Download Glossary Accommodation zone—Accommodation zone is the zone

Glossary
Accommodation zone—Accommodation zone is the zone between two overlapping faults
where offset is transferred from one fault to another (Rosendahl 1987).
Active continental stage of the transform margin development—During this stage, the
transform fault accommodates a displacement between the two continental plates (Wilson
1965; Freund 1974). It is linked with neighbor rift zones or pull-apart terrains via horse-tail
structures (Christie-Blick & Biddle 1985). This stage ends at different times at different
locations along the transform. The end takes place when the continental plate segment
becomes juxtaposed with oceanic plate undergoing active accretion.
Active continental-oceanic stage of the transform margin development—During this
stage the transform fault accommodates the displacement between oceanic and continental
plates (Mascle & Blarez 1987). This stage ends at different times at different locations along
the transform. The end takes place when the continental plate segment gets laterally cleared
by migrating spreading ridge and becomes juxtaposed with an oceanic plate older than the
time of this clearance.
Active contraction-controlled uplift—A contraction component of the displacement during
active continental and continental-oceanic stages of the transform, which can locally drive
transform margin uplift (Mascle & Blarez 1987; Nemčok et al. 2012b). This scenario is rather
rare but it can control a significant uplift where present (Nemčok et al. 2012b).
Active rifting—It is characterized by active asthenosphere upwelling in response to
buoyancy instability, which drives the rifting. This upwelling exhibits broad spatial wave
lengths, in the order of 200-500 km, in proportion to asthenospheric thickness (Ruppel 1995;
Barnouin-Jha et al. 1997). Tensional stresses at the base of the lithosphere are driven by the
ascending convecting material (Turcotte & Emerman 1983). The lithosphere is thermally
thinned by heating and adsorption into the asthenosphere, in addition to necking in response
to extension. Hence, the volume of asthenosphere rising into the lithosphere exceeds the
volume of lithosphere displaced laterally by extension (e.g., Olsen & Morgan 1995).
Advancing side of the strike-slip fault tip—Advancing side of the strike-slip fault
propagation tip is the area affected by the mean stress magnification and the maximum
principal stress σ1 rotation towards a more acute angle with the principal displacement zone
(Homberg et al. 1997). Advancing sides of large strike-slip faults frequently contain folds
and thrusts.
Anastomosing—Pertaining to a network of branching, twining and rejoining surfaces or
surface traces. It is often used to describe braided fault systems (Biddle & Christie-Blick
1985).
Antithetic shear—Defined as a shear that has its displacement sense opposite to those of the
synthetic shear and principal displacement zone (e.g., Christie-Blick & Biddle 1985 and
references therein).
Associated horse-tail structure—Faults splaying off the tip of a transform fault, which are
coeval with and are directly geometrically and kinematically linked to controlling normal
faults of the neighbor rift zone (Nemčok et al., 2016 – this volume). This case takes place
where the rift zone and transform fault are located one after another in the direction of the
break-up propagation.
Breached relay ramp—Relay ramp that has been cut by a fault, transforming it from a softlinked to a hard-linked overlap structure.
Breakaway blocks – Defined as crustal blocks of various sizes, which separate detachment
fault systems that form in distal rifted margins. Similar to extensional allochthons and
hanging-wall blocks, breakaway blocks consist of pre-rift sediments. In contrary to
extensional allochthons and hanging-wall blocks, breakaways blocks are bounded on both
sides by detachment (exhumation) faults (Massini et al. 2013).
Break-away fault—This term can be used for two different things. It can either represent a
fault that forms the upper boundary of the extensional domain of the gravity glide (e.g.,
Stewart & Reeds 2003) or it can represent a fault that divides stretching and necking domains
of the continental margin from each other (Manatschal et al. 2007; Mohn et al. 2012; Sutra et
al. 2013).
Break-up unconformity— Characterized as an unconformity caused by the initiation of seafloor spreading (Embry & Dixon 1990). The age of the breakup unconformity in theory is
synchronous with the initiation of sea-floor spreading in the adjacent ocean basin (Falvey
1974). In reality, this unconformity is typically mistaken for end-of-the-rifting unconformity,
which has a different age in the stretching zone, thinning zone, necking zone, hyper-extended
domain, proximal margin and distal margin. An exact definition of the break-up
unconformity would be to use one, which was formed in response to the break-up of the last
lithospheric layer in the process of the break-up of the lithospheric multilayer, right next to
the initial oceanic crust.
Brittle deformational environment—Environment controlled by its brittle strength, where
deformation takes place by fracturing, frictional sliding and cataclastic flow. The brittle
strength is calculated using the relationship of Ranalli & Murphy (1987) modified from the
relationship of Sibson (1974), i.e. (1 - 3) = gz(1 - ), where 1 and 3 are the maximum
and minimum principal compressional stresses,  is a constant (3 for thrusting, 1.2 for strikeslip faulting and 0.75 for normal faulting, assuming a friction of 0.75),  is the rock density, g
is the acceleration due to gravity, z is the depth and  is the assumed hydrostatic/lithostatic
pressure ratio.
Buoyancy-driven fluid flow— Vertical and lateral fluid flows are likely to occur when any
lateral fluid density gradient exists. Therefore, it is reasonable to assume that buoyancydriven fluid flow is rather common in the extending upper crust in scenarios, which do not
allow topography-driven flow to prevail. These include: (a) extensional scenarios without
sub-aerial exposure such as sea-floor spreading ridges and their surroundings (Lowell 1975;
Lowell & Rona 1985; Fisher & Becker 1995; Yang et al. 1998; Schardt et al. 2006), (b)
submarine environments in extensional basins (Person & Garven 1992; Lampe & Person
2002; Yang et al. 2006), and (c) special sub-aerial extensional scenarios lacking distinct
topographic gradients such as deep isolated permeable aquifers (Gvirtzman et al. 1997).
Compaction-driven fluid flow— Upon decreasing porosity due to compaction, pore fluids
are expelled from the shallow-seated sediments (Shi & Wang 1986 and references therein).
The compaction-driven fluid flow is then controlled by the permeability structure of the basin
fill, which is usually anisotropic (Bethke 1985). The low permeability layers of the
permeability stratigraphy tend to be dominated by fluid perpendicular to bedding planes,
while the high permeability layers tend to undergo a fluid flow parallel to their bedding
planes.
Compression— (1) A certain orientation of stress generating forces that result in the
shortening or decrease of the volume of a substance (Biddle & Christie-Blick 1985; modified
from Bates & Jackson 1980). Compression can be uniaxial or triaxial. The uniaxial
compression is characterized by one principal compressive stress of non-zero value. In the
case triaxial compression, all three principal stresses are of non-zero value (Means 1976).
Compressive principal stress can also occur with one or more tensile principal stresses. (2) A
state of strain where compressive stress shortens material lines (Aydin & Nur 1985; Biddle &
Christie-Blick 1985).
Conjugate Riedel shear—Also referred to as R′ Riedel or antithetic shear. See Riedel shear,
synthetic, and antithetic faults (Cloos 1928; Riedel 1929; Tchalenko & Ambraseys 1970 and
references therein; Biddle & Christie-Blick 1985).
Continent-Ocean boundary—It is characterized as the inboard edge of unequivocal oceanic
crust (Direen et al. 2013).
Continent-Ocean transition zone – A region on the continental margin connecting the
outboard edge of highly attenuated, unequivocal continental crust, and the inboard edge of
unequivocal oceanic crust. It includes sedimentary and magmatic components that vary both
along and across the margin, and may include areas of failed sea-floor spreading (Direen et
al. 2013).
Continental break-up— The continental break-up affects a lithospheric multilayer in several
stages (Manatschal 2004; Huismans & Beaumont 2005, 2008, 2011; Lavier & Manatschal
2006). The synchronous crustal and lithospheric mantle break-up is just one of the possible
scenarios. The end-member scenarios are represented by the first one with crustal break-up
followed by mantle lithosphere break-up and the second one with lithospheric mantle breakup followed by crustal break-up (Huismans & Beaumont 2011).
Continental ribbon— A block of strong, buoyant material embedded into the oceanic
lithosphere (Moresi et al. 2014).
Contraction—A type of strain that results in volume reduction (e.g. thermal contraction), or
reduction of length (e.g. by use of a contraction fault; Norris 1958; McClay 1981). Currently,
term contraction is often used as the general strain term associated with compressive stress.
However, Biddle & Christie-Blick (1985) stated that the use of this word as a general stress
term may be misleading, as it implies a change in volume. They proposed that the better term
for general use would be the word shortening, as defined by Hobbs et al. (1976).
Contractional strike-slip bridge—Synonymous with restraining overstep (Gamond 1987).
This bridge tends to magnify the mean stress and rotate the maximum principal stress σ1
almost towards parallelism with bridge trend. It may also have a tendency for fluid expulsion
(Nemčok et al. 2002).
