Ruhr-Universität Bochum. Institute of Geology, Mineralogy and Geophysics. Structural analysis of inversion features of the Barents Sea

Transcription

1 Ruhr-Universität Bochum Institute of Geology, Mineralogy and Geophysics Structural analysis of inversion features of the Barents Sea Doctoral thesis by Muhammad Armaghan Faisal Miraj Bochum, 2017

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3 Ruhr-Universität Bochum Institute of Geology, Mineralogy and Geophysics Structural analysis of inversion features of the Barents Sea Doctoral thesis by Muhammad Armaghan Faisal Miraj Faculty of Geoscience Ruhr-Universität Bochum Prof. Dr. Christophe Pascal Prof. Dr. A. Immenhauser

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5 Acknowledgement I owe a debt of gratitude to my supervisor Prof. Christophe Pascal for his supervision and valuable discussions. Throughout the research work he provided me sound advices, good teaching and constructive ideas. I am greatly thankful to Prof. Adrian Immernhauser and Prof. Rebecca Harrington for their reviewes. I would also like to thank the Department of Geosciences, Univesity of Oslo, Norway for providing me seismic data. In particular, many thanks to Prof. Roy Helge Gabrielsen and Prof. Jan Inge Faleide for valuable discussions and encouragements. Special thanks to DAAD (Deutscher Akademischer Austauschdienst) and HEC (Heigher Education Commision) Pakistan for funding. Thanks to my friends and Colleagues especially Richard, Caroline, Nicole, Henrick, Kathi and Sara for their moral support. It is difficult to find adequate words to express my many thanks and tremendous gratitude to my parents and specially my wife who rode side by side with me in this tiresome journey of completing my studies at Ruhr University Bochum. Finally, I would like to thank all the teaching and administration staff of the Institut of GMG, RUB. Muhammad Armaghan Faisal i

6 Abstract Basin inversion is a common phenomenon worldwide. Inversion/reactivation of normal fault systems within a basin has a significant impact on its final structure and modifies reservoirs and fluid paths. The study of inversion is prone to shed lights on second or third order tectonic phenomena that have escaped the overall framework of plate tectonics theory until now. In addition, the understanding of inversion structures is of prime importance for e.g. oil exploration. However, in most cases the causes and mechanisms related to inversion remain enigmatic. The present work aims to address these issues by means of studying the particular case of the western Barents Sea. Inversion structures including folds, reverse faults are observed along the Bjørnøyrenna Fault Complex and the Ringvassøy Loppa Fault Complex in the western Barents Sea, although both fault complexes are extensional in origin and developed in mid-jurassic to Early Cretaceous. Subsidence along the fault complexes was interrupted in Early Cretaceous (Valanginian to early Barremian) because of syn-rift localized tectonic inversion, itself related to the uplift of the Loppa High. The Early Cretaceous inversion caused dextral transpression along the boundary faults adjacent to the Loppa High. The second phase of inversion is interpreted to be Late Cretaceous (mid-cenomanian) in age, coeval to the deposition of the Kolmule Formation in the Bjørnøyrenna Fault Complex and the Ringvassøy Loppa Fault Complex. The later phase of compression is of regional significance and related to NW-SE directed far field stresses in Late Cretaceous which caused head-on inversion in the study area. The results of structural restoration of Cretaceous inversion events in the Bjørnøyrenna Fault Complex, western Barents Shelf, are presented. The aim of the study is to identify the structures related to inversion (anticlines, reverse faults) by means of identifying and locating null point positions. 2D MOVETM (structural modeling and analysis software by Midland Valley Exploration Ltd) is used to restore three key seismic profiles located in the central and northern segments of the Bjørnøyrenna Fault Complex. Key profiles 1 and 2 reveal null point positions at the base of the Cretaceous (Hekkingen Formation). Null point positions show progressive compressional inversion of syn-rift Early Cretaceous deposits (Knurr Formation). Below and above null points the geometries of the restored faults show normal and reverse faulting respectively. The results of the restored key profiles 1 and 2 confirm reverse faulting at the Lower Cretaceous triggered by inversion of the study area. The restored sections also show positive inversion features associated with folding of the hangingwall of the base of the ii

7 Upper Cretaceous (Kolmule Formation). The reconstruction of the amount of eroded material on the footwall block also suggests reverse faulting of the base of the Upper Cretaceous. In key profile 3 the footwall block is eroded up to the base of the Upper Cretaceous (Kolmule Formation) due to the uplift of the Loppa High. The corresponding restored section shows a compressional anticline associated with both Early and Late Cretaceous inversion events. The results of numerical modeling of inversion of faults induced by Late Triassic to Miocene tectonic stress fields in the western Barents Sea are presented. The aim is to test the potential for fault reactivation under such circumstances. A finite-element numerical code (ANSYS ) is used to simulate stress and fault slip patterns based on four 2-D thin plate modeling setups. Following previous works, four major regional inversion events are assumed: Late Triassic to Early Jurassic (E-W contraction, Model 1), Late Cretaceous (NW-SE contraction, Model 2), dextral megashear plate margin in Early Eocene (Model 3) and NW-SE Atlantic ridge push starting in Miocene (Model 4). Model 1 confirms the potential for compressional conditions in the western Barents Sea and, hence, contractional reactivation of master fault systems like the Thor Iversen Fault and Troms-Finnmark Fault complexes. Compressive regimes in the Måsøy and Hoop fault complexes favor the development of inversion structures in the study area during Late Triassic to Early Jurassic. Simulated stress patterns in Model 2 (inducing a NW-SE compressional stress) suggest a clockwise stress rotation in the Bjørnøyrenna Fault Complex and the Ringvassøy Loppa Fault Complex and pronounced stress deflections in the Asterias Fault Complex. These modeled stress deflections support tectonic inversion during Late Cretaceous in the corresponding fault complexes. The analyses suggest that significant strike-slip is to be expected to have occurred along some segments. The results obtained in Model 3 suggest that the interior of the western Barents Sea was not severely influenced by Early Eocene North Atlantic opening/shearing. The results suggest that Early Eocene sea floor spreading caused stress partitioning along the Senja Fracture Zone. The observed inversion structures in previous studies may be related to local effects. The results of Model 4 appear to be in agreement with the observed NW-SE contraction, expressed as folds and reverse faults in the study area (e.g. Ringvassøy Loppa, Bjørnøyrenna, Leirdjupt and Asterias fault complexes). The results of the four models suggest the presence of compressive structures along the major fault complexes of the western Barents Sea during iii

8 Late Triassic to Miocene but do not favor the development of inversion structures during Eocene. iv

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10 Contents 1. Introduction The western Barents Shelf Aims and thesis statement Thesis outline and framework Basin inversion tectonics Inversion tectonics Geometry and kinematics of basin inversion Driving forces Compression-related inversion (examples) Strike slip related inversion Transpressive inversion Basin inversion in the western Barents Sea Regional geological setting Inversion phases and evidences in the western Barents Sea Methodology Numerical modeling Finite Element Method ANSYS Seismic reflection survey Kingdom Suit Structural restoration MOVETM Cretaceous inversion of the western Barents Shelf: integrated seismic interpretation of the Bjørnøyrenna and the Ringvassøy Loppa fault complexes Introduction Geological setting Data and methodology Seismic interpretation Snadd Formation (Upper Triassic) Fruholmen Formation (Base Jurassic) Tubåen Formation (Lower Jurassic) Nordmela Formation (Upper Jurassic) Fuglen Formation (upper Middle Jurassic) vi

11 4.4.6 Hekkingen Formation (Base Cretaceous) Knurr Formation (Lower Cretaceous) Kolmule Formation (Base Upper Cretaceous) Results and discussions Early Cretaceous inversion Late Cretaceous inversion Conclusions Structural restoration of Cretaceous inversion events in the Bjørnøyrenna Fault Complex, western Barents Shelf Introduction Geometry and structural evolution of the Bjørnøyrenna Fault Complex Data and Methodology Structural restorations Results Restoration of key profile Restoration of key profile Restoration of key profile Discussions Conclusions Numerical modeling of multi stage basin inversion in the western Barents Shelf Introduction Numerical modeling The Finite Element Method (ANSYS ) Model set up Boundary conditions Results Late Triassic Early Jurassic Late Cretaceous Early Eocene Miocene Discussion Origin of the Cenozoic stress field Modeled stress patterns compared to observations Conclusion vii

12 7- SUMMARY References CURRICULUM VITAE Declaration List of published work viii

13 1. Introduction 1.1 The western Barents Shelf The western Barents Shelf represents the area of the epicontinental Barents Sea extending N-S between the archipelago of Svalbard and mainland Norway and E-W from the NorwegianRussian border to the Atlantic Ocean (Fig. 1.1). The area involves a mosaic of basins, numerous fault complexes and intra-basinal highs, which formed in response to various Late Paleozoic to Cenozoic tectonic events (Doré et al. 1995). The structure of the central and eastern parts of the western Barents Shelf is dominated by NE-SW to ENE-WSW-striking fault complexes, whereas the western part incorporates mostly NNW-SSE and N-S structural elements (Gabrielsen et al. 1990; Fig. 1.1). Several fault complexes of the western Barents Shelf were tectonically inverted to varying degrees from Mesozoic to Cenozoic (Faleide et al. 1988, 1993, Gabrielsen et al. 1997, Grunnaleite 2002, Bergh and Grogan 2003, Fitryanto 2011). Reactivation of pre-existing major fault zones was directly or indirectly emphasized by early investigators (Rønnevik and Motland 1979, Rønnevik et al. 1982, Faleide et al. 1984, Rønnevik and Jacobsen 1984) and was later on attributed to wrenching (Riis et al. 1986, Brekke and Riis 1987, Gabrielsen and Færseth 1988, Faleide et al. 1993a, 1993b) or head-on inversion (Gabrielsen et al. 1992, 1997). Strike-slip related inversion was dated to Mesozoic and Cenozoic (Gabrielsen et al. 1997, 2011). Inversion structures including upright open folds, deformed fault planes, reverse faults and deformation of footwall blocks were reported, in particular, from Turonian throughout Late Cretaceous and into Early Cenozoic (Gabrielsen et al. 1997). Grunnaleite (2002) conducted a regional seismic interpretation study of inversion structures on the Norwegian Shelf, including the western Barents Sea and suggested reactivation and inversion of pre-existing extensional faults during Mesozoic (Cretaceous) and Cenozoic (Eocene and Miocene). Inverted normal faults are distributed in the whole region accompanied by reverse faulting and folding. Thus inversion seems to be common and associated to most of the significant fault systems of the western Barents Sea (Gabrielsen et al. 1997). 1

14 Figure 1.1. Regional setting and major structural elements of the study area (modified from Google Maps and NPD fact maps (AFC = Asterias Fault Complex, BB = Bjørnøya Basin, BFC = Bjørnøyrenna Fault Complex, BP = Bjameland Platform, CFZ = Central Fault Zone, COB = Continental Oceanic Boundary, FSB = Fingerdjupet Sub-Basin, HB = Harstad Basin, HfB = Hammerfest Basin, HFC = Hoop Fault Complex, HrFC = Hornsund Fault Complex, KFC = Knølegga Fault Complex, LFC = Leirdjupt Fault Complex, LH = Loppa High, MB = Maud Basin, MFC = Måsøy Fault Complex, MH = Mercurius High, NB = Nordkapp Basin, NFC = Nysleppen Fault Complex, NH = Norsel High, OB = Ottar Basin, PSB = Polhem Sub-Platform, R-LFC = Ringvassøy Loppa Fault Complex, SB = Sørvestsnaget Basin, SFZ = Senja Fracture Zone, SH = Stappen High, SR = Senja Ridge, TB = Tromsø Basin, T-FFC = Troms-Finnmark Fault Complex, TIFC = Thor Iversen Fault Complex, VH = Veslemøy High, VVP = Vestbakken Volcanic Province). 1.2 Aims and thesis statement The goal of the present research is to investigate the causes and effects of Mesozoic and Cenozoic inversion events in the western Barents Shelf. For this purpose, four main scenarios have been advanced for orientations and timings of stress states responsible for fault reactivation in the western Barents Shelf. These include: (1) westward motion of Novaya Zemlya in Late Triassic-Early Jurassic (Buiter and Torsvik 2007), (2) Late Cretaceous Alpine inversion (Gabrielsen et al. 1997, Vågnes et al. 1998), 3- NW-SE Early Eocene North Atlantic 2

15 opening (Tsikalas et al. 2002, Doré et al. 2008) and 4- NW-SE directed North Atlantic ridge push (Doré and Lundin 1996). The present thesis work consists in three different research phases involving seismic interpretation, structural restoration and numerical modeling. Phase I considered seismic interpretation of 142 2D seismic lines along the Bjørnøyrenna Fault Complex and the Ringvassøy-Loppa Fault Complex. The aim of this study was to detect and interpret eventual inversion structures (folds and reverse faults) of Early and Late Cretaceous. IHSTM Kingdom 8.8 seismic and geological interpretation software was used. In Phase II, 2D structural restoration of Cretaceous inversion structures present in the Bjørnøyrenna Fault Complex was carried out. Three seismic key-profiles with different orientations (WNW-ESE and ENE-WSW) were selected for this purpose. The main objective of Phase II (i.e. kinematic modeling) was to restore the seismic sections backward, to locate null point positions and to investigate Early Cretaceous and Late Cretaceous inversion events. For this purpose 2D MOVETM (structural modeling and analysis software by Midland Valley Exploration Ltd) was used and the 2D kinematic module with different restoration techniques (i.e. 2D unfolding, flexural slip and 2D move on fault, simple shear) was dopted. In general, two main results can be obtained from restoration or backward modeling of a particular structure. The approach can validate the interpreted geometry in cross section and can provide information about the processes linked to regional progressive deformation. In Phase III, numerical modeling techniques were used to investigate inversion induced by Late Triassic to Miocene tectonic stress fields. The aim is to test the potential for fault reactivation under such specific tectonic stress fields. A finite-element numerical code (ANSYS ) was used to simulate stress patterns and fault slip based on four 2-D thin plate modeling setups. Following previous works, four major regional inversion events were assumed: Late Triassic to Early Jurassic (E-W contraction, Model 1), Late Cretaceous (NWSE contraction, Model 2), dextral megashear in Early Eocene (Model 3) and NW-SE Atlantic ridge push from Miocene to present-day (Model 4). 3

16 1.3 Thesis outline and framework The thesis consists in seven chapters including theoretical background, methodology and results of the research. Chapter 2, Basin Inversion Tectonics, gives an account on geometry and kinematics of basin inversion, main driving forces responsible for basin inversion and examples of inverted basins in world. The last sections of the chapter address previous studies of Mesozoic and Cenozoic inversion in the western Barents Sea. Chapter 3, Methodology, introduces the different techniques used in this thesis (i.e. numerical modeling, seismic interpretation and structural restoration) and presents the chosen computer tools (i.e. AnsysTM, Kingdom suit and MOVETM). Chapter 4, Mesozoic inversion in the western Barents Shelf: Integrated seismic interpretation of the Bjørnøyrenna Fault Complex and the Ringvassøy Loppa Fault Complex, presents the results of the 2D seismic interpretation. Chapter 5, Structural Restoration of Cretaceous inversion events in the Bjørnøyrenna Fault Complex, western Barents Shelf, details the results of the structural restoration of the Bjørnøyrenna Fault Complex, western Barents Shelf. Chapter 6, Numerical modeling of multi stage basin inversion in the western Barents Shelf, includes results of the numerical modeling of Mesozoic and Cenozoic inversion events. Four different models were constructed in order to predict stress patterns and to explore the conditions for tectonic inversion during four specific tectonic configurations spanning from Late Triassic to Miocene. Chapter 7 gives a brief summary of the main results of the three individual studies. 4

17 2. Basin inversion tectonics 2.1 Inversion tectonics The term inversion is refes to be reversal of sedimentary basin record in the sense of motion during different stages of basin evolution (Glennie and Boegner 1981, Mitra 1993, Coward 1994). Basin inversion can be defined as the process of shortening of extensional basins which is accommodated by compressional reactivation of pre-existing normal faults (Turner & Williams 2004). According to Cooper et al. (1989) A basin controlled by a fault system that has been subsequently compressed-transpressed producing uplift is defined as basin inversion. Figure 2.1. Schematic diagram (not to scale) showing basin inversion (modified after Williams et al. 1989) Inversion can be caused by various mechanisms and can have different origins. Many authors believe that the main cause for the compressional reactivation of faults are external horizontal stresses related to plate movements (Coward 1994) e.g. continent - continent or arc-continent collision, that can cause compression, uplift and reactivation of pre-existing faults or ocean5

18 continent convergent margins where the changes in the rate and dip of subduction may cause basin inversion. Isostatic rebound of sediments or caused by removal of glacial overburden, flexural and thermal mechanisms and salt tectonics may also cause basin inversion (Voigt 1962). Basin inversion can occur at differen scales (e.g. from basin to sub-basin scale) and is widely documented in different tectonic settings, e.g. continental rifts, rifted continental margins, backarcs, orogenic foredeeps, intracratonic basins, and in regions of strike slip faulting (e.g. Cámara and Klimowitz 1985, Gillcrist et al. 1987, Tucker and Arter 1987, Daly et al. 1989, Letouzey et al. 1990, Turner and Hancock 1990, Boldreel and Andersen 1993, Roure and Colletta 1996, Guiraud and Bosworth, 1997, Butler 1998, Underhill and Paterson1998, Gasperini et al. 2001, Morley et al. 2001, Benkhelil et al. 2002, Turner et al. 2003). 2.2 Geometry and kinematics of basin inversion Glennie & Boegner (1981) documented tectonic inversion as positive (uplift) and negative (subsidence) relative to a fault system. Positive inversion occurs when extensional faults reverse their sense of motion during compressional tectonics which causes the basin to turn inside out and to become a positive feature (Fig. 2.1 and 2.2; Williams et al. 1989). As a result each fault may show net extension at deep levels and depicts net contraction associated with an anticline in the upper portion of the faulted rocks. In positive inversion the areas which initially underwent subsidence were uplifted afterwards. 6

19 a b Figure 2.2. Schematic diagram showing classical positive inversion structure. a) Extension and b) basin inversion (modified after Williams et al. 1989). Negative inversion is the extensional reactivation of existing contractional faults. It gives a useful and relevant concept to understand syn- and post orogenic extension (Williams et al., 1989, Turner et al. 2003). The geometry of inversion structures is very much influenced by the stratigraphic built up of the extensional basin. The pre-rift sequence is deposited before any fault activity and can be recognized by the equal thickness of stratigraphic units on hangingwalls and footwalls. The syn-rift sequence is deposited during extensional faulting and stratigraphic thickness changes from footwall to hangingwall i.e. growth faulting. The post-rift sequence is deposited when extensional faulting ceased. This latter sequence may also be deposited on top of a marked break-up unconformity reflecting erosion or non-deposition (Fig. 2.3). 7

