediff(:,counter)=dt; % timediff is a new matrix with all dt, first column is layer zero (not given in input file) counter=counter+1; end;

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[l,w]=size(timediff);

% This loop filters meaningless interval times, i.e. at those points, where the upper layer has a NaN time entry for i=1:l; nulls=isnan(timediff(i,:)); if 1|isnan clear nanindex; else nanindex=find(nulls==1); timediff(i,nanindex(1)+1:length(timediff(i,:)))=NaN; clear nulls; clear nanindex; end; end;

% Depth conversion of dt dz(l,1)=zeros; % Depth to base of layer under consideration

% Constants: for j=1:w; Vconst=fitdata(j,3); k=fitdata(j,2); % Slope of linear velocity law V0=fitdata(j,3); % Velocity constant q=1-fitdata(j,2)^-1; % Al-Chalabi's q (PL exponent) t=timediff(:,j); % Interval time if fitdata(j,4)==1 & j==1 dz(:,j)=Vconst*t; % Constant velocity elseif fitdata(j,4)==2 dz(:,j)=((V0/k)*(exp(t*k)-1))+dz(:,j-1).*exp(t*k); % Al-Chalabi, eq. 2 else dz(:,j)=(dz(:,j-1).^q+(q*fitdata(j,3)*t)).^(1/q); % Al-Chalabi eq. 6 end; end; formation=data(:,1:2); % X, Y data in col. 1 and 2 for g=1:w; formation(:,3)=dz(:,g); % Depth to base of layer, col. 3 if g==1 % Depth to top of layer, col. 4 formation(:,4)=0; else formation(:,4)=dz(:,g-1); end;

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if g==1 formation(:,5)=dz(:,g); % Thickness of layer, col. 5

else formation(:,5)=dz(:,g)-dz(:,g-1); end;

%The following loop checks whether a row in a given formation has zero entries alone, i.e. zero depth to layer bottom, zero depth to layer top, and zero layer thickness. If this is so, all these entries are converted into NaNs.

[Lf,Wf]=size(formation); for i=1:Lf; if sum(find(formation(i,:)==0))==12; formation(i,3:5)=NaN; else end; end;

saveascii(formation,['Formation_',num2str(g-1),'.txt'],5,' '); end; saveascii(dz,['DZ.txt'],5,' '); saveascii(timediff,['Intervaltime.txt'],5,' '); cd ..;

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Appendix 4

The following contour plots are isopach maps derived from the depth conversion of interpreted seismic horizons presented in Chapter 2. The colour scale on the right margin of the diagrams represents the thickness of each formation in meters. Formation thicknesses derived from boreholes are included on individual diagrams for comparison. The X/Y coordinates (in meters) are generated with a Lambert Conic Conformal map projection.

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Appendix 5

Individual well input files used in 1D backstripping (Petroprob code). The files include the depth to the bottom of each formation (Dewing & Embry, 2007), the adjusted depths including estimated eroded formations, the age of the top and bottom of each formation and the average relative percentages of lithologies. Additional footnotes include information on the removal of igneous intrusions, adjusted depths of formation tops picked by Dewing & Embry (2007), assumptions made on the projection of eroded thicknesses, etc.

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Balaena D-58 Bottom depth age of age of % sand % silt % shale % estimated eroded of formations formation formation limestone thicknesses (Dewing & Embry, 2007) bottom top 16 KANGUK FM 250 95 69 0 3070 0 250 HASSEL FM 350 103 95 60 3010 0 100 CHRISTOPHER FM 950 113 103 24 1560 0 600 Walker Island eroded 1050 116 113 75 025 0 100 WALKER ISLAND MBR 337 1150 126 116 75 0 25 0 1050 total eroded RONDON MBR 390 1203 128 126 20 40 40 0 PAT. ISLAND MBR 910 1723 137 128 65 15 20 0 DEER BAY FM 1222 2035 153 137 5 20 75 5 AWINGAK FM 1269 2082 157 153 60 0 40 0 RINGNES FM 1326 2139 161 157 0 10 90 0 MCCONNELL FM 1361 2174 172 161 0 0 95 5 SANDY POINT FM 1368.5 2181.5 173 172 50 0 50 5 JAMESON BAY FM 1462 2275 184 173 0 10 90 0 REMUS MBR 1470 2283 187 184 30 0 70 0 FOSHEIM MBR 1774 2587 202 187 70 5 25 0 ROMULUS MBR 1874 2687 203 202 35 20 45 0 END