Convergent bend—A bend in a strike-slip fault that causes crustal shortening (convergent
motion; EA Fig. 1) in its vicinity (Biddle & Christie-Blick 1985). Eventually, it may result in
a positive flower structure (EA Fig. 2). It is also referred to as restraining bend (Crowell
1974a). Activity timing of restraining bends tends to match the timing of releasing bends,
providing a displacement accommodation balance along the strike-slip fault (Mann 2013).
Convergent hydrocarbon migration—It takes place in a convex reservoir horizon or fault
component (Hindle 1997). It tends to focus petroleum pathways and concentrate
hydrocarbons.
Convergent (transpressional) strike-slip or wrench fault—A strike-slip or wrench fault,
which contains a convergent motion component perpendicular to the fault direction, along
with the lateral motion component. It results in the shortening transverse to the fault direction
(Wilcox et al. 1973; Biddle & Christie-Blick 1985).
Coupling point—a point, where the first brittle faults cross-cut the entire crust and penetrate
the mantle (Péron-Pinvidic et al. 2013). It tends to be close to or coincide with the taper
break.
Damage zone— Damage zone is the middle of the three fault zone components (Chester &
Logan 1986; Smith et al. 1990). It can include small faults, veins, fractures, cleavage and
folds (Caine 1999).
Decoupling—It is the deformation scenario where the upper part of the multilayer deforms in
a style different from the style of the lower part and/or if the structures developed at these
two levels are not directly kinematically linked.
Depositional model of extensional margin—Model characterized by a low-dip slope,
numerous possibilities of sediment catchment on wide shelf and slope, by the relatively long
sediment transport distance from shelf to basin floor and potential for intra-slope basins
(Towle et al. 2012a, b; Addis et al. 2013a, b). The prograding shelf connection with slope
channel-levee complexes is common. Fairly long compensationally-stacked, slope channellevee complexes are typical. They are wide and rather long. Narrow and highly sinuous
channels with extensive levees are also typical.
Depositional model of transform margin— Model characterized by a steep-dip slope, fairly
limited possibilities of sediment catchment on narrow shelf and slope, by the relatively short
sediment transport distance from the shelf to basin floor and fairly limited potential for intraslope basins (Towle et al. 2012a, b; Addis et al. 2013a, b). Vertically stacked, slope channellevee complexes are rather narrow and not very long. Robust distributary complexes on the
basin floor are also typical.
Detachment—Low-angle or horizontal fault separating the faults of the upper plate from the
lower plate.
Deviatoric stress—The difference between the total stress and the mean stress.
Dextral—Rightward in motion (dextral slip is a right slip; EA Fig. 1), pertaining to the right
side (as defined by Biddle & Christie-Blick 1985).
Differential stress—The difference between the maximum and minimum principal stresses.
Dip separation—Characterized as “separation measured parallel to the dip of a fault”
(Biddle & Christie-Blick 1985; modified from Crowell 1959; Bates & Jackson 1980).
Dip-slip—Slip component measured parallel to the dip of a fault (Crowell 1959; Bates &
Jackson 1980; Biddle & Christie-Blick 1985).
Dip-slip fault—Fault structure, where the majority of displacement is accomplished by a dipslip (Biddle & Christie-Blick 1985; modified from Bates & Jackson 1980).
Disorganized sea-floor spreading—Disorganized sea-floor spreading takes place for several
years after continental break-up. An example of the disorganized spreading comes from
gradual, tortuous and mostly sub-aerial spreading such as the one that takes place in Afar
(Rosendahl 2004, pers. comm.). Another example comes from the oceanic crust adjacent to
the western segment of the Falklands transform margin (Edwards et al. 2013). Here, the cold
edge effect of the continental crust has resulted in either poorly developed or missing oceanic
crust right next to the transform margin, as spreading centers were laterally clearing it during
the stage of disorganized spreading. Further oceanward a spreading ridge reorganization took
place, developing an organized spreading ridge system. Yet another example of the
disorganized spreading, this time from the magma-poor setting, is at least 6 million year-long
exhumation of the continental lithospheric mantle, responsible for a 70-73 km wide corridor
of the so-called proto-oceanic crust in front of the Sergipe-Alagoas conjugate margins
(Rosendahl et al. 2005; Nemčok et al. 2012a).
Distal margin— Part of the rifted margin, separated from the proximal part by a necking
zone (EA Fig. 3). Depending on the study area and varying terminology, the distal margin
can correspond with or include the proximal and distal ocean-continent transition, transitional
domain and/or the zone of exhumed continental mantle. Regularly the distal domain (or distal
margin) is referenced as the hyper-extended domain (with basement that is thinned down to
less than 10 km because of the low-angle detachment faults). The thinned continental crust
contains no remaining ductile layers, thus allowing faults to cut from the surface into the
mantle. The basement of the distal domain can be composed of upper or lower continental
crust (or both), exhumed and variously serpentinized mantle or embryonic oceanic crust
(Péron-Pinvidic & Manatschal 2009; Mohn et al. 2010; Péron-Pinvidic et al. 2013).
Compared to the older proximal margin, the distal margin reflects a more complex evolution
associated with the development of younger, low-angle detachment faults (Massini et al.
2013).
Divergent bend—A bend in a strike-slip fault that causes crustal extension (divergent
motion; EA Fig. 1) in its vicinity (Biddle & Christie-Blick 1985). Eventually, it may result in
a negative flower structure (EA Fig. 2) or a pull-apart basin. It is also referred to as releasing
bend and extensional bend (Crowell 1974a). Similarly to restraining bends, the activity
timing of releasing bends tends to match the timing of restraining bends, providing a
displacement accommodation balance along the strike-slip fault (Mann 2013). Bends form
from sidewall rip-outs (Swanson 2005; Mann 2007) or when faults intersect pre-existing
basement structures (Mann 2007). They can also form linkages of R and P shears (Crowell
1974b; Mann 2007)
Divergent hydrocarbon migration— It takes place in a concave reservoir horizon or fault
component (Hindle 1997). It tends to disperse petroleum pathways and dilute hydrocarbons.
Divergent overstep – see releasing overstep.
Divergent (transtensional) strike-slip or wrench fault—A strike-slip or wrench fault,
which contains a lateral motion component along with a divergent motion component
perpendicular to the fault direction. It results in an extension transverse to the fault direction
(Wilcox et al. 1973; Biddle & Christie-Blick 1985; Harding et al. 1985).
Downlap—A downward termination of initially inclined strata against an initially horizontal
or inclined base-surface (EA Fig. 4; Mitchum 1977; Biddle & Christie-Blick 1985).
Drag fold—(1) A folding of material found in close distance to a fault. It is created by
movement along a fault (see normal drag and reverse drag). The term “fold” in this case does
not refer to normal types of folds that form by the ductile bending of rock material during
plate movement or collision (Hobbs et al. 1976). (2) A minor fold that forms in a less
competent bed between more competent beds as a result of the competent beds’ movement in
opposite direction relative to one another (Bates & Jackson 1980; Biddle & Christie-Blick
1985).
Drape fold—A folding of sedimentary layer that reflects the configuration of underlying
structures, e.g. swells and dips (Friedman et al. 1976; Biddle & Christie-Blick 1985). It is
also referenced as a fold structure formed by differential compaction. This term does not refer
to a normal type of folding that forms by ductile bending of rock material during plate
movement or collision.
Driving forces of stretching— Stretching is driven by (a) deviatoric stresses developed in
stress cycles by block/plate movements, (b) deviatoric stresses developed over upwelling
asthenosphere convection systems, and (c) frictional forces along boundaries of individual
lithospheric layers between themselves and with the asthenosphere. Furthermore, partial
melting driven by adiabatic decompression of the rising lower lithosphere contributes to
lithospheric thinning (McKenzie & Bickle 1988).
Ductile deformational environment— Characterized as the environment controlled by its
ductile strength, where the deformation is continuous at the scale of observation. The ductile
strength is calculated from the power-law creep expression (Kirby 1983), i.e. (1 - 3) =
K1/n*1/nexp(E/nRT) and K = (1/A)1/n, or (1 - 3) = (/A)1/n exp(E/nRT), where K and A are
scaling factors, n is the power law exponent, E is the activation energy,  is the strain rate, R
is the universal gas constant and T is the temperature, taken from the temperature-depth
profile.
Dynamic analysis— Kinematic and kinetic studies, which relate strains to the evolution of
stress fields (Biddle & Christie-Blick 1985).
Echelon— According to the definition of Biddle & Christie-Blick (1985), it is a step, which
can be referred to as overstepping faults – e.g. echelon faults (Clayton 1966; Segall & Pollard
1980).
Effective elastic thickness—The effective elastic thickness (EET) of the lithosphere
represents a long-term integrated strength of multilayer, the layers of which vary from being
coupled to decoupled (e.g., Watts 2001; Burov & Watts 2006; Huismans & Beaumont 2008).