20 Figure 2.3. Schematic diagram showing accumulation of pre-rift, syn-rift, and post-rift stratigraphic units before, during and after extensional faulting. A null point is an apparently unfaulted point on a fault plane (Fig. 2.4). In positive inversion, the downward position of the null point evidences the progressive compressional inversion of an extensional syn-rift sequence. Figure 2.4. Schematic diagram showing the apparently unfaulted point on a fault plane, i.e. the null point (modified after Cramez and Letouzey 2014). Below the null point the fault is normal while above it the fault is reverse (Fig 2.5). The position of the null point on a fault plane depends on the amount of inversion, i.e.the higher the null point on the fault plane the smaller the inversion. The position of the null point at the bottom of the fault plane shows total inversion. 8

21 Figure 2.5. Downward movement of the null point evidencing positive inversion (modified after Cramez and Letouzey 2014). 2.3 Driving forces Far-field stresses transmitted within tectonic plates can cause tectonic inversion (Coward 1994, Lowell 1995). Basins can be inverted by compression, strike-slip or combination of both (e.g. transpression) Compression-related inversion (examples) The orientation of horizontal forces responsible for inversion ranges from 0 to 90 with respect to pre-existing faults. Inversion caused by compression at 90 to existing faults is highly effective (Letouzey et al. 1990), in case of reverse stress regimes (Fig. 2.6a). a) b) Figure 2.6. Schematic diagram depicting the orientation of the three principal stresses (σ1, σ2 and σ3) and related stress regimes: (a) reverse and (b) strike slip (modified from 9

22 The world class example of tectonic inversion caused by direct compression are the Atlas Mountains in Morocco, where the ENE-WSW Triassic-Jurassic Atlas Rift inverted due to NNW-SSE compression caused by Miocene convergence of Africa and Iberia (Fig 2.7). The direct compression resulted into the formation of low-angle thrusts on both sides of the Atlas (Bennett et al. 1992, Brede et al. 1992). Figure 2.7. Map of morocco showing the direction of maximum principal compressive stress (σ1) (Modified after Lowell 1995). Another example of inversion due to almost direct compression is the Uinta Mountains in northeast Utah which are situated in the foreland of the Wyoming-Utah-Idaho thrust and fold belt. An E-W trending Uinta Basin which developed during Proterozoic rifting was inverted during the N-S (Gries 1982) or NE-SW (Stone 1989) late Laramide (Early-Middle Eocene) compression (Fig. 2.8). The inversion of the Uintas is evidenced by reactivation of normal faults and development of km-scale thrusts. 10

23 Figure 2.8. N-S cross section of the Uinta Mountains showing inversion caused by direct compression (Hansen 1986). (Cz = Cenozoic rocks, MzPz = Mesozoic and Paleozoic) Strike slip related inversion Inversion caused by strike slip require a dominance of lateral motion where σ1 (maximum principal stress) and σ3 (minimum principal stress) lay in the horizontal plane and the σ2 (intermediate principal stress) is vertical (Fig. 2.6b). Strike slip related inversion has also been observed in different parts of the world (e.g. offshore northeast Brazil, western Barents Sea). The offshore Ceara Piaui Basin in northeast Brazil experienced inversion. Aptian rift sediments were inverted due to the convergent right-lateral movement caused by the separation of South America and Africa (Ponte and Asmus 1978). The inversion resulted in the reactivation of high angle pre-existing normal faults (Fig. 2.9). 11

24 Figure 2.9. Cross section (based on reflection seismic lines) showing the evolution of the northeast Brazilian continental margin under transpression. (a) Aptian rifting between South America and Africa; (b) inversion, reactivation of high angle pre-existing normal faults and truncation of the upper Aptian below a post-aptian unconformity (modified after Lowell 1995). Strike slip related inversion has also been identified in the western Barents Sea and dated to Late Paleozoic, Mesozoic and Cenozoic (Gabrielsen et al. 2011). Inversion structures including upright open folds, reverse faults, deformation of footwall blocks and deformed fault planes were reported from Turonian throughout Late Cretaceous and into Early Cenozoic in particular (Gabrielsen et al. 1997). However issues concerning the exact timings and stress directions responsible for the multiple reactivations of main fault complexes in the western Barents Sea are still under consideration Transpressive inversion Inversion can be caused by the combination of compression and strike slip, i.e. transpression. Although compression at 90 to pre-existing faults is more effecient (Letouzey et al. 1990), but it is not the general case (Fig. 2.10). 12

25 Figure Possible angles of incidence of compression to pre-existing normal faults. (Modified after Lowell 1995). The analysis of the relative contribution of compression and strike slip in an inverted region is always a difficult task. The azimuth of slip can be determined from measurements of strain in the field but precise assessment of fault slip using 2D seismic data is not possible (Lowell 1995). 2.4 Basin inversion in the western Barents Sea Regional geological setting The complex mosaic of platform areas and basins of the Barents Sea formed mainly through continental collisions in Paleozoic, e.g. Caledonian and Uralian orogenies (Doré 1996, Gee et al. 2008), rifting events during Paleozoic and Mesozoic (Smelror et al. 2009, Tsikalas et al. 2012) and opening of the North Atlantic Ocean in Cenozoic (Gabrielsen et al. 1990, Faleide et al. 2008). 13

26 The series of orogenies and rifting events divided the study area into three distinct provinces (Fig. 2.11): (1) the Svalbard Platform to the north, (2) a rifted domain between the Svalbard Platform and the Norwegian coast and (3) the continental margin to the west (Faleide et al. 1993a). The regional geology of these provinces has been published by various authors (Faleide et al. 1984, Gabrielsen et al and the reference therein). The metamorphic basement of the Barents Sea was consolidated during the Caledonian Orogeny which includes closure of the Iapetus Ocean and the consequent collision of Laurentia with Baltica in early Paleozoic (Dengo and Røssland 1992). The overall strike of Caledonian structures in northern mainland Norway is NE-SW (Sturt et al. 1978, Townsend 1987) whereas a NW-SE structural trend predominates in Spitsbergen (Harland 1985, Dengo and Røssland 1992). The orientation of later extensional features, formed by subsequent rifting phases, mimics the trend of pre-existing fracture systems. This stubbing similarity shows that the orientation of younger extensional features was largely controlled by the preexisting structural grain (Gabrielsen 1984, Gabrielsen et al. 1990, Dengo and Røssland 1992). The N-S to NNW-SSE and WNW-ESE to NW-SE structural strike in Svalbard and northern Norway formed due to the Archean to late Precambrian deformation (Harland 1969, Harland et al. 1974, Beckinsale et al. 1975, Kjøde et al. 1978, Berthelsen and Marker 1986, Rider 1988). In contrast, the Caledonian deformation resulted in ENE-WSW to NE-SW striking structural features (Roberts 1971, 1972, Worthing 1984, Fig. 2.11). According to Gabrielsen et al. (1990), the structural trend of all features in the western Barents Sea cannot be directly linked to the Caledonian Orogeny because most of the major structural trends may have been shaped by Devonian tectonics. The rifting events responsible for the development of the post - Caledonian geological setting of the western Barents Sea includes: Late (?) Devonian - Carboniferous, Middle JurassicEarly Cretaceous and Early Cenozoic events (Faleide et al. 1993a). All of these major rifting phases involved several tectonic pulses. 14

27 Figure Regional setting and major structural elements of the study area (modified from google map and NPD fact maps (AFC = Asterias Fault Complex, BB = Bjørnøya Basin, BFC = Bjørnøyrenna Fault Complex, BP = Bjameland Platform, CFZ = Central Fault Zone, COB = Continental Oceanic Boundary, FSB = Fingerdjupet Sub-Basin, HB = Harstad Basin, HfB = Hammerfest Basin, HFC = Hoop Fault Complex, HrFC = Hornsund Fault Complex, KFC = Knølegga Fault Complex, LFC = Leirdjupt Fault Complex, LH = Loppa High, MB = Maud Basin, MFC = Måsøy Fault Complex, MH = Mercurius High, NB = Nordkapp Basin, NFC = Nysleppen Fault Complex, NH = Norsel High, OB = Ottar Basin, PSB = Polhem Sub-Platform, R-LFC = Ringvassøy Loppa Fault Complex, SB = Sørvestsnaget Basin, SFZ = Senja Fracture Zone, SH = Stappen High, SR = Senja Ridge, TB = Tromsø Basin, T-FFC = Troms-Finnmark Fault Complex, TIFC = Thor Iversen Fault Complex, VH = Veslemøy High, VVP = Vestbakken Volcanic Province). The change in stress system, from compressional to extension in Late Devonian to Early Carboniferous caused the formation of Bjørnøya Basin, Fingerdjupet Basin, Hammerfest Basin, Maud Basin, Nordkapp Basin, Ottar Basin and Tromsø Basin in the western Barents Sea (Dengo and Røssland, 1992). According to Harland (1969), Faleide et al. (1984), Rønnevik & Jacobsen (1984) and Ziegler (1988) the first rifting event (Late Devonian to Early Carboniferous) in the western Barents Sea initiated along the sinistral strike-slip fault and along a conjugate dextral strike-slip fault in the central Barents Sea. Dengo and Røssland 15

28 (1992) argued and suggested that the major structural elements in the Barents Sea mainly developed due to dip-slip normal faulting with little evidence of strike-slip components. Deformation along the western parts of the Barents Sea continued throughout the Mesozoic and the Cenozoic whereas the eastern and northeastern parts remained tectonically less active since the Late Carboniferous (Gabrielsen et al. 1990). The Permian is considered to be a period of thermal subsidence in the Barents Sea (Dengo and Røssland, 1992). Major structural elements which controlled the subsequent structural architecture of the Barents Sea may have been established by the end of the late Paleozoic (Gabrielsen et al. 1990). In the eastern part of the Barents Sea, closure of the Uralian Sea took place from Late Permian to Early Triassic and the Barents Sea is assumed to be a distal foreland basin of the Uralian Orogeny (Dengo and Røssland 1992). The eastern parts of the Barents Sea experienced subsidence during Triassic and Early Jurassic whereas the western parts remained tectonically quiet; however the Stappen High and the Loppa High were submitted to tilting (Gabrielsen et al. 1990). The Middle Jurassic to Early Cretaceous rifting phase is believed to be the most significant one in the western Barents Sea. It resulted in the formation of major basins and highs. The tectonic process caused high rates of subsidence in the western part of the Bjørnøya Basin and the Tromsø Basin in Early Cretaceous is merely complex, while evidences of local inversion along the Ringvassøy-Loppa Fault Complex and Bjørnøyrenna Fault Complex are also recorded (Gabrielsen et al. 1990, 1997). The study area was also affected by Late Cretaceous-Early Cenozoic tectonic inversion (Gabrielsen et al. 1997). Minor folds and thrust faults at the base of the Upper Cretaceous in the central segment of the Bjørnøyrenna Fault Complex were interpreted by Gabrielsen et al. (1997). During and after Early Cenozoic rifting and breakup (earliest Eocene), the western margin of the Barents Sea was subject to tectonic dextral shear and associated folding with NW-SE-striking fold axes (Faleide et al. 1996). The western Barents Sea continental margin developed at the Paleocene-Eocene transition (~ Ma) as a result of continental breakup and opening of Norwegian-Greenland Sea which was linked by the regional megashear system (De Geer Zone) to the Arctic Eurasian Basin (Faleide et al. 1996, 2008). According to Libak et al. (2012), the De Geer Zone is a large dextral continental strike-slip zone and was located between the western Barents Sea/Svalbard and northeast Greenland, and extended from northern Norway to the Arctic Ocean (Fig. 2.12). 16

29 The Northeast Atlantic Basin and the Eurasian Basin were linked by this transform margin (Doré 1995). Initial spreading started along the Aegir and Mohns ridges in Early Eocene (Talwani & Eldholm 1977, Czuba et al. 2011) and relative plate motions were parallel to the strike-slip system (Libak et al. 2012, Fig. 2.12). As a matter of fact, the breakup did not propagate into the southwestern Barents Sea, but shear motions along the De Geer Zone were developed and this relative plate movement generated different margin segments, i.e. shear and rifted ones on the western Barents Sea margin (Libak et al. 2012). To the south, the Senja Fracture Zone (SFZ) marks the southern segment of the purely sheared margin (Fig.2.12), developed due to the opening of the Norwegian-Greenland Sea during the Eocene (Faleide et al. 2008). The generation of the Senja Fracture Zone (SFZ) was initially related to continent-continent shear followed by continent-ocean shear and has been passive since Oligocene imes (Faleide et al. 2008). In the central part of the western Barents Sea continental margin, the strike-slip system changed and resulted into a pull-apart setting (Faleide et al. 1993, Breivik et al. 1998, Ryseth et al. 2003) which caused generation of a rifted segment associated with volcanism, i.e. the VVP (Vestbakken Volcanic Province). The rifted margin (VVP) linked sheared margin segments to the north and south (Fig. 2.12). A number of buried mounds interpreted as buried volcanoes of Early Eocene to Early Oligocene age are reported in the northwestern parts of the province (Faliede et al. 1988, Libak et al. 2012). The significant intrusion of dense magmatic material at the VVP (Vestbakken Volcanic Province) was due to the transtension which caused thinning of the crust (Sundvor and Eldholm 1979, Eiken and Austegard 1987, Eldholm et al. 1987, Czuba et al. 2011). The Cenozoic evolution of the VVP (Vestbakken Volcanic Province) includes several tectonic and volcanic events (Faleide et al. 1988, Richardsen et al. 1991, Sættem et al. 1994, Eidvin et al. 1998, Jebsen 1998, Ryseth et al. 2003). In Early Eocene, only the western parts of the province experienced volcanic activity, but later effected by erosion (Jebsen 1998, Faleide et al. 1988). The presence of middle late Eocene sediments above the volcanic flows in the VVP proves that the area was a major sedimentary basin at that time. The main source of these sediments was the uplifted Stappen High in the northeast (Fig. 2.11, Richardsen et al. 1991, Ryseth et al. 2003, Libak et al. 2012). 17

30 Figure Schematic diagram (not to scale) showing tectonic setting of the Arctic region (a) Early Eocene breakup and (b) Post Eocene extension. AR = Aegir Ridge, DGZ = De Geer Zone, HFC = Hornsund Fault Complex, MR = Mohns Ridge, SFZ = Senja Fracture Zone, VVP = Vestbakken Volcanic Province. (Modified after Doré et al. 2008) The Cenozoic development of the VVP (Vestbakken Volcanic Province) was mainly controlled by two major fault zones including; the Knølegga Fault Complex (KFC) (Gabrielsen et al. 1990) and Central Fault Zone (CFZ). The N-S to NNW-SSE Knølegga Fault Complex (KFC) lies in the east of the VVP and marks the western boundary of the Stappen High (Fig. 2.11). The Central Fault Zone (CFZ) lies on the eastern limit of the VVP (Fig. 2.11) and marks a boundary between Eocene seabed sediments in the east and PliocenePleistocene seabed sediments in the west (Sættem et al. 1994). In contrast, Faleide et al. (1988), observed volcanic products also in the east of the CFZ (Central Fault Zone). Many previous authors (Eldholm et al. 1987, Faleide et al. 1988) believed that the VVP is underlain by thick oceanic crust and the COB (Continent-ocean boundary) is located close to the CFZ (Central Fault Zone). Faleide et al. (1991) suggested that the outer parts of the province contain thick oceanic crust while the inner parts are composed of continental crust which covered by sediments and volcanics. Later studies (Ryseth et al. 2003) showed that the eastern parts of the province represent stretched continental crust covered by pre-breakup sediments. Most recent studies (Breivik et al. 1999, Engen et al. 2008, Czuba et al. 2011) of the VVP (Vestbakken Volcanic Province) suggest that the whole province is underlain by stretched 18

31 continental crust. From middle Miocene to middle Pliocene the western Barents Sea and the VVP experienced regional uplift (Libak et al. 2012). To the north, a margin segment developed along the Hornsund Fault Complex (Faleide et al. 1993, Faleide et al. 2008, Libak et al. 2012) which is affected by continent-ocean and oblique continent-continent shearing with both transpressional and transtensional components during Eocene (Gorgan et al. 1999, Berg and Grogan 2003). The restraining bend along northnorthwest trending faults between the northeast Greenland and Svalbard caused transpression and as a result the Spitsbergen fold and thrust belt was formed (Czuba et al. 2011), while the releasing bend between the Hornsund Fault Complex and the Senja Fracture Zone facilitated by Oligocene rifting (Faleide et al. 1993). This rifting caused reactivation of NE trending normal faults in the Sørvestnaget Basin (Jebsen 1998). In Early Oligocene spreading ceased in the Labrador Sea Baffin Bay (Talwani & Eldholm 1977, Mosar et al. 2002) and the spreading direction between Greenland and Eurasia changed from NNW-SSE to NW-SE (Oakey 2005, Faleide et al. 2008). The change in relative plate movement caused the progressive development of the Mid Atlantic Ridge towards north (Czuba et al. 2011). The propagation of the spreading axis into the Spitsbergen Shear Zone resulted into the formation of an obliquely spreading and asymmetric the Knipovich Ridge (Czuba et al. 2011). The opening of the Fram Strait and a deep water connection to the Arctic Ocean basin in Early Miocene are also outcomes of the change in plate motion (Jakobsson et al. 2007, Engen et al. 2008). Cenozoic inversion in the western Barents Sea is assumed to be caused by North Atlantic ridge push (Ranalli and Chandler 1975, Stephansson 1988, Talbot and Slunga 1989, Spann et al. 1991). According to Doré and Lundin (1996), the renewed direction of plate motion caused compression in the area due to the counterclockwise shift in the poles of rotation in the North Atlantic, during A13 A7 chrons (35 25 Ma). NW-SE transfer of stress (ridge push) in Miocene also caused development of inversion structures including anticlines and reverse faults in the western Barents Sea (Doré and Lundin 1996). Vågnes et al. (1998) suggested that the ridge push force will not be affected by shift in plate motion. According to Srivastava and Tapscott (1986), no such changes in spreading axis are noticed in the North Atlantic in the Early Oligocene. 19

32 Inversion phases and evidences in the western Barents Sea Most of the fault complexes in the western Barents Sea were inverted due to different tectonic processes from Late Triassic to Miocene (Eldholm 1977, Myhre and Eldholm 1988, Faleide et al. 1988, 1993, Richardson 1992, Gabrielsen et al. 1990, 1997 and references therein, Grunnaleite 2002, Bergh and Grogan 2003, Fig. 2.13). Figure Inversion direction in the western Barents Sea (modified from Grunnaleite 2002) Reactivation of pre-existing major fault zones was evidenced by early investigators in the form of wrenching (Riis et al. 1986, Brekke and Riis 1987, Gabrielsen and Færseth 1988, Faleide et al a, b) or head-on inversion (Gabrielsen et al. 1992, 1997). Strike slip related inversion has also been recognized and dated to Late Paleozoic, Mesozoic and Cenozoic (Gabrielsen et al. 2011). Grunnaleite (2002) conducted a regional study of inversion structures including the greater part of the western Barents Sea. In the study, all the major fault complexes, including the Knølegga Fault Zone, Bjørnøyrenna Fault Complex, Leirdjupet Fault Complex, Ringvassøy Loppa Fault Complex and Hoop Fault Complex were found to show signs of inversion. Inverted normal faults are distributed in the whole region accompanied with reverse faulting and folding. 20