*** Assuming that Romulus Mbr is 200 m thick (Cape Allison C-47 and Char G-07) *** Removed 237 m of water

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Cape Allison C-47 bottom age bottom age top % sand % silt % shale % lmst est. thickness 14 KANGUK FM 250 95 69 0 30 700 250 HASSEL FM 350 103 95 60 30 100 100 CHRISTOPHER FM 950 113 103 24 15 600 600 WALKER ISLAND MBR 1150 126 113 75 5 200 200 RONDON MBR 1200 128 126 30 40 300 50 PAT. ISLAND FM 772.1 1707.5 137 128 80 0 20 0 1200 total eroded DEER BAY FM 1165.0 2100.4 153 137 0 0 100 0 AWINGAK FM 1222.9 2158.3 157 153 50 0 50 0 RINGNES FM 1288.0 2223.4 161 157 0 0 100 0 MCCONNELL FM 1344.0 2279.4 172 161 0 0 100 0 SANDY POINT FM 1371.0 2306.4 173 172 10 20 70 0 JAMESON BAY FM 1469.5 2404.9 184 173 0 5 95 0 HEIBERG FM 2016.0 2951.4 204 184 65 5 30 0 BARROW FM 2100.1 3035.5 210 204 25 20 55 0 END

*** Barrow assumed to be 250 m thick (King Christian N-06 and Sutherland O-23) *** Assuming Paterson complete *** Eroded units from 400-1000 m from Dewing & Sanei

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Char G-07 bottom age bottom age top % sand % silt % shale % lmst est. thickness 17 KANGUK FM 250 95 69 0 30 700 250 HASSEL FM 350 103 95 60 30 100 100 CHRISTOPHER FM 675 113 103 25 15 600 325 walker island eroded 275 700 115 113 100 0 0 0 25 WALKER ISLAND MBR 450 874 126 115 100 0 0 0 700 total eroded RONDON MBR 504 928 128 126 10 20 70 0 PAT. ISLAND MBR 976 1400 137 128 60 20 20 0 DEER BAY FM 1314 1738 153 137 0 30 70 0 AWINGAK FM 1415 1839 157 153 40 20 20 0 RINGNES FM 1439 1863 161 157 10 0 90 0 MCCONNELL FM 1477 1901 172 161 0 0 100 0 SANDY POINT FM 1498 1922 173 172 40 0 60 5 JAMESON BAY FM 1561 1985 184 173 0 10 90 0 REMUS MBR 1577 2001 187 184 70 0 30 0 FOSHEIM MBR 1838 2262 202 187 50 0 50 0 ROMULUS MBR 1982 2406 204 202 70 0 30 0 BARROW FM 2179 2603 214 204 0 10 90 0 END

*** Assuming that Barrow is 250 m (King Christian N-06 and Sutherland O-23) *** Removed 276m of water *** Assume Dewing and Embry forgot McConnell (1439 from Panarctic Oil Ltd.)

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Dome Bay P-36 bottom age bottom age top % sand % silt % shale % lmst est. thickness 16 KANGUK FM 350 95 69 0 30 700 350 HASSEL FM 600 103 95 60 30 100 250 CHRISTOPHER FM 1400 113 103 25 15 60 0 800 WALKER ISLAND MBR 1600 126 113 60 20 20 0 200 RONDON MBR 1650 128 126 30 10 5010 50 Pat. Island eroded 1775 130 128 65 25 10 0 125 PAT. ISLAND MBR 425.8 2200.8 137 130 65 25 10 0 1775 total eroded DEER BAY FM 852.8 2627.8 153 137 0 10 90 0 RINGNES FM 1177.1 2952.1 161 153 10 0 90 0 MCCONNELL FM 1564.8 3339.8 172 161 15 0 85 0 SANDY POINT FM 1575.2 3350.2 173 172 80 0 20 0 JAMESON BAY FM 1857.5 3632.5 184 173 20 10 70 0 KING CHRISTIAN FM 2036.1 3811.1 190 184 80 0 20 0 LOUGHEED ISLAND 2100.0 3875.0 193 190 5 45 50 0 MACLEAN STRAIT FM 2375.6 4150.6 197 193 60 10 30 0 GROSVENOR ISLAND 2453.7 4228.7 200 197 10 20 70 0 END