It reflects the integrated strength of the lithosphere that responds to long-term (> 105 yr)
loading by flexure and controls the depth of necking in the rift process. EET has been defined
as the combined effect of thicknesses of detached strong layers in the lithospheric multilayer
(Burov & Diament 1995): EET = (Σni=1Δhi3)1/3, where n is the number of layers with
thicknesses h1, h2 and so on.
En echelon—A step-like arrangement of relatively short, consistently overlapping or
underlapping structural elements like faults or folds that are approximately parallel to each
other but at the same time oblique to the linear or relatively narrow zone in which they occur
(EA Fig. 1; preferred definition of Biddle & Christie-Blick 1985; modified from Campbell
1958; Harding & Lowell 1979). En echelon structures can be preserved in both map view and
cross section (Shelton 1984; Aydin & Nur 1985).
Entry window—Fault overlap with the source horizon for the hydrocarbon migration (Knipe
1993).
Erosion— An effect of exogenic elements and processes (such as water and wind) on rocks,
which removes material from one location on the Earth's crust, and transports it to another
location where it is deposited (Sklar & Dietrich 2004).
Erosional truncation— Termination of strata against an overlying erosional surface (Emery
& Myers 1996).
Erosional unloading-controlled uplift—Uplift driven by erosional unloading of the
transform margin (Basile & Allemand 2002). This uplift model requires a high flexural
strength of the transform margin, which is not usually the case.
Euler rotation parameters— Defined as four main mathematical constructs used to
represent the attitude of a rigid body in 3-D space, including: (1) the rotation matrix, (2) three
Euler angles, (3) the unit quaternion and (4) the rotation vector (Diebel et al. 2006).
Exhumation— Exhumation is a mechanism that causes the ascent of material points with
respect to reference surface, causing an upward heat advection (England & Molnar 1990).
Exhumed mantle zone— Characterized as a transitional zone located between the oceanic
and continental crusts, which formed during continental rifting (Sibuet et al. 2007).
Exit window—Fault overlap with the sink horizon for the hydrocarbon migration (Knipe
1993).
Extension—A type of strain that results in an increase in length (Biddle & Christie-Blick
1985).
Extension fault—A fault that results in the lengthening of the material it intersects,
commonly (but not necessarily) bedding. It can also refer to a fault of any dip and is
synonymous to term normal fault (Christie-Blick 1983; Biddle & Christie-Blick 1985; Suppe
1985).
Extension fracture— Forms when effective stresses are tensile (for example when porefluid pressure exceeds lithostatic pressure). In strike-slip systems, these kinds of fractures
form at about 45° to the master fault and their formation is caused by a simple shear (Biddle
& Christie-Blick 1985). Extension fractures were also defined as fractures that show no
motion in the plane of crack, and as cracks of the 1st mode, which form as a consequence of a
negative lithostatic load (Lawn & Wilshaw 1975; Biddle & Christie-Blick 1985). They are
partly synonymous with tension fractures (T fractures of Tchalenko & Ambraseys 1970).
Extension localization—Extension localization is controlled by lithospheric weakening in
response to extension as elevated isotherms reduce the strength of lithospheric multilayers
which have temperature-dependent rheology. This results in strain localization and stress
concentration into a thinned lithosphere (e.g., Davis & Kusznir 2002).
Extensional bend – see divergent bend.
Extensional allochthon— Characterized as a fault block/blocks with the width of a few
kilometres to a couple of hundred metres, which is/are systematically truncated by
a detachment fault at the base (Manatschal 2004). Nice examples of extensional allochthons
can be found on the Iberian and East Indian margins (Manatschal 2004; Sinha et al. 2015).
Extensional margin—An extensional margin is the result of successful orthogonal rifting
that reached continental break-up and subsequent oceanic crust accretion. It went through the
development stage including: (a) orthogonal rifting of continental lithosphere, (b) the
continental break-up and development of sea-floor spreading centers parallel to the
extensional margin, (c) early drifting and (d) advanced drifting. It is characterized by a
relatively wide shelf and slope (Towle et al. 2012a, b; Addis et al. 2013a, b). Extensional
margins are usually characterized by the sediment entry point system that takes a relatively
long time after the breakup to develop. It keeps evolving during the remaining passive margin
development history. Examples come from the Santos (Mohriak et al. 2008) and Nova Scotia
margins. The same applies to the shelf break location that usually undergoes distinct shifts.
All these sedimentary characteristics are related to the occurrence of the post-breakup uplift
of this margin.
Extensional overstep – See releasing overstep.
Extensional strike-slip bridge— Synonymous with releasing overstep (Gamond 1987). This
bridge tends to reduce the mean stress and rotate the maximum principal stress σ1 almost
towards being perpendicular to the bridge trend. It may also have a tendency for fluid
accumulation (Nemčok et al. 2002).
Fault-angle depression—Area parallel to the trace of fault with an oblique-slip, which is
subsiding (Ballance 1980; Biddle & Christie-Blick 1985b).
Fault core—Fault core is the innermost of the three fault zone components (Chester & Logan
1986; Smith et al. 1990). It can include a single-slip surface, unconsolidated clay-rich gouge
zone, brecciated and geochemically altered zone or highly indurated cataclasite zone (Caine
1999).
Fault-flank depression—A depression formed between the subsidiary folds of a strike-slip
fault system (Crowell 1976; Biddle & Christie-Blick 1985).
Fault rupturation (propagation) — The fault rupturation and propagation is triggered by
preceding stress buildup (e.g., Yielding et al. 1981; Rockwell et al. 1988; Klinger &
Rockwell 1989; Philip et al. 1992; Treiman 1995). It is accompanied by elastic unloading.
Fault seal—There are several different types of fault seals including: (a) shale smear fault
seal, (b) cementation fault seal, (c) grain-size reduction fault seal, and (d) juxtaposition fault
seal (Adams & Dart 1998).
Fault-slice ridge—Characterized as the linear topographic high that forms at the boundary of
a fault-bounded uplifted block within a fault zone (Crowell 1974b; Biddle & Christie-Blick
1985). It is also referred to as a pressure ridge (Tchalenko & Ambraseys 1970).
Fault splay—Defined as a subsidiary fault that is genetically related to the more prominent
fault it merges with. They are common near the termination of a major strike-slip fault, unless
they intersect with another strike-slip fault (Biddle & Christie-Blick 1985).
Fault strand—A single individual fault from a set of closely spaced, parallel or subparallel
faults in a fault system (Biddle & Christie-Blick 1985).
Fault-wedge basin—Basin formed by means of extension at a releasing junction between
two predominantly strike-slip faults with the same sense of offset (as defined by Crowell
1974a; Biddle & Christie-Blick 1985). Also referred to as a wedge graben (Freund 1982).
Fault zone architectural styles—There are several fault zone architectural styles including:
(a) single-fracture fault, (b) localized deformation zone, (c) distributed deformation zone, and
(d) composite deformation zone (Caine 1999).
Fault zone permeability structures—Fault zone permeability structures include a localized
conduit, localized barrier, distributed conduit and combined conduit-barrier (Caine 1999).
Flexural rebound— Happens in response to the mechanical unloading of the lithosphere
during extension (Petit et al. 2007). It may couse permanent uplift of rift flanks.
Flexural strength— A parameter that describes the elastic response of a multilayer, treating
its flexural response as one elastic layer of a given thickness. One of its implications for
lithosphere is that lithospheric loads are compensated flexurally rather than in an isostatic
manner. Temporal changes in lithosphere rheology induced by rifting process and subsequent
cooling cause significant changes in flexural response of the lithosphere. For example, the
flexural strength (rigidity) of the transform margin during the active continental-oceanic and
subsequent early passive margin stages is low due to high temperatures (Karner & Watts
1982; Holt & Stern 1991; Kooi et al. 1992). As a result they usually lack any distinct flexural
uplift or isostatic rebound (Nemčok et al. 2012b).
Flow lines—Flow lines (Rosendahl 2004, pers. comm.) are linear features in gravity maps
that are parallel to oceanic fracture zones and perpendicular to the spreading ridges, typically
used as a spreading vector indicator.
Footwall— The rock volume below the dipping fault. A schematic example of a footwall
block is shown in EA Fig. 5.
Graben— An area of structural low, represented by an elongated, relatively depressed block,
bounded by two controlling normal faults of opposite dips (Bates & Jackson 1980; Biddle &
Christie-Blick 1985).
Gravitational instability— Differential erosion commonly leads to gravitational instability
by exposing the ends of a set of beds along valley sides (EA Fig. 6; Park 1997). The
consequences of a gravitational instability resulting from such an exposure range from
comparatively minor bending of the strata close to the ground surface to gravitational
collapse structures, which can be hundreds of meters long. Instability produced by a
topographic slope may be greatly accentuated if the beds dip towards the slope. This scenario
often results in the slipping of competent layers that are resting on weak material towards the
topographic low. Percolating ground water may also reinforce this process considerably.