33 Asterias Fault Complex The Asterias Fault Complex also known as the Southern Loppa High Fault System (Faleide et al. 1984, Gabrielsen et al. 1984, Berglund et al. 1986), is located between N, 20 E and N, 24 E and separates the Hammerfest Basin from Loppa High (Fig. 2.11). The western limit of the Asterias Fault Complex connects to the Ringvassøy - Loppa Fault Complex (Gabrielsen et al. 1990). According to Gabrielsen et al. (1984, 1990), the E-W trending Asterias Fault Complex is a first- or second- order basement-involved extensional structure initiated in Triassic to Jurassic (Gabrielsen et al 1984). Presence of inversion structures including reverse faults, half flower structures and local domes at the Jurassic Cretaceous boundary are evidenced at the western segment of fault complex (west of E, Fig. 2.14) and at its junction with the Ringvassøy - Loppa Fault Complex (Berglund et al. 1986, Brekke and Riis 1987). A (dextral) strike-slip fault forming half flower structures at the end of Jurassic time is also suggested by Rønnevik and Jacobsen (1984). According to Indrevær et al. (2016), head-on contraction is suggested in the early Barremian early Aptian and early Barremian middle Albian along the Asterias Fault Complex. The formation of inversion structures during the above mentioned time periods was caused by uplift of the Loppa High due to space accommodation problems (Indrevær et al. 2016). 21

34 Figure Uninterpreted and interpreted seismic lines (BSS and NBR07RE ) crossing the Hammerfest Basin, from the Finnmark Platform in the south and the Loppa High in the north (modified after Indrevær et al. 2016). See figure 2.11 for location. Several models have been proposed for the development of inversion structures along the Asterias Fault Complex including e.g. regional strike slip movement (Rønnevik et al. 1982, Rønnevik and Jacobsen 1984, Riis et al. 1986, Gabrielsen et al. 2011), gravity induced dextral shear of the Hammerfest Basin sedimentary fill (Ziegler et al. 1986), largescale horizontal rotation of the Hammerfest Basin relative to the Loppa High (Gabrielsen and Færseth 1988) and uplift and clockwise rotation of the Loppa High (Indrevær et al. 2016) Bjørnøyrenna Fault Complex The Bjørnøyrenna Fault Complex strikes mainly NE-SW and is situated between 72º N, 19' E and 73º 15' N, 22º E. Rønnevik and Jacobsen (1984) described the fault complex as the southeastern boundary fault of the Bjørnøya Basin whereas Gabrielsen et al. (1984), defined it as the northeastern extension of the Ringvassøy - Loppa Fault Complex (Fig. 2.11). In general the fault complex marks the boundary between the Bjørnøya Basin and the Loppa High in the southwest and it separates the Loppa High from the Fingerdjupet Subbasin in the north east (Rønnevik et al. 1975, Hinz and Schlüter 1978, Rønnevik and Motland 1981, Gabrielsen et al. 1990). The Bjørnøyrenna Fault Complex is of extensional origin with sets of normal faults having large throws and was active in the Late Jurassic to Early Cretaceous. Signs of tectonic inversion including domal features, deformed fault planes and reverse faults are observed affecting Cretaceous to Cenozoic sedimens along the fault complex (Gabrielsen et al. 1997). Gabrielsen et al. (1997) suggested that the Early Cretaceous reactivation of the normal faults wass caused by dextral shear (Fig. 2.15a) and that they experienced NW-SE compressional inversion in Late Cretaceous Early Cenozoic (Fig. 2.15b). 22

35 N a b Figure Structural development of the Bjørnøyrenna Fault Complex indicating (a) Early Cretaceous dextral shear and (b) Late Cretaceous-Early Cenozoic NW-SE compression (modified from Gabrielsen et al. 1997) Hoop Fault Complex The Hoop Fault Complex is located between 72º 50 N, 21º 50 E and 74º N, 26º E (Fig. 2.11) and is considred to be an old zone of weakness which cuts across the Bjarmeland Platform and Loppa High (Gabrielsen et al. 1990). It also separates the Mercurius High from the Maud Basin in the SW of the fault complex. According to Gabrielsen et al. (1990), the NE-SW to NS trending fault complex is characterised by normal faulting and further subdivided into three main segments. Its northern segment strikes N-S and consists of swarm of normal faults that cut the Bjarmeland Platform. The NE-SW central segment of the fault complex is related to the development of the Maud Basin and the Svalis Dome. The southern NE-SW segment of the Hoop Fault Complex comprises of narrow graben which is part of minor grabens arranged in an en echelon pattern in the northern Loppa High. The arrangement of the system defines the transition between the Hoop Fault Complex and the Bjørnøyrenna Fault Complex. 23

36 Figure Inversion structure (minor fold) at the level of the MT (Middle Triassic) on the southern segment of the Hoop Fault Complex (modified after Fiytriano 2011). MT (Mid Triassic), ET (Early Triassic), P (Permian). See figure 2.11 for location. Activity along the fault complex started in Late Carboniferous to Permian and was followed by thermal subsidence in Early-Late Permian. Growth faulting was active during EarlyMiddle Triassic and was followed by mild inversion (minor folds) due to head-on contraction (Fig. 2.16) in Middle-Late Triassic (Gabrielsen et al. 2016). Subsidence along the fault complex occurred during Early-Middle Jurassic, which was interrupted by the Late Jurassic - Early Cretaceous NW-SE rifting. The Late Cretaceous was marked by regional uplift and erosion, which was followed by Paleogene subsidence and Neogene glaciation uplift and erosion (Fiytriano 2011) Knølegga Fault Complex The Knølegga Fault Complex is part of the Hornsund Fault Complex and defines the western boundary of the Stappen High (Sundvor and Eldholm 1976, Myhre et al. 1982, Gabrielsen et 24

37 al. 1984). The fault complex strikes NNE-SSW to NS (Fig and 2.13) and is referred to as the Bjørnøya-Sørkapp fault zone by Faleide et al. (1988) and Myhre and Eldholm (1988). According to Gabrielsen et al. (1990) the Knølegga Fault Complex has a listric geometry and the main phase of movement was in Cenozoic times. The contractional structures observed by Ur-Rehman (2012) along the Knølegga Fault Complex including synclines and anticlines (Fig. 2.17) are suggested to be the result of compression in Oligocene. The intra-oligocene, intra-miocene and intra-pliocene reflectors are eroded on the anticline (Fig. 2.17). Ridge push direction changed from NW-SE to WNW-ESE at the Eocene-Oligocene boundary, as response to the adjustment of the rotation poles in the North Atlantic (Boldreel and Andersen 1993 in Vågnes et al. 1998). The contractional stresses have deformed both the hanging wall and the footwall. This contraction is suggested to be older than the PliocenePleistocene glacial sediments as no contractional structures are observed in the younger Pliocene-Pleistocene sediments along the fault complex. Figure Observed inversion structures (syncline and anticline) at the southern segment of the Knølegga Fault Complex (modified after Ur-Rehman 2012, see figure 2.11 for location). UN (Upper Neogene), IP (Intra Pliocene), IM (Intra Miocene), IO (Intra Oligocene), IE (Intra Eocene), NBE (Near Base Eocene). 25

38 Leirdjupet Fault Complex The N-S Leirdjupet Fault Complex extends from the Loppa High towards the Stappen High between 'N at 21 E (Fig. 2.11). The fault marks the transition zone that divides the Bjørnøya Basin into deep western and shallow eastern (Fingerdjupet Subbasin) parts (Rønnevik and Jacobsen 1984, Gabrielsen et al. 1990). The fault complex is believed to had been an extensional feature and was active during different time periods but the main tectonic activity took place in Carboniferous, Mid Jurassic and Early Cretaceous (Gabrielsen et al. 1990). A change in structural style along strike was avidenced by Gabrielsen et al. (1990) and the fault complex is further subdivided into three segments from south to north. The southern part (segment 3) has in general a NW-SE trend and consists of several rotated smaller normal faults. The central part (segment 2) is characterised by a single normal fault (N-S) with a large throw towards Bjørnøya Basin and the NE-SW northern part (segment 1) is composed of horst and graben structures (Fig. 2.11). The Leirdjupet Fault Complex is considered to be a northern continuation of the Bjørnøyrenna Fault Complex (Gabrielsen et al. 1990). The Leirdjupet Fault Complex experienced a phase of Early Cretaceous dextral shear and Late Cretaceous-Early Cenozoic NW-SE contraction (Gabrielsen et al. 1997, Bjørnestad 2012). Inversion structures (folds) are reported by earlier investigators (Gabrielsen et al. 1990, Bjørnestad 2012) along the central and northern parts of the fault complex. The orientation of the observed folds with axes almost parallel to the fault strike suggests head-on inversion in Late Cretaceous Early Cenozoic (Fig. 2.18). 26

39 Figure Folding along the Leirdjupet Fault Complex indicating inversion in Late Cretaceous Early Cenozoic (modified after Bjørnestad 2012). See figure 2.11 for location; BUC (Base Upper Cretaceous), EC (Early Cretaceous), BC (Base Cretaceous), UMJ (Upper Mid Jurassic) Måsøy Fault Complex The Måsøy Fault Complex is the southern marginal fault of the Nordkapp Basin and is located between N, E and N, E (Fig. 2.11). This NE-SW fault complex marks the structural division between the Nortkapp Basin and the Finnmark Platform and is considered to be an extensional structure with an en echelon pattern and mainly dip slip components (Gabrielsen et al. 1990). Tectonic activity is observed during Early Carboniferous but also recorded in Mesozoic and Cenozoic (Gabrielsen et al. 1990). The hanging wall of the major fault is reported to be severely damaged; showing minor folds, and may indicate inversion at the Jurassic-Cretaceous transition (Gabrielsen & Faerseth 1989). Glørstad-Clark et al. (2010) also suggested inversion related to (dextral) strike-slip movements during Early Middle Jurassic. 27

40 Ringvassøy Loppa Fault Complex The N-S Ringvassøy-Loppa Fault Complex is an extensional fault complex involving old zones of weakness. The northern part of the fault complex defines the western boundary of the Loppa High and in the south it merges into the Troms-Finnmark Fault Complex. It separates the Tromsø Basin in the west from the Hammerfest Basin in the east (Fig. 2.11). The southern part of this fault complex is dominated by normal faulting (Øvrebø and Talleraas 1977, Faleide et al. 1984, Gabrielsen 1984, Berglund et al. 1986). The geometry of the complex is interpreted as two levels of detached listric normal faults and a possible deeper zone of weakness (Gabrielsen 1984). The main subsidence along the southern part of the complex was from Mid Jurassic time to Early Cretaceous (Gabrielsen, 1990). This was due to large-scale rifting (Talleraas 1979). According to Braut (2012) and Zalmstra (2013) the Ringvassøy-Loppa Fault Complex was reactivated during the Late Cretaceous but Cenozoic strata have also been affected by faulting. Compression is suggested during the Late Cretaceous period and caused the formation of the observed inversion structures along the fault complex (Fig. 2.19). 28

41 Figure Inversion along the western margin of the Ringvassøy Loppa Fault Complex affecting the Cretaceous (after Zalmstra 2013). IP = Intra Permian, IC = Intra Cretaceous, BT = Base Tertiary. The summary of main tectonic events in the western Barents Sea is given in figure

42 Figure Summary of the main tectonic events in the western Barents Sea 30

43 3. Methodology 3.1 Numerical modeling A number of techniques have been applied to predict stress/strain distribution and fracture patterns (Bourne and Willemse 2001, Maerten et al. 2002, Lunn et al. 2008). In general, widely used approaches for numerical modeling are the finite element method (FEM) and discrete element method (DEM) Finite Element Method A very basic concept of the finite element method (FEM) is to divide the structural body into elements which are connected by nodes. In that way, equilibrium equations for each element can be calculated and solved simultaneously. The system of equilibrium equations can be written as: K D = F Where D is the displacement vector which contains displacements with all degrees of freedom. F is the force vector and K is rigidity matrix. FEM computations begin after having created geometry (surface bodies) and after having assigned material properties and boundary conditions. Bodies are then divided into elements and equilibrium equations are set. The K matrix is built according to material properties and elements geometries and the nodal displacements d for each element are solved. For each element, displacement fields u can be calculated by using an interpolation method u = N d. The interpolation functions in N are called shape functions. The solver then can calculate strain fields according to the constituve laws related to the selected material properties. Material parameters i.e Young s modulus (E) and Poisson s ratio (v) can be used for linear elastic materials to describe the stress-strain relation. The FEM can calculate stresses and strains for heterogeneous structures with very complex geometries ANSYS ANSYS is a multiphysics FEM program based on advanced engineering simulation technologies involving various analysis systems, e.g. fluid dynamics, structure mechanics, thermal, electromagnetics. With the purpose of calculating horizontal stress in this thesis 31

44 work, 2D linear elastic models were generated using the ANSYS Workbench. For the detailed mathematical description of the modeling approach, the reader is referred to ANSYS Mechanical APDL Theory Reference, Release 15.0, Inc., The basic workflow involved in the analyses is given in the following Analysis system A number of analyses (e.g. Electric, Fluid Flow, Rigid Dynamics, Static Structural, ThermalElectric, and Transient Structural) are available (Fig. 3.1) and can be performed by using any component system (e.g. Mechanical APDL, Fluent, Polyflow). Each analysis includes the individual components of the analysis, e.g. geometry and model properties. Static structural analysis which determines the displacements, stresses, strain and forces in structures caused by loads is adopted here. Figure 3.1. List of available analysis systems in ANSYS ANSYS Mechanical APDL Theory Reference, Release 15.0, Inc.,

45 Engineering data A very important parameter during analysis is defining the engineering data. As the response of each analysis result is determined by the assigned material properties. Depending on the application and analysis type, material properties can be linear or nor-linear, as well as temperature dependent. Linear material properties can be constant and isotropic (Fig. 3.2). A part from defined ones, user can also assign material properties (elastic, inelastic etc.) for each analysis. Figure 3.2. List of engineering data and material properties Geometry creation or attachment There are two ways to work on geometry in the ANSYS workbench, either to import an existing mesh file (.cdb) or create a new geometry using DesignModeler. It is designed to be used as a geometry editor of existing CAD models. The application allows the user to draw 2D lines, arcs and splines (Fig. 3.3) and convert into 3D models. 33

46 Figure 3.3. Example of geometry construction Coordinate systems When a model is imported into the ANSYS workbench, the coordinate system object and its sub section object called Global Coordinate System is automatically added to the working tree with default location of 0, 0, 0. For solid parts and bodies, Global Coordinate System is used by default but the user can apply a local coordinate system to any part or body Material properties After creation of geometry and coordinate systems, the user can choose predefined material for the simulation or can add new material properties. Different options including e.g. create a new material definition, import a material, edit the characteristics of the current material or assign a material from the list of available materials can be selected (Fig. 3.4). 34

47 Figure 3.4. Screen shot showing different options for assigning material properties Define connections Once the material properties of the model have been assigned, one should apply connections to the bodies in the model so that they are connected as a unit while sustaining the applied loads for the analysis. The contact defines an area where two or more bodies are in contact with each other. That connection could be e.g. bonded, frictional, frictionless, rough or no separation, according to the analysis need and required results. During the analysis, the application should prevent the two bodies from passing through each other. When the application does so, it is said to enforce contact compatibility (Fig. 3.5). Different contact formulations e.g. pure penalty, augmented Lagrange, MPC and Normal Lagrange can be used in order to enforce compatibility at the contact interface. 35

48 Figure 3.5. Schematic diagram showing penetration of contacts due to non-enforcement Meshing During meshing, the geometry is divided into elements and nodes depending on the mesh method (triangular or quadrilateral). The meshed body along with the material properties shows the mass distribution and stiffness of the structure. The analysis allows the user to define the mesh sizing and refinement along certain body contacts (Fig. 3.6). The default element size depends on different factors e.g. model size, body curvature and complexity of the geometry. Figure 3.6. Example of triangular mesh with refinement along the contact between bodies regions. 36

49 Setting up boundary conditions Boundary conditions include applied loads and support types which depend on the analysis. For static structural stress analysis, pressure and force for loads, and displacement for supports can be applied. These boundary conditions constrain or act upon the model by exerting forces, rotations or by fixing the model in such a way that it cannot be deformed (Fig. 3.7). Figure 3.7. Assumed boundary conditions and applied force Analysing results For structural analysis, equivalent stress, total displacement (Fig. 3.8) and vector principal stress can be reviewed. The results can be visualised in the form of contour, vector, probe, chart or animation. 37

50 Figure 3.8. Model results showing total displacement. 3.2 Seismic reflection survey Seismic reflection survey is based on the principle that the sound waves reflect off the interfaces between layers within the earth (Fig. 3.9). The earth is composed of different layers having different physical properties (i.e. porosity, density). When sound waves travel through the earth and found change in physical properties of layers, they penetrate into the earth or reflect back to the surface. The physical properties responsible for reflection of the waves are density and seismic velocity. The seismic data acquired in the field contains noise and has to be processed before final interpretation. The seismic data processing is basically consists of chain of operations to refine the raw data. A pre-defined program is used to eliminate noises for better data presentation. After processing the data has to be interpreted. The seismic interpretation process involves its geological expression. Seismic reflection is displayed in two way travel time (TWT) and the interpretation is basically the transformation of the data into a geological structure using time depth conversion (Dorbin and Savit 1976). 38

51 Figure 3.9. Cartoon showing basic seismic reflection methodology (geosphere inc. n.d. 12 May 2017 < Kingdom Suit 8.82 Kingdom Suite is easy to use fully integrated geoscience software which includes different modules including geophysical interpretation 2d/3dPAK, geological interpretation EarthPAK and Geosteering etc. The software package is used for seismic analysis including generation of horizons and fault on seismic lines and slices in both time and depth domains. It can also produce seismic-based interpretation maps by combined utilization of horizon and fault picking tools2. 2 The KINGDOM Suit, Seismic Micro-Technology, Inc.,

52 3.3 Structural restoration3 Deformation is assumed to neither create nor destroy rock volume; therefore reconstruction of deformed structure to its pre-deformed state is possible. Structural restoration is basically the validation of an interpreted section through back stripping of depositional and tectonic events by applying specific geometric rules. The process of restoration includes different techniques i.e. removal of faulting effects, folding associated with faulting and flexural slip and volume loss caused by compaction or erosion. According to Chamberlin (1910, 1919), the foremost use of balancing cross section was to estimate the depth to the décollement underlying concentric folds. The section can be restored back in time to place the beds into their depositional and pre-deformed position. The process can link the deformed and undeformed positions of beds and finite strain analysis can be performed which can predict fracture distribution and orientation. The outcome of the restored model can be used as a key to validate the seismic interpretation and can give a better idea to understand the geological history of the area and can enhance the quality of the work. There are certain rules for balancing which have to be followed when performing the restoration techniques. It assumes that there is in general conservation of rock volume during deformation (Hossack 1979, Goguel 1952). To build an accurate restored model of a particular area, the model must account for erosion, sediment compaction (Sanderson 1976, Wood 1974), over all tectonic compaction (Wood 1974), pressure solution (Plessmann 1964) and elongations along orogenic strike (Ramsay and Wood 1973). There are number of recently developed techniques for 2D/3D modeling of geological structures including interpolation methods based on Geostatistical approaches (Goovaerts 1997, Chilès and Delfiner 1999, Wellmann et al. 2010), interface and orientation approaches like in Calcagno et al. (2008), implicit function described in Frank et al. (2007) etc. The results of these techniques are available in the form of 2D/3D modeling software packages e.g. MOVETM, GOCAD, GeoSec, EarthVision. The main motivation of all these approaches is to generate a digital model showing subsurface structural geometry and to perform tests using different modules (according to the needs of the project). 3 Help manual of Move TM 2016 Midland Valley Exploration Ltd 40