*** Removed intrusions from McConnell Fm *** Grosvenor assumed complete at 180 m thick (Mocklin Point D-23)

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Dumbells E-49 bottom age bottom age top % sand % silt % shale % lmst est. thickness 14 KANGUK FM 400 95 69 0 3070 0 400 HASSEL FM 650 103 95 60 3010 0 250 Christopher eroded 1450 112 103 0 2080 0 800 CHRISTOPHER FM 76.2 1526.2 113 112 0 20 80 0 1450 total eroded WALKER ISLAND MBR 225 1675 126 113 60 20 20 0 RONDON MBR 365 1815 128 126 0 10 90 0 PAT. ISLAND MBR 1254 2704 137 128 60 13 27 0 DEER BAY FM 2112.3 3562.3 153 137 3 3 93 2 RINGNES FM 2575.6 4025.6 161 153 1 2 95 3 MCCONNELL FM 2768.2 4218.2 172 161 30 0 70 0 SANDY POINT FM 2780.7 4230.7 173 172 80 0 20 0 JAMESON BAY FM 3114.5 4564.5 184 173 0 10 90 0 HEIBERG FM 3408.9 4858.9 194 184 60 0 40 0 added Heiberg 3688.9 5138.9 204 194 60 0 40 0 END

*** Heiberg est. to be 1000 m thick (Hoodoo Dome J-12) *** Thicknesses for Isachsen members from seisimics (correlate well with logs)

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Grenadier A-26 bottom age bottom age top % sand % silt % shale % lmst est. thickness 18 KANGUK FM 200 95 69 0 30 70 0 200 HASSEL FM 275 103 95 60 30 10 0 75 CHRISTOPHER 825 113 103 25 15 60 0 550 WALKER ISLAND MBR 1000 126 113 100 0 0 0 175 RONDON MBR 1050 128 126 10 20 70 0 50 PAT. ISLAND MBR 645 1522 137 128 30 40 30 0 1050 total eroded DEER BAY FM 849 1726 153 137 0 30 70 0 AWINGAK FM 910 1787 157 153 60 20 20 0 RINGNES FM 929 1806 161 157 0 20 80 0 MCCONNELL FM 939 1816 172 161 0 30 70 0 SANDY POINT FM 955 1832 173 172 70 0 30 0 JAMESON BAY FM 981 1858 184 173 0 20 70 10 HEIBERG FM 1134 2011 204 184 50 0 50 0 BARROW FM 1589 2466 217 204 0 20 80 0 HOYLE BAY FM 1671 2548 226 217 0 50 50 0 GORE POINT MBR 1751 2628 228 226 5 0 35 60 CHADS POINT MBR 1835 2712 233 228 0 60 20 20 MURRAY HARBOUR FM 1991 2868 245 233 0 70 30 0 END

*** Reorganized McConnell and Ringnes a bit (McConnell removed from GSCC tops) *** Removed 173 m of water *** Assumed Pat Island complete