Gravity gliding— May be defined as gravity-driven lateral extension and vertical
contraction, regardless of basal slope and coherence of the body. However, in most cases a
more detailed description should be used in addition to (or instead of) this one to capture the
behavior of rock masses when deforming under gravity (Schultz-Ela 2001). Also referred to
as 'gravity spreading'.
Gravity-driven fluid flow— If precipitation and infiltration of water in regions of high rift
and passive margin-related elevations are sufficient to recharge the water table, a continuous
supply of groundwater is available to maintain flow (Deming 1994 and references therein).
Topographically driven fluid flow in such a system occurs nearly everywhere, where there is
topographic relief.
Growth strata— Can be defined as strata with thickness variations across faults (Pochat et
al. 2009).
Half-graben— Characterized as a structural low, controlled by one controlling normal fault.
Hanging wall—Defined as a rock volume above the dipping fault.
Hanging-wall block – Forms between conjugate normal faults that define necking zones. It
consists of a more-or-less undeformed pre-rift successions of upper crust (and possibly lower
crust). It is only preserved in failed rifts. In successful rifts, these blocks are delaminated
during extensional deformation, creating a residual hanging-wall block and several
delaminated blocks, which are all preserved as extended continental blocks on one of the two
conjugate margins, and they correspond to extensional allochthons and breakaway blocks.
The residual hanging-wall block is bounded by the necking zone on it proximal side and by
younger exhumation faults on its distal side (Masini et al. 2013). A schematic example of a
hanging-wall block is shown on EA Fig. 5.
Hard link—At least one fault that is mappable at the scale of observation, connecting the
overlapping faults through their overlap zone.
Helicoidal geometry of Riedel shears— Helicoidal geometry is the true 3-D geometry of
Riedel shears. It is characterized by fault traces at progressively higher acute angles from the
principal displacement zone as they approach it. Their geometry in the cross section is
slightly listric before they join the principal displacement zone.
High-relief accommodation zone— If the opposing non-overlapping half grabens are active
roughly in the same time, there is no space problem to be solved in the linkage area. This area
tends to be “left behind” as a relatively unsubsided structure, called high-relief
accommodation zone (Reynolds 1984; Burgess et al. 1988; Rosendahl 1987).
Hinge line—The line of the maximum curvature of the hanging wall, located on the opposite
side of the half-graben from its controlling normal fault.
Horsetail splay (structure)—A fault splay from a set of curved fault splays near the end of a
strike-slip fault (EA Fig. 1). The whole set forms an array crudely resembling a horse’s tail
(Biddle & Christie-Blick 1985). It serves as a fault displacement-dissipation structure.
Horst—Structural high controlled by two controlling normal faults of opposite dips on its
opposite sides.
Hyper-extension— Extreme extension of a continental crust that thins it down to less than
10 km. Often used in relation to a hyper-extended margin or distal margin, or in relation to
the last rifting stage before break-up (Péron-Pinvidic & Manatschal 2009; Mohn et al. 2010;
Péron-Pinvidic et al. 2013).
Hyper-extended zone— Part of a rifted margin that underwent hyper-extension, thinning the
crust to less than 10 km (EA Fig. 3). It tends to be referred to as a distal margin. It is
separated from an unextended or weakly extended crust of a proximal margin by a necking
zone (Péron-Pinvidic & Manatschal 2009; Mohn et al. 2010; Péron-Pinvidic et al. 2013).
Hyper-extended margins are known in the South Atlantic, Australia-Antarctica, Gulf of
Mexico, Iberia-Newfoundland, Arctic region or in the Lower Austro-Alpine nappes of southeast Switzerland and Northern Italy. In Iberia-Newfoundland, the distal (hyper-extended)
domains are more assymetrical than their proximal, symmetrical counterparts (Massini et al.
2013).
Inter-basin high in the pull-apart basin terrain—The high developed between individual
sub-basins of the pull-apart basin that is detached along the ductile detachment (Sims et al.,
1999). The pull-apart basin above a thick ductile detachment horizon commonly centers
along a dominant Riedel shear that directly links propagation tips of main bounding strikeslip faults. It controls a simple flip-flop basin asymmetry. The basin is relatively thin. Its
development can be followed by development of neighbor basins divided from the initial
pull-apart basin by inter-basin highs, which do not usually contain syn-rift sediments because
they form areas “left behind” as relatively unsubsided structures, characterized by high-relief.
Integrated lithospheric yield strength profile—It is a strength profile that can be estimated
using the extrapolation of laboratory failure criteria (Goetze & Evans 1979; Brace &
Kohlstedt 1980; Kirby 1985; Carter & Tsenn 1987; Ranalli & Murphy 1987). It is the lower
of the two strength curves calculated according to functions described in brittle deformational
behavior and ductile deformational behavior that controls the deformational behavior of the
respective layer. The integrated strength profile becomes weaker with the decreasing strain
rate, with weaker rheologies representing individual layers, with increasing temperature and
with increasing fluid pressure. It is also controlled by the stress regime, being strongest in the
thrust regime and weakest in the normal faulting regime.
Intermediately strong relatively young stable lithosphere—Represents a lithosphere,
which became stable during the Phanerozoic eon (Huismans & Beaumont 2005). The
intermediately strong relatively young stable lithosphere contains a weak lower crustal layer,
which results in its integrated yield strength profile indicating four-layer strength distribution
characterized by brittle upper crustal and upper mantle layers separated by ductile lower crust
and lower mantle layers (e.g., Brun 1999; Huismans & Beaumont 2005). The rifting of this
kind of lithosphere is characterized by viscous flow in the lower crust, separating brittle
upper crust and upper mantle layers. The cause behind the progressive loss of the rift
asymmetry is the decrease in coupling between upper crust and upper mantle along the lower
crust, which contributes into a more distributed style of the upper crustal deformation.
Intra-basin high in the pull-apart basin terrain— The high developed in the center of the
pull-apart basin that is detached along the brittle detachment (Sims et al., 1999). The pullapart basin of this detachment scenario is represented by a single basin with dominant normal
faults controlling the basin geometry and subsidence. Unlike in the basin above ductile
detachment, this intra-basin high is minimal.
Irrotational deformation (strain) – See pure shear.
Isostatic rebound—Isostatic rebound is the positive vertical movement of the area in
reaction to the overburden removal. In the case of normal fault-related footwall uplift, it is the
lateral removal of the hanging wall (Rosendahl 1987). In this case, the maximum uplift
resides in the footwall next to the maximum displacement of the controlling normal fault. In
the map view, this uplift diminishes to zero along the fault trace towards the propagation tips
of the normal fault. Such uplift distribution in the rift zone controlled by unlinked normal
faults controls the development of variable depositional environments. Potential reservoirrock prone sediments enter the rift zone through the relay ramps between controlling normal
faults while depocenters are potentially prone to source rock deposition.
Joining horse-tail structure— Faults splaying off the tip of a transform fault, which are
younger than and not directly geometrically and initially not kinematically linked to
controlling normal faults of the neighbor rift zone (Nemčok et al., 2016 – this volume). This
case takes place where the propagating transform fault encounters an already developed rift
zone and this zone is located in the direction of the break-up propagation, and where the rift
zone is older than the one on the opposite side of the transform fault where the transform
propagation started.
Kinematic analysis— Displacement-based study of a movement pattern, disregarding force
or stress (Biddle & Christie-Blick 1985; modified from Spencer 1977).
Lateral hydrocarbon migration— Capability of oil and gas to transport in the lateral
direction (EA Fig. 7). Lateral (and vertical) hydrocarbon migration is often caused by two
factors – buoyancy and gravity. Buoyancy-driven migration tends to result from an existing
lateral fluid density gradient (Nemčok et al. 2005 and references therein). Gravity-driven
migration includes migration that results from compaction, when the rock porosity decreases
and the hydrocarbon is being pushed out (Shi & Wang 1986). Compaction- and gravitydriven migration is further constrained by topography (faults, and the presence and dip of
pervious sedimentary layers). Lateral hydrocarbon migration tend to progress along
unconformities and through sandy, silty, and limestone units (Doligez, 1987).
Left-hand overstep (stepover)—Defined as an overstep, in which every following fault or
fold segment occurs to the left of the segment from which it is being viewed (EA Fig. 1;
Campbell 1958; Wilcox et al. 1973; Biddle & Christie-Blick 1985). In cross sections, the
direction from which the overstep is being viewed needs to be specified. Also referred to as
left-stepping overstep (stepover).
Left-lateral—A fault offset where the far side is apparently displaced to the left (EA Fig. 1)
when compared to the near side (Biddle & Christie-Blick 1985).
Left separation—Refers to the separation of blocks in strike-slip faulting where the far side
of a fault is apparently displaced to the left when compared to the near side (Biddle &
Christie-Blick 1985).
