Salt The following notes are abbreviated versions of Mark Rowan’s three-day industry short course on salt tectonics.

Introduction Diapirs and diapirism Diapir initiation during differential Distribution and origin of salt basins loading Diapir initiation during extension or Mechanics of salt contraction Active and passive diapirism Salt withdrawal structures and welds Reactivation of diapirs during Turtle structures extension or contraction Expulsion rollovers Diapir interiors and margins Welds Near-diapir deformation - folding and faulting Tectonic styles of salt deformation Thick-skinned extension Collisional mountain belts

1 INTRODUCTION There has been an enormous revolution in our understanding of salt tectonics in the past decade or so Structural restoration. The technique of cross- (see Jackson, 1995, for an excellent history of salt was first applied to salt tectonics research). The beginnings of the revolution date structures in the late 1980s (e.g., Worrall and back a little farther within some of the exploration Snelson, 1989). In the past decade, numerous companies, but their ideas only became public starting in people have used restoration to reconstruct the about 1989. Our increased understanding of the history of salt movement and associated geometry and evolution of salt bodies and associated deformation of surrounding strata. strata is due in large part to the fortuitous convergence of advances in four areas: Field studies. Armed with new ideas, various researchers have reexamined exposed salt basins Seismic imaging. There has been a steady improvement throughout the world, leading to improved in seismic data acquisition and processing over the years. But with the advent of such techniques as pre- understanding of the processes of salt-related stack depth migration, images of salt bodies became deformation. much clearer, with improved pictures of the bases of salt In this course, we will concentrate on the new sheets and overhangs and the steep flanks of many ideas about salt tectonics. Many of the illustrations diapirs (e.g., Ratcliff, 1993). used here are examples of the four areas of Experimental and numerical modeling. Attempts to advance listed above. Much of the work has been model salt deformation have been made for many concentrated in the northern Gulf of Mexico, but decades (e.g., Nettleton, 1934), but until fairly recently, the impact of the advances has spread to salt both salt and its overburden were modeled as viscous basins worldwide. Thus, we will also examine the fluids. Starting in the late 1980s, however, B. Vendeville and coworkers started modeling the overburden as a geometries and structural styles of salt from such brittle material, more in keeping with the known places as the North Sea, the Red Sea, offshore mechanical behavior. The results demonstrated salt’s West Africa, offshore Brazil, the Precaspian more passive role of reacting to, rather than causing, Basin, and onshore Mexico. deformation (e.g., Vendeville and Jackson, 1992a, b) and fundamentally changed the ways most people look at salt deformation. 2 DISTRIBUTION AND ORIGIN OF SALT BASINS SE oriented transfer faults. The geometry, coupled with the rate of sedimentation and the relative time of Salt basins are found throughout the world (Fig. 1), but a evaporite formation, determines the areal extent and quick look will show that they occur primarily in rift thickness distribution of evaporites. Salt may be restricted basins and along passive margins, as well as in their to individual half- (Fig. 5 it may be regionally deformed counterparts, such as the Alpine/Himalayan tabular (Fig. 6), or it may have an intermediate geometry, system. According to a review by Jackson and Vendeville with a regional distribution but significant thickness (1994), many salt deposits were formed during the early variations (Fig. 7). postrift phase, including the basins of offshore Brazil, offshore West Africa, the U.S. Gulf Coast, and the Red Salt commonly occurs in paired basins on either side of Sea (Fig. 2). Others were formed either during rifting or oceanic crust, such as across the Gulf of Mexico, the South during lulls between distinct rift episodes, for example Atlantic, and the North Atlantic. Thus, it has generally many of the basins on either side of the northern Atlantic been thought that evaporite deposition occurs only on Ocean (Fig. 2). Finally, a few salt basins appear to be older continental crust, with subsequent oceanic spreading than rifting, namely those in the North Sea and Persian separating a once contiguous basin into two parts (Fig. 8). Gulf (Fig. 2). Many passive margins have seaward-dipping reflectors, or SDRs (Fig. 9) – wells show that these consist of subaerial Many rift basins have a similar history, one that is basalts that are considered the initial expression of oceanic conducive to evaporite deposition. They form during spreading. Autochthonous salt in parts of offshore Brazil extension of the continental crust, and grabens are initially occurs at a stratigraphic level above the SDRs, leading to a filled with nonmarine clastics because of the high heat model in which salt deposition postdates the onset of flow and associated regional uplift during rifting. oceanic spreading (Fig. 10). Salt in the South Atlantic Subsidence of the grabens, either during rifting or, more (offshore West Africa or Brazil) is interpreted to occur typically, during postrift thermal and loading subsidence, above the breakup unconformity and thus is underlain by a leads to marine incursion. If the climatic conditions are combination of landward continental crust and basinward appropriate, it is during the transition from nonmarine to oceanic crust (Fig. 11). There is then a transition to marine environments that evaporites are formed (often shallow-water carbonates and then deeper-water facies. episodically). In the typical (but not universal) scenario, The interbedded nature of the salt-carbonate transition and continued subsidence leads to true marine conditions. the similar seismic velocities means that there is typically Rift basins have a distinct basement architecture made a low acoustic impedance contrast and thus no good top- up of grabens and half-grabens segmented by transverse salt reflector for the autochthonous salt layer. In contrast, structures such as accommodation zones or transfer faults the salt is usually in contact with underlying clastic (Fig. 3). An example from the Brazilian margin is shown redbeds or basement so that there is a good base-salt in Figure 4, where the rift system includes a series of NW- reflector. 