Salt TectonicsThe following notes are abbreviated versions of Mark Rowan’s

three-day industry short course on salt tectonics.

Diapirs and diapirismDiapir initiation during differential

loadingDiapir initiation during extension or

contractionActive and passive diapirismReactivation of diapirs during extension or contractionDiapir interiors and marginsNear-diapir deformation - folding and

faultingTectonic styles of salt deformation

Thick-skinned extensionCollisional mountain belts

Introduction

Distribution and origin of salt basins

Mechanics of salt deformation

Salt withdrawal structures and weldsTurtle structuresExpulsion rolloversWelds

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INTRODUCTION

There has been an enormous revolution in our understanding of salt tectonics in the past decade or so (see Jackson, 1995, for an excellent history of salt tectonics research). The beginnings of the revolution date back a little farther within some of the exploration companies, but their ideas only became public starting in about 1989. Our increased understanding of the geometry and evolution of salt bodies and associated strata is due in large part to the fortuitous convergence of advances in four areas:

Seismic imaging. There has been a steady improvement in seismic data acquisition and processing over the years. But with the advent of such techniques as pre-stack depth migration, images of salt bodies became much clearer, with improved pictures of the bases of salt sheets and overhangs and the steep flanks of manydiapirs (e.g., Ratcliff, 1993).

Experimental and numerical modeling. Attempts to model salt deformation have been made for many decades (e.g., Nettleton, 1934), but until fairly recently, both salt and its overburden were modeled as viscous fluids. Starting in the late 1980s, however, B. Vendevilleand coworkers started modeling the overburden as a brittle material, more in keeping with the known mechanical behavior. The results demonstrated salt’s more passive role of reacting to, rather than causing, deformation (e.g., Vendeville and Jackson, 1992a, b) and fundamentally changed the ways most people look at salt deformation.

Structural restoration. The technique of cross-section restoration was first applied to salt structures in the late 1980s (e.g., Worrall andSnelson, 1989). In the past decade, numerous people have used restoration to reconstruct the history of salt movement and associated deformation of surrounding strata.

Field studies. Armed with new ideas, various researchers have reexamined exposed salt basins throughout the world, leading to improved understanding of the processes of salt-related deformation.

In this course, we will concentrate on the new ideas about salt tectonics. Many of the illustrations used here are examples of the four areas of advance listed above. Much of the work has been concentrated in the northern Gulf of Mexico, but the impact of the advances has spread to salt basins worldwide. Thus, we will also examine the geometries and structural styles of salt from such places as the North Sea, the Red Sea, offshore West Africa, offshore Brazil, the Precaspian Basin, and onshore Mexico.

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DISTRIBUTION AND ORIGIN OF SALT BASINS

Salt basins are found throughout the world (Fig. 1), but a quick look will show that they occur primarily in rift basins and along passive margins, as well as in their deformed counterparts, such as the Alpine/Himalayan system. According to a review by Jackson and Vendeville (1994), many salt deposits were formed during the earlypostrift phase, including the basins of offshore Brazil, offshore West Africa, the U.S. Gulf Coast, and the Red Sea (Fig. 2). Others were formed either during rifting or during lulls between distinct rift episodes, for example many of the basins on either side of the northern Atlantic Ocean (Fig. 2). Finally, a few salt basins appear to be older than rifting, namely those in the North Sea and Persian Gulf (Fig. 2).

Many rift basins have a similar history, one that is conducive to evaporite deposition. They form during extension of the continental crust, and grabens are initially filled with nonmarine clastics because of the high heat flow and associated regional uplift during rifting. Subsidence of the grabens, either during rifting or, more typically, during postrift thermal and loading subsidence, leads to marine incursion. If the climatic conditions are appropriate, it is during the transition from nonmarine to marine environments that evaporites are formed (often episodically). In the typical (but not universal) scenario, continued subsidence leads to true marine conditions.

Rift basins have a distinct basement architecture made up of grabens and half-grabens segmented by transverse structures such as accommodation zones or transfer faults (Fig. 3). An example from the Brazilian margin is shown in Figure 4, where the rift system includes a series of NW- 3

SE oriented transfer faults. The rift geometry, coupled with the rate of sedimentation and the relative time ofevaporite formation, determines the areal extent and thickness distribution of evaporites. Salt may be restricted to individual half-grabens (Fig. 5 it may be regionally tabular (Fig. 6), or it may have an intermediate geometry, with a regional distribution but significant thickness variations (Fig. 7).

Salt commonly occurs in paired basins on either side of oceanic crust, such as across the Gulf of Mexico, the South Atlantic, and the North Atlantic. Thus, it has generally been thought that evaporite deposition occurs only on continental crust, with subsequent oceanic spreading separating a once contiguous basin into two parts (Fig. 8). Many passive margins have seaward-dipping reflectors, orSDRs (Fig. 9) – wells show that these consist of subaerial basalts that are considered the initial expression of oceanic spreading. Autochthonous salt in parts of offshore Brazil occurs at a stratigraphic level above the SDRs, leading to a model in which salt deposition postdates the onset of oceanic spreading (Fig. 10). Salt in the South Atlantic (offshore West Africa or Brazil) is interpreted to occur above the breakup unconformity and thus is underlain by a combination of landward continental crust and basinward oceanic crust (Fig. 11). There is then a transition to shallow-water carbonates and then deeper-water facies. The interbedded nature of the salt-carbonate transition and the similar seismic velocities means that there is typically a low acoustic impedance contrast and thus no good top-salt reflector for the autochthonous salt layer. In contrast, the salt is usually in contact with underlying clastic redbeds or basement so that there is a good base-salt reflector.