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Helicopter J-12 bottom age bottom age top % sand % silt % shale % lmst est. thickness 14 KANGUK FM 300 95 69 0 30 70 0 300 HASSEL FM 500 103 95 60 30 10 0 200 CHRISTOPHER FM 1250 113 103 0 20 80 0 750 Walker Island eroded 1350 120 113 40 25 35 0 100 WALKER ISLAND MBR 75.0 1425.0 126 120 40 25 35 0 1350 total eroded RONDON MBR 125.0 1475.0 128 126 0 10 80 0 PAT. ISLAND MBR 1353.9 2703.9 137 128 65 20 14 1 DEER BAY FM 2307.4 3657.4 153 137 0 1 99 0 **removed 60 m intrusion RINGNES FM 2974.9 4324.9 161 153 0 0 90 10 **removed 300 m intrusion MCCONNELL FM 3334.5 4684.5 172 161 0 10 90 0 **removed 70 m intrusion SANDY POINT FM 3358.6 4708.6 173 172 0 40 60 0 JAMESON BAY FM 3698.8 5048.8 184 173 0 2 98 0 HEIBERG FM 3813.7 5163.7 188 184 0 90 10 0 added Heiberg 4341.7 5691.7 200 188 0 90 10 0 END

*** Geology of Rondon and Patterson updated from well logs *** Removed igneous material from Deer Bay, Ringnes and McConnell *** Heiberg assumed to be 1100 m thick; original depth to 3813.7 m but added material to make up to 200 Ma for modelling

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Hoodoo Dome H-37 bottom age bottom age top % sand % silt % shale % lmst est. thickness 16 KANGUK FM 250 95 69 0 30 700 250 HASSEL FM 350 103 95 60 30 100 100 Christopher eroded 600 107 103 24 15 600 250 CHRISTOPHER FM 411.5 1011.5 113 107 24 15 60 1 600 total eroded WALKER ISLAND MBR 600 1200 126 113 75 5 20 0 RONDON MBR 652 1252 128 126 30 40 30 0 PAT. ISLAND MBR 1146.4 1746.4 137 128 70 20 10 0 DEER BAY FM 1752.6 2352.6 153 137 15 20 60 5 RINGNES FM 1960.2 2560.2 161 153 0 5 90 5 MCCONNELL FM 2090 2690 172 161 0 5 90 5 SANDY POINT FM 2125.4 2725.4 173 172 30 10 60 5 JAMESON BAY FM 2301 2901 184 173 5 10 85 0 REMUS MBR 2353.1 2953.1 186 184 80 5 15 0 FOSHEIM MBR 2787.4 3387.4 202 186 65 10 25 0 ROMULUS MBR 3246.2 3846.2 204 202 35 20 45 0 BARROW FM 3374.8 3974.8 217 204 30 30 40 0 END

*** Barrow assumed to be complete

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Hoodoo N-52 bottom age bottom age top % sand % silt % shale % lmst est. thickness 12 KANGUK FM 300 95 69 0 30 700 300 HASSEL FM 375 103 95 60 30 100 75 Christopher eroded 925 110 103 20 10 700 550 CHRISTOPHER FM 263.0 1188.0 113 110 20 10 70 0 925 total eroded WALKER ISLAND MBR 716.9 1641.9 126 113 60 20 20 0 RONDON MBR 759.9 1684.9 128 126 45 10 45 0 PAT. ISLAND MBR 948.0 1873.0 137 128 50 20 30 0 DEER BAY FM 1012.0 1937.0 153 137 10 10 80 0 RINGNES FM 1159.0 2084.0 161 153 5 5 90 0 no deposition 1160.0 2085.0 184 161 30 30 30 0 REMUS MBR 1183.0 2108.0 187 184 40 0 60 0 FOSHEIM MBR 1495.1 2420.1 202 187 60 10 30 0 END

*** Assuming no erosion?? Added 1m for period of little to no deposition *** Assuming Romulus is 460 m thick (Hoodoo Dome H-37)

