Geology and Geophysics

PhD Project: A Comparison Between Active and Ancient Rift Processes (2005)

This project was started in October 2001, supervised by Dr Nicky White and Professor James Jackson at the Bullard Labs, University of Cambridge. I was granted the PhD in July 2005. Professor Dan McKenzie was the main driving force behind the chapter on the elastic strength of extensional basins. This project was sponsored by NERC and Shell.

Dissertation Abstract:

Actively extending sedimentary basins provide a snapshot of the way in which basins evolve. However, observations from such regions are temporally limited: they provide little information on the long-term processes involved even though they yield excellent insight into the short-term motions (10⁰—10⁴ years). The converse is true of tectonically dead basins where the deformation has ceased and only the end product can be observed. This dissertation compares the inactive Late Jurassic extensional East Shetland Basin (ESB) in the northern North Sea with the actively deforming basins in central Greece and the Basin and Range Province, western USA.

In central Greece, faulting appears to migrate basinward with time, on a million-year timescale. Analysis of five fault systems in the ESB suggests that, on a 10 Myr timescale, all the faults were simultaneously active during the Late Jurassic. Syn-rift sediment geometries are very similar in each fault-bounded half-graben, as are the log data from well boreholes. In addition, the rate of extension in the ESB, derived from fault-heave summation, was approximately an order of magnitude slower than present-day strain rates in central Greece. Extension in the ESB occurred at a rate comparable to the current stretching of western North America. This result is consistent with previous strain rates obtained from subsidence modelling. The effective elastic thicknesses of the North Sea, Greece and the Basin and Range Province have been estimated using gravity modelling and are found to be low — less than 5 km in each region. The elastic thickness of the North Sea does not appear to have increased since the end of rifting, probably because post-rift sediments insulate the lithosphere and keep it hot and thus weak. The similarity in elastic thickness between the three basins is likely to be the reason why half-graben in each province have similar wavelengths, between 10 — 40 km.

Rotations of fault blocks about vertical axes are important in the deformation of central Greece. Similar rotations are inferred to have occurred at the Tern-Eider Ridge in the ESB in response to gradients in Late Jurassic extension along the ridge-bounding faults. The structure of the Tern-Eider Ridge is compatible with Late Jurassic east-west extension in the northern North Sea, rather than strike-slip tectonics. A re-evaluation of the structural evolution of the Tern-Eider Ridge by comparison with analogous processes in central Greece can therefore enhance our understanding of how ancient rifts evolved

Click here for a brief geological history of the North Sea.

Click here for details of the North Sea straigraphy used for this project.

Structural Evolution of the Tern-Eider Ridge

Left Location of the Tern-Eider (black box) and North Cormorant (red box) seismic volumes in the ESB. A regional map of the Base Cretaceous Unconformity is shown. Some major fault systems are labelled. UK quadrant numbers are in grey. The Tern-Eider Ridge is a horst block in the East Shetland Basin about 30 km long and 7 km wide. It trends northeast-southwest and lies between the Cormorant fault block and the East Shetland Platform. The ridge is bounded by three major faults which were active throughout the Late Jurassic rift episode in the northern North Sea (168–140 Ma).

Right Map of the Top Brent surface (168 Ma) in the Tern-Eider region. The map is in depth and approximates the pre-rift structure. The original interpretation was depth-converted using well checkshot surveys from c20 wells that penetrate the Top Brent horizon in the Tern-Eider region. The ridge is defined by three major normal faults (offsets of c1 km): the Northern Fault (NF), Southern Fault (SF) and Eastern Fault (EF). Smaller normal faults (offsets of c100s metres) are labelled 1—5. The smallest faults (offsets of c10s metres) are extensional in nature, but are not labelled. A thrust fault is present in the southwest of the ridge. Intra-ridge faulting is organised north-south and is predominantly oblique to the trend of the ridge.