3 The evolution for the northern Gulf of Mexico is shown in The Precaspian Basin example nicely shows the kinds of Figure 12. Initial rifting during the Early Jurassic resulted vertical and lateral facies variations commonly found in in the SSW movement of Yucatan away from N. America, evaporite basins. Although salt dominates the basin center, whereas oceanic spreading resulted in southern movement there is also interbedded anhydrite that becomes more and rotation about a nearby pole. This means that dominant towards the northern margin. There can also be basement structures will have different orientations in the significant components of both carbonates and thinned continental crust and the oceanic crust. The siliciclastics that are concentrated along basin margins. boundary between the two is interpreted to be in the Another example is provided by the central North Sea, approximate position of the present-day shelf edge, so that where non-salt facies in the proximal part of the basin most of the deepwater province is underlain by oceanic disappear towards the center (Fig. 16). Of course, the crust (Fig. 13). So instead of one salt basin on continental presence of other lithologies means that there can be crust that was subsequently split by spreading (Fig. 12), significant, even coherent, reflectivity within the evaporite there were most likely two salt basins with the downdip layer (Fig. 17). 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). 4 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- (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 11 seafloor spreading) (Pindell et al., 2000). 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 14 Figure 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 MECHANICS OF SALT DEFORMATION overburden also behaved as a weak viscous fluid, but in Rock salt is a very different material from other, more fact, the overburden is a brittle material with real strength typical sedimentary rock types. There are several key (Figs. 18 and 20). The under-consolidated nature of factors that play roles in dictating its behavior. First, salt is shallow sediments in places like the northern Gulf of much weaker than other lithologies under both Mexico may suggest a viscous nature, but this is incorrect. and compression (Fig. 18). Even overpressured shale In fact, there are plenty of scarps at the sea floor that almost always has more strength than salt. In fact, the attest to the brittle style of deformation of even very young curve for wet salt in Figure 18 falls essentially on the axis sediments. of zero strength. The reason is that salt deforms as a Vendeville and Jackson (1992a) argued that the strength viscous material that effectively flows, with flow rates up and brittle nature of the overburden means that the density to 15m per year in exposed diapirs in Iran (Talbot et al., contrast plays only a limited role in salt tectonics. Instead, 2000). Flow is by a combination of Poiseuille flow due to they showed that salt should be viewed as a pressurized overburden loading and Couette flow due to overburden fluid and that it is the differential fluid pressure that drives translation (Fig. 19). The viscous nature of salt means that salt flow (Fig. 22). There are three key messages in this it forms a constant-strength, albeit very weak, layer figure: (1) density is only a second-order factor – the between normal sedimentary layers whose strength scenario is similar whether the overburden is of lesser, increases with depth (Fig. 20). Thus, salt serves as an equal, or greater density than the salt; (2) if there is a excellent detachment surface into which faults sole. differential load on the fluid (salt) and there is a place for Second, salt has a constant density of about 2.2 g/cm3, it to go, salt will flow; and (3) conversely, the salt cannot irrespective of burial depth (it actually get less dense with just push its way into the overburden because of the depth due to temperature effects). This is in contrast to overburdens strength. In other words, for salt withdrawal other sedimentary strata, such as sands and shales, that and diapirism to occur: (1) there must be an open path to a become compacted during burial and thus become more near-surface salt body; (2) the overburden must be thin and dense (Fig. 21). Thus, salt is more dense than its weak (faulted) enough for the differential fluid pressure to surrounding strata when it is near the surface, but is less overcome the overburden strength, in which case it will dense once it is buried beneath 1000-1500m of sediment. deform the thin overburden; or (3) some other process (e.g., tectonic) must create space. Put another way, salt The relatively low density of buried salt was historically does not drive salt tectonics; instead, salt merely used as a rationale for models in which salt punches its facilitates and reacts passively to external forces. This is a way through more dense overburden until it reaches its key point that underlies almost all subsequent discussion. 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.