53 3.3.1 MOVETM 2D MOVETM is mainly used for structural restoration and balancing of cross-sections. The product suite belongs to Midland Valley Exploration Ltd which is designed for complete geological structural modeling and analysis4. The suite comprises a full digital environment which provides better results in structural modeling and reduces the risk and uncertainty in geological models. It also provides a platform for integration and interpretation of data, construction of cross-section, 2D/3D kinematic modeling, 2D/3D restoration and validation D Kinematic modeling There are several modules available in MOVETM including 2D kinematic modeling, 3D kinematic modeling, goemechanical modeling, fracture modeling, stress analysis and fault analysis. The 2D kinematic module comprises of a comprehensive range of tools (e.g. block restoration, simple shear (unfolding and move on fault), flexural slip (unfolding), trishear (planar and non-planar faults), fault-parallel flow (move on fault) and fault bend folding (move on fault). It also includs sedimentation, erosion and salt movements along with decompaction, thermal subsidence and isostasy4. All these modules are based on different algorithms which allow the user to fold/unfold and fault/unfault geological models resulting in the pre-deformed tectonic setting of the study area. Structural restoration or backward modeling has mainly two goals. (1) To validate the geometry of the cross section and (2) to provide maximum information about the tectonic processes that caused any kind of deformation. The workflow used in 2D structural restoration involves application of different tools and algorithms. The details of some basic tools from the MOVE product and help manual , 4, are given in the following. 4 Product manual of Move TM 2016 Midland Valley Exploration Ltd. 41

54 D Unfolding 2D unfolding allows the user to unfold (restore) the geological beds to their pre-deformed position. The pre-deformed position could be any target horizon or an assumed regional datum. The beds can be unfolded by using either simple shear, flexure slip or line length unfolding algorithm. Simple Shear Unfolding3 The simple shear unfolding tool is used to unfold the horizons and the algorithm is best for flattening a regional dip. The only limited factor with this algorithm is that line length is not persevered (Fig. 3.11). The upper bed (blue horizon) is to be restored to a horizontal datum (Green line). Vertical shear vectors are used to restore the upper bed to target datum (3.11A). The restored geometry (3.11B) of the upper and lower beds showed that the original length of the upper bed (before restoration) is greater than the restored bed and also the original length of the lower bed (red horizon) is greater before restoration. The principle of the simple shear unfolding algorithm is that neither length nor area is preserved. Line length in the unfolding directions and surface areas of beds vary before and after deformation. Line length loss mainly depends on the dip, i.e. steep dip of restored bed causes greater line length loss. 3 Help manual of Move TM 2016 Midland Valley Exploration Ltd. 42

55 Figure Folded horizon restoration using the simple shear unfolding tool (modified from help manual of MoveTM 2016 Midland Valley Exploration Ltd.). Flexural Slip Unfolding3 The flexural slip unfolding tool can be used for concentric, layer-parallel folds. The basic principle of the flexural slip unfolding algorithm uses a pin and a slip-system parallel to the template bed to control the unfolding. The procedure involves the rotation of limbs of a fold to a datum. The effects of flexural slip component are then removed by applying the layer parallel shear to the rotated fold limbs. Unlike the simple shear unfolding algorithm, this tool allows the user to maintain bed thickness between the template horizon and other passive objects. The algorithm is built to maintain the line length of the template horizon in the direction of unfolding and maintain the area of the fold and the model. It can be used to validate complex thrust deformations. Sequence of cross-sections showing how the flexural slip unfolding works can be seen in figure Help manual of Move TM 2016 Midland Valley Exploration Ltd. 43

56 Figure Basic workflow for flexural slip unfolding. A) Folds with thickness variations. B) Construction of a slip system parallel to the template bed using dip domain bisector. C) Unfolded/restored template bed and passive beds about the pin. (modified from help manual of MoveTM 2016 Midland Valley Exploration Ltd.) D Move-on-Fault3 The move on fault module can be used for both forward models and to restore deformation. It gives guidance and new ideas for seismic structural interpretation and allows the user to model pre- and syn- tectonic successions and different displacements. The module is capable to test and modify fault geometries within the stratigraphic framework and can directly create balanced interpretation. There are different methods (simple shear, fault parallel flow, fault bend fold, fault propagation, trishear, detachment fold) available within the capacity of the module and can be used to control the deformation along the fault. Most of the algorithm tools can be used for both compressional and extensional models, using either positive or negative displacement3. 3 Help manual of Move TM 2016 Midland Valley Exploration Ltd. 44

57 Simple Shear3 The simple shear algorithm is used to model the relationship between the hanging wall deformational features and fault geometry. It extends the deformation along the hanging wall rather than generating discrete slip between beds (i.e. flexural slip) 3. The hanging wall is folded and maintains bed area during modeling. The tool is useful in extensional systems e.g. half graben or rollover anticline developed on nonplanar normal fault. The algorithm can also be used for forward modeling of restoration of growth faults and inverted basins. According to Yamada and McClay (2003), the simple shear algorithm in MOVETM can only be used to restore inverted basins and proposed to use a shear angle of 32 with respect to vertical D Decompaction3 The 2 D decompaction module allows for modeling change in rock volume due to porosity loss as a function of depth. The algorithm can also be used for isostatic effects and burial history of any area. There are different methods within the range of the module to calculate porosity change with depth, e.g. Sclater and Christie (1980), Baldwin and Butler (1985), and Dickinson (1953). According to Sclater and Christie (1980), porosity decreases with increasing depth (compaction) and increases with decreasing depth (decompaction) and can be expressed as: f = f0 (e -cy) where: 3 f is present-day porosity f0 is porosity at the surface c is the porosity-depth coefficient (km-1) y is depth Help manual of Move TM 2016 Midland Valley Exploration Ltd. 45

58 4. Cretaceous inversion of the western Barents Shelf: integrated seismic interpretation of the Bjørnøyrenna and the Ringvassøy Loppa fault complexes 4.1 Introduction The Barents Sea represents a large part of the Arctic, it is located in the northwestern corner of the Eurasian plate and bounded by the Norwegian Greenland Sea to the west, Svalbard and Franz Josef Land to the north, Novaya Zemlya to the east and the Norwegian Russian mainland in the south (Fig. 4.1). It covers a tectonically extended shelf which consists of basins, highs and fault complexes (Gabrielsen et al. 1990, Indrevær et al. 2016). Most of these fault complexes strike NE-SW and N-S in the eastern and central parts of the western Barents Sea and developed as a result of multiple extensional events (throughout the Carboniferous to Eocene) after the collapse of the Caledonian orogen (Faleide et al. 1984, 1993, 2008, Gabrielsen et al. 1990, Gudlaugsson et al. 1998). The rifting episodes terminated with the opening of the North Atlantic and Arctic oceans and the final stages of rifting characterized by the transition from a rift system in the south to a dextral transform one connecting the North Atlantic rift to the Arctic rift (Faleide et al. 2008, Indrevær et al. 2016). In addition to several phases of extension, a number of authors have reported late Paleozoic, Mesozoic and Cenozoic events of inversion related to strike-slip movements and head-on compression which affected most of the fault complexes in the western Barents Sea (Ziegler 1978, Rønnevik et al. 1982, Riis et al. 1986, Berglund et al. 1986, Sund et al. 1986, Brekke & Riis 1987, Wood et al. 1989, Gabrielsen & Færseth 1989, Gabrielsen et al. 1990, 1997, 2011, Vågnes et al. 1998, Grogan et al. 1999, Henriksen et al. 2011, Glørstad-Clark et al. 2011, Faleide et al. 2015, Indrevær et al. 2016). The different phases of inversion included 1) NW-SE directed head-on inversion in the late Cretaceous Palaeocene caused by far field stresses (Gabrielsen et al. 1997, Vågnes et al. 1998), 2) dextral shearing due to Early Cenozoic (earliest Eocene) rifting and breakup (Faleide et al. 1996,) and 3) NW-SE contraction related to ridge push in Miocene (Gabrielsen & Færseth 1989, Engen et al. 2008, Faleide et al. 2015, Gac et al. 2016). The current research focused on the Early and Late Cretaceous inversion events which affected the Bjørnøyrenna Fault Complex and the Ringvassøy-Loppa Fault Complex (Gabrielsen et al. 1990, 1997, Vågnes et al. 1998, Indrevær et al. 2016). Although Cretaceous inversion has long been discussed, the exact timing and main source of driving force (s) are 46

59 not fully constrained. In the present study, inversion structures of the Early and Late Cretaceous are interpreted along with the timing and mechanisms of development of these structures Geological setting The Phanerozoic evolution of the Barents Sea involves series of orogenic events, subsequent collapses and rifting (Gabrielsen et al. 1990, Worsley 2008, Henriksen et al. 2011, Gernigon et al. 2014). The major tectonic phases responsible for the development of the geological framework of the Barents Sea include the Timanian, Caledonian, and Uralian orogenies (Gee et al. 2008) from Late Proterozoic to Late Paleozoic (Doré 1991) and are followed by protoatlantic rifting events during Mesozoic (Smelror et al. 2009, Tsikalas et al. 2012) and opening of the North Atlantic Ocean along the western margin of the shelf during Cenozoic (Gabrielsen et al. 1990, Faleide et al. 2008). The Barents Shelf is generally subdivided into two major geological provinces. The eastern province was mainly affected by tectonics pertaining to Novaya Zemlya, the Timan-Pechora Basin and the Uralian Orogeny (Worsley 2008), while the western province was mainly controlled by major post-caledonian rifting episodes (Fig. 4.1). The Late Paleozoic structures of the western Barents Sea reflect mainly WNW-ESE and N-S to NE-SW structural grains inherited from the Timanian (i.e. Ediacaran) and the Caledonian orogenies respectively (Ritzmann and Faleide 2007, Gernigon and Brönner 2012, Gernigon et al. 2014). 47

60 Figure 4.1. Barents Sea major structural features including major faults, basins, structural highs and platform areas. BF, Baidsratsky Fault Zone; BFC, Bjørnøyrenna Fault Complex; BFZ, Billefjorden Fault Zone; HFZ, Hornsund Fault Zone; KFZ, Knølegga Fault Zone; KHFZ, Kongsfjorden Hansbreen Fault Zone; LFC, Leirdjupet Fault Complex; MFC, Masøy Fault Complex; RLFC, Ringvassøy Loppa Fault Complex; SJZ, Senja Fracture Zone; SKZ, Sørkapp Fault Zone; SRFZ, Sredni Rybachi Fault Zone; TIFC, Thor Iversen Fault Complex; TFFC, Troms Finnmark Fault Complex; TKFZ, Trollfjorden Komagelva Fault Zone (modified from Marello et al. 2013). With respect to present-day geography, the Caledonian orogeny was characterized by east- to southeast-directed convergent movement from Late Cambrian to Early Devonian (i.e. ~410 Ma) and involved both the closure of the Iapetus Ocean and the collision between Laurentia and Baltica in the region under scope (Roberts 2003, Gee et al. 2006, Ritzmann and Faleide 2007, Gasser 2014, Gernigon et al. 2014). The post-caledonian geological history of the 48

61 western Barents Sea was dominated by large-scale regional sinistral shear promoting both transtension and transpression in Late Devonian Early Carboniferous (Faleide et al. 1984). Late Paleozoic tectonics resulted in the formation of three major sedimentary basins located in the east of the studied region (Fig. 4.2): the ENE-WSW Nordkapp Basin, the NE-SW Ottar Basin, located between the Loppa High and the Norsel High, and the NE-SW Maud Basin located between the Loppa High and the Mercurius High (Jensen and Sørensen 1992). In Early Triassic, the western Barents Sea was affected by rifting. This rifting phase is recorded in many parts of the Arctic and North Atlantic regions (Tsikalas et al. 2012, Gernigon et al. 2014). The Early Triassic extension continued until late Anisian-early Ladinian and was characterized by normal faulting, tilting of fault blocks and erosion. According to Gudlaugsson et al. (1998), Early Triassic faulting occurred mainly along N-S striking structures in the western Barents Shelf. This was followed by Middle to Late Triassic (post-rift) thermal subsidence in the North Atlantic and Arctic basins (Gernigon et al. 2014). However, during the corresponding period of time the Barents Shelf was subject to progressive uplift of its northern, eastern, and southern regions (Worsley 2008). From Late Triassic to Early Jurassic, Late Paleozoic structural elements of the Barents Sea were inverted following an E-W compressive regime (Otto and Bailey 1995, Buiter and Torsvik 2007). Westward motion of Siberia has been advanced as a potential cause for this compressional event (Buiter and Torsvik 2007). Otto and Bailey (1995) also interpreted inversion structures on the eastern margin of the South Barents Sea Basin dated to Late Triassic-Early Jurassic. Some other investigators (e.g. Gabrielsen et al. 1990, Gudlaugsson et al. 1998, Fitryanto 2011, Gernigon et al and references therein, Gabrielsen et al. 2016) confirmed reactivation of major fault complexes (e.g. Troms-Finnmark, Måsøy, Thor Iversen and Hoop fault complexes) in the western Barents Sea during the same period of time. The Middle Jurassic to Early Cretaceous regional rifting affected mainly the western Barents Shelf (Faleide et al. 2008). The Jurassic rifting was responsible for the development of major deep basins, e.g. Hammerfest and Tromsø basins and fault complexes, e.g. RingvassøyLoppa, Bjørnøyrenna, Leirdjupt and Asterias fault complexes (Faleide et al. 1993a, b). The N-S-striking Ringvassøy-Loppa Fault Complex and its southern extension coincide with the transition zone between the Hammerfest and Tromsø basins (Øvrebø and Talleraas 1977, Gabrielsen et al. 1990) and merges with the Troms-Finnmark Fault Complex to the south. To the north, the fault complex defines the western boundary of the Loppa High and defines the 49

62 southeastern margin of the Bjørnøyrenna Fault Complex (Fig. 4.2). The main subsidence along the southern segment of the Ringvassøy-Loppa Fault Complex started in Mid Jurassic and culminated in Aptian to Albian (Gabrielsen et al. 1990, Gudlaugsson et al. 1998). Figure 4.2. Regional setting and major structural elements including major faults, basins, structural highs and platform areas of the study area (modified from and google map and NPD fact maps The NE SW Bjørnøyrenna Fault Complex is the northern continuation of the Ringvassøy Loppa Fault Complex and defines the western margin of the Loppa High, separating it from the Bjørnøya Basin (Fig. 4.2). Rønnevik and Jacobsen (1984) described the fault complex as the south-eastern boundary fault of the Bjørnøya Basin whereas Gabrielsen et al. (1984) defined it as the north-eastern extension of the Ringvassøy - Loppa Fault Complex. In general the fault complex defines the boundary between the Loppa High and the Bjørnøya Basin in the southwest and in the northeast it separates the Loppa High from the Fingerdjupet Subbasin (Rønnevik et al. 1975, Hinz and Schlüter 1978, Rønnevik et al. 1982). The area was affected 50

63 by tectonic inversion and presence of folds, reverse faults point to NW-SE directed Late Cretaceous to Early Cenozoic contraction (Gabrielsen et al. 1997). During and after Early Cenozoic rifting and breakup in earliest Eocene, the western margin of the Barents Sea was subject to tectonic dextral shear and associated folding with NW-SE striking fold axes (Faleide et al. 1993a). Cenozoic inversion in the western Barents Sea is assumed to be caused by North Atlantic ridge push causing post Miocene shortening and initiating folds with NE-SW-striking fold axes (Gabrielsen et al. 1990). Neogene uplift and glaciations resulted in deep erosion of the western Barents Shelf (Faleide et al. 1996). 4.3 Data and methodology The data used in the present study includes 142 2D seismic lines with different orientations (N-S, E-W, NE-SW, NW-SE, Fig. 4.3). These seismic lines belong to different surveys (BJRE, BJSY, NBR06, NBR07, NBR08, NBR10, TTR74R1 and TTR83R1) and are of variable quality and variable depth resolution. The data were provided by the Department of Geosciences, University of Oslo, Norway. A total 162 wells have been drilled in the Barents Sea (NPD fact page, and most of them are located in the Hammerfest basin (Fig. 4.3). Well control was provided by ten exploration wells on the Loppa High (wells 7120/1-1 and 7120/2-1), the Hammerfest Basin (wells 7120/7-2, 7120/8-1, 7120/8-2 and 7120/8-4) and along the eastern margin of the Bjørnøya Basin (wells 7219/8-1, 7219/9-1, 7220/5-1 and 7220/6-1). 51

64 Figure 4.3. Base map of the study area showing the respective locations of 2D seismic lines and wells. Red lines and dots indicate key profiles and exploration wells respectively. IHS TM Kingdom 8.8 (seismic and geological interpretation software) is used for current work. Kingdom Suit1 is easy to use fully integrated geoscience software which involves different modules including geophysical interpretation 2d/3dPAK, geological interpretation Earth PAK and Geosteering. 4.4 Seismic interpretation The first step while doing the seismic interpretation across the entire data sets (BJRE, BJSY, NBR06, NBR07, NBR08, NBR10, TTR74R1 and TTR83R1) was to choose key reflectors of different ages. Eight reflectors were chosen because of their continuity, prominence and geological importance. The seismic tie to well 7219/9-1, located in the eastern margin of the Bjørnøya Basin is shown in figure 4.4. The next step is to extend these reflectors along the major fault complexes (i.e. Bjørnøyrenna Fault Complex and Ringvassøy Loppa Fault Complex). It was a particularly challenging task to interpret the reflectors in all parts of the study area due to intense faulting. A detailed description of interpreted reflectors and stratigraphy is given in the following. 1 The KINGDOM Suit, Seismic Micro-Technology, Inc.,

65 Figure 4.4. Seismic tie to well 7219/9-1, located in the eastern margin of the Bjørnøya Basin (Data courtesy of TGS and Spectrum). 53