Left slip—Slip component measured parallel to the strike of a fault, where the far side of the
fault is displaced to the left (EA Fig. 1) when compared to the near side (Biddle & ChristieBlick 1985).
Left-stepping overstep (stepover)— See left-hand overstep.
Listric fault—A curved, spoon-shaped fault, which is generally concave in the upper part
and flattening downward. It can be characterized by normal or reverse separation (Biddle &
Christie-Blick 1985).
Lithospheric multilayer—A multilayer is the most usual representation of the rheologic
distributions in the lithosphere, dictated by the analog material and numerical modeling need
for a simplified representation of various types of the continental lithosphere. Provided that
one has the seismic velocity and density data on the lithospheric multilayer, the lithology of
individual lithospheric layers can be implied from rock tables containing seismic velocity and
density measurements from rock specimens (Christensen & Mooney 1995). The amount of
rheologically specified layers of this multilayer varies among types of lithospheres. See
strong old stable lithosphere, intermediately strong relatively young stable lithosphere and
weak young thickened lithosphere.
Low-angle fault—A fault that dips less than 30o.
Low-relief accommodation zone—If the opposing overlapping half-grabens are active
roughly in the same time, the space problem in the linkage area results in a development of
the positive structure known as low-relief accommodation zone (Rosendahl 1987; McClay et
al. 2002). The subsidence of these hinged highs pales in comparison with the depocenters at
graben-bounding faults, but they do subside.
Lower crustal decoupling— Brittle faults often detach in a region of plasticity. The depth of
such a detachment can be referred to as the depth of necking (Braun & Beaumont 1989;
Weissel & Karner 1989). It is assumed that the necking depth relates to a depth-dependent
crustal-decoupling zone, as described by Driscoll & Karner (1998) and Karner & Driscoll
(1999). Crust above the decoupling zone is referred to as the upper plate and the crust and
lithospheric mantle below the decoupling zone is referred to as the lower plate.
Lower crustal ductile flow-controlled uplift—Lateral flow of the ductile lower crust to the
region underneath the marginal ridge of the transform margin from the surrounding regions
undergoing higher pressure drives a local uplift (Sage 1994). Understanding the magnitude of
this process in a transform margin setting requires further studies.
Magma-dominated margin – See magma-rich margin.
Magma-poor margin— Defined by a range of features including: wide continent-ocean
transition zone (generally 100-200 km), peridotite rocks of unroofed mantle exposed at the
paleo-sea floor, relatively flat-lying, serpentinized detachment faults and low volumes of
magmatic rocks (Boillot & Froitzheim 2001; Whitmarsch et al. 2001; Huismans & Beaumont
2011). In the past it was referred to as a non-volcanic margin (Mutter et al. 1988) or magmastarved margin (Sawyer et al. 2007).
Magma-rich margin— Defined mostly by larger volumes of magmatic rocks than would be
expected based on the passive upwelling of “normal” asthenosphere beneath the thinning
crust. Magmatic material is present in the form of thick igneous crust, including the
underplated material. It is said to be a result of magmatism with potential temperature
(temperature it would have if brought to surface rapidly) of 1300°C. The total thickness often
exceeds 10 km. In places, the emplacement of seaward-dipping lavas is accompanied by large
landward-dipping faults (Reston & Manatschal 2011). In the past it was referred to as a
volcanic margin (Mutter et al. 1988) or magma-dominated margin (Sawyer et al. 2007).
Magma-starved margin – See magma-poor margin.
Mantle exhumation stage— The third and last stage of rifting, follows after the stretching
and thinning stages. This stage is characterized by downward concave faulting that reaches to
the rising asthenosphere and generates large fault offsets (more than 10 km). Subcontinental
mantle, representing an exhumed fault zone (also called top-basement detachment fault or
sometimes extraction fault) is exhumed during this stage (Manatschal 2004; Lavier &
Manatschal 2006). The type and origin of the exhumed mantle rocks depend on the geometry
of the detached fault (Manatschal 2004). The deeper mantle levels are only exhumed if this
fault roots in the asthenosphere. If it roots at the top of the mantle, it can logically only
exhume the uppermost mantle levels. The type of mantle rocks present also depends on their
distance from the continent. Exhumed mantle present near the continent is formed by spinel
peridotite mixed with garnet-pyroxenite layers, which equilibrated at lower temperatures. On
the other hand, mantle rocks present further from the continent are formed by pyroxenitepoor peridotite that equilibrated in the plagioclase stability field (Müntener & Picardo 2003).
According to Péron-Pinvidic et al. (2013), the stage of mantle exhumation is not mandatory
in the evolution of a rifted margin, in contrast to the hyperextension phase.
Marginal ridge—Not present at all transform margins, the marginal ridge occupies their
most oceanward location, the outer corner (e.g., Le Pichon & Hayes 1971; Scrutton 1979;
Mascle & Blarez 1987; Todd & Keen 1989; Basile et al. 1993, 1996). It is more-or-less
parallel to the margin. There are several models explaining its development. The first of them
is the continental sliver model (Le Pichon & Hayes 1971). The second model is the
transpressional origin of the ridge (Mascle & Blarez 1987; Huguen et al. 2001; Attoh et al.
2004). The third model is the thermal expansion-driven uplift model where the heat source is
oceanic crust together with the spreading center passing the continent during the active
continental-oceanic transform stage (Scrutton 1979; Mascle & Blarez 1987; Todd & Keen
1989; Lorenzo et al. 1991; Lorenzo & Vera 1992; Basile et al. 1993; Gadd & Scrutton 1997;
Vagnes 1997). The fourth model is the flexural uplift model where uplift can be responding
to tectonic denudation or erosional removal of a rock column (e.g., Clift & Lorenzo 1999;
Basile & Allemand 2002). The fifth model is based on the ductile flow of either lower crust
(Sage 1994) or those lithospheric layers which reach ductile behavior (Reid 1989; Vagnes
1997). The facts that transforms undergo various kinematic adjustments during their
continent-oceanic stage discussed in this article and some transforms have marginal ridge
while others do not indicates that different combinations of the aforementioned models must
have developed those ridges that exist. It is probable that a few more factors constrain their
development, such as the (a) location of the break-up trajectory with respect to the preexisting geometry of the pull-apart floor in the case of brittle detachment that contains a
distinct polarity flip and intra-basin high, (b) location of the break-up trajectory with respect
to the location of narrow inter-basin highs in the case of the pull-apart basin with ductile
detachment (see Sims et al. 1999 for different pull-apart structural grains), (c) location of the
break-up trajectory with respect to the complex topography of the horse-tail structure
constrained by the uneven distribution of the isostatic rebound due to various kinematic
behavior of involved faults.
Master fault—A dominant fault in a fault system (Wilcox et al. 1973; Rodgers 1980; Biddle
& Christie-Blick 1985). The term is almost synonymous with principal displacement zone
defined by Tchalenko & Ambraseys (1970).
Mean stress—The arithmetic mean of the principal stresses.
Mechanical stratigraphy—Mechanical stratigraphy is the information about the rheology of
all the layers in the deforming multilayer. Mechanical stratigraphy of the lithosphere controls
the rift style (Huismans & Beaumont 2005, 2008, 2011). The strong old cratonic lithosphere
starts to rift using conjugate shear zones until one of them is preferred by shear strain
softening. This shear progressively develops into the upward-convex shallow-dip fault zone,
which is one of the main accommodators for the crustal and upper mantle necking and the
ascent of the lower mantle and asthenosphere. The developed rifts are narrow and
characterized by deep syn-rift basins and pronounced flexural uplifts on one or both flanks.
The rifting type is the narrow asymmetric rifting of the whole lithosphere. The intermediately
strong relatively young stable lithosphere starts to rift using conjugate shear zones until one
of them is preferred by shear strain softening. In this case the extension in the upper mantle is
transferred into the upper crust via viscous drag in the ductile lower crust. The coupling is
sufficient to transfer the localized strain into the overlying upper crust. Less coupling
between the upper mantle and upper crust, however, allows for a wider rift mode and the
reduction of the rift flank uplift due to the reduction of the flexural stresses. The rifting style
is the wider asymmetric rifting of the upper lithosphere combined with narrow symmetric
rifting of the lower lithosphere. The weak young thickened lithosphere starts to rift using
conjugate shear zones until one of them is preferred by shear strain softening inside the upper
lithospheric mantle. In this case the decoupling between the upper crust and upper mantle
along the ductile lower crust is very advanced. As a result, the extension transfer to the upper
crust results in a wide region as the upper mantle drives almost pure shear deformation of the
upper crust via simple shear drag in the lower crust. Minimum coupling between the upper
mantle and upper crust allows for a wide rift mode and the suppression of the rift flank uplift
due to a suppression of the flexural stresses. The rift width is determined by the integrated
strength of the decoupling lower crustal layer. The rifting style is the wide symmetric rifting
of the crust combined with narrow symmetric rifting of the mantle lithosphere.