19 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 Paleogene, on the order of 100 million years after the salt once the weld forms, the flanks, which are still underlain was deposited. Two maps from a nearby area show how by salt, may collapse, forming new, flanking depocenters salt was inflated from the Late Jurassic through the and inverting the original depocenter into a turtle Oligocene (reds and yellow in Fig. 31a) and then collapsed structure (Figs. 25 and 26). It is uncertain why some to form turtle structures in the Miocene (blues in Fig. 31b). minibasins collapse symmetrically to form turtles and The age of a turtle structure is defined here as the age of others don’t collapse at all. It may have to do with the touchdown of the initial basin and the start of flank geometry of the wedge of strata above the flanking salt: if collapse (i.e., the yellow horizon in Fig. 28). This age can it is relatively long and thin, it is mechanically easy to vary significantly across a basin, even between adjacent and thus collapse; if it is a short, thick wedge, it will not turtles. The timing is dependent on when subsidence fold as easily. Another possibility is that extension plays a began, the sedimentation rate, and the initial salt thickness. role. Experimental models with no lateral movement Thus, one cannot simply correlate between turtles. resulted in simple minibasins, whereas extension helped Moreover, the presence of a turtle tells us nothing about flank collapse and the formation of turtle structures (Fig. facies: there are, of course, turtle structures that have 27). reservoir-quality sands (e.g., Thunder in the GoM), A classic turtle structure from the Precaspian Basin shows but there are also turtles around the world where the strata the weld (with some remnant salt), the initial depocenter forming the turtle are composed dominantly of shale, that is now inverted, the flanking depocenters, and the carbonate, or even evaporites. Generally, once a basin is adjacent salt diapirs (Fig. 28). Note that crestal faulting subsiding into salt, it will keep subsiding regardless of the due to bending of the strata is a common secondary depositional setting and the minibasin will be filled with feature. Subsidence in this example started right after salt whatever sediment is available, including slumps off the deposition, as shown by the variable thickness of the oldest adjacent highs. suprasalt 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 21 Figure 25. Formation of a turtle structure by flank collapse and of a minibasin whose center has touched down and formed a weld. 22 Figure 26. Experimental model of a turtle structure (courtesy of B. Vendeville).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 Crestal faulting Flank collapse Flank collapse Salt Salt Initial depocenter

Weld

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 SW NE 20 km

meters Figure 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 (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. When progradational 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 Allochthonous Suprasalt salt Thickening onto weld/salt