66 4.4.1 Snadd Formation (Upper Triassic) The deepest reflector interpreted in the study area is the top Snadd Formation (Upper Triassic). The Snadd Formation is of Ladinian to early Norian age and composed of basal grey shales which coarsen up into shales with interbeds of grey siltstones and sandstones. The lower and middle part of the formation consists of limestone and calcareous interbeds, while thin coaly lenses are also developed in the upper part (Dalland et al. 1988). The reflector exhibit strong amplitude and has been interpreted between ms TWT (Fig. 4.4) Fruholmen Formation (Base Jurassic) The base Jurassic reflector is interpreted as the top of the Fruholmen Formation between ms TWT in the study area (Fig. 4.4). The formation is part of the Realgrunnen Subgroup. The dominant lithologies of the group are sandstones, shales and coals of Late Triassic to Mid Jurassic (early Norian to Bajocian, Dallmann 1999). The sandstones were deposited in coastal plain and deltaic through shallow marine environments (Worsley et al. 1988). The base of the formation belongs to the Late Triassic (early Norian) and the top corresponds to the Triassic/Jurassic transition. The formation consists of basal grey to dark grey shales which gradually pass upwards into interbedded sandstones, shales and coals. The lowermost Akkar Member (Norian) consists of open marine shales which pass up into a coastal and fluvial sandstone dominated sequence of the middle Reke Member (Norian -? Rhaetian). Marine shales dominate the uppermost Krabbe Member (Rhaetian, Dalland et al. 1988, Larssen et al. 2002) Tubåen Formation (Lower Jurassic) The Lower Jurassic reflector is interpreted as top of the Tubåen Formation of the Realgrunnen Subgroup which consists of various lithologies including sandstones and subordinate shales with minor coals. The upper and lower part of the formation is dominated by a sand-rich unit whereas a shaly interval exists in the middle part. The sand unit is believed to represent stacked series of high energy marginal marine deposits (tidal inlet dominated, barrier complex and/or estuarine) while marine shale characterizes more distal environments. Coal was deposited in protected back-barrier lagoonal environments to the south-east (Dalland et al. 1988). The age of the formation is Late Triassic to Early Jurassic (late Rhaetian to early Hettangian) and its top is found between 1790 and 2000 ms TWT (Fig. 4.4). 54

67 4.4.4 Nordmela Formation (Upper Jurassic) The third formation of the Realgrunnen subgroup is the Nordmela Formation. In the present study, the Nordmela Formation is interpreted as an Upper Jurassic reflector between 1750 and 2000 ms TWT (Fig. 4.4). The formation is relatively thick when measured in the reference well (7119/12-2; 71 00'51.81"N, 19 58'20.81"E) but in contrast the thickness is only 62 m in the type well (Fig. 4.3, 7121/5-1; 71 35'54.88"N, 21 24'21.78"E). The lateral thickness variation between type and reference well suggests a southwest thickening wedge evidencing early Kimmeridgian subsidence of the Ringvassøy-Loppa Fault Complex. The upper part of the formation is mainly composed of sandstones but in general the formation consists of interbedded siltstones, sandstones, shales and claystones with modest amounts of coal. The variation in lithologies shows that the formation was deposited in tidal flat to flood plain environments. Individual sandstone sequences represent estuarine and tidal channels that dissected this low-lying area (Dalland et al. 1988). The Nordmela Formation is Early Jurassic to early Mid Jurassic in age (Sinemurian - late Pliensbachian to Aalenian) Fuglen Formation (upper Middle Jurassic) The upper Mid Jurassic reflector corresponds to the Fuglen Formation between ms TWT on seismic section (Fig. 4.4). The formation belongs to the Adventdalen Group which includes sediments from Middle Jurassic (Bathonian) to Lower Cretaceous (Cenomanian). The group is further subdivided into the Hekkingen, Knurr, Kolje and Kolmule formations. The thickness of the group varies from m on Svalbard to m on the Barents Sea Shelf. The dominant lithologies of the group are dark marine mudstones with some deltaic and shelf sandstones along with carbonates of Late Jurassic to Early Cretaceous in age. The group contains major hydrocarbon source rocks (Fuglen and Hekkingen formationa) in the Upper Jurassic successions (Larssen et al. 2002). The Fuglen Formation is composed of pyritic mudstone with interbedded thin limestones and was deposited in marine environments during a high stand system tract. A relatively thick (48 m) unit of the formation was deposited in the western parts of the Hammerfest Basin (7119/12-1). It thins to less than 10 m on the central highs of the basin. The age of the formation is Late Callovian to Oxfordian (Dallmann 1999, Dalland et al. 1988). 55

68 4.4.6 Hekkingen Formation (Base Cretaceous) The top of the Hekkingen Formation is represented in the study area by a reflector between 1680 and 1780 ms TWT (Fig. 5.4). The Hekkingen Formation is relatively thick (359 m) in the type well (7120/12-1) but its thickness decreases down to 113 m in the reference well (7119/12-1) showing northwards thinning to the Hammerfest Basin. The depositional pattern shows the development of semi-grabens along basin margins. The formation consists of brown-grey to very dark grey shales and claystones with thin interbeds of limestones, dolomite, siltstones and sandstones (Larssen et al. 2002). The suggested age of the formation is late Oxfordian/early Kimmeridgian to Ryazanian. The Hekkingen Formation is further subdivided into the Alge and Krill members. The lower Alge Member consists of black shales rich in organic material and the base of the member is defined by a transition from carbonate cemented and pyritic mudstone to poorly consolidated shales deposited in restricted shelf environments. The age of the Alge Member is late Oxfordian - Kimmeridgian. The upper Krill Member of the Hekkingen Formation consists of brown-grey to very dark grey shales and mudstone with some thin beds of limestone, dolomite, siltstone and sandstone. The overall lithology of the Krill Member represents an open to restricted shelf environment. The age of the member is Kimmeridgian Volgian (Dalland et al. 1988) Knurr Formation (Lower Cretaceous) The top of the Knurr Formation is interpreted between 1600 and 1650 ms TWT in the study area (Fig. 4.4). Based on dinoflagellates and foraminifera the age of the formation is Ryazanian / Valanginian to early Barremian. The formation consists of dark grey to greybrown claystone with thin limestone and dolomites interbeds. A small amount of sandstone is also present in the lower part of the unit but disappears laterally into the Hammerfest Basin. The upper part of the formation consists of red to yellow brown claystones. The lithology of the formation shows that the sediments were deposited in open and generally distal marine environments with local restricted bottom conditions (Dalland et al. 1988). The thickness of the formation recorded in the type well (7119/12-1) is 56 m and 285 m in the reference well (7120/12-2). 56

69 4.4.8 Kolmule Formation (Base Upper Cretaceous) The formation is mainly composed of dark grey to green claystone and shale. Limestone and dolomite stringers along with minor thin siltstone beds are also present (Dalland et al. 1988). Traces of glauconite and pyrite also occur in some places reflecting open marine environments. The age of the formation is Aptian to mid-cenomanian (Dalland et al. 1988). 4.5 Results and discussions Early Cretaceous inversion The study area was influenced by tectonic inversion (e.g. dextral strike slip) during Early Cretaceous (Gabrielsen et al. 1997). The Bjørnøyrenna Fault Complex strikes mainly NE-SW and is basically extensional, but the master fault plane is occasionally over-steepened. NNESSW-striking structural highs (positive flower structures?) in the Bjørnøyrenna Fault Complex affecting the assumed Lower Cretaceous (Hauterivian-Aptian?) sequence have been interpreted to represent an event of dextral wrenching at that time (Riis et al. 1986, Brekke and Riis 1987, Gabrielsen and Færseth 1988, Gabrielsen et al. 1992, 1997, Faleide et al. 1993a, 1993b, Vågnes et al. 1998). In the present study, a lens-shaped structure affecting the Lower Cretaceous sediments is detected in the central segment of the Bjørnøyrenna Fault Complex and interpreted as a positive inversion one. The interpreted seismic section (Fig. 4.5), located at the central segment of the Bjørnøyrenna Fault Complex (Fig. 4.2), depicts an inverted structure (red square) affecting the EC (Early Cretaceous; Valanginian to early Barremian in age) and the BUC (Base Upper Cretaceous; Aptian to Cenomanian). Both horizons have been subsequently compressed and uplifted. The fault retains extension at a deeper level and experienced net contraction associated with an anticline in the upper portion. 57

70 Figure 4.5. Interpreted inversion structure at the Early Cretaceous (Top Knurr Formation) level. See figure 4.2 for location (Data courtesy of TGS and Spectrum). The northernmost part of the central segment of the Bjørnøyrenna Fault Complex also shows inversion structures affecting the Lower Cretaceous and Late Cretaceous sediments. Interpreted seismic section (Fig. 4.6) showing inversion structure affecting LC (Lower Cretaceous; Valanginian to early Barremian) and the BUC (Base Upper Cretaceous; Aptian to Cenomanian) sediments on the northern segment of the Bjørnøyrenna Fault Complex. Small wavy undulations can be imaged between the Base Cretaceous and the Lower Cretaceous reflectors. The anticline imaged at the Lower Cretaceous level as well as the slight bulge on the BUC (Base Upper Cretaceous) indicates Early and Late Cretaceous inversion phases. Such positive structures are formed when extensional faults reverse their sense of motion during compressional tectonics causing the basin to turn inside out and to become a positive feature (Williams et al. 1989). 58

71 Figure 4.6. Interpreted inversion structure at the Early Cretaceous (Top Knurr Formation) and Base Upper Cretaceous (Top Kolmule Formation) levels. See figure 4.2 for location (Data courtesy of TGS and Spectrum). Gabrielsen et al. (1997) and Hameed (2012) advance dextral strike-slip movements coeval to Early Cretaceous inversion (Fig. 4.7 and 4.8). The contractional structures are overlain by a set of extensional faults which affect the upper part of the section. According to Gabrielsen et al. (1984, 1997), signs of tectonic inversion in the fault complex, including deformed fault planes and reverse faults were dated to Cretaceous and Cenozoic. These structures are developed above over-steepened fault branches in the hanging wall and are associated with mild inversion (Fig. 4.7). The downward-steepening faults define structures resembling positive (half) flower structures developed due to dextral strike-slip movement (Gabrielsen et al. 1997). 59

72 ESE s WNW twt 1 BUC ILC rn the r No of nt e m seg na ren y ø ør Bj x ple m o lt C u Fa 2 km Figure 4.7. Seismic line (NH ) cutting the central segment of the Bjørnøyrenna Fault Complex and depicting minor fold trains associated with assumed thrust faults. See Figure 4.2 for location. ILC (Intra Lower Cretaceous), BUC, (Base Upper Cretaceous). Modified after Gabrielsen et al Hameed (2012) reported inversion structures in the central segment of the Bjørnøyrenna Fault Complex. Based on the strike of the axis of the imaged fold, which is oblique to the strike of the master fault, he interpreted the inversion to be caused by dextral strike-slip movement (Fig. 4.8). Indrevær et al. (2016) also interpreted inversion structures of early Barremian to mid-albian age (ca Ma) along the margins of the Loppa High and concluded that these inversion structures developed due to uplift of the Loppa High along its inclined boundary fault (e.g. Bjørnøyrenna Fault Complex). 60

73 Figure 4.8. Interpreted seismic section showing inversion structure (minor fold) at Lower Cretaceous level in the central segment of the Bjørnøyrenna Fault Complex. See figure 4.2 for location (modified after Hameed 2012). The southern part of the Ringvassøy-Loppa Fault Complex is dominated by normal faulting (Øvrebø and Talleraas 1977, Gabrielsen 1984, Faleide et al. 1984, Berglund et al. 1986). The geometry of the complex is interpreted as two levels of detached listric normal faults and a possible deeper zone of weakness (Gabrielsen 1984). Braut (2012) and Zalmstra (2013) proposed that the two inversion events of the Ringvassøy-Loppa Fault Complex, during Cretaceous (i.e. Early and Late Cretaceous), are of regional significance. The present study also evidenced inversion in the eastern margin of the Ringvassøy-Loppa Fault Complex, where a snake-head like structure affecting the Lower Cretaceous (Valanginian to early Barremian) sediments in the hanging wall is observed and interpreted as an inversion structure (Fig. 4.9). The interpreted seismic line (Fig. 4.9) shows inversion structures in the eastern margin of the Ringvassøy Loppa Fault Complex at lower and Late Cretaceous levels. Folding at the Lower Cretaceous (Valanginian to early Barremian) and 61

74 Base Upper Cretaceous (Aptian to Cenomanian) sediments favors inversion in the Ringvassøy Loppa Fault Complex. Figure 4.9. Interpreted inversion structures at the Early and Late Cretaceous level along the Ringvassøy Loppa Fault Complex. See figure 4.2 for location (Data courtesy of TGS and Spectrum) Late Cretaceous inversion The study area was also affected by Late Cretaceous tectonic inversion (Gabrielsen et al. 1997). In the present study, an open fold affecting the Base Upper Cretaceous (Aptian to Cenomanian) reflector is observed in the hanging wall of the central segment of the 62

75 Bjørnøyrenna Fault Complex (Fig. 4.5 and 4.6) and interpreted as an inversion structure. It is assumed that the tilted and eroded Late Cretaceous horizon in the footwall was also involved in folding or reverse faulting (Fig. 4.5). Inversion is also proposed for the eastern margin of the Ringvassøy-Loppa Fault Complex at the base of Upper Cretaceous (Aptian to Cenomanian) horizon. A Fold is detected affecting the Base Upper Cretaceous horizon and is interpreted as an inversion structure (Fig. 4.9). It is suggested that the tectonic event that affected the Bjørnøyrenna Fault Complex during Late Cretaceous is also responsible for the inversion of the Ringvassøy Loppa Fault Complex. Minor fold trains associated with assumed thrust faults at the base of the Upper Cretaceous in the central segment of the Bjørnøyrenna Fault Complex (Fig.4.7) were previously interpreted by Gabrielsen et al. (1997) who suggested that the compressional event started to affect the depositional geometry after the establishment of an intra-cenomanian reflector (BUC; Kolmule Formation) and continued throughout the Late Cretaceous and into the Early Cenozoic. This compressional event caused inversion of local depocentres and development of folds and reverse faults along the Bjørnøyrenna Fault Complex. The local erosion of Late Cretaceous deposits is due to the thrusting which also affected the depositional pattern in the area (Gabrielsen et al. 1997). The inversion features striking parallel to the fault complex suggest that the Late Cretaceous inversion resulted from compression perpendicular to the strike of the Bjørnøyrenna Fault Complex, i.e. SE-directed and presumably related to far field stresses. Riis et al. (1986) identified E W-trending fold axes affecting the Cenomanian along the southeastern margin of the Bjørnøya Basin and suggested that these structures were related to wrenching. Braut (2012) suggested inversion, based on the interpretation of folds in the hanging wall block of the central segment of the Ringvassøy Loppa Fault Complex where the base of the Upper Cretaceous is affected by compression (Fig. 4.10). 63

76 Figure Interpreted seismic section showing inversion structures (minor folds) at BUC (Base Upper Cretaceous) level in the central segment of the Ringvassøy Loppa Fault Complex. See Figure 4.2 for location (modified after Braut 2012). 64

77 4.6 Conclusions Kingdom 8.8 was used for the current research and 142 2D seismic lines were interpreted. The main aim was to identify and interpret Cretaceous inversion structures in the Bjørnøyrenna and the Ringvassøy Loppa fault complexes. The results of the present study suggest two different episodes of inversion which affected the fault complexes. The Early Cretaceous inversion phase is related to strike-slip (dextral?) movement and the Late Cretaceous relates to NW-SE far field stresses. 65

78 5. Structural restoration of Cretaceous inversion events in the Bjørnøyrenna Fault Complex, western Barents Shelf. 5.1 Introduction The Barents Sea consists of a large epicontinental sea bounded by young passive continental margins in the north and west (Faleide et al. 1984) and covers an area of approximately 1.4 million km2. It is bounded by Svalbard archipelago in the north and the Norwegian and Russian coasts in the south. The Norwegian Greenland Sea lies to the west and Novaya Zemlya forms the eastern boundary of the Barents Sea (Fig. 5.1). The Barents Sea contains some of the deepest basins in the world which developed due to different regional tectonic events from Paleozoic to Cenozoic within the North Atlantic Artic region (Faleide et al. 1993a). Its western part constitutes the northern Norwegian continental shelf and is located between the mainland Norway and Svalbard, informally termed as southwestern Barents Sea (Gabrielsen et al. 1997). Most of the fault complexes in the southwestern Barents Sea have an overall NE-SW to ENE-WSW structural trend in its eastern and central parts, whereas the western part of the area consists of NNW-SSE to N-S fault complexes (Gabrielsen et al. 1990, 1997; Fig. 5.1). These fault complexes formed from late Proterozoic to Cenozoic in response to several rifting and collision events and indication of inversion structures were reported by earlier investigators (Rønnevik and Motland 1979, Rønnevik et al. 1982, Faleide et al. 1984, Rønnevik and Jacobsen 1984, Faleide et al. 1993a, Gabrielsen et al. 1997, Vågnes et al. 1998). The development of inversion structures has been suggested to be related to strike-slip tectonics (Riis et al. 1986, Brekke and Riis 1987, Gabrielsen et al. 2016) or head-on inversion (Gabrielsen et al. 1992, 1997). The aim of present study is to identify and decipher eventual Cretaceous inversion structures in the Bjørnøyrenna Fault Complex by means of structural restoration. To these aims, 2D MOVETM, a structural modeling and analysis software by Midland Valley Exploration Ltd, is used and three key seismic lines crossing the central and northern segments of the Bjørnøyrenna Fault Complex are restored. 66

79 5.2 Geometry and structural evolution of the Bjørnøyrenna Fault Complex The NE-SW striking Bjørnøyrenna Fault Complex (Gabrielsen et al. 1990) is located in the western Barents Sea between 72º N, 19 E and 73º 15' N, 22º E (Fig. 5.1) and is considered to be the northern continuation of the Ringvassøy-Loppa Fault Complex (Gabrielsen et al. 1984). The fault complex marks the boundary between the platform-like Loppa High to the southeast and deep Cretaceous basins to the northwest (Hinz and Schlüter 1978 in Gabrielsen et al. 1997). Figure 5.1. Regional setting and major structural elements of the study area (modified from Google Maps and NPD fact maps The Bjørnøyrenna Fault Complex is an extensional structural feature (Gabrielsen et al. 1990, 1997) which lies over a crustal zone of weakness (Gabrielsen et al. 1984) and displays complex geometries due to multiple phases of deformation, including reactivation (Grunnaleite 1991) and inversion phases (Gabrielsen et al. 1997). Based on fault geometry and structural trend, Gabrielsen et al. (1997) subdivided the fault complex into four major 67

80 segments (Fig. 5.2). The southern segment (NNE-SSW and NE-SW), central segment (NESW) and northern segment (NNE-SSW) are separated from the northwestern shoulder of the Loppa High (segment 4), which consists of narrow horsts and grabens (Gabrielsen et al. 1997). According to Gabrielsen et al. 1992, these grabens have flower-like structures (segment 4; Fig. 5.2). Figure 5.2. Subdivision of the Bjørnøyrenna Fault Complex based on structural trend and geometry. Modified after Gabrielsen et al. (1997). Regional extension and minor strike-slip adjustments along old lineaments in the western Barents Sea started again in Middle Late Jurassic (Faleide et al. 1993a). At regional scale, the Jurassic extensional structures of the Barents Sea belonged to the larger Arctic-North Atlantic rift system (Doré 1991). The Middle Late Jurassic extensional phase created the Hammerfest and Bjørnøya basins following pre-existing structures and caused block faulting (Faleide et al. 1993a). This was also coeval with renewed subsidence in the Tromsø and Bjørnøya basins (Faleide et al. 1993a). The Bjørnøyrenna Fault Complex is of extensional origin (Gabrielsen et al. 1990, 1997) and presents listric fault geometries (Faleide et al. 68