Microcontinent—A microcontinent is a block of continental crust of variable size
surrounded by oceanic crust, at variable distance from the parent continent. It has a
bathymetric high, positive Free-air anomaly, magnetically quite area, prograding sedimentary
wedges on its flanks, continental crust architecture, a crustal age older than that of the
surrounding oceanic crust, and a heat flow regime different from that of the surrounding
oceanic crust (Rey et al. 2003). Microcontinent release mechanisms include (Müller et al.
2001; Collier et al. 2008; Péron-Pinvidic & Manatschal 2010; Nemčok et al. 2012a, 2015b):
(a) plume refocusing, (b) competing wrench faults, (c) competing horse-tail structure
elements, (d) effect of consecutive tectonic events controlled by different stress regimes and
(e) competing rift zones.
Monocline—A sub-cylindrical fold that has one limb formed by the horizontal reference line
and the other one inclined.
Multiple overstep—Refers to a series of discontinuities between approximately parallel
overlapping or underlapping strike-slip faults, as defined by Biddle & Christie-Blick (1985).
Necking— A process, during which pinch-and-swell structures (areas of extremely different
thickness) get created (Fossen 2010). In relation to continental crust, during necking the
continental crust thins down from unextended or weakly extended to extremely extended
crust. This gradual change in thickness occurs in a so-called necking zone.
Necking zone— Separates the proximal and distal parts of a rifted margin. It represents an
area where the crustal thinning was most significant, relates to a specific wedge shape of the
crust where the Moho (seismically) is said to “define an inflection point associated with a
drastic crustal thinning from ± 30 km to less than 10 km” (Lavier & Manatschal 2006; Mohn
et al. 2010; Péron-Pinvidic & Manatschal 2009). It contains an area of the margin that is
characterized by a basinward increase in total accommodation space (Sutra & Manatschal
2012). Also, the necking zone records the localization, migration, and depth of deformation
that followed the formation of proximal basin (Péron-Pinvidic et al. 2013). EA Fig. 3 shows
an example of necking zone separating a little extended proximal margin from a
hyperextended distal margin.
Negative flower structure—A flower structure with mainly normal, upward-diverging fault
splays (EA Fig. 2). This structure tends to be associated with a prominent synformal
structure/structures that is/are located in the strata above the faults, or cut by the faults
(Harding 1983, 1985; Biddle & Christie-Blick 1985; Harding et al. 1985). Also referred to as
normal flower structure.
Non-volcanic margin—See magma-poor margin.
Normal flower structure—See negative flower structure.
Oblique margin—The oblique margin develops when oblique rifting is followed by oblique
oceanic accretion. Typical examples are the Northern Exmouth and Canning-Browse
continental margin segments of Western Australia. Magnetic stripe anomalies of the Argo
Abyssal Plain in front of them are at about 45o to both segments. Oceanic transforms of the
Argo Plain are perpendicular to stripe anomalies and do not indicate any continental stage of
their development.
Oblique slip—Slip component, which has characteristics of both dip-slip and strike-slip
(Biddle & Christie-Blick 1985).
Oblique-slip fault— Fault structure where displacement is accomplished by a combination
of strike-slip and dip-slip (Biddle & Christie-Blick 1985). Also referred to as a slip-oblique
fault.
Ocean-Continent transition – Transition from the distal continental margin to the first
oceanic crust (Manatschal 2004). Localization with respect to other terms is shown on EA
Fig. 3.
Oceanic accretion— A process by which new material is added to the oceanic crust, making
the oceanic crust grow – may be identified with the formation of new oceanic crust.
Oceanic crust (normal oceanic crust)—Normal oceanic crust is a result of organized seafloor spreading (EA Fig. 3; Rosendahl 2004, pers. comm.). See organized sea-floor spreading
for further explanation. The oceanic crust in East India is 3.9-6.8 km thick, having an average
thickness of 5.4 km. Oceanic crust thickness in Gabon varies between 4 and 6.5 km. These
values are very similar to thickness values reported for oceanic crust in offshore Cameroon
and Equatorial Guinea (Rosendahl et al. 2005; Nemčok & Rosendahl 2006). They fall within
the published values for the normal oceanic crust of the Atlantic Ocean (Rosendahl &
Groschel-Becker 1999). The global average thickness, however, is 7.08 km (White 1992).
Unlike in Gabon, the thickness of the oceanic crust in offshore East India varies even in
relatively small regions. The top of the oceanic crust is imaged by a strong continuous
reflector. The base of the oceanic crust is indicated by a fairly continuous reflector, which
does not reach the acoustic impedance contrast typical for the boundary between top crust
and overlying sedimentary cover. The image of the oceanic crust itself is relatively
transparent, with more-or-less horizontal weak reflector patterns or is sometimes seismically
unlayered and structureless. It does not show any internal deformations, apart from zones
around oceanic fracture zones, which represent the only complexity in the oceanic crust
image. Some of these are characterized by the offset of the Moho surface.
Oceanic fracture zone - Oceanic fracture zones represent 50-150 km wide zones of ridges
and troughs (Rosendahl 2004, pers. comm.). While axes of these zones are perpendicular to
the spreading ridge active at their time of formation, trends of individual ridges and troughs
can be parallel to or at a very acute angle to the axis (e.g., Nemčok et al. 2015a).
Offlap— A progressive diminution of conformable strata in the lateral extent in passing
upwards from older to younger strata (EA Fig. 4), so that each stratum leaves a portion of the
underlying one exposed (http://findwords.info/term/offlap). The term was first used by
Grabau (1913).
Onlap—An upward termination of initially horizontal or inclined strata against an initially
inclined surface (EA Fig. 4; Biddle & Christie-Blick 1985; modified from Mitchum 1977).
Organized sea-floor spreading—The organized sea-floor spreading takes place after several
millions of years of disorganized spreading that takes place after the continental break-up in
the magma-rich settings or continental lithospheric mantle exhumation in magma-poor
settings. Example of the disorganized spreading comes from gradual, tortuous and mostly
sub-aerial spreading such as takes place in Afar while an example of the organized spreading
comes from the Red Sea (Rosendahl 2004, pers. comm.). The example of their most likely
transition comes from the oceanic crust adjacent to the western and central segments of the
Falklands transform margin (Edwards et al. 2013). In the west, the cold edge effect of the
continental crust has resulted in either poorly developed or missing oceanic crust right next to
the transform margin, as spreading centers were laterally clearing it during the stage of
disorganized spreading. Further oceanward a spreading ridge reorganization took place
causing a 30 km northward jump of the spreading center. This was associated with an
abandonment of a crustal sliver adjacent to the margin and the development of more
organized spreading ridge system.
Orogenic collapse—Orogen collapsing under its own weight represents the situation that
occurs when gravitational forces created by the top of the orogen exceed the strength of the
orogenic wedge (e.g., England 1983; Platt 1986; Dewey 1988; England & Houseman 1988).
Overlap—(1) Characterized as the distance between the ends of overlapping parallel faults.
It is measured parallel to the faults (Rodgers 1980; Mann et al. 1983; Aydin & Nur 1985;
Biddle & Christie-Blick 1985) and is usually applied to strike-slip faults in map view. It is
almost synonymous with the term separation of Segall & Pollard (1980); (2) Defined as a
relationship between two superimposed stratigraphic units onlapping a certain surface where
the upper unit extends beyond the line of pinch-out in the lower unit (Biddle & Christie-Blick
1985).
Overstep—(1) Characterized as the discontinuity between two approximately parallel faults
that are overlapping or underlapping (EA Fig. 1). Overstep is observable in both a map view
and cross section, usually it is applied to strike-slip faults in map view. However, it can be
observed on both strike-slip and dip-slip faults (Aydin & Nur 1985; Biddle & Christie-Blick
1985). Overstep is further divided into solitary oversteps, multiple oversteps, releasing
oversteps, and restraining oversteps. It is also referred to as stepover (Aydin & Nur 1982a, b,
1985); (2) Also defined as the case where the eroded edge of older, generally tilted or folded
sedimentary strata is unconformably overlaid by one or more stratigraphic units (Biddle &
Christie-Blick 1985).
P shear—A single fault from the set that generally develops in a simple shear after the
formation of Riedel shears. P shears form at an angle to the principal displacement zone, and
have the same sense of movement as the Riedel R shears. Compared to P-faults, the principal
displacement zone is approximately of the same magnitude but of an opposite sign
(Skempton 1966; Tchalenko & Ambraseys 1970; Biddle & Christie-Blick 1985). Also
referred to as secondary synthetic strike-slip fault.
Parallel hydrocarbon migration— Takes place in a planar reservoir horizon or fault
component (Hindle 1997). It tends to maintain petroleum pathways and hydrocarbon
concentration.