The Northern and Southern faults have strong gradients in extension along their strike (left). Each panel shows the Top Brent horizon from each major fault hanging wall and footwall projected onto a plane parallel to the fault. The distance between the two lines is the throw across the fault. In the Northern Fault the extension increases to the northeast, and to the southwest in the Southern Fault. Extension was fairly constant along the Eastern Fault. The Eastern and Southern Faults are separate with no linkage between them visible in the seismic data.

A thrust fault formed at the southern end of the ridge during the Late Jurassic (right). This fault (TF) was active, as shown by the growth in the syn-rift deposits (the Heather Formation, bounded by the Top Brent Group (TBG) and Base Kimmeridge Clay (BKC), and the Kimmeridge Clay Formation, between the BKC and Base Cretaceous (BX) horizons). The thrust fault is separate from the Northern Fault (NF). TB is Top Basement.

The principle structural observations from the Tern-Eider Ridge are: i) the ridge-bounding faults have gradients in extension along their length; ii) the intra-ridge faults were active simultaneously with the major faults, are extensional in nature, and are organised north-south; iii) the ridge was in compression at its southwestern end during the rift period. These observations can be explained by a simple kinematic model (left). In a. the right side of a thin, rigid plate is moved right (arrow) whilst holding the left side fixed. The plate then deforms as in b. (black lines denote faults). Application of this model to the Tern-Eider Ridge is shown in c. and d.. The diagram is schematic and extension and contraction directions are approximate. The geometry of the faults can be explained by extension in an approximately WNW-ESE direction. Compression in the southwest is a consequence of rotation in a clockwise direction due to asymmetric extension across the ridge. The intra-ridge faults form in response to the ridge being in east-west tension.

The inspiration for interpreting the Tern-Eider ridge in the above way was similar work done in central Greece by Goldsworthy (2001). Strong gradients in extension along the major half-graben structures in Greece (e.g. the Gulfs of Corinth and Evia) also result in the rotation of the intervening semi-rigid continental blocks. The thrust fault in the Tern-Eider region can be understood as a consequence of inter-block motions, rather than a change in extension direction during rifting (cf Thomas and Coward, 1995).

Timing of Faulting in the East Shetland Basin

Studies of the active tectonics of central Greece (i.e. the Gulf of Corinth and Northern Gulf of Evia regions) have demonstrated that normal faulting has migrated basinward over the past 5 Ma on a timescale of ~1 Myr (Goldsworthy et al. 2001). The aim of this section was to see if a similar progression of faulting occured in the ESB during the Late Jurassic. Five major normal fault systems were examined: the Dunlin-Hutton-Ninian fault system; South Tern-Eider; North Tern-Eider; North Cormorant; and Brent (shown left with a map of the Top Brent horizon in the ESB). The activity on each fault system was determined for three time intervals within the Humber Group: the Heather Formation, lower Kimmeridge Clay Formation and upper Kimmeridge Clay Formation. Each of these sediment packages represents approximately 10 Myr of deposition. Fault activity was determined by looking at sediment geometries using isopach maps and seismic sections: divergent sediment packages indicate the fault was active during the time of their deposition. Each of the the three sediment intervals was decompacted using the methodology of Sclater and Christie (1980) in order to recover the sediment geometry at the time of deposition.

Above From left to right: Dunlin, Hutton, Ninian, South Tern, North Tern, and North Cormorant fault systems. Each figure shows: top-left, map of the Top Brent horizon in depth (location given in the first figure of this section); top-right, decompacted isopach map of the Heather Formation; bottom-left, decompacted isopach map of the lower Kimmeridge Clay Formation; bottom-right, decompacted isopach map of the upper Kimmeridge Clay Formation. Each isopach map is decompacted such that the top of the respective sediment package is at the sea bed. The main point of these isopach maps is to show that each exhibits growth into each fault indicating that all faults were active between the top of the Brent Group and the Base Cretaceous Unconformity. As each sediment interval lasts for ~10 Myr, this conclusion is only good at the 10 Myr timescale. Any migration of fault activity on a smaller timescale than this is not resolvable.