81 1993a). The main phase of subsidence along the fault complex started in Callovian (Gabrielsen et al. 1990, Faleide et al. 1993a). The more complex structuration occurred in the western Barents Sea at the end of Jurassic due to the development of regional scale extensional fault blocks and the influence of shear movements in parts of the North Atlantic (Håkansson and Stemmerik 1984, Riis et al. 1986). In Early Cretaceous, subsidence went on in the western Barents Sea and, in particular, rapid subsidence occurred along the Bjørnøyrenna and Ringvassøy-Loppa Fault complexes (Gabrielsen et al. 1997). Faleide et al. (1993a) described three tectonic rifting phases affecting the major basins of the western Barents Sea (e.g. Hammerfest, Tromsø and Bjørnøya Basin) during Early Cretaceous. The first two phases (Berriasian/Valanginian and Hauterivian/Barremian) strongly affected the Tromsø and Bjørnøya basins. The last phase included thermal subsidence in the Tromsø Basin which is evidenced by gradually increased thickness of the Barremian sediments (i.e. Kolje Formation) westwards in the Hammerfest Basin and into the Ringvassøy-Loppa Fault Complex. In the western Hammerfest Basin, uplift and thinning of the Aptian sequences (i.e. Kolmule Formation) towards the RingvassøyLoppa Fault Complex evidence further the occurrence of an Aptian tectonic event in the area (Faleide et al. 1993a). In general, the development of Early Cretaceous structures in the western Barents Sea were coeval with the opening of the Amerasian Basin and the North Atlantic rifting and mainly characterized by extensional faults with large downthrow to the west with minor wrench component. Faleide et al. (1993a) described sinistral transtensional strike-slip along the Bjørnøyrenna Fault Complex which caused formation of the Senja Ridge and the Veslemøy High as positive structural elements. Riis et al. (1986) also suggested sinistral shear which caused compressional faulting and folding in the Senja Ridge during Early Cretaceous. Gabrielsen et al. (1997) suggested that the local thinning of the Hauterivian Aptian sequence along the Bjørnøyrenna Fault Complex is associated with mild inversion of local depocentres and reverse faults associated with minor hanging wall folds. Gabrielsen and Færseth (1988) also suggested dextral transpressional strike-slip movement during Early Cretaceous. According to Gabrielsen et al. (1992) the fault pattern in the central segment of the Bjørnøyrenna Fault Complex are similar to positive (semi)-flower structures. Indication of local Early Cretaceous inversion is also observed along the Ringvassøy Fault Complex and its junction with the Asterias Fault Complex (Gabrielsen et al. 1990). During Late Cretaceous, opening of the Labrador Sea started and regional subsidence occurred along the North Atlantic rift basins. These deep and broad basins ended at the Dee 69

82 Geer Zone where they were prolonged by pull-apart basins as a response to dextral oblique slip (Faleide et al. 1993a). Most of the basins in the western Barents Sea (i.e. Tromsø and Sørvestsnaget basins) continued to subside during the Late Cretaceous. Although extension was dominant at regional scale during the Late Cretaceous, compressional deformation evidenced by reverse faults and folds was also observed along the Bjørnøyrenna Fault Complex and the Ringvassøy Fault Complex (Gabrielsen et al. 1990, 1997). Seismic stratigraphy analysis suggested that compression started to act after the deposition of the Intra-Cenomanian reflector and the Late Cretaceous strata, which was also involved in folding and reverse faulting was eroded and uncomformably sealed by the Cenozoic strata (Gabrielsen et al. 1997). The compressional phase continued into Cenozoic and early Cenozoic sequences also experienced head-on inversion with NW-SE compression direction (Figure 11c in Gabrielsen et al. 1997). The Late Cretaceous inversion phase is characterised by e.g. upright open folds, close to tight inclined to recumbent folds and compressional footwall shortcuts. In contrast, Rønnevik and Jacobsen (1984) and Riis et al. (1986) suggested sinistral shear movement due to the opening of the North Atlantic Ocean, causing reactivation of the Bjørnøyrenna Fault Complex in the Late Cretaceous and Cenozoic. The Ringvassøy Fault Complex was also reactivated during the Late Cretaceous (Brekke and Riis 1987, Gabrielsen et al. 1990). 5.3 Data and Methodology We used 2D MOVETM (structural modeling and analysis software by Midland Valley Exploration Ltd) for structural restoration of Cretaceous inversion structures in the Bjørnøyrenna Fault Complex, western Barents Sea. Three key seismic profiles with different orientations (Fig. 5.3) have been selected for this purpose. These seismic lines belong to two different surveys (NBR08, NBR10) and are of variable quality and variable depth resolution. The data were provided by the Department of Geosciences, University of Oslo, Norway. The main motivation for the modeling is to restore the sections backwards, to locate null point positions and to isolate inversion events at Lower and Late Cretaceous levels. To these aims the 2D kinematic module has been used and different restoration techniques (i.e. 2D unfolding or flexural slip, 2D move on fault or simple shear) have been adopted. In general, two objectives can be obtained from restoration or backward modeling of a particular structure. The structural restoration can validate the interpreted geometry in cross section and can provide information about the processes of progressive deformation in the region. Detailed 70

83 accounts on structural restoration, 2D kinematics and theory of the used modules are given in section 3.3. Figure 5.3. Base map of the study area showing locations and orientations of the restored seismic profiles and locations of used well. In order to identify inversion events and to locate null points the models need to be decompacted and restored. The workflow used in the current restored models includes digitizing of seismic cross sections. Different horizons including sea bottom and top basement were digitized and polygons representing the different sediment packages were created. The details of all digitized horizons, age and thickness (according to Well 7219/9-1) are given in table 5.1. A brief description of all stratigraphic units can be found in section 5.4. The 2D depth-conversion tool was used to convert the seismic sections into depth. The 2D decompaction tool was used to decompact the rock units and footwall blocks were unfolded using the flexure slip method. The hanging wall blocks were moved upwards along the fault to restore the latest pre-compressional sedimentary layer, i.e. Hekkingen Formation, to its original position. 71

84 Table 5.1. Digitized horizons representing colour and age. 5.4 Structural restorations Structural restoration allows for validating an interpreted section through backstripping of depositional and tectonic events by applying certain geometric rules. The robustness of the restoration is checked by means of identifying space problems during unfolding restoration. The 2D unfolding tool, involving flexural slip unfolding, to restore footwall block and the move-on-fault technique for hanging wall block were selected in the present study. The flexural slip unfolding algorithm maintains bed thickness between the template horizon and other passive objects. The algorithm is also built to maintain the line length of the template horizon in the direction of unfolding and maintain the area of the fold and the model4. A detailed description of the 2D unfolding procedure is given in section The move-on-fault module is used to restore deformation and allows the user to model pre and syn tectonic successions and different displacements. Simple shear maintains the relationship between fault geometry and hanging wall deformational features. 72

85 It diffuses deformation throughout the hanging wall instead of partitioning it into discrete slip between beds (i.e. flexural slip). A detailed description of the 2D move-on-fault procedure is given in section Results Based on the interpretation of seismic cross-sections, calibrated by exploration wells, the structural restoration permits to investigate the Early and Late Cretaceous inversion events in the western Barents Sea. Total three key seismic profiles located in the central and northern segments of the Bjørnøyrenna Fault Complex have been used for this purpose (Fig. 5.3) Restoration of key profile 1 The WNW-ESE seismic cross-section crossing the central segment of the Bjørnøyrenna Fault Complex (Fig. 5.3), was imported in MOVETM (Fig. 5.4). Ten key horizons (Table 5.1) and two faults were digitized using horizon and fault options (Fig. 5.5). Stratigraphy and rock properties were assigned for each sedimentary package according to data available from well 7219/9-1 (Table 5.2). The polygons of the different sediments packages, ranging in age from (Late Triassic to Late Cretaceous, were created using the create auto polygon tool. The seismic image then converted into depth using 2D depth conversion tool. The tool used for conversion of a seismic section from time to depth or vice versa. 4 Product manual of Move TM 2016 Midland Valley Exploration Ltd. 73

86 The equation implemented is: Z= V0 (ekt - 1)/k Where: Z = depth in meters v0 = initial velocity (m/s) k = rate of change in velocity with increasing depth t = one-way travel time(s) If k is equal to zero then the formula becomes: Z= V0 t Stratigraphy Top Time depth (TWT) Porosity Depth Velocity Thickness coefficient (m/s) (m) (m) Sea bottom Kolmule Knurr Formation Hekkingen Formation Formation Fuglen Formation Nordmela Formation Tubaen Formation Fruholmen Formation Snadd Table 5.2. Rock properties of sedimentary packages adopted from well 7219/

87 Figure 5.4. Interpreted seismic cross-section of the central segment of the Bjørnøyrenna Fault Complex. See figure 5.4 for location (Data courtesy of TGS and Spectrum). All the sediment packages (polygons) were then decompacted using Sclater and Christie s laws (Sclater and Christie 1980). In present study, most of the used decompacted parameters (e.g. surface porosity and depth coefficient) are given by the program (Table 5.2). A detailed description of the 2D decompaction procedure is given in section The top-most horizon (Earth surface) was deleted after decompaction. 75

88 Figure 5.5. Digitized key profile 1; seismic cross-section showing sediment packages and major faults (see figure 5.4 for location). The footwall block is then unfolded using the flexural slip method (Fig. 5.6). The base Cretaceous reflector (top Hekkingen Formation) is used as a template bed and unfolded to an arbitrary depth of ~1550 m. The rest of all horizons are kept as passive features which follow the template bed and preserve interbed volumes. The resulted restored section shows the unfolded footwall block with anticlockwise rotation and the hanging wall block on the left side remains unaltered (Fig. 5.6). The restored section resulted in ~ m gap between the two fault blocks. 76

89 Figure 5.6. Unfolded footwall block at Base Cretaceous level (top of Hekkingen Formation). The hanging wall block is moved along the fault to connect the base Cretaceous horizon (Fig. 5.7). A simple shear method is adopted and a shear angle of 32 with respect to vertical (Yamada and McClay 2003) is applied. The hanging wall block was moved ~300m upwards and the resultant horizontal elongation or restored shortening is ~ m. All the sediment packages up to the Base Cretaceous level are well connected with each other (Fig. 5.8). The thickness variations of Lower Cretaceous sediment package (Knurr Formation) from the hangingwall block and to the footwall fault block indicate positive inversion. The extensional fault reversed the sense of motion during compression which caused Lower Cretaceous sediments to turn inside out and to become a positive feature (Fig. 5.8). As a result the fault retains net extension at a deeper level and net contraction associated with an anticline in the upper portion. 77

90 The resultant horizontal elongation or restored shortening is ~ m. The hanging wall block moved ~300m upward. Figure 5.7. Hanging wall block is moved along the fault up to the base Cretaceous level. The restored section shows the null point at the base of the Cretaceous (Hekkingen Formation, Fig. 5.8). The position of the null point evidences the progressive compressional inversion of Lower Cretaceous syn-rift sequence (Knurr Formation). Below the null point, the geometry of the restored fault shows normal faulting while above the null point the geometry points to reverse. The eroded part of Lower Cretaceous (Knurr Formation) is interpolated which suggests reverse faulting. The restored section shows also positive inversion as a fold in the hangingwall at the base of the Upper Cretaceous (Kolmule Formation). The eroded Kolmule Formation is interpolated from the footwall towards the hanging wall (Fig. 5.8). The interpolation suggests reverse faulting of the base of the Upper Cretaceous. The restored section is slightly modified and space issues (~ m) between the hanging wall and the footwall are solved (Fig. 5.8). 78

91 Assumed eroded part of the Early Cretaceous (Knurr Formation). Position of null point at the base Cretaceous (Hekkingen Formation). Assumed eroded part of the base upper Cretaceous (Kolmule Formation). Figure 5.8. Position of null point indicates positive inversion affecting Lower and Upper Cretaceous sediment packages Restoration of key profile 2 The WNW ESE seismic cross section (Fig. 5.9) located north of key profile 1 and passing through the central segment of the Bjørnøyrenna Fault Complex (see figure 5.4 for location) is used for structural restoration. Eight horizons including top Fruholmen Formation (Base Jurassic), top Tubåen Formation (Early Jurassic), top Nordmela Formation (Mid Jurassic), top Fuglen Formation (Upper Mid Jurassic), top Hekkingen Formation (Base Cretaceous), top Knurr Fomation (Lower Cretaceous), top Kolmule Fomation (Base Upper Cretaceous) and seafloor are digitized (Fig. 5.10). 79

92 Figure 5.9. Interpreted seismic cross- crossing the central segment of the Bjørnøyrenna Fault Complex (for location see figure 5.4; Data courtesy of TGS and Spectrum). 80

93 Figure Digitized key profile 2 (for location see figure 5.4). After creating polygons of all sediment packages the seismic image is converted into depth using the 2D depth conversion tool. The footwall is then unfolded using the flexural slip method (Fig. 5.11). A similar procedure to the one used in the restoration of key profile 1 is applied and the base Cretaceous reflector is used as a template bed. The footwall is unfolded to the 1600 m depth level. The other horizons are kept as passive features. The restored section shows an unfolded footwall with anticlockwise rotation and the hanging wall remains unaltered (Fig. 5.11). The restored section resulted in ~ m gap between the two fault blocks. The base of the Upper Cretaceous (Kolmule Formation) shows clockwise tilting on the footwall as a response to the underlying folded Knurr Formation (Fig. 5.11). 81

94 Fault gap (~ m) between two fault blocks. Unfolded to datum ~ 1600m (The presumed depth of base Cretaceous used as reference). Figure D unfolding (flexure slip method) using the Base Cretaceous (top of Hekkingen Formation) as template bed. The hanging wall is moved ~700m along the fault in order to the base Cretaceous horizon in the footwall block (Fig. 5.12). A simple shear method is adopted and a shear angle of 32 from the vertical is applied as proposed by Yamada and McClay (2003). The resultant horizontal elongation or restored shortening is ~500m. The eroded part of Lower Cretaceous (Knurr Formation) and the Upper Cretaceous (Kolmule Formation) are interpolated. The eroded sediment packages on the footwall (Knurr Formation and Kolmule Formatio) are extended to the hangingwall (Fig. 5.12). In the upper part of the Lower Cretaceous sequence, the extensional fault inverted its sense of motion during compression and, as a result, the restored section suggests a positive inversion structure (anticline, reverse fault) at the level of Lower Cretaceous (Fig. 5.12). The restored section also indicates a compressional structure (reverse fault) at the base of the Upper Cretaceous (Kolmule Formation). 82

95 The resultant horizontal elongation or restored shortening is ~ 500m. Position of null point at the base Cretaceous level. The hanging wall block moved ~ m upward. Figure Hanging wall block is moved along the fault up to the base Cretaceous (top of the Hekkingen Formation). 83

96 Restoration of key profile 3 The WSW-ENE seismic cross section (Fig. 5.13) crossing the northern segment of the Bjørnøyrenna Fault Complex (figure 5.4) was used for structural restoration. Figure Interpreted seismic cross-section of the northern segment of the Bjørnøyrenna Fault Complex (see figure 5.4 for location; Data courtesy of TGS and Spectrum). Seven horizons including top Tubåen Formation (Early Jurassic), top Nordmela Formation (Mid Jurassic), top Fuglen Formation (Upper Mid Jurassic), top Hekkingen Formation (Base Cretaceous), top Knurr Fomation (Lower Cretaceous), top Kolmule Fomation (Base Upper Cretaceous) and seafloor are digitized (Fig. 5.14). 84

97 Figure Digitized key profile 3 (for location see figure 5.4). Auto polygon tool is used and polygons are created representing different sediment packages ranging in age from Early Jurassic to Late Cretaceous. A similar procedure to the one used for the depth conversion and decompaction of profile 1 and 2 is applied. In key profile 3, the footwall block is eroded upto the base of the Upper Cretaceous (Kolmule Formation) due to the uplift of the Loppa High, located in the east of the Bjørnøyrenna Fault Complex. After depth conversion and decompaction, 2D unfolded technique is applied on hanging wall block. The base Cretaceous horizon (Hekkingen Formation) is used as a template bed and the section is unfolded at the depth of ~1430m. The resulted restored section shows compressional structure (anticline) associated with Early and Late Cretaceous inversion events. Both the Lower Cretaceous (Knurr Formation) and the base of the Upper Cretaceous (Kolmule Formation) moved down ward showing compression. A minor gap (~50-75m) occurred between the hanging wall block and the Loppa High due to unfolding (Fig. 5.15). The base of the restored section is slightly modified (20-30m) to avoid further gaps. 85

98 Fault gap (~ 50-75m). Figure D unfolding (flexure slip method) using the Base Cretaceous horizon (top of Hekkingen Formation) as templet bed. The restored section shows downward movement of hangingwall depicting inversion structure (anticline) at Lower Cretaceous (Knurr Formation) and at the base upper Cretaceous (Kolmule Formation). 86

99 5.6 Discussions The Bjørnøyrenna Fault Complex is defined as extensional by origin (Gabrielsen et al. 1990, 1997) with listric fault geometries (Faleide et al. 1993). The main phase of subsidence along the fault complex started in late Middle Jurassic i.e. Callovian (Gabrielsen et al. 1990, Faleide et al. 1993a). The subsidence along the Bjørnøyrenna Fault Complex is locally interrupted in (Hauterivian-Aptian) which is evidenced by local thinning of Lower Cretaceous (HauterivianAptian) sediments (Gabrielsen et al. 1997). The inversional event is marked as a strike-slip origin (dextral wrenching) on the basis of interpreted basin ward tilt of strata which define half-flower-like structures (Riis et al. 1986, Brekke and Riis 1987, Gabrielsen and Færseth 1988, Gabrielsen et al. 1992, 1997). Indrevær et al also interpreted inversion structures of early Barremian to mid-albian age (ca Ms) along the margins of the Loppa High and concluded that these inversion structures developed due to uplift of the Loppa High along its inclined boundary fault (e.g. Bjørnøyrenna Fault Complex). The uplift of the Loppa High may be due to the contemporaneous extreme lithospheric thinning going on in the Tromsø and Bjørnøya Basin and somehow the heat created from this thinning caused the Loppa High to be uplifted more than its surroundings and hence cause local tectonic inversion in the area as it forced itself upwards like a wedge (Indrevær et al. 2016). The restored seismic sections along the Bjørnøyrenna Fault Complex also showed inversion in Early Cretaceous. The unfolding of footwall block helped to mark the position of null point in key profile 1 and 2, and restored eroded parts of Lower Cretaceous (Knurr Formation). The position of null pint is showing the progressive compressional inversion of extensional syn-rift of Lower Cretaceous sequence (Knurr Formation). Below the null point, the geometry of the restored fault showing normal faulting while above the null point the geometry is reverse. The Bjørnøyrenna Fault Complex also affected by Late Cretaceous tectonics and was reactivated in compression in the Cenomanian, and again in the Late Cretaceous (Gabrielsen et al. 1997, Vågnes et al. 1998). The Late Cretaceous sediments were also involved in folding and reverse faults and were eroded and diconformably overlain by the early Cenozoic strata. The Late Cretaceous inversional phase is resulted from head-on, southeasterly-directed contraction and characterised by minor fold trains associated with assumed thrust faults at the base of the Upper Cretaceous (Kolmule Formation) in the central segment of the Bjørnøyrenna Fault Complex (Gabrielsen et al. 1997). Restoration of key profiles also suggested folds and reverse faults associated with Late Cretaceous inversion event. The results also suggested horizontal elongation or restored shortening ~ m in the central 87