The evolution for the northern Gulf of Mexico is shown in Figure 12. Initial rifting during the Early Jurassic resulted in the SSW movement of Yucatan away from N. America, whereas oceanic spreading resulted in southern movement and rotation about a nearby pole. This means that basement structures will have different orientations in the thinned continental crust and the oceanic crust. The boundary between the two is interpreted to be in the approximate position of the present-day shelf edge, so that most of the deepwater province is underlain by oceanic crust (Fig. 13). So instead of one salt basin on continental crust that was subsequently split by spreading (Fig. 12), there were most likely two salt basins with the downdip edges defined by the incipient spreading center (Fig. 13).

Although most salt basins are closely associated with rifting, salt deposits can form anytime conditions are appropriate for evaporite formation. Thus, any restricted basin with an arid climate is a potential salt basin. The modern examples are the sabkha deposits of the Arabian peninsula, but these are rare in the geologic record and, similarly, the types of salt basins common in the past are not observed today (Warren, 1999).

One ancient setting for evaporite deposition is a basin with an open marine connection that gets closed off. This will lead to the development of various evaporite facies as the basin essentially dries up (Fig. 14). Examples include basins with narrow entrances that become emergent during major sea-level drops, such as the Mediterranean (Messinian salinity crisis) and the Red Sea. Another example is when plate tectonic motions close off a basin, as in the case of the Precaspian Basin during the Permian (Fig. 15).

The Precaspian Basin example nicely shows the kinds of vertical and lateral facies variations commonly found inevaporite basins. Although salt dominates the basin center, there is also interbedded anhydrite that becomes more dominant towards the northern margin. There can also be significant components of both carbonates andsiliciclastics that are concentrated along basin margins. Another example is provided by the central North Sea, where non-salt facies in the proximal part of the basin disappear towards the center (Fig. 16). Of course, the presence of other lithologies means that there can be significant, even coherent, reflectivity within the evaporite layer (Fig. 17).

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Figure 1. Locations of salt basins around the world (Jackson and Talbot, 1991). 5

Figure 3. 3-D block diagram of basement rift architecture with offset grabens separated by an accommodation zone (Stonely, 1981). Note that more asymmetric rift geometries are also possible. 6

Figure 5. Example from east Africa where the salt is isolated within a single half-graben (Malek-Aslani, 1985).

7

Figure 6. Example from west Africa, where salt is high enough in the section to form a regionally continuous, tabular salt body (Perrodon, 1981). 8

Figure 7. Model for the onshore northern Gulf of Mexico, where the salt is continuous but of highly variable thickness because it partially fills in the rift-basin architecture (Adams, 1989). 9

Figure 11. Regional cross section across Angolan margin showing salt deposition above both continental and oceanic crust (Jackson et al., 2000).

10

Figure 12. Present-day geometry of the Gulf of Mexico and three reconstructions to the Early Jurassic (beginning of rifting), Late Jurassic (onset of seafloor spreading), and Early Cretaceous (end of seafloor spreading) (Pindell et al., 2000). 11

Figure 13. Map showing distribution of oceanic and continental crust in the Gulf of Mexico – note that most of the deepwater province is underlain by oceanic crust (Reitz, 2001). 12

13

14Figure 16. Facies distribution within the Zechstein of the central North Sea (Stewart and Clark, 1999).

Figure 17. Undeformed and deformed examples of the Zechstein showing interbedded facies and variations in mobility (Stewart and Clark, 1999). 15

level of neutral buoyancy. This might be valid if the overburden also behaved as a weak viscous fluid, but in fact, the overburden is a brittle material with real strength (Figs. 18 and 20). The under-consolidated nature of shallow sediments in places like the northern Gulf of Mexico may suggest a viscous nature, but this is incorrect. In fact, there are plenty of fault scarps at the sea floor that attest to the brittle style of deformation of even very young sediments.

Vendeville and Jackson (1992a) argued that the strength and brittle nature of the overburden means that the density contrast plays only a limited role in salt tectonics. Instead, they showed that salt should be viewed as a pressurized fluid and that it is the differential fluid pressure that drives salt flow (Fig. 22). There are three key messages in this figure: (1) density is only a second-order factor – the scenario is similar whether the overburden is of lesser, equal, or greater density than the salt; (2) if there is a differential load on the fluid (salt) and there is a place for it to go, salt will flow; and (3) conversely, the salt cannot just push its way into the overburden because of the overburdens strength. In other words, for salt withdrawal and diapirism to occur: (1) there must be an open path to a near-surface salt body; (2) the overburden must be thin and weak (faulted) enough for the differential fluid pressure to overcome the overburden strength, in which case it willdeform the thin overburden; or (3) some other process (e.g., tectonic) must create space. Put another way, salt does not drive salt tectonics; instead, salt merely facilitates and reacts passively to external forces. This is a key point that underlies almost all subsequent discussion.