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King Christian N-06 bottom age bottom age top % sand % silt % shale % lmst est. thickness 23 KANGUK FM 250 95 69 0 30 70 0 250 HASSEL FM 350 103 95 60 30 10 0 100 CHRISTOPHER FM 950 113 103 25 15 60 0 600 WALKER ISLAND MBR 1150 126 113 85 5 8 0 200 RONDON MBR 1200 128 126 50 50 0 0 50 PAT. ISLAND MBR 1675 137 128 70 10 20 0 475 Deer Bay eroded 1775 141 137 10 15 70 5 100 DEER BAY FM 310.3 2085.3 153 141 10 15 70 5 1775 total eroded AWINGAK FM 355.1 2130.1 157 153 50 0 40 10 RINGNES FM 419.7 2194.7 161 157 5 0 80 15 MCCONNELL FM 478.5 2253.5 172 161 0 10 75 15 SANDY POINT FM 486.2 2261.2 173 172 0 90 10 0 JAMESON FM 612.0 2387.0 184 173 0 15 85 0 KING CHRISTIAN FM 702.0 2477.0 190 184 100 0 0 0 LOUGHEED ISLAND FM 716.0 2491.0 193 190 50 50 0 0 MACLEAN STRAIT FM 956.2 2731.2 197 193 50 20 30 0 GROSVENOR ISLAND FM 970.8 2745.8 200 197 0 60 40 0 SKYBATTLE FM 1107.0 2882.0 204 200 40 30 30 0 BARROW FM 1603.0 3378.0 217 204 20 20 60 0 HOYLE BAY FM 2330.0 4105.0 226 217 0 15 80 5 ROCHE POINT FM 2358.0 4133.0 240 226 0 10 0 90 MURRAY HARBOUR FM 2396 4171.0 245 240 0 10 80 10 BLIND FIORD FM 2889 4664.0 251 245 0 15 80 5 END

*** Blind Fiord assumed to be 2800 m thick (GSC paper and Sutherland O-23) *** removed 170 m of igneous material for Murray Harbour ***removed 300 m from Blind Fiord *** 600 to 1200 m estimated eroded units from Dewing & Sanei, 2009

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Kristoffer Bay B-06 bottom age bottom age top % sand % silt % shale % lmst est. thickness 20 KANGUK FM 350 95 69 0 3070 0 350 HASSEL FM 600 103 95 60 3010 0 250 CHRISTOPHER FM 1400 113 103 25 1560 0 800 WALKER ISLAND MBR 1600 126 113 85 58 0 200 RONDON MBR 1650 128 126 50 500 10 50 Pat. Island eroded 1825 131 128 70 1020 0 175 PAT. ISLAND MBR 375.8 2200.8 137 131 70 0 30 0 1825 total eroded DEER BAY FM 897.3 2722.3 153 137 2 3 95 0 RINGNES FM 1027 2852 161 153 0 0 100 0 MCCONNELL FM 1105 2930 172 161 10 0 95 5 SANDY POINT FM 1217.1 3042.1 173 172 10 20 70 0 JAMESON BAY FM 1418 3243 184 173 0 2 97 1 KING CHRISTIAN FM 1559.4 3384.4 190 184 95 0 5 0 LOUGHEED ISLAND FM 1588 3413 193 190 60 30 10 0 MACLEAN STRAIT FM 1863 3688 197 193 70 20 10 0 GROSVENOR ISLAND 1976 3801 200 197 0 5 95 0 SKYBATTLE FM 2038.5 3863.5 204 200 0 60 40 0 BARROW FM 2790 4615 217 204 10 20 70 0 PAT BAY FM 2886.7 4711.7 220 217 0 60 40 0 HOYLE BAY FM 3624.9 5449.9 226 220 0 8 90 2 END

*** Assuming that Hoyle Bay is complete (King Christian N-06 and Sutherland O-23 < Kristoffer Bay B-06) *** Removed 300 m of intrusions from Barrow, Pat Bay and Hoyle Bay