100 segment of the Bjørnøyrenna Fault Complex (Key profile 1) and ~500m shortening in the northern part of the central segment (Key profile 2). 5.7 Conclusions Inversion in the Bjørnøyrenna Fault Complex is evidenced by thickness vatiations of Lower Cretaceous (Hauterivian-Aptian) sediments and by folds associated with assumed reverse faults in Late Cretaceous. The restored seismic sections also confirmed two inversion events i.e. Early and Late Cretaceous. The thickness variations of the Lower Cretaceous sediment package (Knurr Formation) in the hangingwall block and footwall fault block (key profile 1 and 2) indicate positive inversion. The extensional fault reversed the sense of motion during compression which caused Lower Cretaceous sediments to turn inside out and became positive feature. As a result fault retains net extension at a deeper level and got net contraction associated with anticline in the upper portion. 88

101 6. Numerical modeling of multi stage basin inversion in the western Barents Shelf 6.1. Introduction The western Barents Shelf covers the area between Svalbard in the north, mainland Norway in the south, Novaya Zemlya in the east and the Norwegian Greenland Sea in the west (Fig. 6.1). It is bounded by the Norwegian Greenland Sea in the west and the Eurasian Basin in the north, which developed during the final continental breakup in Cenozoic (Faleide et al. 1993a). During the time of crustal breakup the western Barents Sea consisted in two megalineaments including the North Atlantic rift zone between the present Charlie Gibbs and Senja Fracture Zones and a shear zone, the De Geer Zone (Harland 1969, Faleide et al. 1993a). The overall setting of the area showed a transcurrent to transform notion during the Late Cretaceous - Paleogene rifting and early opening between Greenland and Eurasia. The study area contains some of the deepest basins of the world separated by intra-basinal structural highs and deep seated fault complexes (Gabrielsen et al. 1990, Faleide et al. 1993a). Different regional tectonic events from Paleozoic to Cenozoic within the North Atlantic Arctic region are responsible for the development of these structures (Faleide et al. 1993a, Gernigon et al. 2014). The overall structural trend in the western Barents Sea is NE-SW to ENE-WSW in its eastern and central parts, whereas the western part of the area consists of NNW-SSE to N-S fault complexes (Gabrielsen et al. 1990; Fig. 6.1). The existence of contractional structures (anticlines, synclines and reverse faults) of Cretaceous Cenozoic age and reactivation of major fault complexes in the western Barents Sea were confirmed by numerous investigators (Rønnevik and Motland 1979, Rønnevik et al. 1982, Rønnevik and Jacobsen 1984, Faleide et al. 1984, Faleide et al. 1988, 1993a, b, Richardson 1992, Vågnes et al. 1998, Bergh and Grogan 2003, Gabrielsen et al. 1990, 1997). The development of inversion structures in the Barents Sea is suggested to be related to strikeslip (Riis et al. 1986, Brekke and Riis 1987, Gabrielsen et al. 1990, 1997) or head-on inversion (Gabrielsen et al. 1992, 1997). Possible mechanisms for the formation of Cenozoic inversion structures in the Atlantic region could be: rifting and subaerial sea-floor spreading (Vågnes et al. 1998), ridge push and mantle drag (Bott 1991, Boldreel and Andersen 1993, Wilson 1993, Doré and Lundin 1996, Gölke and Coblentz 1996, Bird 1998, Vågnes et al. 89

102 1998, Doré et al. 2008, Ziegler et al. 2001) and differential sediment loading (Stuevold et al. 1992). Figure 6.1. Regional setting and major structural elements of the SW Barents Shelf (modified from and google map and NPD fact maps AFC = Asterias Fault Complex, BB = Bjørnøya Basin, BFC = Bjørnøyrenna Fault Complex, BP = Bjameland Platform, CFZ = Central Fault Zone, COB = Continental Oceanic Boundary, FSB = Fingerdjupet Sub-Basin, HB = Harstad Basin, HfB = Hammerfest Basin, HFC = Hoop Fault Complex, HrFC = Hornsund Fault Complex, KFC = Knølegga Fault Complex, LFC = Leirdjupt Fault Complex, LH = Loppa High, MB = Maud Basin, MFC = Måsøy Fault Complex, MH = Mercurius High, NB = Nordkapp Basin, NFC = Nysleppen Fault Complex, NH = Norsel High, OB = Ottar Basin, PSB = Polhem Sub-Platform, RLFC = Ringvassøy Loppa Fault Complex, SB = Sørvestsnaget Basin, SFZ = Senja Fracture Zone, SH = Stappen High, SR = Senja Ridge, TB = Tromsø Basin, T-FFC = Troms-Finnmark Fault Complex, TIFC = Thor Iversen Fault Complex, VH = Veslemøy High, VVP = Vestbakken Volcanic Province. The aim of the research is to investigate the causes and effects of Mesozoic and Cenozoic inversion events from Late Triassic to Miocene in the western Barents Shelf through numerical modeling. In order to model Mesozoic and Cenozoic stress pattern in the western 90

103 Barents Sea, ANSYS (Work bench) was used and four different 2-D thin plate modeling setups were made Numerical modeling The Finite Element Method (ANSYS ) A multiphysics FEM program, namely ANSYS, an advanced engineering simulation tool that is aimed at the analysis of various physical systems and problems like fluid dynamics, structure mechanics, thermomechanical systems and electromagnetics was used. The tool can also be used in the simulation of stress and strain. For detailed mathematical description the reader is referred to section 3. Considering the specific geometry of the problem and the data at hand, a 2-D thin-plate approach was adopted, assuming plane strain conditions. In the models linear elasticity was used and contact elements were introduced to simulate major faults. 6.3 Model set up Four different models were constructed in order to predict stress patterns and to explore the conditions for fault reactivation and eventually tectonic inversion during four specific and regional tectonic events spanning from Late Triassic to Miocene. The Late Triassic to Early Jurassic (i.e. Model 1), the Late Cretaceous (i.e. Model 2), Early Eocene (Model 3) and Miocene to recent (Model 4) inversion events. Model 1 involved ~70250 triangular solid elements with mid-nodes and ~ 2970 surface to surface body contact elements. Model 2, 3 and 4 included ~ triangular solid elements and approximately 3000 contact elements. In the most refined parts of the models (i.e. along the simulated faults) element minimum size is ~20 m and average size is ~100 m (Fig. 6.2). 91

104 Figure 6.2. Numerical meshes with refinement along simulated faults. (a) Model 1 involved ~70250 triangular solid elements with mid-nodes and ~ 2970 surface to surface body contact elements. (b) Model 2, 3 and 4 included ~ solid elements and ~3000 contact elements. Elastic and isotropic elastic behavior was prescribed to the models and two domains with distinct properties were defined: (1) an oceanic one with relatively high Young's modulus (E=100 GPa) (Model 4) and (2) a continental one (E=60 GPa). The latter value remains reasonable for continental crust (e.g. Pascal and Gabrielsen 2001). A standard Poisson's ratio value of 0.25 was chosen for the whole modeled area. The major fault zones introduced into the models represent the major structures bordering the sedimentary basins of the western Barents Sea (Faleide et al. 1993a, b, Gabrielsen et al. 1990, 1997; Table. 6.1). In order to explore what fault segments are the most likely to be reactivated during the four studied inversion events, friction coefficient values of 0.1 to 0.6 were tested and 0.1 is set for all fault segments in models because the modeled stress magnitudes are too low to generate displacement along faults with standard friction coefficients. Normal stiffness (FKN) values from 0.01 to 1 were tested. The lowermost FKN values ( ) resulted in bending and contact penetration problems. No difference was noted for FKN values from 0.1 to 1. A value of 1 was finally retained for the modeling. 92

105 Model 1 Young Modulus (E) = 60 GPa Model 2 Young Modulus (E) = 60 GPa Model 3 Young Modulus (E) = 60 GPa Model 4 Young Modulus (E) = 100 GPa Poisson s ratio = 0.25 Poisson s ratio = 0.25 Poisson s ratio = 0.25 Poisson s ratio = 0.25 Asterias Fault System 0.1 Bjørnøyrenna Fault Complex Fault zones Hoop Fault System Hornsund Fault Complex Leirdjupt Fault Complex Måsøy Fault Complex Nysleppen Fault Complex Ringvassøy Loppa Fault Complex Senja Fracture Zone Friction coefficient (µ) Troms-Finmark Fault Complex 0.1 Thor Iversen Fault Complex 0.1 Table 6.1. Major fault zones introduced into the models with their material properties. 93

106 6.4 Boundary conditions The imposed boundary stresses were taken in agreement with stress directions suggested in the literature (Doré and Lundin 1996, Gabrielsen et al. 1997, Vågnes et al. 1998, Mosar et al. 2002, Tsikalas et al. 2002, Buiter and Torsvik 2007, Doré et al. 2008). Different stress magnitudes were tested (10, 25 and 50 MPa) for each model. Finally, a stress magnitude of 50 MPa was selected (Fig. 6.3). In Model 1, Late Triassic to Early Jurassic E-W compression, associated with westward motion of Novaya Zemlya (Buiter and Torsvik 2007), was assumed (Fig. 6.3a). The western and northern boundaries of Model 1 were kept fixed in the x and y directions respectively. For Model 2, NW-SE directed compression in Late Cretaceous was assumed (Fig. 6.3b) with fixed western (in the x direction) and northern (in the y direction) boundaries. In Model 3, NW-SE Early Eocene North Atlantic opening (Tsikalas et al. 2002, Doré et al. 2008) was introduced (Fig. 6.3c). In the present modeling 1 km of opening was imposed. Such a value allows for keeping the model numerically stable while simulating stress directions. The northern boundary of the model was kept fixed according to y and the eastern and western boundaries according to x. In addition, ~ 20 nodes on the Barents Shelf side along the Senja Fracture Zone are fixed to prevent artefacts due to physical spreading of the material towards the gap created during the opening. As a far field stresses, 50 MPa stress is induced from the southern boundary. For Model 4, NW-SE Atlantic ridge push from Miocene to present-day was introduced (Fig. 6.3d). All four models cover ~ 2000 km2 but only ~ 900 km2 represent the area of interest. In order to avoid potential edge effects, the external boundaries of the models were placed relatively far away (i.e. ~ 550 km) from the area of interest (Fig. 6.3). 94

107 Figure 6.3. Geometry and assumed boundary conditions used in the modeling (acronyms as in caption of fig. 6.1). See text (section 6.4) for explanation. 95

108 6.5 Results Late Triassic Early Jurassic Simulated stress patterns are presented for the Late Triassic-Early Jurassic in figure 6.4. The results indicate local stress rotation in the northern segment (1) of the Thor Iversen Fault Complex. The NE-SW to ENE-WSW fault was reactivated as a dextral one in the model implying that the maximum principal stress rotated counterclockwise (Fig. 6.4). The angle between the strike of the fault and the regional maximum principal stress is suggestive of transpression. Slight counterclockwise rotation of ϭhmax is also modeled in the central segment (2) of the Thor Fault Complex. The change in direction for the maximum principal stress is due to the NE-SW orientation of the fault. The simulated results show dextral reactivation in particular for fault segment (2). The southern segment (3) of the Thor Iversen Fault Complex, which lies parallel to the main applied stress, does not show any rotation due to E-W alignment of the fault strike (Fig. 6.4). Figure 6.4. Simulated stress patterns for Late Triassic-Early Jurassic 96

109 In concert with the observations (Glørstad-Clark et al. 2010) reactivation as a dextral fault and subsequent stress rotation is also predicted for the Måsøy Fault Complex (4) similarly to the Thor Iversen Fault Complex. Large counterclockwise σ1 rotations were also modeled in the southern tip of the Måsøy Fault Complex (5) depicting reactivation. A very slight stress rotation is also observed in the southern part of the Hoop Fault Complex (6) where maximum principal stresses are simulated with a counterclockwise rotation (Fig. 6.4). The simulated stress pattern shows dextral reactivation which is due to the NNE-SSW strike of fault. The modeled maximum principal stress pattern shows rotation in the northern part (N-S) of the Troms-Finnmark Fault Complex (7), where counterclockwise rotations at its southern tip and clockwise rotations at its northern tip point to dextral reactivation. No change in maximum stress pattern is modeled in the northern segment of the Hoop Fault Complex (8) which lies perpendicular (N-S) to the applied stress direction (E-W). Also along the Norwegian mainland (9) and on the western side of the model 1 (10) the simulated stress pattern is almost regular due to the modeled uniform rheological properties Late Cretaceous Simulation of stress patterns (Model 2) is presented in Figure 6.5, which was designed to account for Late Cretaceous tectonic situation. Modeled stress rotations are observed in most of the major fault complexes located in the study area (Bjørnøyrenna Fault Complex, Ringvassøy Loppa Fault Complex, Leirdjupet Fault Complex, Hoop Fault Complex, Asterias Fault Complex, Nysleppen Fault Complex, Måsøy Fault Complex and TromsFinnmark Fault Complex). 97

110 Figure 6.5. Simulated stress patterns for the Late Cretaceous. Major stress rotation were modeled in the eastern side of the study area, where the NE-SW the Nysleppen Fault Complex, shows a change in stress pattern (1). The southern segment of the Nysleppen Fault Complex experienced clockwise stress rotation, depicting reactivation during Late Cretaceous (Fig. 6.5). A change in stress pattern is suggested in the northern part (2) of the Troms -Finnmark Fault Complex (Fig. 6.5) where the maximum principal stress (regional strike NW-SE) is indeed modeled with a N-S strike. A clockwise stress rotation at the northern tip (2) of the segment suggests reactivation of the Troms -Finnmark Fault Complex. Moderate stress rotation was simulated for the NE-SW-striking the Hoop Fault Complex (Fig. 6.5). The simulated maximum principal stress pattern becomes parallel in the northern segment (NNW-SSE) of the Hoop Fault Complex (3). In contrast, simulated stress directions in the southern segment (NE-SW) of the fault rotate counterclockwise (4). The change in 98

111 direction for the maximum principal stress is due to the alignment of the fault segments and suggests reactivation of both segments of the Hoop Fault Complex (Fig. 6.5; 3, 4). The modeled Late Cretaceous tectonic situation shows clockwise stress rotations in the eastern segment (5) of the Asterias Fault Complex. The direction of maximum horizontal stress ( Hmax) becomes perpendicular along the strike of the fault (ENE-WSW) in the central segment (5) and rotated counterclockwise in the western segment of the Asterias Fault Complex. Notable clockwise rotation in the eastern tip and counterclockwise stress rotations in the western tip of the Asterias Fault Complex indicates reactivation of the segment. The simulated stress pattern close to the central segment of the Asterias Fault Complex depicts inversion (Fig. 6.5). It is noted that Model 2 displays a stress situation that would be compatible with a compressional regime in the Asterias Fault Complex in the Late Cretaceous. Model 2 suggest stress rotations for the Leirdjupet Fault Complex (6). The simulated results show a change in Hmax direction in the northern segment (6) of the Leirdjupet Fault Complex which is oriented NNE-SSW (Fig. 6.5). Hmax rotated clockwise and aligned itself with the fault (6). The maximum stress pattern shows reactivation and inversion of the Leirdjupet Fault Complex. Clockwise stress rotations were simulated for the Bjørnøyrenna Fault Complex and the Ringvassøy Loppa Fault Complex in Model 2 (Fig. 6.5). A change in direction for the maximum principal stress was modeled in the central segment (7) of the Bjørnøyrenna Fault Complex (Fig. 6.5). The results show that the direction of Hmax changed from NW-SE to NNW-SSE (clockwise) depicting reactivation. The northern segment of the Bjørnøyrenna Fault Complex (7) also shows a change in Hmax direction (Fig. 6.5). Maximum principal stresses become almost perpendicular to the fault strike (NE-SW) suggesting inversion of the northern segment of the fault. Compressional stress was also modeled for the Ringvassøy Loppa Fault Complex (Fig. 6.5). The counterclockwise stress rotation in the northern part (8) here is in agreement with the presence of contractional structures (e.g. folds and reverse faults) observed by Gabrielsen et al. (1997) and Hameed (2012) and the findings of the seismic interpretation in this thesis. Drastic Hmax rotations were also simulated in the Veslemøy High and the Senja Ridge (Fig. 6.5, 9). Triple junction stress rotations are due to the closely located edges of the structural highs. Pronounced clockwise maximum principal stress rotation was modeled in the central 99

112 part of the Dee Geer Zone (10) where Hmax became parallel to the strike of this latter zone (NW-SE). The uniform stress pattern orientation on the Finnmark Platform (11) and the western part (12) of the Model 2 is a consequence of to choice of uniform rheological properties for these areas Early Eocene Figure 6.6 shows the simulated stress pattern caused by dextral shear between Greenland and the Barents Sea Shelf in Eocene. Along the Senja Fracture Zone (Fig. 6.6a, 1), ϭhmax becomes parallel to the strike of the fault (i.e. NW-SE). The modeled stress pattern shows shearing along the southern segment of the De Geer Zone (Fig. 6.6b). A slight change in stress direction on the western side of the southern segment is due to the imposed boundary conditions (Fig. 6.6a and 6.6b). The simulated stress pattern shows rifting in the central segment (2) of the De Geer Zone (Fig. 6.6b). The development of the rift segment follows the NW-SE applied stress direction which acted perpendicularly to the NE-SW striking central segment (Fig. 6.6b). Along the northern segment of the zone, the results of the modeling suggest shearing and rifting (3). ϭhmax was modeled parallel to the fault strike (i.e. NW-SE) in the northern part of the segment, depicting shearing/strike-slip movement (4). Along the southern part, just above the rifted segment, the simulated stress pattern suggests moderate extension and shearing (Fig. 6.6b, 3). A slight change in the stress pattern on the western side of the model is due to the boundary conditions. On the eastern side (5) of the model 3, the stress pattern is regular due to the applied far-field stresses. The results suggest that Early Eocene sea floor spreading caused stress partitioning along the Senja Fracture Zone and did not influence significantly the eastern side of it (Fig. 6.6). 100

113 Figure 6.6. Simulated Eocene stress patterns Miocene Figure 6.7 shows the Miocene simulated stress pattern (Model 4). Stress rotations of ~30-40 are modeled at the continent-ocean boundary (1) because of the imposed contrasting rheologies between the two modeled crustal domains. Pronounced stress rotations are simulated along the boundary segments which are oriented obliquely to the applied stress. More drastic stress deflections are visible in the Hornsund Fault Complex (2), where ϭhmax becomes perpendicular to fault strike (Fig. 6.7). Large stress rotations occurred at the tips of the fault zone, where ϭhmax becomes parallel or perpendicular to the fault. A notable counterclockwise rotation at the northern tip and clockwise stress rotations in the south of the Hornsund Fault Complex indicate dextral reactivation of the segment. The southern part of the continental oceanic boundary (i.e. Senja Fracture Zone; 3) also shows stress deflections (Fig. 6.7). Maximum principal stresses become perpendicular to the strike of the fault. The simulated maximum principal stress pattern becomes parallel to the NW-SE northern segment (4) of the Knølegga Fault Complex (Fig. 6.7). In contrast, simulated stress directions in the southern segment of fault (5), striking N-S and NNE-SSW, are perpendicular. The change in direction for the maximum principal stress is due to the alignment of the fault segments and may cause reactivation of the Knølegga Fault Complex. 101