MECHANICS OF SALT DEFORMATION

Rock salt is a very different material from other, more typical sedimentary rock types. There are several key factors that play roles in dictating its behavior. First, salt is much weaker than other lithologies under both tension and compression (Fig. 18). Even overpressured shale almost always has more strength than salt. In fact, the curve for wet salt in Figure 18 falls essentially on the axis of zero strength. The reason is that salt deforms as a viscous material that effectively flows, with flow rates up to 15m per year in exposed diapirs in Iran (Talbot et al., 2000). Flow is by a combination of Poiseuille flow due to overburden loading and Couette flow due to overburden translation (Fig. 19). The viscous nature of salt means that it forms a constant-strength, albeit very weak, layer between normal sedimentary layers whose strength increases with depth (Fig. 20). Thus, salt serves as an excellent detachment surface into which faults sole.

Second, salt has a constant density of about 2.2 g/cm3, irrespective of burial depth (it actually get less dense with depth due to temperature effects). This is in contrast to other sedimentary strata, such as sands and shales, that become compacted during burial and thus become more dense (Fig. 21). Thus, salt is more dense than its surrounding strata when it is near the surface, but is less dense once it is buried beneath 1000-1500m of sediment.

The relatively low density of buried salt was historically used as a rationale for models in which salt punches its way through more dense overburden until it reaches its

16

Figure 18. Strength of various rock types in both tension and compression (Jackson and Vendeville, 1994). Note that wet salt falls effectively on the axis of zero strength because it is a viscous material.

17

Figure 20. Simple 3-layer model of the crust with a weak, but constant-strength, salt layer between two brittle layers whose strength increases with depth (Vendeville and Jackson, 1993). 18

SALT-WITHDRAWAL STRUCTURES AND WELDS

Differential loading induces a differential fluid pressure that drives salt withdrawal and minibasin formation. The initial load may have been induced by a variety of means: extension of the salt, contraction, emplacement of a depositional lobe, etc. – basin initiation will be discussed later. In any case, salt displaced from beneath the “protobasin” moves laterally into flanking areas (Fig. 23). The dynamic nature of the salt flow results in bathymetric highs adjacent to the minibasin. This sets up a feedback process where the minibasin is a low that receives further sediments; this in turn increases the differential pressure, driving further salt withdrawal and flow into the flanking highs (Fig. 23). The process continues until the suprasalt minibasin touches down on the subsalt strata, forming a salt weld (indicated by pairs of dots).

As the minibasin grows, the load increases compared to the thin overburden above the flanking highs. Thus, withdrawal and subsidence rates gradually increase through time (Fig. 24). Note that subsidence is not driven by the sediment added during any short time interval, but by the net differential load that exists at any time. Thus, the existing basin drives salt flow – active sedimentation just adds to the differential load and is thus a second-order effect. Eventually, as the source layer thins, viscous drag forces inhibit lateral flow of salt, and subsidence rates slow even though there is a large differential load. Minibasin subsidence ceases once the weld forms.

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Figure 23. Evolution of a minibasin subsiding into salt: (a) initiation due to extension, shortening, deposition of a sand lobe, etc.; (b) subsidence and growth of basin even in the absence of sediment input because of the differential load on the salt; and (c) cessation of subsidence once weld forms. 20

Turtle structures

Although the center of a minibasin will stop subsiding once the weld forms, the flanks, which are still underlain by salt, may collapse, forming new, flanking depocenters and inverting the original depocenter into a turtle structure (Figs. 25 and 26). It is uncertain why someminibasins collapse symmetrically to form turtles and others don’t collapse at all. It may have to do with the geometry of the wedge of strata above the flanking salt: if it is relatively long and thin, it is mechanically easy to fold and thus collapse; if it is a short, thick wedge, it will not fold as easily. Another possibility is that extension plays a role. Experimental models with no lateral movement resulted in simple minibasins, whereas extension helped flank collapse and the formation of turtle structures (Fig. 27).

A classic turtle structure from the Precaspian Basin shows the weld (with some remnant salt), the initial depocenter that is now inverted, the flanking depocenters, and the adjacent salt diapirs (Fig. 28). Note that crestal faulting due to bending of the strata is a common secondary feature. Subsidence in this example started right after salt deposition, as shown by the variable thickness of the oldestsuprasalt strata. Moreover, the turtles in this basin are linear features along and parallel to the basin margin (Fig. 29), suggesting that extension may have played a role. In other cases, there may be a prekinematic section that represented condensed sedimentation on an inflated salt high before collapse and formation of the initial basin. An example from the northern GoM (Fig. 30) shows that initial minibasin subsidence did not start until probably the

Paleogene, on the order of 100 million years after the salt was deposited. Two maps from a nearby area show how salt was inflated from the Late Jurassic through the Oligocene (reds and yellow in Fig. 31a) and then collapsed to form turtle structures in the Miocene (blues in Fig. 31b).

The age of a turtle structure is defined here as the age of touchdown of the initial basin and the start of flank collapse (i.e., the yellow horizon in Fig. 28). This age can vary significantly across a basin, even between adjacent turtles. The timing is dependent on when subsidence began, the sedimentation rate, and the initial salt thickness. Thus, one cannot simply correlate between turtles. Moreover, the presence of a turtle tells us nothing aboutfacies: there are, of course, turtle structures that have reservoir-quality sands (e.g., Thunder Horse in the GoM), but there are also turtles around the world where the strata forming the turtle are composed dominantly of shale, carbonate, or even evaporites. Generally, once a basin is subsiding into salt, it will keep subsiding regardless of the depositional setting and the minibasin will be filled with whatever sediment is available, including slumps off the adjacent highs.