Thinning Subsalt onto weld

Weld

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 31 Figure 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 An exposed vertical weld in La Popa Basin, Mexico is along the original, autochthonous salt layer when the illustrated in Figures 44 and 45. It is 25 km in length but is subsiding overburden comes into contact with the subsalt a true weld only over the southeastern half. To the west, it strata (Fig. 37) – a nice example of a so-called ‘primary’ consists of continuous remnant evaporite 100-200 m thick, weld is shown in Figure 38 – or they can be inclined due to and in between, it consists of patchy evaporite separated evacuation above a dipping base salt (Fig. 39). by true welds. However, this would not be known from a seismic profile because of the difficulty in imaging steep Welds are not always obvious, as shown in Figures 40 and structures. 41 – they are identified by discontinuous, high-amplitude reflectors, often with angular discordance between strata Another exposed weld, this time from the Flinders Ranges on either side (but this can also reflect an unconformity or in South Australia, is shown in Figure 46. It extends onlap surface). The seismic character reflects the fact there upward from a triangular diapir pedestal formed above the are pods of remnant salt along the weld, which are there autochthonous salt layer, and separates two minibasins because the top and base of salt do not have perfectly with very different thicknesses and facies. The weld has matching geometries. Ultimately, correctly identifying and remnant sandstones along it that were originally deposited interpreting welds requires a good mental image of the within the evaporite sequence, and is bordered on one side three-dimensional salt geometry and its evolution over by a shale sheath. 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.

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).

34 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 A variation of reactive diapirism occurs in wrench-fault models suggest that many diapirs are initiated and grow settings. Pull-apart basins form where there is a releasing during regional extension. The models show that diapirs bend or step-over in strike-slip faults. These are the sites of may go through three evolutionary stages: reactive, active, very rapid thinning and are ideal locations for the and passive diapirism (Fig. 50). At this stage, we will generation of reactive diapirs (Fig. 54). An example of a examine only reactive diapirism. shallow diapir at a releasing bend on the edge of an allochthonous sheet is shown in Figure55. When a salt layer is buried by constant-thickness strata, nothing will happen (even if the overburden is more Contraction can also initiate diapirism. When shortening dense) until some external force is applied. In the case of occurs above a salt décollement, the overburden in extension, the overburden is lengthened and thinned, growing detachment folds may be uplifted, faulted, and which is accommodated by graben formation at the surface eroded enough so that salt can break through to the surface and an “inverse graben” at the salt-sediment interface and subsequently grow as a diapir (Fig. 56). It is important (large-scale ). Salt reacts to the extension by to that this mechanism of diapir initiation, like overall thinning and by filling the space in the inverse progradation and unlike extension, can only happen when graben (Fig. 51). The result is a triangular diapir (elongate there is a thin overburden. Otherwise, the differential load in the strike direction) that has flanking growth faults that is insufficient and the thickness and strength of the get younger toward the diapir crest. The diapir grows in overburden is too great and only a fold will form (Fig. 57). 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 a subhorizontal 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.

41 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 and Vendeville, 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 arbitrary pin A Cover Regional Salt Basement

arbitrary erosion surface B

C

Figure 56. Formation of a diapir by erosion and salt breakthrough in a contractional (Coward and Stewart, 1995). 45 Active and passive diapirism In most of the models for diapir initiation that we have It is increasingly clear that passive diapirism is the examined, a salt body starts inflation or growing at some dominant style of diapir growth in basins throughout the point after it has been buried beneath an overburden. Thus, world. Diapirism may get triggered by various means, for two more stages of diapirism are typically involved in example extension (reactive diapirism), contraction, or subsequent growth. First, once the overburden becomes differential loading. But this usually happens early in the thin and weak enough and/or the pressure differential in history, beneath a relatively thin overburden, and salt soon the salt is great enough, the diapir will break through to breaks through to the surface. Thus, there is typically an the surface as an active diapir. This is equivalent to the initial phase of diapir formation, a very brief stage of classic model of piercement diapirism long thought to be active diapirism, and then a long-lived history of passive dominant in salt tectonics, but is a brief episode in the diapirism (as long as there is still adequate salt in the diapir history that occurs only when the overburden is thin source layer to maintain diapir growth at the sea floor). 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.