Passive rifting— Passive rifting is characterized by passive asthenosphere upwelling in
response to overlying layers separation controlled by regional tectonic extension (Keen 1985;
Ruppel 1995; Kincaid et al. 1996). During early rifting, the lithosphere is thinned only in
response to extension. The passive asthenospheric upwelling drives many secondary
processes such as decompression-driven melting, crustal/lithospheric magma underplating,
eruption of continental flood basalts, the onset of secondary convection and development of
large thermal gradients between extended and unextended regions (e.g., McKenzie & Bickle
1988; Ruppel 1995; Huismans et al. 2001).
Passive transform margin stage of the transform margin development— During this
stage, the transform fault is no longer active. One of its sides is formed by a continental
margin and the other one by a progressively cooling oceanic plate (Mascle & Blarez 1987).
Permeability stratigraphy—Stratigraphic division of rocks based on differing values of
rock permeability.
Piercing points—Defined as the points of intersection of previously adjacent linear features
(real or constructed, e.g. pinchout lines of sedimentary wedges, offset streams, or facies
boundaries combined with structure contours) on opposite sides of a fault. Piercing points
enable the net slip on the fault to be determined (Crowell 1959; Biddle & Christie-Blick
1985).
Plane strain—Strain where the intermediate strain axis Y remains unchanged. It is
characterized by the shortening along the Z axis compensated by extension along the X axis.
Plateau (deep-sea continental plateau)—About 30% of transform margins contain deep-sea
continental plateaus (Loncke et al. 2013). Most of them appear in boundary areas of oceanic
systems of different ages (e.g., the Guinea Plateau, the Demerara Plateau, and the Exmouth
Plateau). Their development requires several subsequent rift events. Their bounding segments
contain both rifted and transform segments.
Polarity flip— Polarity flip can occur in the case of a rift system. It is associated with a fact
that individual half-grabens forming rift zones can be linked in various ways, which can be
grouped into linkages of half-grabens with opposing and similar polarities. Polarity flip can
also occur in the case of breaking-up plate segments. In the case of the strong old stable
lithosphere, the decision about footwall versus hanging wall sides is made in relatively late
stages of the rift-drift transition (Huismans & Beaumont 2005). Natural case studies indicate
that this decision takes place during the hyper-extension stage (Manatschal et al. 2013). The
final result is a narrow asymmetric rifting of the whole lithosphere with obvious footwall and
hanging wall sides. In the case of the intermediately strong relatively young stable
lithosphere, the decision is apparently made during the early stages of the rift-drift transition
(Huismans & Beaumont 2005). The final result is a narrow asymmetric rifting of the upper
lithosphere combined with narrow symmetric rifting of the lower lithosphere, where footwall
and hanging walls are also obvious. Finally, in the case of the weak young thickened
lithosphere, it is fairly difficult to distinguish between footwall and hanging wall (Huismans
& Beaumont 2005). The final result is a wide symmetric rifting of the crust combined with
narrow symmetric rifting of the mantle lithosphere, where the knowledge of the crustal
architecture alone may not be sufficient for determining the footwall and hanging wall sides.
Pop-up—Characterized as a relatively uplifted block located between thrusts verging in
opposite directions. The term originally applied to structures in thrust and fold belts (Butler
1982; Biddle & Christie-Blick 1985).
Positive flower structure—A flower structure with mainly reverse, upward-diverging fault
splays (EA Fig. 2). This structure tends to be associated with a prominent antiformal
structure/structures that is/are located in the strata above the faults, or cut by the faults
(Harding 1983; Biddle & Christie-Blick 1985; Harding et al. 1985). Nice examples of the
positive flower structures were formed by transpressional reactivation of Aptian strike-slip
faults of the Romanche transform fault zone (Davison et al. 2015). Also referred to as reverse
flower structure.
Post-break-up uplift—See rift shoulder uplift for all potential driving mechanisms.
Numerical models (e.g., Huismans & Beaumont 2005) indicate that the post-break-up uplifts
are most prominent in the case of the strong old stable lithosphere where conjugates are
almost symmetric in the case of strong lithospheric mantle and strong lower crust. This
symmetry is progressively lost with the mantle lithosphere becoming weaker. The post-breakup uplifts can occur in the case of the intermediately strong relatively young stable
lithosphere although they are subdued. The same applies to their asymmetry. The uplifts do
not develop in the case of the weak young thickened lithosphere. On the contrary, both
conjugate margins have a tendency to subside.
Post-rift unconformity—The post-rift unconformity is the result of erosion that post-dates
the rifting accommodated by normal faults and predates the thermal subsidence event. This
unconformity can be amalgamated with younger unconformity at both flanks of the sag basin
developed over the failed rift units, where the younger one becomes developed as associated
with the migrating-away flexural bulging.
Post-transform unconformity—The post-transform unconformity is the result of erosion
that post-dates the transform activity. It is diachronous, younging oceanward, along the
transform margin (Basile 2015).
Pressure ridge— See fault-slice ridge.
Principal displacement zone—Characterized as a relatively narrow zone, in which most of
the slip on a given fault takes place (Tchalenko & Ambraseys 1970; Biddle & Christie-Blick
1985).
Progradation—Defined as the outward building of sediment in the transport direction.
Usually the progradation progresses from a shoreline towards a body of water (Biddle &
Christie-Blick 1985).
Proto-oceanic crust—Proto-oceanic crust is the very first type of oceanic crust produced in a
magma-poor setting (Rosendahl 2004, pers. comm.). It is developed by continental
lithospheric mantle unroofing (see Whitmarsh et al. 1996; Manatschal 2004). It tends to
develop in the settings involving old stable lithosphere, which undergoes break-up in the
crust-first mantle-second scenario. Its development follows the stretching, necking and hyperextension stages, and precedes the mantle break-up. It is usually composed of rock suite
representing the upper lithospheric mantle (Rosendahl et al. 2005), although it may contain
slivers of lower crust, as indicated by forward gravity modeling (Meyers et al. 1998) and the
scattered bodies of mafic magmatic rocks (Manatschal 2004). The maximum thickness of the
proto-oceanic crust in offshore Gabon reaches 10 km (Nemčok et al. 2012a). Its thickness in
East Coast India ranges in thickness from 4.5 to 11.2 km, having an average thickness of 8.7
km (Nemčok et al. 2012a). These values are similar to values reported in offshore Cameroon
and Equatorial Guinea (Rosendahl et al. 2005; Nemčok & Rosendahl 2006). While the top of
the proto-oceanic crust is imaged by a higher reflectivity zone, the base in most cases does
not have any distinct signature, which represents one of the main differences between seismic
images of oceanic and proto-oceanic crusts. Further difference is in the highly contorted
reflector patterns representing the proto-oceanic crust and not just deformation along the
oceanic fracture zones, documenting its faulted and deformed character. Fracture zones are
defined by large faults that sometimes dissect the entire crust and usually coincide with large
offsets of the Moho surface.
Proximal margin— Part of the rifted margin that corresponds to platforms - inboard
continental crust stretched by none or weak extensional forces). Also called proximal domain
(of the rifted margin). It is separated from the distal margin by a necking zone (EA Fig. 3).
According to Froitzheim & Eberli (1990), proximal margins developed as a result of an older
phase of rifting than that of the distal margins, because their syn-tectonic sediments are older
than the distal domain syn-rift deposits. They are characterized by graben and half-graben
basins filled with wedge-shaped syn-tectonic sedimentary units. Typical for proximal
margins are the high-angle normal listric faults related to fault-bounded rift basins, which
affect the brittle upper crust. Crustal thinning tends to be moderate, which is associated with
β-factors that are typically less than 2 (Keen & de Voogd 1988; Lau et al. 2006; Mohn et al.
2012, Péron-Pinvidic et al. 2013). Compared to the distal margin, only modest amounts of
accommodation space are created in proximal margin during syn- and post-rifting phase. This
is due to the moderate extension in the area (Péron-Pinvidic et al. 2013).
Pull-apart basin—(1) A basin that forms as a result of crustal extension at a releasing bend,
releasing overstep along a strike-slip fault zone or its terminating horse-tail structure
(preferred definition; Burchfiel & Stewart 1966; Crowell 1974b; Mann et al. 1983; Biddle &
Christie-Blick 1985). It is nearly synonymous with the term rhomb graben; (2) Any kind of
basin, which forms as a result of crustal extension (Klemme 1980; Bois et al. 1982; Biddle &
Christie-Blick 1985). In comparison to rift basins, extension intervals of pull-apart basins are
short. They are frequently characterized by fast post-rift subsidence due to lateral heat flow
(Pitmann & Andrews 1985).
Pull-apart basin with brittle detachment—Pull-apart basins developed in the continental
crust, which was thicker and cooler during the early stages of stretching (Nemčok et al.
2012b). The pull-apart basin of this detachment scenario is represented by a single
rhombohedral basin with dominant normal faults controlling the basin geometry and
subsidence, which are clearly distinguishable from controlling strike-slip faults (Sims et al.