21

Figure 25. Formation of a turtle structure by flank collapse and inversion of a minibasin whose center has touched down and formed a weld.

22

courtesy of B. VendevilleFigure 26. Experimental model of a turtle structure (courtesy of B. Vendeville).

23

(a) (b)

Figure 27. Model results of minibasin subsidence showing no turtle formation in the absence of extension (a) and turtle formation triggered by extension (b) (courtesy of B. Vendeville). 24

courtesy of B. Vendeville

Weld

Crestalfaulting

Initial depocenter

Flank collapseFlank collapseSalt Salt

Figure 28. Example of a turtle structure from the Precaspian Basin, with an inverted central depocenter above a salt weld and flanked by younger depocenters and adjacent diapirs. Note the crestal faulting and erosion. Also note that subsidence began immediately after salt deposition.

25

20 kmSW NE

metersFigure 30. Example of a turtle structure from the northern Gulf of Mexico with the same features as the previous figure, except that there is a prekinematic sequence representing condensed sedimentation above a salt high prior to eventual collapse and minibasin formation (data courtesy of GX Technology). 26

Expulsion rollovers

Another type of withdrawal structure is the so-called expulsion rollover, which is essentially a half-turtle. In expulsion rollovers, the initial basin touches down and welds out, just as in turtles, but flank collapse is then asymmetric, so that the depocenter shifts progressively in one direction, forming a growth monocline (Fig. 32). As the weld grows in length, salt is displaced basinward, where it inflates and lifts a condensed overburden.

An example of an expulsion rollover from the Precaspian basin is shown in Figure 33 – the yellow horizon marks when the weld first formed and minibasin growth shifted from the central depocenter to the flank collapse as salt was driven into an allochthonous body. Another purported example is provided by the famous Cabo Frio structure in offshore Brazil (Fig. 34), whose proposed evolution is shown in Figure 35. However, regional patterns of extension and contraction in the Santos Basin suggest that the Cabo Frio structure is actually a landward-dipping normal fault, where the footwall moved basinward to create the accommodation.

An important point is that progradation must occur early in the basin history if it is to drive salt movement. Whenprogradational geometries form above a thin prekinematic section, the differential load is great enough to drive subsidence and inflation (Fig. 36a). If progradation is late, however, the net load differential is not as large and the overburden is thicker and stronger, so that no deformation takes place (Fig. 36b).

27

Figure 32. Model of expulsion rollover structure, in which progradational deposition results in a progressive shift of depocenters and the underlying weld, while salt is displaced into a distal,inflating salt plateau (Ge et al., 1997). 28

Allochthonoussalt

Weld

Thinningonto weld

Thickeningonto weld/salt

Suprasalt

Subsalt

Presalt

Figure 33. Expulsion rollover structure in the Precaspian Basin, showing the characteristic change from thinning onto the weld to thickening onto the weld/salt. The transition marks the timing of initial basin touchdown and welding of the salt layer. 29

Figure 34. Two sections through the Cabo Frio “fault” zone, offshore Brazil, and plots of dip variation supporting interpretation as a salt-withdrawal feature (Ge et al., 1997).

30

31Figure 35. Model evolution of the Cabo Frio structure: progressive evacuation inflates distal salt, which eventually evolves into a diapir that gets buried (Ge et al., 1997).

Welds

Welds may have variable geometries. They can form along the original, autochthonous salt layer when the subsiding overburden comes into contact with the subsalt strata (Fig. 37) – a nice example of a so-called ‘primary’weld is shown in Figure 38 – or they can be inclined due to evacuation above a dipping base salt (Fig. 39).

Welds are not always obvious, as shown in Figures 40 and 41 – they are identified by discontinuous, high-amplitude reflectors, often with angular discordance between strata on either side (but this can also reflect an unconformity oronlap surface). The seismic character reflects the fact there are pods of remnant salt along the weld, which are there because the top and base of salt do not have perfectly matching geometries. Ultimately, correctly identifying and interpreting welds requires a good mental image of the three-dimensional salt geometry and its evolution over time.

Welds can also be vertical (or ‘secondary’), formed by the squeezing of a salt wall in response to updip extension (Fig. 42). A GoM shelf example is shown in Figure 43, where a vertical weld is indicated by a teepee structure beneath the landward edge of an allochthonous (so-called ‘tertiary’) weld. However, there are many cases where vertical welds are over-interpreted. Teepee structures can form along strike-slip faults or normal faults that have been reactivated during shortening, and many apparent vertical welds are simply migration artifacts below the edges of overlying salt bodies and minibasins.

An exposed vertical weld in La Popa Basin, Mexico is illustrated in Figures 44 and 45. It is 25 km in length but is a true weld only over the southeastern half. To the west, it consists of continuous remnant evaporite 100-200 m thick, and in between, it consists of patchy evaporite separated by true welds. However, this would not be known from a seismic profile because of the difficulty in imaging steep structures.

Another exposed weld, this time from the Flinders Ranges in South Australia, is shown in Figure 46. It extends upward from a triangular diapir pedestal formed above the autochthonous salt layer, and separates two minibasins with very different thicknesses and facies. The weld has remnant sandstones along it that were originally deposited within the evaporite sequence, and is bordered on one side by a shale sheath.

32

Figure 37. Formation of primary salt weld as the overburden subsides into the autochthonous salt and comes into contact with subsalt strata (Jackson and Cramez, 1989).