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 minibasins subside 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 The tops of many diapirs have a zone of caprock (Fig. 69), from the flow of the evaporite and any interbedded typically consisting of anhydrite, calcite, and minerals lithologies (Figs. 66 and 67). Typical features include such as pyrite and barite. This is usually interpreted as the vertical lineations, isoclinal folds, pinch-and-swell insoluble residue after the halite has been removed by structures, etc. Exposures in such areas as La Popa basin dissolution. A plot of caprock thickness versus diapir (Mexico), the Flinders Ranges (Australia), and the Zagros burial depth shows that caprock is best developed in the Mts. (Iran) show that exotic clasts are comprised upper 3000-5000 ft. of the earth’s surface (Fig. 70). This is exclusively of lithologies originally deposited within or because dissolution is a consequence of the circulation of intruded into the evaporite layer (Fig. 68). There are no meteoric water, which dissolves and carries the salt away blocks of strata plucked from the walls of diapirs rooted in and is constantly replenished. Deeper diapirs do not feel the autochthonous layer (we will discuss allochthonous this circulation. Likewise, caprock is rare to absent in salt later). Although very interesting, we will not address deepwater environments. One explanation is that the the internal characteristics of diapirs any further in this surrounding water is already very saline; another is that course. Instead, we are concerned here more with the condensed muds covering diapirs inhibit the transport external geometry of salt bodies and the deformation in away of any dissolved salt (Fletcher et al., 1995). Having surrounding sediments. said that, there are brine pools known in the deepwater GoM, 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 above diapirs, and ends up as shale sheath as the diapir flank collapses during minibasin growth (Fig. 72). 51 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). 52 Figure 67. Cross section through onshore Texas diapir showing complex folding of internal stratigraphy (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.

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

57 Figure 95. Modified time-slice through two diapirs with beautiful radial fault patterns (courtesy of WesternGeco and Kerr-McGee). 58 Figure 98. Intersecting fault trends that generate a pseudo-radial pattern above a buried salt diapir (Sealy, 1962).

59 TECTONIC STYLES OF SALT DEFORMATION In this section of the course, we will look at the impact of footwall of the deeper normal faults because they salt in different tectonic provinces, from thick-skinned preferentially form at the updip hinge of the suprasalt extension to collisional mountain belts to passive margins. drape folds. The coupling or decoupling of supra- and subsalt deformation has been modeled by various researchers. Thick-skinned extension They showed that the most important parameters are the Depending on the relative timing of salt layer formation thickness of the salt, the thickness and strength of the and rifting, extensional salt tectonics may be thick-skinned overburden, the rate of fault slip, and the magnitude of the or thin-skinned. If the salt layer is prerift, as in the North fault (Fig. 108). If the salt is very thin or the slip is very Sea, extension affects both the supra- and subsalt section rapid, for example, deformation is coupled with little, if (thick-skinned). In such cases, salt tends to decouple the any, salt flow. At the other extreme, thick salt and slow deformation, so that the structural styles above and below slip result in significant decoupling. A gradual increase in the salt can be quite different. Typically, the overburden salt thickness away from the basin margin thus can have an is draped over subsalt faults, with salt separating and impact on the structural style, as shown schematically in accommodating the different styles (Fig. 105). Figure 109. There can also be dramatic changes in the degree of decoupling over time. A schematic evolutionary Diapirs are commonly located above subsalt normal faults, model of the central North Sea (Fig. 110) shows decoupled with the idea being that differential loading between the deformation during the first (Triassic) rift event, when the graben and the induces salt flow (Fig. 106). salt layer was thick, but largely coupled deformation However, there are many faults that do not have overlying during the subsequent (Upper Jurassic) rifting because the diapirs and many diapirs that do not have underlying salt layer had thinned during prior diapiric flow. faults. In fact, basement faulting and diapirism are not necessarily coupled at all (Fig. 107). The basement can There is an important added complication. Thick-skinned extend by domino faulting, but the overburden lengthens extension usually results in rotation of fault blocks, which by forming symmetrical graben that are unrelated to induces thin-skinned gravitational deformation of the deeper faults. The salt separates the two levels of overburden above the salt (Fig. 111). In three dimensions, deformation and forms reactive diapirs, which may movement vectors can be highly varied because of the subsequently evolve into passive diapirs. More typically, complex rift architecture (Fig. 112). An example of linked, however, there is a link between subsalt normal faults and thin-skinned extension and contraction from the central overlying diapirs, with diapirs typically located above the North Sea is shown in Figure 113. 60 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. 62 Figure 109. Role of increasing salt thickness in determining the structural style of thick-skinned extension in the North Sea (Stewart and Clark, 1999). 63 Collisional mountain belts Salt occurs in many contractional foldbelts around the It should be understood that it is usually possible to world. Because of its relative weakness, salt forms a interpret a given fold geometry several different ways. For detachment level overlain by folds and thrusts that sole example, the fold in Figure 117 can be interpreted as a into the salt. There are three main contractional settings fault-bend fold cored by a duplex (top section). But such (Fig. 114): (1) in the frontal portions of collisional fold- an interpretation is purely a geometrical construct when and-thrust belts; (2) in rift basins that are inverted during salt is present; a better interpretation that is consistent with collisional tectonics; and (3) at the toe of gravitational the mechanics of salt shows a detachment fold cut by collapse systems. The locations of many of these foldbelts minor reverse faults on both flanks (bottom section). are shown in Figure 115. In this next section, we will briefly examine those formed during crustal-scale collision of tectonic plates. The reason for the lack of in salt-related folds is the weakness of salt. Critical wedge taper theory (Davis et Salt plays a dominant role in determining the structural al., 1983) suggests that the external geometry and internal style (Fig. 116). Foldbelts without salt are dominated by deformation of foldbelts is a function of, among other thrust faults and asymmetric fault-related folds with a factors, the coefficient of sliding friction along the basal consistent vergence toward the foreland. In contrast, those detachment (Fig. 118). In the case of salt, where the underlain by salt are characterized by a regular wavetrain detachment is relatively frictionless, the foldbelt has a of narrow, usually symmetrical separated by narrower cross-sectional taper, a wider belt of broad, flat-bottomed . Thrust faults are less deformation, and more symmetrical structures. An common with no preferred vergence. The folds are example of this is seen in the Sierra Madre Oriental of detachment folds cored by the salt and sometimes cut on northeastern Mexico (Figs. 119 and 120). To the south and one or both limbs by steep reverse faults. Fault-bend folds north of Monterrey, the foldbelt is relatively narrow, with are generally rare, except where there is enough shortening a broad taper, and the deformation is dominated by that detachment folds can no longer accommodate the asymmetric thrust faults and ramp anticlines (Fig. 120b). strain. 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).