Above Seisimic sections, from left: Dunlin, Hutton, Ninian, South Tern, North Tern, North Cormorant. For locations see Top Brent maps above. TKC: Top Cromer Knoll Group; BX: Base Cretaceous Unconformity; IKC: Intra-Kimmeridge Clay; BKC: Base Kimmeridge Clay Formation; TBG: Top Brent Group. Below each section is a depth-section showing sediment horizons (BX:red; IKC:black; BKC:green; TBG: blue) decompacted such that the top of the Kimmeridge Clay Formation is at the sea bed (i.e. this is a section showing the sediment geometries at Base Cretaceous time, assuming no subsequent deformation). These diagrams show how the three Humber Group sediment intervals each show growth into each of the six fault systems. The overlying Cromer Knoll Group is not internally divergent, but characterised by parallel layering that onlaps the Base Cretaceous surface. The Brent Group is also made up of parallel reflections. The Brent fault system has been shown by Mcleod et al (2000) to have been active throughout deposition of the Humber Group. Thus faulting was confined to the time represented by the Humber Group and did not migrate from one fault system to another, on a timescale of 10 Myr.

Rates of Faulting in the East Shetland Basin

1D and 2D subsidence modelling suggests back-arc basins, such as central Greece, extend an order of magnitude faster than intra-cratonic basins, such as the North Sea (Newman and White, 1997, 1999; Bellingham and White, 2000; Hanne, 2002). In this section, fault displacements are used to estimate the strain rate in the East Shetland Basin during the Late Jurassic, and compared with present-day strain rates from central Greece derived from GPS data.

Fault throw profiles along the Ninian-Hutton-Dunlin fault system (far left), North Cormorant fault (centre left) and Brent fault (near left, redrawn from McLeod et al., 2000). The maximum extension across each fault was estimated by measuring the maximum throw and the average fault dip. The total maximum extension across the Northern and Southern Faults of the Tern-Eider Ridge, the North Cormorant Fault, the Ninian Fault and the Brent Fault is 6.85±1.4 km over a distance of 65 km. For a rift period of 28 Myr this equates to a strain rate of 3.8±0.6 Ga^-1. This assumes all the extension was accommodated on those five fault systems. In reality, these faults account for about 80% of the total extension (Scholz & Cowie, 1990, using a maximum fault length of 50 km, a minimum mapped fault length of 10 km, a value for C of 1.34 (Yielding et al., 1996), and assuming that the total displacement on a fault is linearly proportional to its total length). Fault block rotation is also ignored, but accounts for about 1.1 km of the extension (Sclater & Cèlèrier, 1989). Thus, allowing for an error in the radiometric ages of ±6 Myr, the Late Jurassic strain rate across the ESB was 5±2 Ga^-1.

This strain rate is very comparable to strain rates obtained by 2D subsidence modelling of the ESB: Bellingham & White (2000) estimated a peak strain rate of about 6 Ga^-1; Hanne(2002) estimated it at 4.4 Ga^-1. Subsidence modelling ignores the faults, and thus the estimates of strain rates obtained using this method and from the fault heaves are independent. The present-day strain rate across central Greece is between 100-150 Ga^-1, using data from GPS (Goldsworthy et al., 2002; Mattei et al., 2004). Therefore the extension in the ESB during the Late Jurassic was about an order of magnitude slower than the present-day deformation in Greece. There is an obvious difference in the timescale of the observations from the two regions in reaching this conclusion. In Greece this strain rate can only be demonstrated to have remained constant over the past 350 kyr. In the ESB the strain rate has been averaged over the entire rift period. However, there is a well established difference in the strain rates of intra-cratonic basins compared to back-arc basins using subsidence modelling. The Basin and Range, an intra-cratonic basin active today, is stretching at strain rates more comparable to the ESB than to Greece (9—16 Ga^-1; e.g. Dixon et al., 1995; Bennet et al., 1998).