33

Figure 38. Time-migrated seismic profile from offshore Brazil showing the primary weld (Mohriak et al., 1995).

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Figure 40. Time-migrated 3-D seismic from the Louisiana shelf with an allochthonous weld indicated by discontinuous, high-amplitude reflectors and structural discordance.

35

Figure 41. Interpretation of Figure 28 showing weld geometry.36

Figure 42. Model for the development of a vertical (secondary) weld as extension on the normal fault is accommodated by squeezing of the salt wall (Jackson and Cramez, 1989).

37

DIAPIRS AND DIAPIRISM

It used to be thought that density contrast alone was responsible for the initiation and growth of diapirs – the idea was that salt, once it became buried deeply enough to create a density inversion, first bulged into a salt pillow and then punched through to the surface. Instead, we now know that density is a secondary factor and that diapirs are triggered by a variety of mechanisms.

Diapir initiation during differential loading

Simple differential loading can lead to the formation of diapirs. We have already seen how the formation of turtle structures and expulsion rollovers inflates adjacent salt and triggers diapirism (e.g., Figs. 25 and 27). Another, similar process is shown in Figure 47, where progradational loading causes localized inflation where the salt thins over basement steps. A possible example may be in southeastern Mississippi Canyon, where many of the diapirs are located above apparent basement steps (Figs. 48 and 49) and thus may have been triggered by differential loading and local inflation.

38

Figure 47. Model in which the original salt thins abruptly over basement steps; these serve as nucleation points for salt inflation and diapir growth (Ge et al., 1997). 39

WSW ENE

Figure 48. Strike line in southeastern Mississippi Canyon (GoM) showing the Louann salt stepping down, presumably over basement steps (Rowan et al., 2000a; data courtesy of WesternGeco).

40

Diapir initiation during extension or contraction

Analyses of salt basins combined with experimental models suggest that many diapirs are initiated and grow during regional extension. The models show that diapirs may go through three evolutionary stages: reactive, active, and passive diapirism (Fig. 50). At this stage, we will examine only reactive diapirism.

When a salt layer is buried by constant-thickness strata, nothing will happen (even if the overburden is more dense) until some external force is applied. In the case of extension, the overburden is lengthened and thinned, which is accommodated by graben formation at the surface and an “inverse graben” at the salt-sediment interface (large-scale boudinage). Salt reacts to the extension by overall thinning and by filling the space in the inversegraben (Fig. 51). The result is a triangular diapir (elongate in the strike direction) that has flanking growth faults that get younger toward the diapir crest. The diapir grows in size as extension progresses. An example from the northern GoM is shown in Figure 52 – again, note the triangular shape, the overlying seafloor graben, and the increase in fault age down the diapir flanks. The diapir’s position at the landward (extensional) edge of asubhorizontal salt tongue is also characteristic. Restoration shows how this diapir formed as the tongue overburden moved basinward above the salt, pulling away from the deeper minibasin to the north (Fig. 53). The width of the salt at any given stratigraphic level shows how much extension is hidden in the salt.

A variation of reactive diapirism occurs in wrench-fault settings. Pull-apart basins form where there is a releasing bend or step-over in strike-slip faults. These are the sites of very rapid thinning and are ideal locations for the generation of reactive diapirs (Fig. 54). An example of a shallow diapir at a releasing bend on the edge of anallochthonous sheet is shown in Figure55.

Contraction can also initiate diapirism. When shortening occurs above a salt décollement, the overburden in growing detachment folds may be uplifted, faulted, and eroded enough so that salt can break through to the surface and subsequently grow as a diapir (Fig. 56). It is important to stress that this mechanism of diapir initiation, likeprogradation and unlike extension, can only happen when there is a thin overburden. Otherwise, the differential load is insufficient and the thickness and strength of the overburden is too great and only a fold will form (Fig. 57).

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Figure 50. Model of diapir initiation and growth during extension (Vendeville and Jackson, 1992a): (1) necking down of the overburden creates a reactive diapir; (2) once the diapir is tall enough and the overburden is thin and weak enough, the diapir will actively punch through to the surface; and (3) once at the surface, the diapir will grow passively as surrounding minibasins subside into and displace the salt in the source layer. 42

Figure 51. Steps in the evolution of a reactive diapir with no synkinematic sedimentation (Jackson andVendeville, 1994). The diapir widens and gets taller with time, and the faults get younger toward the crest of the diapir (note that it is a passive diapir by step (e)). 43

Figure 52. Example of a reactive diapir from the northern GoM, showing the characteristic triangular shape (it is elongate in the 3rd dimension), overlying graben, and increasing age of faults down the diapir flanks (Rowan, 1995). 44

Salt

Basement

Cover RegionalA

arbitrarypin

arbitrary erosion surface

B

C

Figure 56. Formation of a diapir by erosion and salt breakthrough in a contractional detachment fold (Coward and Stewart, 1995). 45

Active and passive diapirism

In most of the models for diapir initiation that we have examined, a salt body starts inflation or growing at some point after it has been buried beneath an overburden. Thus, two more stages of diapirism are typically involved in subsequent growth. First, once the overburden becomes thin and weak enough and/or the pressure differential in the salt is great enough, the diapir will break through to the surface as an active diapir. This is equivalent to the classic model of piercement diapirism long thought to be dominant in salt tectonics, but is a brief episode in thediapir history that occurs only when the overburden is thin and weak.