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Linckens Island P-46 bottom age bottom age top % sand % silt % shale % lmst est. thickness 18 KANGUK FM 200 95 69 0 30 70 0 200 HASSEL FM 275 103 95 60 30 10 0 75 CHRISTOPHER FM 875 113 103 25 15 60 0 600 WALKER ISLAND MBR 1075 126 113 100 0 0 0 200 RONDON MBR 1125 128 126 10 20 70 0 50 Paterson Island eroded 1525 136 128 60 20 20 0 400 PATERSON ISLAND MBR 70.1 1595.1 137 136 60 20 20 0 1525 total eroded DEER BAY FM 554.1 2079.1 153 137 0 30 70 0 SLIDRE MBR 576 2101 155 153 90 0 10 0 HOT WEATHER MBR 628 2153 157 155 0 30 70 0 CAPE LOCKWOOD MBR 732.7 2257.7 159 157 80 0 20 0 RINGNES FM 786.4 2311.4 161 159 20 0 75 5 MCCONNELL ISLAND FM 846.7 2371.7 172 161 0 0 100 0 SANDY POINT FM 870.8 2395.8 173 172 30 0 70 0 JAMESON BAY FM 979.9 2504.9 184 173 0 40 60 0 REMUS MBR 1010 2535 187 184 60 0 40 0 FOSHEIM MBR 1462.4 2987.4 202 187 30 30 40 0 ROMULUS MBR 1569.7 3094.7 204 202 70 10 20 0 END

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Louise O-25 bottom age bottom age top % sand % silt % shale % lmst est. thickness 16 KANGUK FM 350 95 69 0 30 70 0 350 HASSEL FM 550 103 95 60 30 10 0 200 CHRISTOPHER FM 1300 113 103 25 15 60 0 750 WALKER ISLAND MBR 1550.0 126 113 60 20 20 0 250 RONDON MBR 1600.0 128 126 30 10 50 10 50 PAT. ISLAND MBR 2175.0 137 126 65 25 10 0 575 Deer Bay eroded 2275.0 140 137 0 10 90 0 100 DEER BAY FM 477.3 2752.3 153 140 0 10 90 0 2275 total eroded RINGNES FM 709.0 2984.0 161 153 10 0 90 0 MCCONNELL FM 1210.0 3485.0 172 161 15 0 85 0 SANDY POINT FM 1216.0 3491.0 173 172 80 0 20 0 JAMESON BAY FM 1527.0 3802.0 184 173 20 10 70 0 KING CHRISTIAN FM 1743.5 4018.5 190 184 80 0 20 0 LOUGHEED ISLAND FM 1758.7 4033.7 193 190 5 45 50 0 MACLEAN STRAIT FM 2164.1 4439.1 197 193 60 10 30 0 GROSVENOR ISLAND 2280.8 4555.8 200 197 10 20 70 0 END

*** Geology and eroded thicknesses taken from Dome Bay P-36 *** No pdf on strata, taken from Dome Bay P-36 *** Grosvenor assumed complete (Kristoffer Bay B-04)

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Mocklin Bay D-23 bottom age bottom age top % sand % silt % shale % lmst est. thickness 19 KANGUK FM 350 95 69 0 30 70 0 350 HASSEL FM 600 103 95 60 30 10 0 250 CHRISTOPHER FM 1400 113 103 25 15 60 0 800 WALKER ISLAND MBR 1600.0 126 113 60 20 20 0 200 RONDON MBR 1650.0 128 126 30 10 50 10 50 Pat. Island eroded 1900 131 128 65 25 10 0 250 PAT. ISLAND MBR 292.6 2192.6 137 131 65 25 10 0 1900 total eroded DEER BAY FM 905.3 2805.3 153 137 0 10 90 0 RINGNES FM 1094.5 2994.5 161 153 10 0 90 0 MCCONNELL FM 1292.7 3192.7 172 161 15 0 85 0 SANDY POINT FM 1325.3 3225.3 173 172 80 0 20 0 JAMESON BAY FM 1566.7 3466.7 184 173 20 10 70 0 KING CHRISTIAN FM 1713.6 3613.6 190 184 80 0 20 0 LOUGHEED ISLAND FM 1853.2 3753.2 193 190 5 45 50 0 MACLEAN STRAIT FM 1955.3 3855.3 197 193 60 10 30 0 GROSVENOR ISLAND 2046.7 3946.7 200 197 10 20 70 0 SKYBATTLE FM 2141.2 4041.2 204 200 10 20 70 0 BARROW FM 2721.0 4621.0 217 204 60 20 20 0 PAT BAY FM 2802 4702.0 219.5 217 0 60 40 0 END

*** Eroded thicknesses taken from Thor P-38 and Kristoffer Bay B-06 *** assumed Pat Bay is 100 m (Kristoffer Bay B-06) *** Geology from Dome Bay P-36 and Kristoffer Bay B-06 (for Barrow only)

