Formation of continental margins

Passive continental margins form by extension of the lithospheric plate. The growth of margins determine the stretching factor. The amount and rate of stretching determine the temporal and spatial variation of crustal and lithospheric thinning across the margin, which in turn controls variations in heat flow, subsidence, extensional faulting and decompression melting. During thinning, the base of the lithosphere is advected upwards and heat flow to the surfacerapidly increases. At the same time, hotter asthenophere is passively moved upwards, and generates a thermal anomaly. When stretching stops, the thermal anomaly. Within the upper crust, thinning is shown by rapid, fault-controlled subsidence. When stretching stops, the thermal anomaly decays exponentially with time, heat flow decreases and a phase of thermally driven subsidence occurs.

From a thermal perspective we can divide margins into three broad categories. Note! We understand under the hot margins –Volcanic rifted margins, and under cold margins - Non-volcanic rifted margins.

The first, are ‘hot’ margins, where lithospheric thinning has taken place over a upwelling mantle plume. The best examples occur in the North Atlantic Ocean on either side of the Iceland Plume. But ‘Hot’ margins present difficulties for deep-water exploration. The presence of high-velocity lava flows, which scatter acoustic energy, considerably impedes our ability to image underlying sedimentary strata. It is also difficult to accurately constrain thermal histories because the spatial and temporal distribution of hot molten rock, which advects heat, is not easy to determine with accuracy.

The second are ‘cold’ margins, where there is little evidence for magmatism until new oceanic crust has formed. The best-example margin occurs west of the Iberian Peninsula. This type of margin is more widespread than ‘hot’ ones.

The third category of margin lies between these two mentioned hereinbefore. At these margins, there is no convincing evidence in favour of a mantle plume, even though high velocities consistent with magmatic under plating occur in the lower crust. These ‘warm’ margins occur along the east coast of North America.

Subsidence history

For determing the oil and gas potential of a basin and estimation reservoir porosities can be used subsidence history when we plot subsidence/uplift and paleowater depth as a function of time. Subsidence history curves from a number of locations can also be used to construct paleostructure maps at specific time slices. Combined with information on thermal maturity, this can be a powerful tool in evaluating the timing of oil migration and likely migration pathways in relation to the development of suitable traps.

Fig. 3

When generating a subsidence curve for a basin, it is important to remember that the basin floor and its fill will have subsided in response to more processes than just tectonic stretching and the thermal decay that follows lithospheric thinning. Tectonic subsidence will be added to in particular by

1) the weight of the sediments replacing water or air in the basin, which pushes the basin floor down, and

2) the weight of the sediments in the basin causing compaction of the older sediments below them.

Analytical techniques that take account of these factors are called decompaction and backstripping.

Fig. 4 Fig. 5

Using data from drill cores, the porosities of various types of sediments can be observed to decrease with increasing burial. This is partly because, like a sponge being squeezed, sediments become compacted as the weight of material on top of them pushes more and more pore fluid out of the space between the solid grains, or clasts, that they are also made up of. Usually the decrease becomes milder with depth, as clasts start to press together and support the weight of the overburden, and as diagenetic processes lead to cements forming between them. (Fig. 2)

Decompaction is a mathematical process that conceptually undoes this squeezing, restoring lost pore space and thus increasing the volumes of sedimentary units.

Porosity-depth curves (Fig.3) have exponential shapes. They can be described mathematically using an expression of the form

So λ and c are numbers that can be used to describe the rate at which exponential compaction occurs. φ₀ is the porosity at the surface, and φ the porosity at depth z.

W e define c simply so as not to have an unwieldy term in the brackets after the exponential function.

We have just seen that porosity

T o figure out the total volume of pore space in a sediment unit with top and bottom surfaces at burial depths z₁ and z₂, we need to integrate the porosity expression between those depths:

This evaluates as

Fig.4 Fig.5

In a rock column of unit area, this is equal to the thickness made up of pore space, z_pores.

Since the total thickness of this sediment (z₁-z₂) is made up of

s ome clasts and some pores, the thickness made of clasts is

We can use these expressions for z_clasts and z_pores to model the changes in the basin and its fill that occur when material is removed from above it: this kind of modelling is a decompaction analysis.

In each step of a decompaction consideration, the top package of sediments is removed from a basin. The basin has a new top surface and so the depths of all the underlying units become shallower. Also, for each unit, z_pores increases as the reduced lithostatic pressure allows water back into the sediment. The new depths of the top and bottom of the unit are z’₁ and z’₂.

T he decompacted thickness is thus equal to the unchanging clast thickness plus the increased total pore thickness:

The decompaction analysis gives a good picture of how sediments have accumulated within the basin, but its results are not immediately interpretable in terms of tectonic subsidence.

Look what an effect sedimentation has. First, let’s think of the basin floor as starting off at sea level before stretching happens.

As we have seen, stretching often causes subsidence. Not all of the sediment that accumulates is due to this subsidence because the weight of accumulated sediments depresses the lithosphere into the fluid mantle below. We can take account of this effect by assuming that the basin is either in local isostatic equilibrium, or that the sediment load is supported by lithospheric flexure, or some combination of the two.

Fig.6 Fig.7

Now let’s think about what happens if the same amount of stretching occurs, but when the basin floor was well below sea level before that stretching started. Some accommodation space exists before stretching.

So here, even after an isostatic or flexural correction for the sediment loading of the lithosphere during the stretching phase, it is wrong to assume that all of the sediment that accumulated during tectonic stretching did so owing to stretching. To take this into account, some estimate of paleobathymetry must be applied as a correction for the decompaction analysis. This estimate is usually made on the basis of paleontological analyses.

A final correction to make is one for the effects of global sea level changes (eustatic changes) during the sedimentation of the basin fill. Eustatic changes may have many causes, including global warming (which causes the ocean water to expand), ice cap fluctuations, and seafloor spreading rate changes (which affect how much of the ocean floor globally is shallow or deep).

T

Fig. 8

he details of eustatic variations are poorly understood, and usually only the long-term changes are applied as a correction.

Interestingly, eustatic curves like these ones are based on studies of sediment accumulation in basins for which the stretching factor and paleobathymetry are thought to be very tightly constrained. The thermal subsidence phase of a basin’s life is most useful, as this should have a smooth predictable shape for subtracting from accumulation signals seen in borehole data but also in seismic stratigraphic features like marginal onlap and offlap. With compaction, isostatic and paleobathymetric effects also removed, the remaining accumulation signal can be related to eustatic sea level changes.

T he curves on the right are based on data from three separate wells in basins on the NE coast of the USA.

Once corrections have been made for compaction, isostasy/flexure, paleobathymetry, and long-timescale eustatic changes, the subsidence profile that remains can more confidently be referred to as the result of tectonic stretching, and a β factor determined.