If the salt does break through to the seafloor, it then enters the stage of passive diapirism, in which it continues to grow as long as there is adequate salt in the source layer. This is equivalent to the old concept of “downbuilding”(Barton, 1933), where a diapir keeps its crest essentially at the sea floor as the surrounding strata subside into the source layer (Figs. 58 and 59). It is important to note that once the source layer is depleted, the diapir ceases to grow and is buried by further sedimentation. This is because, despite the density contrast, there is no longer a differential fluid pressure to drive salt flow.

It is increasingly clear that passive diapirism is the dominant style of diapir growth in basins throughout the world. Diapirism may get triggered by various means, for example extension (reactive diapirism), contraction, or differential loading. But this usually happens early in the history, beneath a relatively thin overburden, and salt soon breaks through to the surface. Thus, there is typically an initial phase of diapir formation, a very brief stage of active diapirism, and then a long-lived history of passivediapirism (as long as there is still adequate salt in the source layer to maintain diapir growth at the sea floor).

46

Figure 58. Passive diapir from the northern GoM growing at the sea floor with effectively no overburden and little deformation of adjacent strata (Rowan, 1995).

47

Figure 59. Restoration of a passive diapir showing how it stays at the sea floor as flanking minibasinssubside into the source salt layer (Worrall and Snelson, 1989).

48

Figure 61. Serial cuts through a model showing different stages in the collapse of a diapir(Vendeville and Jackson, 1992b).

49

Figure 63. Examples of two diapirs with significant thinning and upturn of flanking strata, interpreted as a result of shortening of the diapirs at the toe of an allochthonous salt sheet (Rowan, 1995).

50

Diapir interiors and margins

Diapirs have complex internal deformation resulting from the flow of the evaporite and any interbedded lithologies (Figs. 66 and 67). Typical features include vertical lineations, isoclinal folds, pinch-and-swell structures, etc. Exposures in such areas as La Popa basin (Mexico), the Flinders Ranges (Australia), and the Zagros Mts. (Iran) show that exotic clasts are comprised exclusively of lithologies originally deposited within or intruded into the evaporite layer (Fig. 68). There are no blocks of strata plucked from the walls of diapirs rooted in the autochthonous layer (we will discuss allochthonous salt later). Although very interesting, we will not address the internal characteristics of diapirs any further in this course. Instead, we are concerned here more with the external geometry of salt bodies and the deformation in surrounding sediments.

The tops of many diapirs have a zone of caprock (Fig. 69), typically consisting of anhydrite, calcite, and minerals such as pyrite and barite. This is usually interpreted as the insoluble residue after the halite has been removed by dissolution. A plot of caprock thickness versus diapir burial depth shows that caprock is best developed in the upper 3000-5000 ft. of the earth’s surface (Fig. 70). This is because dissolution is a consequence of the circulation of meteoric water, which dissolves and carries the salt away and is constantly replenished. Deeper diapirs do not feel this circulation. Likewise, caprock is rare to absent in deepwater environments. One explanation is that the surrounding water is already very saline; another is that condensed muds covering diapirs inhibit the transport away of any dissolved salt (Fletcher et al., 1995). Having said that, there are brine pools known in the deepwaterGoM, so at least some dissolution is occurring.

Another known feature of diapir margins is the so-called shale sheath that is found as a flanking skirt around the deeper portions of some diapirs (Fig. 71). This typically consists of a thin zone of overpressured shale that is older than onlapping, more shallow-water facies. It was traditionally interpreted as fault gouge formed as the diapir punched through the overburden, but we now recognize it as a remnant of the condensed section found on top of many salt bodies. This overburden is condensed and mud-rich because it forms on the bathymetric highs abovediapirs, and ends up as shale sheath as the diapir flank collapses during minibasin growth (Fig. 72).

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Figure 66. Aerial photograph of exposed salt diapir in the Great Kavir desert of Iran, showing complex internal deformation of evaporite layers and folding of surrounding strata (Jackson et al., 1990).

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Figure 67. Cross section through onshore Texas diapir showing complex folding of internalstratigraphy (Talbot and Jackson, 1987). 53

L

M

Figure 68. Photograph of El Papalote diapir in northeastern Mexico, consisting of gypsum caprock with exotic blocks of Upper Jurassic limestone (L) and metaigneous rocks (M) (courtesy of R. Goldhammer). 54

Figure 72. Explanation of shale sheaths as simply representing preserved portions of the original condensed section that was deposited on top of the salt body when it was at the sea floor (Worrall and Snelson, 1989). 55

Faulting

Diapirs have classically been associated with radial faults (Figs. 94). However, this interpretation of near-diapir deformation was driven in large part by the idea of active intrusion of diapirs. Although modern seismic data and field exposures show that there are certainly radial faults (Figs. 95 and 96), the situation is more complicated. Many radial faults actually curve to become roughly tangent to the diapir face, producing cuspate salt outlines (Fig. 97). The irregular, cuspate-lobate plan-view geometry of diapirs increases with depth where there is greater displacement on the salt-intersecting faults. Moreover, many faults patterns around diapirs are dominated by one or more fault trends (Figs. 98 and 99).