Elastic Thickness of Extensional Sedimentary Basins

The elastic thickness (T_e) of the Earth's lithosphere is a contentious issue in geophysics today. The best review is provided by Watts (2001). The strength of the lithosphere can be estimated using signal processing methods on the topography and gravity field of the Earth. In this section the method of McKenzie (2003) is used to estimate T_e in three extensional sedimentary basins: the Basin and Range Province, USA, Greece, and the North Sea. The admittance method of McKenzie (2003) is preferred over the coherence method of Forsyth (1985) because it returns a true estimate of T_e rather than an upper bound.

To find out how T_e is estimated from topography and gravity, click here.

Far left Topography (GTOPO30) and free-air gravity anomaly (Hittleman et al., 1994) of the Basin and Range Province. Left Admittance and coherence between the topography and free-air gravity anomaly in the Basin and Range. Best-fit T_e is 3.4 km (solid line) ±0.6 km (dashed and dot-dashed lines). Right Misfit of the flexural model as a function of T_e and F2 (F2 is the fraction of the load which is internal and has no surface expression). Best fit model has F2 = 0.08. T_e is well constrained, but F2 is not. Densities used were 2.67 Mg m^-3 (load and upper crust); 2.9 Mg m^-3 (lower crust) and 3.3 Mg m^-3 (mantle). Upper crustal thickness was 3.2 km and total crustal thickness 20 km.

Far left Topography and free-air gravity anomaly (D. Fairhead, written communication) of Greece and surrounding region. Left Admittance and coherence between the topography and free-air gravity anomaly in Greece. Best-fit T_e is 3.4 km (solid line); max T_e is 5 km (dot-dashed line), min T_e is 2.4 km (dashed line). Right Misfit of the flexural model as a function of T_e and F2. Best fit model has F2 = 0.345. Densities used were 2.0 Mg m^-3 (load and upper crust); 2.9 Mg m^-3 (lower crust) and 3.3 Mg m^-3 (mantle). Upper crustal thickness was 5 km and total crustal thickness 25 km.

In the North Sea, there is no significant topographic signal. The load is an internal load with no topohgraphic expression. However, the geometry of this load, i.e. the sediment infill of the North Sea basin, is well constrained by the geophysical exploration of the region. In this case the admittance is measured between the load thickness and the free-air gravity anomaly, rather than between the surface expression of the load and the gravity. F2 cannot be estimated for this region because the entire gravity field is used in estimating the admittance, rather than just that part of it that is coherent with the load. This is because there is noise present in both the load and the gravity data.

Far left Cenozoic sediment load (redrawn from Surlyk et al., 2003) and free-air gravity anomaly (Sandwell & Smith, 1997) of the North Sea. The sediment thickness was reduced to a solid sediment thickness using the empirical porosity relationship of Sclater & Christie (1980). Left Admittance and coherence between the Cenozoic solid-sediment load and free-air gravity anomaly in the North Sea. Best-fit T_e is 4.6 km (solid line), but may vary between 0 km (dashed) and 6.6 km (dot-dashed). Right Misfit of the flexural model as a function of T_e. Densities used were 1.03 Mg m^-3 (water layer), 2.675 Mg m^-3 (load and upper crust), 2.9 Mg m^-3 (lower crust) and 3.3 Mg m^-3 (mantle). The upper crustal thickness was set to 7 km and the total crustal thickness to 25 km.

The elastic thickness of the North Sea since during the Cenozoic is therefore of the order of 5 km. Modelling of geotherms enables a stress profile of the North Sea lithosphere to be estimated and thus the depth at which the lithosphere begins to deform by ductile as opposed to brittle deformation:

Left Geotherms in the North Sea for two different sediment models (solid=northern North Sea; dot-dash=central North Sea, each ~5 km thick) for three different surface heat flows (in mW m^-2). Dashed lines are Moho depths. These geotherms are calculated on the basis that heat flow is 1D and steady-state. Thermal conductivity was allowed to vary with depth, being a function of porosity in the sediments, 2.5 W m^-1 K^-1 in the crystalline crust, and 3.1 W m^-1 K^-1 in the mantle. Radiogenic heat production was included: 1.5 μW m^-3 for sediments; 1.0 μW m^-3 for the crust; 0.02 μW m^-3 for the mantle. Above Right Stress profile for the North Sea using the geotherm from left with a surface heat flow of 70 mW m^-2. The upper crust is characterised by Byerlee's Law until the stress required to deform rock is low enough for ductile deformation to occur by creep processes. Three different flow laws are shown for the upper crust for three different wet quartzites (solid black line — Gleason & Tullis, 1995; dashed line — Luan & Paterson, 1992; dotted line — Kronenberg & Tullis, 1984). The lower crust is characterised by wet diabase flow law (Caristan, 1982) and the mantle by two wet dunites (Chopra & Paterson, 1984). The rock is assumed to have zero pore fluid pressure and the strain rate was set to 1 x 10^-15 s^-1. The lithosphere ceases to be strong at about 10 km depth, which is approximately the depth to the base of the elastic layer in the North Sea if the sediments do not contribute to the strength of the lithosphere and the elastic strength is concentrated in a single layer in the crust. However, the errors in the geotherm models and the experimental flow laws means that the stress profile is only consistent with T_e being a single layer in the crust.

Conclusions

These are my conclusions, lifted straight from my dissertation:

A comparison of Late Jurassic rift structures within the northern North Sea with active tectonics of central Greece has revealed both important similarities and differences between these two basins. Here, my principal conclusions are summarised, and some avenues for further research are proposed.

The first similarity between the ESB and central Greece is the role rotations about vertical axes play in the structural development of the basin. Faults, and thus the blocks they bound, rotate clockwise in central Greece in order for the regional velocity field to be accommodated by dip-slip motion on east-west striking normal faults. In addition to these rotations, gradients in extension along strike of the faults on the south side of the Gulf of Corinth means that the entire block between this gulf and the northern Gulf of Evia rotates clockwise, relative to the Peloponnese. The study of the Tern-Eider Ridge in the ESB has shown that similar rotations can be inferred to have occurred in mature extensional basins. Even in the absence of active tectonics and palaeomagnetic analyses, such rotations can manifest themselves in the geological structure of a basin. In the case of the Tern-Eider Ridge, the strong gradients of extension along the ridge-bounding faults, which caused a small rotation about a vertical axis to occur, is one model that can explain the observed structures. The Tern-Eider Ridge is not similar to the Gulf of Corinth in terms of configuration; the length scales of both are different (~30 km vs. ~100 km), as is the magnitude of rotation relative to adjacent blocks (2—5^o for the Tern-Eider Ridge compared to ~10^o in central Greece). Nevertheless, it is the principle behind the comparison that really matters: there may not be similarity of configuration between the Tern-Eider Ridge and central Greece, but there is in terms of process. Therefore, the observations of increasing extension westward along the Gulf of Corinth and the occurrence of block rotations about a vertical axis provided the impetus for explaining the structure of the Tern-Eider Ridge. Such a model does not require the vector of extension in the northern North Sea to change during the Late Jurassic, since structures previously attributed to basin inversion or strike-slip tectonics can be more simply understood as local effects arising from the interaction of fault blocks due to vertical axis rotations.

As well as rotations about vertical axes, the ESB and central Greece are similar in terms of inter-fault spacing and elastic thickness. These similarities also extend to the Basin and Range Province. A quantitative analysis of the spacing between normal faults in these three regions has demonstrated that they are similar. In each basin, there are fault-bounded blocks that have rotated coherently about a horizontal axis. The range in width of these blocks in the ESB is similar to that in central Greece (10—30 km), and overlaps with the range observed in the Great Basin (10—50 km). That there is such uniformity in block widths between these three basins is unsurprising given that the control on half-graben width has been attributed to elastic thickness. T_e in all three regions is very similar, less than 5 km. However, the controls on fault spacing are not fully understood. Although the similarity in T_e between the three basins can account for the restricted range of half-graben widths, it does not explain why half-graben form in some areas, but not others. In the ESB, central Greece and the Great Basin there are regions, up to 100 km wide, where there is little faulting and no well-developed fault blocks. The reason why these areas did not break up into regularly spaced half-graben structures is unknown, although the role of pre-existing crustal structures cannot be discounted.