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Noice D-41 bottom age bottom age top % sand % silt % shale % lmst est. thickness 17 KANGUK FM 350 95 69 0 30 70 0 350 HASSEL FM 600 103 95 60 30 10 0 250 CHRISTOPHER FM 1400 113 103 25 15 60 0 800 WALKER ISLAND MBR 1600.0 126 113 60 20 20 0 200 RONDON MBR 1650.0 128 126 30 10 50 10 50 Pat. Island eroded 1900 134 128 65 25 10 0 250 PAT. ISLAND MBR 307.0 2207.0 137 134 65 25 10 0 1900 total eroded DEER BAY FM 827.0 2727.0 153 137 0 10 90 0 RINGNES FM 1152.0 3052.0 161 153 10 0 90 0 MCCONNELL FM 1385.0 3285.0 172 161 15 0 85 0 SANDY POINT FM 1409.0 3309.0 173 172 80 0 20 0 JAMESON BAY FM 1674.0 3574.0 184 173 20 10 70 0 KING CHRISTIAN FM 1838.0 3738.0 190 184 80 0 20 0 LOUGHEED ISLAND FM 2005.9 3905.9 193 190 5 45 50 0 MACLEAN STRAIT FM 2076.0 3976.0 197 193 60 10 30 0 GROSVENOR ISLAND 2216.0 4116.0 200 197 10 20 70 0 SKYBATTLE FM 2347.0 4247.0 204 200 60 20 20 0 END

*** Skybattle assumed to be 100 m, so complete (e.g. Mocklin Point D-23) *** Geology from Dome Bay P-36

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Sutherland O-23 bottom age bottom age top % sand % silt % shale % lmst est. thickness 23 KANGUK FM 250 95 69 0 3070 0 250 HASSEL FM 350 103 95 60 3010 0 100 CHRISTOPHER FM 1200 113 103 25 1560 0 850 WALKER ISLAND MBR 1400 126 113 85 58 0 200 RONDON MBR 1450 128 126 50 500 0 50 PAT. ISLAND MBR 1925 137 128 70 1020 0 475 Deer Bay eroded 2225 150 137 10 1570 5 300 DEER BAY FM 73.2 2323.2 153 150 5 5 90 0 2225 total eroded AWINGAK FM 113.7 2363.7 157 153 20 5 75 0 RINGNES FM 171 2421 161 157 5 5 90 0 MCCONNELL FM 237.7 2487.7 172 161 10 20 65 0 SANDY POINT FM 250.5 2500.5 173 172 0 100 0 0 JAMESON BAY FM 387.1 2637.1 184 173 0 95 0 5 KING CHRISTIAN FM 497 2747 190 184 100 0 0 0 LOUGHEED ISLAND FM 518 2768 193 190 60 40 0 0 MACLEAN STRAIT FM 646.2 2896.2 197 193 100 0 0 0 GROVESNOR ISLAND 686.4 2936.4 200 197 10 10 80 0 SKYBATTLE FM 804 3054 204 200 60 0 40 0 BARROW FM 1200 3450 217 204 20 20 60 0 HOYLE BAY FM 1708 3958 226 217 0 17 80 3 GORE POINT MBR 1793 4043 228 226 10 0 10 80 MURRAY HARBROUR FM 2002 4252 245 228 15 20 60 20 BLIND FIORD FM 4457.1 6707.1 251 245 15 35 50 0 END