114 Clockwise stress rotations are modeled in the Leirdjupet Fault Complex. Maximum principal stresses are parallel throughout the main fault (6) depicting a strike-slip component (Fig. 6.7). Stress rotations are modeled by the Bjørnøyrenna Fault Complex. The modeling results in the northern segment (7) of the Bjørnøyrenna Fault Complex show that σ1 is simulated perpendicular to the fault (NE-SW), whereas it appears clockwise rotated and almost parallel (NW-SE) to fault complex in its central part (8) (Fig. 6.7). The change in stress pattern suggests reactivation and inversion in the particular segments of the fault and suggests the development of strike-slip related inversion along the central and head-on inversion in the northern segment of the fault complex. The southern part of the fault exhibits inversion, as the modeled clockwise rotation of the maximum principal stress suggests (Fig. 6.7). The maximum horizontal stresses are deflected and become almost parallel to the N-S southern segment of the Ringvassøy Loppa Fault Complex (Fig. 6.6, 9). Change in stress directions were also modeled along the northern segment, where maximum principal stresses rotated clockwise to become almost parallel to the fault. The simulated stress pattern suggests strike-slip movement in the southern segment and head-on inversion in the northern segment of the fault complex. The rotation of maximum principal stresses is due to the N-S and NESW orientation of the fault segments. The northern segment of the Ringvassøy Loppa Fault Complex and the central part of the Bjørnøyrenna Fault Complex showed stress deflection, where maximum principal stresses aligned parallel to the NNE-SSW strike of the faults. The stress pattern is consistent with strike-slip. Major stress rotations were modeled in the eastern side of the study area, where the NE-SW Hoop Fault Complex (10) and the Nysleppen Fault Complex (11) are associated with changes in stress pattern. The southern segment of the Hoop Fault Complex experienced clockwise stress rotation, suggesting reactivation during Miocene. The southern tip of the Nysleppen Fault Complex was also reactivated like suggested by the clockwise deflection of the maximum principal stresses (Fig. 6.7). 102

115 Figure 6.7. Simulated Miocene stress patterns. The Asterias Fault Complex, located almost by the middle of the study area (12) shows slight stress rotations. The change in stress pattern suggests reactivation and inversion in the western segment of the fault complex. Maximum principal stresses were simulated parallel to the Måsøy Fault Complex (13), which is the southern boundary fault of the Nordkapp Basin (Fig. 6.1). Stress rotation suggests reactivation. Clockwise rotation is modeled in the southern segment of the fault. Along the Norwegian mainland (14) the simulated stress pattern is almost regular due to the modeled uniform rheological properties there (Fig. 6.7). 6.6 Discussion Several models have been advanced to explain the probable causes for the reactivation of different fault complexes and development of inversion structures in the western Barents Sea including effect of tectonic forces, gravity loading and sliding, differential sediment loading, 103

116 mantle drag and plate driving forces (Bott 1991, Stuevold et al. 1992, Wilson 1993, Ziegler 1993, Bird 1998, Vågnes et al. 1998, Ziegler et al. 2001, Mosar et al. 2002, Doré et al. 2008). In concert with the regional model of Gabrielsen & Færseth (1988), Gabrielsen et al. (1997) Vågnes et al. (1998) and Doré et al. (2008), the current research favors the influence of different regional tectonic forces and proposes that the Late Triassic Miocene stress development was due to far-field (plate-margin) stress and was responsible for the reactivation of pre-existing faults and inversion in the study area. According to Letouzey et al. (1990) compression perpendicular to existing faults is highly efficient for inversion to occur, but it is seldom to find compression exactly normal to preexisting normal faults. Analyzing the relative contribution of compression and strike-slip in an inverted region is always difficult. The azimuth of slip can be determined from measurements in the field but exact slip determination using 2D seismic data is difficult (Lowell 1995). Experience shows that reactivation of normal faults occur where strike slip is dominant whereas development of new contractional structures (reverse faults) is mainly related to compressional components (Lowell 1995). The stress pattern predicted for Model 1 (Late Triassic to Early Jurassic) shows counterclockwise rotation of the maximum principal stress, σ1 having the tendency to become perpendicular to the strikes of the major fault complexes (Troms-Finnmark Fault Complex, Thor Iversen Fault Complex, Hoop Fault Complex and Måsøy Fault Complex). The simulated stress pattern also favors the presence of inversion structures interpreted by earlier investigators (Berglund et al. 1986, Gabrielsen et al. 1990, Glørstad-Clark et al. 2010, Fitryanto 2011, Gabrielsen et a. 2016). A counterclockwise rotation of σ1 in the Hoop Fault Complex suggests the presence of inversion structures in Late Triassic Early Jurassic. The results are in agreement with the findings of Fitryanto (2011), who interpreted structures related to mild inversion (minor folds) in the southern side of the Hoop Fault Complex dating back to mid-late Triassic. In agreement with Glørstad-Clark et al. (2010), reactivation as a dextral fault and subsequent stress rotation is also predicted for the Måsøy Fault Complex. The area has been mostly under E-W extension throughout its tectonic history. However, the modelling suggests that this general stress situation was possibly interrupted by inversion from Late Triassic to Early Jurassic, presumably associated with the westward push of the Novaya Zemlya (Buiter and Torsvik 2007), as indicated on interpreted seismic sections (Fiytriano 2011, Otto and Bailey 1995, Glørstad-Clark et al. 2010; Fig. 6.8). 104

117 Figure 6.8. Comparison between the numerical modeling results (a) and an interpreted seismic section (b; modified after Fiytriano 2011). The seismic interpretation confirms inversion at MT (Middle Triassic) to early Late Triassic levels on the southern segment of the Hoop Fault Complex. See figure 6.1 for location. For Model 2 a clockwise stress rotation is simulated by the major fault complexes. The modeling result reveals that NNW-SSE re-orientation of the maximum principal stress occurs at the Bjørnøyrenna Fault Complex and the Ringvassøy Loppa Fault Complex. The Bjørnøyrenna Fault Complex is of extensional origin but with signs of tectonic inversion including deformed fault planes and reverse faults affecting Early Cretaceous to Early Cenozoic strata (Gabrielsen et al. 1997). Earlier investigators (e.g. Rønnevik and Jacobsen 1984, Gabrielsen et al. 1984, 1997, Riis et al. 1986, Faleide et al. 1993, Braut 2012, Hameed 2012, Indrevær et al. 2016) suggested that the Early Cretaceous reactivation of normal faults was caused by dextral shear and Late Cretaceous Early Paleogene. The Ringvassøy Loppa Fault Complex was also reactivated during the Early and Late Cretaceous and Cenozoic strata have also been affected by faulting. Results of seismic interpretation also confirm inversion in the Bjørnøyrenna Fault Complex and the Ringvassøy Loppa Fault Complex during the early and Late Cretaceous (Chapter 4). A comparison of seismic and numerical modeling results for 105

118 the Bjørnøyrenna Fault Complex and the Ringvassøy Loppa Fault Complex is presented figure 6.9. Moderate stress deflections are also modeled for the Asterias Fault Complex (Fig. 6.5, 5). The E-W trending Asterias Fault Complex is an extensional structure initiated from Triassic to Jurassic (Gabrielsen et al. 1984, 1990) but reactivated during the Late Jurassic Cretaceous rifting (Indrevær et al. 2016). The presence of inversion structures include reverse faults, halfflower structures and local doming at the Jurassic Cretaceous transition and is evidenced at the western segment of the Asterias Fault Complex (Berglund et al. 1986, Brekke and Riis 1987, Mongat 2010, Indrevær et al. 2016). Figure 6.9. Comparison of numerical modeling results (a) and interpreted seismic section in central segment of the Bjørnøyrenna Fault Complex (b) and the Ringvassøy Loppa Fault Complex (c) confirming Inversion structure at LC (Lower Cretaceous) and BUC (Base Upper Cretaceous) level. See figure 6.1 for locations of seismic lines (Data courtesy of TGS and Spectrum). BUC (Base Upper Cretaceous), LC (Lower Cretaceous), BC (Base Cretaceous), UJ (Upper Jurassic), BJ (Base Jurassic), UT (Upper Triassic). 106

119 A change in maximum principal stress is modeled in the Troms-Finnmark Fault Complex which is believed to be an extensional feature (Gabrielsen 1984) but inversion connected to sinistral slip in the northeastern segment of fault complex was proposed by Rønnevik et al. (1982) and Rønnevik and Jacobsen (1984). The fault complex runs parallel to the coastline of the Troms and the Finnmark counties (Fig. 6.1). The Troms-Finnmark Fault Complex display listric fault geometry with normal dip-slip while the hanging wall is associated with roll-over anticlines and antithetic faults (Fønstelien and Horvei 1979, Faleide et al. 1984, Gabrielsen et al. 1984, Berglund et al. 1986, Gabrielsen et al. 1990). The modeled stress pattern would initiate reactivation in the northern segment of the fault complex which is also observed by Gabrielsen et al. (1990) and Ahmed (2012). The modeling results support the conclusion of these authors. Change in ϭhmax direction in the southern segment of the Hoop Fault Complex is modeled. The fault complex is characterized mainly by normal faulting (Gabrielsen et al. 1990) but reactivated during different tectonic time periods. The modeling results also suggest compression in the Hoop Fault Complex during Late Cretaceous. The modeled maximum stress pattern suggests inversion in the Leirdjupet Fault Complex, which also experienced a phase of Early Cretaceous dextral shear (transtension?) and Late Cretaceous - Early Cenozoic contraction (inversion) (Bjørnestad 2012). Inversion structures (folds) are observed by earlier investigators in the central and northern parts of fault complex. The simulated results show compression in the northern segment of the Leirdjupet Fault Complex which is oriented NNE-SSW (Fig. 6.5; 6). The resulted stress pattern favors the conclusion of Gabrielsen et al. (1990) and Bjørnestad (2012) Origin of the Cenozoic stress field Two main sources of stresses were advanced to explain Cenozoic inversion in the western Barents Sea: Early Eocene North Atlantic opening (Tsikalas et al. 2002, Doré et al. 2008) and ridge push from the North Atlantic mid-oceanic ridge system (Ranalli and Chandler 1975, Stephansson 1988, Talbot and Slunga 1989, Spann et al. 1991, Doré and Lundin 1996). Breakup in the North Atlantic started in Early Eocene (55 Ma; anomaly A24b, e. g. Talwani and Eldholm 1977, Srivastava and Tapscott 1986, Eldholm et al. 1987, Vågnes et al. 1998). The ocean spread northwards and reached the Barents Sea margin at the time of Anomaly A23 (Eocene, 54 Ma). According to Faleide et al. (1996, 2008) the western Barents Sea 107

120 continental margin developed at the Paleocene-Eocene transition (~ Ma) as a result of continental breakup and opening of Norwegian-Greenland Sea. The breakup and initial seafloor spreading in the Norwegian-Greenland Sea was linked to the Arctic Eurasian Basin by the regional De Geer Zone megashear system (Faleide et al. 2008). During and after Early Cenozoic rifting and breakup in earliest Eocene, the western margin of the Barents Sea was subject to dextral shear and associated folding with NW-SE-striking fold axes (Faleide et al. 1996). Vågnes et al. (1998) suggested that the topographic momentum associated with the early phase of sub-aerial sea floor spreading may have contributed to initiate contractional deformation. Sea floor spreading reached the margin off southern Spitsbergen and a narrow oceanic basin formed between the western Barents Sea and continental margin of NE Greenland at the end of Eocene (Faleide et al. 2008). In Early Oligocene the spreading direction between Greenland and Eurasia changed from NNW-SSE to NW-SE (Faleide et al. 2008). A progressive development of the Mid Atlantic Ridge caused by a change in relative plate movement resulted in the formation of an obliquely spreading Knipovich Ridge (Czuba et al. 2011). The main source of Cenozoic inversion in the western Barents Sea is assumed to be caused by the Knipovich Ridge push (Ranalli and Chandler 1975, Stephansson 1988, Talbot and Slunga 1989, Spann et al, 1991). Doré and Lundin (1996) also favor compression due to the counterclockwise shift in the poles of rotation in the North Atlantic between anomalies A13 and A7 (35 25 Ma). Reactivation of NE-SW trending faults in the Vestbakken Volcanic Province is also due to the change in relative plate motion in Early Oligocene Modeled stress patterns compared to observations In Model 3, the simulated stress pattern caused by dextral between Greenland and the Barents Sea Shelf in Eocene is presented. Maximum principal stresses are NW-SE and aligned with the Senja Fracture Zone. The simulated stress pattern shows shearing along the southern segment (Senja Fracture Zone) of the De Geer Zone. Faleide et al. (2008) marked the Senja Fracture Zone as a pure sheared margin which developed due to the opening of the Norwegian-Greenland Sea during Eocene. The generation of the sheared segment was initially related to continent-continent shear followed by continent-ocean shear and has been passive since earliest Oligocene (Faleide et al. 2008). The simulated stress pattern shows rifting in the central segment of the De Geer Zone. The rifted central segment (VVP; Vestbakken Volcanic Province) of the De Geer Zone formed 108

121 due to NW-SE opening. The modeling results show that the main rifting occurred along NESW striking segment. Many previous authors (e.g. Faleide et al. 1993, Breivik et al. 1998, Ryseth et al. 2003) advanced rifting along the central part of the western Barents Sea continental margin, where the strike-slip system changed and resulted into a pull-apart, which caused generation of a rifted segment associated with volcanism, i.e. the VVP (Vestbakken Volcanic Province). Simulated stresses depict the formation of a sheared and to some extent rifted margin along the northern segment (Hornsund Fault Complex) of the De Geer Zone. The ϭhmax aligned with the fault strike (NW-SE) in the northen part of the northern segment showing shearing/strikeslip movement and along the southern part, just above the rifted segment, the modelted stress pattern suggests modest rifting and shearing. Previous studies (Faleide et al. 1993, Faleide et al. 2008, Libak et al. 2012) support the development of an oblique continent-continent and partly continent-ocean shearing along the northern margin segment (Hornsund Fault Complex) during Eocene (Gorgan et al. 1999, Berg and Grogan 2003). The shearing exhibited both transtension and transpression where the restraining bend along north-northwest trending faults between Svalbard and northeast Greenland caused transpression and, as a result, formation of the Spitsbergen fold and thrust belt (Czuba et al. 2011), while the releasing bend between the Senja Fracture Zone and the Hornsund Fault Complex facilitated Oligocene rifting (Faleide et al. 1993). The present study suggests that Early Eocene sea floor spreading caused stress partitioning along the Senja Fracture Zone and any direct effect of NE Atlantic opening on the studied area is deemed minor. The stresses are concentrated along the shear margins (Senja Fracture Zond and Hornsund Fault Zone) and did not penetrate east of the De Geer Zone (i.e. the interior of the Barents Sea). Inversion structures observed by previous authors (Gabrielsen et al. 1984, Riis et al. 1986, Gabrielsen and Færseth, 1988, Bjørnestad, 2012) along the major fault complexes (Hoop Fault Complex, Troms-Finnmark Fault Complex, Ringvassøy Loppa Fault Complex and Bjørnøyrenna Fault Complex) during Eocene are may be related to other mechanisms (e.g. gravity loading and sliding, underplating and Iceland hotspot influence, differential sediment loading, mantle drag). For Model 4, it was inferred that the NW-SE directed Atlantic ridge push is the main source of stresses since Miocene (Gölke and Coblentz 1996, Czuba et al. 2011, Ranalli and Chandler 1975, Stephansson 1988, Talbot and Slunga 1989, Spann et al 1991, Doré and Lundin 1996). 109

122 The results of Model 4 suggest counter-clockwise rotation of the stress field in the western Barents Sea. The maximum principal stress strikes in general NW-SE but often rotates to WNW-ESE. The stresses rotated along the ocean- continent transition because of the implemented change in rheology (Gölke et al. 1996, Pascal and Gabrielsen 2001). Strong deflections were modeled along the fault complexes which are oblique to the applied boundary stress. Maximum principal stresses become perpendicular to the strike of the major fault complexes (Knølegga Fault Complex, Bjørnøyrenna Fault Complex, Ringvassøy-Loppa Fault Complex and Leirdjupt Fault Complex), potentially causing direct inversion. According to Gabrielsen et al. (1990) the Knølegga Fault Complex was mainly inverted during Cenozoic. The contractional structures observed by previous investigators (Gabrielsen et al. 1990) along the Knølegga Fault Complex, including synclines and anticlines, were suggested to be the result of compression in Oligocene. It was proposed that ridge push direction changed from NW-SE to WNW-ESE at the Eocene-Oligocene boundary, caused by the adjustments of the poles of rotation in the North Atlantic (Boldreel and Andersen 1993, in: Vågnes et al. 1998). Pronounced stress rotations are also modeled in the Senja Ridge and the Veslemøy High. The N-S Leirdjupet Fault Complex which extends from the Loppa High towards the Stappen High in the north (Fig. 6.1) and divides the Bjørnøya Basin into a deep western part and a shallow eastern part (Fingerdjupet Subbasin) (Rønnevik and Jacobsen 1984, Gabrielsen et al. 1990) were also affected by Early Cretaceous dextral shear and Late Cretaceous - Early Paleogene inversion (Bjørnestad 2012). Inversion structures are also observed by earlier investigators along the central and northern parts of the fault complex in Miocene. The modeling results are also in agreement with previous studies. Maximum principal stresses are parallel to the main fault line depicting a strike-slip component (Fig. 6.7). Rotations of maximum principal stresses are modeled by the Bjørnøyrenna Fault Complex depicting inversion. The fault complex is believed to be the southern continuation of the Leirdjupet Fault Complex (Gabrielsen et al. 1997) and lies in the north of the RingvassøyLoppa Fault Complex (Fig. 6.1). It was also reactivated during different tectonic episodes (Gabrielsen et al. 1997, Vågnes et al. 1998). The simulated stress pattern shows a change in ϭhmax direction, confirming head-on/strike-slip inversion of the southern and northern segments of the Ringvassøy Loppa Fault Complex. The N-S to NE-SW trending Ringvassøy-Loppa Fault Complex is an extensional fault complex reactivating old zones of weakness. The fault complex was reactivated during Late 110

123 Cretaceous but also Cenozoic strata have been affected by faulting (Gabrielsen 1984). Vågnes et al. (1998) and Doré and Lundin (1996) suggested NW-SE compressional inversion. The modeled stresses are in agreement with previous studies and favor inversion in Miocene due to ridge push from the Knipovich Ridge. A comparison between the present and previous studies is presented in figure Conclusion A FEM numerical tool (ANSYS workbench) was used in the present research and four 2D linear elastic models were generated to investigate the causes for the development of inversion structures from Late Triassic to Miocene in the western Barents Sea. The results of the modeling suggest that different tectonic stresses affected the study area during the mentioned time period and reorientation of stress patterns at major fault complexes indicate the presence of inversion structures. The modeled principal stresses orientations in Model 1, 2 and 4 are in agreement with findings from previous studies. In particular, Model 3 succeeds to predict the observed deformation field and suggests that the opening of the NE Atlantic during Early Eocene had no direct impact on the observed inversion of study area. 111

124 Figure Comparison of present study with previous studies results showing different fault complexes affected by inversion events in the western Barents Sea. 112