In effect, there are two classes of faults associated with diapirs. The first comprises the large growth faults and associated smaller-scale faults that form due to salt withdrawal or basinward translation of the overburden during gravitational failure of a margin. These have distinct trends that often intersect at diapirs because the salt, as the weakest part of the section, localizes the deformation, and they typically curve to become tangential to the salt edge. The second class comprises the radial faults, which typically extend a relatively short distance away from diapirs. These are a consequence of three-dimensional folding during passive diapirism: where the diapir edge is curved in map view, drape folding requires radial faults, much as flower petals separate as they open up.

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Figure 94. Radial faults adjacent to an onshore Louisiana diapir (Murray, 1966).

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Figure 95. Modified time-slice through two diapirs with beautiful radial fault patterns (courtesy ofWesternGeco and Kerr-McGee).

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Figure 98. Intersecting fault trends that generate a pseudo-radial pattern above a buried salt diapir (Sealy, 1962).

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TECTONIC STYLES OF SALT DEFORMATION

In this section of the course, we will look at the impact of salt in different tectonic provinces, from thick-skinned extension to collisional mountain belts to passive margins.

Thick-skinned extension

Depending on the relative timing of salt layer formation and rifting, extensional salt tectonics may be thick-skinned or thin-skinned. If the salt layer is prerift, as in the North Sea, extension affects both the supra- and subsalt section (thick-skinned). In such cases, salt tends to decouple the deformation, so that the structural styles above and below the salt can be quite different. Typically, the overburden is draped over subsalt faults, with salt separating and accommodating the different styles (Fig. 105).

Diapirs are commonly located above subsalt normal faults, with the idea being that differential loading between thegraben and the horst induces salt flow (Fig. 106). However, there are many faults that do not have overlyingdiapirs and many diapirs that do not have underlying faults. In fact, basement faulting and diapirism are not necessarily coupled at all (Fig. 107). The basement can extend by domino faulting, but the overburden lengthens by forming symmetrical graben that are unrelated to deeper faults. The salt separates the two levels of deformation and forms reactive diapirs, which may subsequently evolve into passive diapirs. More typically, however, there is a link between subsalt normal faults and overlying diapirs, with diapirs typically located above the

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footwall of the deeper normal faults because they preferentially form at the updip hinge of the suprasalt drape folds.

The coupling or decoupling of supra- and subsalt deformation has been modeled by various researchers. They showed that the most important parameters are the thickness of the salt, the thickness and strength of the overburden, the rate of fault slip, and the magnitude of the fault (Fig. 108). If the salt is very thin or the slip is very rapid, for example, deformation is coupled with little, if any, salt flow. At the other extreme, thick salt and slow slip result in significant decoupling. A gradual increase in salt thickness away from the basin margin thus can have an impact on the structural style, as shown schematically in Figure 109. There can also be dramatic changes in the degree of decoupling over time. A schematic evolutionary model of the central North Sea (Fig. 110) shows decoupled deformation during the first (Triassic) rift event, when the salt layer was thick, but largely coupled deformation during the subsequent (Upper Jurassic) rifting because the salt layer had thinned during prior diapiric flow.

There is an important added complication. Thick-skinned extension usually results in rotation of fault blocks, which induces thin-skinned gravitational deformation of the overburden above the salt (Fig. 111). In three dimensions, movement vectors can be highly varied because of the complex rift architecture (Fig. 112). An example of linked, thin-skinned extension and contraction from the central North Sea is shown in Figure 113.

Figure 105. Salt decoupling and the generation of drape folds above subsalt faults during thick-skinned extension along the Norwegian margin (a and b) and in the Gulf of Suez (c and d) (Withjack and Callaway, 2000). 61

Figure 106. Southern North Sea diapir located above a subsalt normal fault and proposed evolutionary model (Remmelts, 1995). The model invokes differential loading and density contrast in driving salt movement, but the differential pressure on the salt is probably insufficient.

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Figure 109. Role of increasing salt thickness in determining the structural style of thick-skinned extension in the North Sea (Stewart and Clark, 1999).

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Collisional mountain belts

Salt occurs in many contractional foldbelts around the world. Because of its relative weakness, salt forms a detachment level overlain by folds and thrusts that sole into the salt. There are three main contractional settings (Fig. 114): (1) in the frontal portions of collisional fold-and-thrust belts; (2) in rift basins that are inverted duringcollisional tectonics; and (3) at the toe of gravitational collapse systems. The locations of many of these foldbelts are shown in Figure 115. In this next section, we will briefly examine those formed during crustal-scale collision of tectonic plates.

Salt plays a dominant role in determining the structural style (Fig. 116). Foldbelts without salt are dominated by thrust faults and asymmetric fault-related folds with a consistent vergence toward the foreland. In contrast, those underlain by salt are characterized by a regular wavetrain of narrow, usually symmetrical anticlines separated by broad, flat-bottomed synclines. Thrust faults are less common with no preferred vergence. The folds are detachment folds cored by the salt and sometimes cut on one or both limbs by steep reverse faults. Fault-bend folds are generally rare, except where there is enough shortening that detachment folds can no longer accommodate the strain.

It should be understood that it is usually possible to interpret a given fold geometry several different ways. For example, the fold in Figure 117 can be interpreted as a fault-bend fold cored by a duplex (top section). But such an interpretation is purely a geometrical construct when salt is present; a better interpretation that is consistent withthe mechanics of salt shows a detachment fold cut by minor reverse faults on both flanks (bottom section).