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Thor P-38 bottom age bottom age top % sand % silt % shale % lmst est. thickness 16 KANGUK FM 350 95 69 0 30 70 0 350 HASSEL FM 600 103 95 60 30 10 0 250 CHRISTOPHER FM 1500.0 113 103 25 15 60 0 900 WALKER ISLAND MBR 1700.0 126 113 85 5 8 0 200 RONDON MBR 1750.0 128 126 50 50 0 10 50 Pat. Island eroded 2050.0 133 128 70 10 20 0 300 PAT. ISLAND FM 234.1 2284.1 137 133 60 10 30 0 2050 total eroded DEER BAY FM 521.2 2571.2 153 137 2 18 80 0 RINGNES FM 734.6 2784.6 161 153 0 8 90 2 MCCONNELL FM 932.7 2982.7 172 161 15 15 70 0 JAMESON BAY FM 1175.6 3225.6 184 173 4 5 90 1 KING CHRISTIAN FM 1332.0 3382.0 190 184 70 0 30 0 LOUGHEED ISLAND FM 1393.2 3443.2 193 190 40 20 40 0 MACLEAN STRAIT FM 1631.3 3681.3 197 193 90 0 10 0 GROSVENOR ISLAND 1738.0 3788.0 200 197 15 0 85 0 SKYBATTLE FM 1800.0 3850.0 204 200 40 0 60 0 END

*** Skybattle assumed complete (from Mocklin Point D-23 and Kristoffer Bay B-06) *** removed 30 m of igneous material from Skybattle

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Wallis A-73 bottom age bottom age top % sand % silt % shale % lmst est. thickness 17 KANGUK FM 250 95 69 0 3070 0 250 HASSEL FM 350 103 95 60 3010 0 100 Christopher eroded 850 109 103 15 2550 10 500 CHRISTOPHER FM 344.0 1194.0 113 109 15 25 50 10 850 total eroded WALKER ISLAND MBR 556.0 1406.0 126 113 95 5 0 0 RONDON MBR 616.0 1466.0 128 126 0 20 80 0 PAT. ISLAND MBR 1178.0 2028.0 137 128 80 10 10 0 DEER BAY FM 1529.0 2379.0 153 137 5 15 80 5 RINGNES FM 1663.0 2513.0 161 153 0 20 80 0 MCCONNELL FM 1800.0 2650.0 172 161 10 40 50 0 JAMESON BAY FM 2027.0 2877.0 184 173 0 60 40 0 KING CHRISTIAN FM 2132.0 2982.0 190 184 50 25 25 0 LOUGHEED ISLAND FM 2210.0 3060.0 193 190 70 0 30 0 MACLEAN STRAIT FM 2414.0 3264.0 197 193 70 0 30 0 GROSVENOR ISLAND 2517.0 3367.0 200 197 0 50 50 0 SKYBATTLE FM 2647.0 3497.0 204 200 60 40 0 0 BARROW FM 2827.0 3677.0 208 204 0 30 70 0 END

** assuming Barrow is 450 m (average of the thickness from Sutherland O-23 and King Christian N-06)

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Wallis K-62 bottom age bottom age top % sand % silt % shale % lmst est. thickness 17 KANGUK FM 250 95 69 0 3070 0 250 HASSEL FM 350 103 95 60 3010 0 100 Christopher eroded 800 109 103 15 2550 10 450 CHRISTOPHER FM 391.0 1191.0 113 109 15 25 50 10 800 total eroded WALKER ISLAND MBR 570.0 1370.0 126 113 95 5 0 0 RONDON MBR 624.8 1424.8 128 126 0 20 80 0 PAT. ISLAND MBR 1085.7 1885.7 137 128 80 10 10 0 DEER BAY FM 1500.0 2300.0 153 137 5 15 80 5 RINGNES FM 1635.0 2435.0 161 153 0 20 80 0 MCCONNELL FM 1755.0 2555.0 172 161 10 40 50 0 JAMESON BAY FM 1956.0 2756.0 184 173 0 60 40 0 KING CHRISTIAN FM 2064.4 2864.4 190 184 50 25 25 0 LOUGHEED ISLAND FM 2122.6 2922.6 193 190 70 0 30 0 MACLEAN STRAIT FM 2319.6 3119.6 197 193 70 0 30 0 GROSVENOR ISLAND 2426.0 3226.0 200 197 0 50 50 0 SKYBATTLE FM 2523.5 3323.5 204 200 60 40 0 0 BARROW FM 2628.9 3428.9 206 204 0 30 70 0 END

** assuming Barrow is 450 m (average of the thickness from Sutherland O-23 and King Christian N-06)

145