The third similarity between the basins is their uniform elastic thickness. That basins should be weak is not unexpected, given that they are heavily fractured. However, the low values of T_e estimated for the Basin and Range and Greece do imply that if normal faulting is responsible for making the lithosphere weak, the lithospheric mantle cannot have much strength. The implication, therefore, is that much of the long-term strength of the lithosphere lies in, or near, the seismogenic layer. The estimate of T_e obtained for the North Sea is more surprising, as it is similar to the values obtained from the active basins. It might be expected that as a basin cools, the lithosphere should recover its strength. Such a recovery has evidently not occurred in the North Sea, which has remained weak since the Late Jurassic, and is probably due to the insulating effect of the sediments deposited during and after rifting. Modelling of geotherms in the North Sea has confirmed that the base of an elastic layer lying beneath the sediments and of the same thickness as T_e is hot enough to deform significantly by creep, although the errors in such an analysis are significant. Hence there may be a thermal control on the strength of continental lithosphere, but it could be masked by the insulating effect of basin sediment infill.

As well as the similarities described above, there are two important differences between the ESB and central Greece that have been highlighted in this dissertation. The first of these is in terms of fault migration and timing. In central Greece, activity on sub-parallel faults is inferred to migrate basinward with time, on a timescale of < 10⁶ years. A similar pattern of fault migration cannot be resolved in the ESB using the available seismic and well data, and the faults examined all appear to have been simultaneously active. However, the data used for this study have a limited resolution in time and depth, and thus the conclusion of synchronous faulting is only good to a resolution of ~10 Myr, i.e. on a 10 Myr timescale, all the faults appear to have been active at the same time. The disparity between the timescales over which the deformation can be constrained is a problem in both basins. In the North Sea, there is a lack of high resolution seismic data and biostratigraphic control, which means that it is difficult to determine the history of Late Jurassic faulting at the sub-million year timescale. Part of this limitation is inherent; the Late Jurassic rift structures are quite deep, and thus cannot be resolved at the same scale as shallower geology, and the Tarbert Formation, for example, is a sand-rich lithology and thus poorly constrained in time by biostratigraphy. In Greece, the problem is reversed: through geodesy and geomorphology, the history of faulting is known at the 10⁵ years timescale, but at longer timescales, the tectonics are less well known. Thus it is not possible to determine if the pattern of fault migration observed today is stable, or if it will be reversed in the future. Similarly, it is not known if those faults which appear to be less active at the moment are in fact 'dead' faults, are merely dormant and awaiting reactivation later on in the rift episode, or are simply still active, but much less so than those faults which currently dominate the geomorphology and seismological record. The answers to these questions await further developments in techniques and data, by which the history of deformation in central Greece can be better constrained, and the short-term processes during Late Jurassic rifting in the ESB be better resolved. Any models of faulting within extensional basins therefore need to explain faults which appear to be active simultaneously on long timescales (~10 Myr) yet migrate basinward on short timescales (~1 Myr).

The second difference between the ESB and central Greece concerns the magnitude of strain rate during extension. Not only does this difference arise if strain rates in the North Sea are obtained by subsidence modelling, but also if fault heaves are used to estimate rates of extension. The discrepancy between strain rates from the two regions is therefore robust, and is about an order of magnitude, the ESB being slower than central Greece (~10 Ga^-1 vs. ~100 Ga^-1). The present-day strain rate across the northern Basin and Range, obtained using GPS and fault-heave summation, is similar to that for the Late Jurassic in the northern North Sea, and therefore fits the pattern established from subsidence modelling, where intra-cratonic basins appear to extend more slowly than those from a back-arc setting. The reasons for this discrepancy in strain rate between extensional basins from different tectonic settings have yet to be explained.

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