The reason for the lack of vergence in salt-related folds is the weakness of salt. Critical wedge taper theory (Davis et al., 1983) suggests that the external geometry and internal deformation of foldbelts is a function of, among other factors, the coefficient of sliding friction along the basal detachment (Fig. 118). In the case of salt, where the detachment is relatively frictionless, the foldbelt has a narrower cross-sectional taper, a wider belt of deformation, and more symmetrical structures. An example of this is seen in the Sierra Madre Oriental of northeastern Mexico (Figs. 119 and 120). To the south and north of Monterrey, the foldbelt is relatively narrow, with a broad taper, and the deformation is dominated by asymmetric thrust faults and ramp anticlines (Fig. 120b). In the Monterrey salient, where salt is present, the foldbelt is wider, with a narrower taper, and the deformation is dominated by upright detachment folds with a regular wavelength (Fig. 120a).

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There are two end-member models for the development of detachment folds: one in which the limbs maintain a constant dip as they lengthen, and the other in which limbs maintain their length but gradually rotate during shortening (Fig. 121). In the latter case, which is favored by almost all researchers, strike-parallel fold propagation means that fold terminations will have more open geometries while culminations, with more shortening, will be tighter. Growth wedges are characterized by strata thatsteepen with depth (Figs. 122 and 123), although there are commonly complications in the form of onlap surfaces and local unconformities due to fluctuations in shortening rate and sedimentation rate.

In many mountain belts, the salt layer was deposited in a passive margin setting and was then involved in later deformation when the ocean basin closed up duringcollisional tectonics. In other cases, the salt layer was deposited in the foreland basin during deformation, so that the main detachment beneath the basement-involvednappes in the hinterland ramps up to the salt detachment in the foreland (Fig. 124). The result is that the structural style changes abruptly, from fault-bend and fault-propagation folds and associated thrusts where there is no salt, to simple detachment folds with only minor faulting where there is salt. An evolutionary model for the southern Pyrenees example shows how buckle folding typically precedes any faulting (Fig. 125). Faulting may be favored

by more rapid shortening rates, but the most common cause is a lack of sufficient salt to fill the growing fold cores. Thus, detachment folds in the center of a salt basin may give way to thrusted folds near the margin where the salt layer thins (Fig. 126). A nice example can be seen in the Sierra Madre Oriental of Mexico.

So far, we have looked at collisional foldbelts where the salt is merely the detachment. In other cases, diapirism may have occurred prior to regional shortening, creating a complex architecture of preexisting weaknesses (thediapirs) and minibasins with differing geometries and strengths. As can be imagined, the results of shortening in such a scenario can be highly complex. Two examples from the Southern Carpathians in Romania are illustrated in Figure 127.

Finally, many collisional mountain belts are inverted rift basins. Because salt is a common component of rift basins, it often influences the structural styles during inversion. Deformation above and below the salt will be partly or wholly decoupled, and any preexisting diapirs or other salt-related structures will impact subsequent deformation (Fig. 128). Examples of inverted rift basins with salt are shown from the Atlas Mountains of northern Africa (Fig. 129).

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Figure 114. Three settings for contractional salt tectonics: (a) foreland portions of collisionalmountain belts; (b) deepwater portions of passive margins that fail gravitationally; and (c) inverted rift basins (Letouzey et al., 1995). 66

Figure 115. Locations of salt-involved contractional tectonics around the world (Letouzey et al., 1995).

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Figure 116. Cartoons showing difference in structural style between foldbelts with (a) and without (b) salt (Jackson and Talbot, 1991). Those with salt have folding dominating over faulting, symmetrical structures with a regular wavelength and no preferred vergence, and high-angle reverse faults instead of low-angle thrusts. 68

Figure 117. Two interpretations of the same fold in the Perdido foldbelt: (a) fault-bend fold cored by a duplex; and (b) salt-cored detachment fold cut by minor high-angle reverse faults on both limbs (Trudgill et al., 1999). The latter interpretation is consistent with known contractional mechanics above a salt layer. 69

Figure 118. Critical wedge taper theory, in which a weak detachment (e.g., salt) results in a narrower taper angle, a wider zone of deformation, and no preferred vergence of faults (Jaumé and Lillie, 1988). 70

Figure 119. Map of northeastern Mexico showing the width of the Sierra Madre foldbelt increasing and the taper angle decreasing in the Monterrey salient due to the presence of salt (Marrett and Aranda-García, 2001). 71

Figure 120. Cross sections through the Sierra Madre foldbelt in (a) the Monterrey salient, where there is a salt detachment; and (b) to the south, where salt is absent (Marrett and Aranda-García, 2001). 72

Figure 121. Models of detachment fold formation with either constant limb dip and variable limb length (Model 1) or variable limb dip and constant limb length (Model 2) (Poblet and McClay, 1996).

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Figure 123. Example of growth strata deposited during limb rotation on an isoclinal detachment fold in the southern Pyrenees (Poblet and Hardy, 1995). 74

Figure 124. Regional and detailed cross sections across the southern Pyrenees, northern Spain (Sans and Vergés, 1995). The basal detachment ramps up to the Eocene Cardona salt layer that was deposited in the foreland basin; resulting structures are simple, symmetrical, salt-cored detachment folds. 75

Figure 125. Evolutionary model for the Pyrenean folds, in which early detachment folds are later modified by minor reverse faults (Sans and Vergés, 1995).

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