64 There are two end-member models for the development of by more rapid shortening rates, but the most common detachment folds: one in which the limbs maintain a cause is a lack of sufficient salt to fill the growing fold constant dip as they lengthen, and the other in which cores. Thus, detachment folds in the center of a salt basin limbs maintain their length but gradually rotate during may give way to thrusted folds near the margin where the shortening (Fig. 121). In the latter case, which is favored salt layer thins (Fig. 126). A nice example can be seen in by almost all researchers, strike-parallel fold propagation the Sierra Madre Oriental of Mexico. means that fold terminations will have more open So far, we have looked at collisional foldbelts where the geometries while culminations, with more shortening, will salt is merely the detachment. In other cases, diapirism be tighter. Growth wedges are characterized by strata that may have occurred prior to regional shortening, creating a steepen with depth (Figs. 122 and 123), although there are complex architecture of preexisting weaknesses (the commonly complications in the form of onlap surfaces and diapirs) and minibasins with differing geometries and local unconformities due to fluctuations in shortening rate strengths. As can be imagined, the results of shortening in and sedimentation rate. such a scenario can be highly complex. Two examples In many mountain belts, the salt layer was deposited in a from the Southern Carpathians in Romania are illustrated setting and was then involved in later in Figure 127. deformation when the ocean basin closed up during Finally, many collisional mountain belts are inverted rift collisional tectonics. In other cases, the salt layer was basins. Because salt is a common component of rift basins, deposited in the during deformation, so that it often influences the structural styles during inversion. the main detachment beneath the basement-involved Deformation above and below the salt will be partly or in the hinterland ramps up to the salt detachment in wholly decoupled, and any preexisting diapirs or other the foreland (Fig. 124). The result is that the structural salt-related structures will impact subsequent deformation style changes abruptly, from fault-bend and fault- (Fig. 128). Examples of inverted rift basins with salt are propagation folds and associated thrusts where there is no shown from the Atlas Mountains of northern Africa (Fig. salt, to simple detachment folds with only minor faulting 129). 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

65 Figure 114. Three settings for contractional salt tectonics: (a) foreland portions of collisional mountain 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).

67 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).

73 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). 76