1999). The subsidence in this basin initiates earlier than the subsidence in basins developed
above ductile detachment. It is because there is no ductile layer, which would take a certain
share of the extensional deformation before the extension in the brittle rock section developed
faults controlling the basin subsidence. Unlike in the basin above ductile detachment, the
within-basin high in the center of the basin is minimal (Sims et al. 1999).
Pull-apart basin with ductile detachment—Pull-apart basins developed in the continental
crust, which was thinner and warmer during the mature stages of stretching (Nemčok et al.
2012b). Thickness of the ductile detachment horizon exerts important control on the basin
geometry (Sims et al. 1999). The reason is that basin geometry is controlled by Riedel shears,
which have geometry controlled by the thickness of the ductile detachment layer. The pullapart basin above a thick ductile detachment horizon commonly centers along a dominant
Riedel shear that directly links propagation tips of main bounding strike-slip faults. It controls
a simple flip-flop basin asymmetry. The basin is relatively thin. Its development can be
followed by the development of neighbor basins divided from the initial pull-apart basin by
between-basin highs, which do not usually contain syn-rift sediments because they form areas
“left behind” as relatively unsubsided structures, characterized by high-relief. The basin
above a thick ductile detachment undergoes subsidence, which takes part later than the
subsidence in the basin above a thin ductile detachment. This is because the ductile layer
takes a certain share of the extensional deformation before the extension in the brittle
overlying rock section develops faults, which control the basin subsidence.
Pure shear— Characterized as homogeneous deformation that does not result in internal
rotation when affected by plane or general strain. Lines of particles, which are parallel to the
principal axes of the strain ellipsoid retain the same orientation after deformation as they had
before (Hobbs et al. 1976). Also referred to as an irrotational deformation or irrotational
strain (Biddle & Christie-Blick 1985). In McKenzie’s (1978) pure shear model for the
extensional basin, the lithosphere is instantaneously thinned in a uniform fashion, producing
gravitational and thermal instabilities. In this model, the equilibration of the gravitational
instability occurs instantaneously, producing the syn-rift subsidence. The thermal
equilibration causes time-dependent thermal subsidence stage. Numerous successful
applications of the uniform stretching model (McKenzie 1978) to initial-stage rift terrains
indicate that most of them were formed by mechanisms dominated by the passive rifting of
the lithosphere. Studies of passive margins, structures which went beyond advanced rifting
development, document that there are practically no passive margins, which have been
developed by homogeneous stretching. They require various modifications of the uniform
stretching model to explain the observed data (e.g., depth dependent-stretching - Royden &
Keen 1980; Beaumont et al. 1982; Roberts et al. 1997; Driscoll & Karner 1998; Davis &
Kusznir 2004; the effect of different rheology development histories of various portions of
the lithospheric multilayer – e.g., Faugére & Brun 1984; Brun 1999; the interaction between
lithospheric rheology and erosion - Burov & Poliakov 2001; the effect of mineral phase
transitions on rifting - O’Connell & Wasserburg 1972; Yamasaki & Nakada 1997;
Artyushkov et al. 2000; Simon & Podladchikov 2006). Active rifting plays a progressively
more important role in advanced rifting stages (e.g., Huismans & Beaumont 2005, 2008,
2011).
Pure strike-slip— A purely lateral displacement on a fault, either a right or a left lateral
shear motion. It results in no net addition or subtraction of an area perpendicular to the
motion direction (Twiss & Moores 2007).
Push-up—Characterized as a block, which was elevated by means of crustal shortening at a
restraining bend or restraining overstep along a strike-slip fault zone (Aydin & Nur 1982a;
Mann et al. 1983; Biddle & Christie-Blick 1985).
Receding (retreating) side of the strike-slip fault tip— Receding (retreating) side of the
strike-slip fault propagation tip is the area affected by the mean stress reduction and the
maximum principal stress σ1 rotation towards an almost perpendicular position with respect
to the principal displacement zone (Homberg et al. 1997). Receding sides of large strike-slip
faults frequently contain a horse-tail structure with oblique-slip and normal faults (e.g.,
Nemčok et al. 2012c).
Relay ramp—Synonymous for strike-slip bridge and stepover (Peacock & Sanderson 1995).
Releasing bend— see divergent bend.
Releasing fault junction—Intersection between two strike-slip faults, which is associated
with crustal extension and the formation of a fault-wedge basin (Christie-Blick & Biddle
1985).
Releasing overstep—Defined as a right overstep between right-slip faults or a left overstep
between left-slip faults. It is associated with crustal extension and the formation of basins
between the faults (Christie-Blick & Biddle 1985). A strike-slip fault relationship to poles of
rotation in the region control the areas of local transtension and transpression (Mann 2013).
Restraining bend—See convergent bend.
Restraining fault junction— Intersection between two strike-slip faults, which is associated
with crustal shortening and uplift between the faults (Christie-Blick & Biddle 1985).
Restraining overstep—Defined as a right overstep between left-slip faults or a left overstep
between right-slip faults. It is associated with crustal shortening and uplift between the faults
(Christie-Blick & Biddle 1985). Similar to releasing overstep, a strike-slip fault relationship
to poles of rotation in the region controls the areas of local transpression and transtension
(Mann 2013).
Retrogradation— Defined as the landward back-stepping of sedimentary units. Usually the
retrogradation progresses from a shoreline, e.g. landward migration of facies belts (Biddle &
Christie-Blick 1985).
Reverse flower structure – See positive flower structure.
Rhombochasm—Gap with parallel sides in continental crust, which is filled by oceanic
crust. For example the Gulf of California (Carey 1958; Biddle & Christie-Blick 1985).
Rhomb graben—Can be characterized as a basin formed by crustal extension at a releasing
bend or releasing overstep in a strike-slip fault zone (Freund 1971; Aydin & Nur 1982b;
Biddle & Christie-Blick 1985). Also referred to as a pull-apart basin, especially sharp or
angular pull-apart basins (as defined by Burchfiel & Stewart 1966; Crowell 1974a, b).
Rhomb horst—Can be characterized as a block raised by crustal shortening at a restraining
bend or restraining overstep in a strike-slip fault zone (Aydin & Nur 1982b). It is almost
synonymous with the term push-up, particularly in the case of push-ups that are angular in the
map view (Aydin & Nur 1982a; Mann et al. 1983).
Ridge jump— A sudden change of location of the spreading ridge axis. Time and distance of
the ridge-jump are the two main parameters describing a ridge jump. Relocation of the ridge
can be caused by hot-spot influenced asymmetric spreading. Mechanisms considered for the
hot-spot influenced ridge jump include “lithospheric tension induced by buoyant and
convecting asthenosphere (Mittelstaedt & Ito 2005), mechanical and thermal thinning of the
lithosphere due to hot flowing asthenosphere (Jurine et al. 2005), and penetration of magma
through the plate (Kendall et al. 2005)” as mentioned by Mittelstaedt et al. (2008). Ridge
jumps were reported from many places along the mid-ocean ridge network – for example
from areas near Shatsky Rise, Ascension, Iceland, Galápagos, or the area of East India-Elan
Bank-Antarctica, where the original ridge between Antarctica and Elan Bank moved between
Elan Bank and East India (Mittelstaedt et al. 2008; Sinha et al. 2015).
Riedel shear—Defined as two sets of shear fractures oriented at Φ/2 (R shears) and 90°- Φ/2
(R′ shears) to the principal displacement zone in simple shear (Φ being the internal
coefficient of friction, commonly about 30o; modified from Tchalenko & Ambraseys 1970).
Also referred to as synthetic fault (shear).
Rift margin—See extensional margin.
Rift shoulder— Uplifted flanks, or shoulders, are common features of rifts. The rift shoulder
is typically asymmetric. It has a steep scarp facing the sea and a gentle slope facing the land.
Rift shoulders have been explained by models based on transient and permanent uplift
mechanisms. Transient uplift models are related to the thermal effects of rifting and include
depth-dependent extension (Royden & Keen 1980; Hellinger & Sclater 1983; Watts &
Thorne 1984; Morgan et al. 1985), lateral heat flow (Steckler 1981; Cochran 1983; Alvarez et
al. 1984; Buck et al. 1988) and secondary convection under rift shoulders (Keen 1985;
Steckler 1985; Buck 1986). These mechanisms, which can create 500-1500 m shoulder
elevation, operate only during the elevated thermal regime of the lithosphere. Created
positive topography will decay over the time period equivalent to the thermal time constant of
the lithosphere, which is roughly 60 Ma. Permanent uplift models are related to the magmatic
underplating (Cox 1980; Ewart et al. 1980; McKenzie 1984; White & McKenzie 1988) and
lithospheric unloading and/or plastic necking (Zuber & Parmentier 1986; Parmentier 1987;
Braun & Beaumont 1989; Issler et al. 1989; Weissel & Karner 1989; Chery et al. 1992).
Subsequently, processes of erosion and deposition started to be incorporated into existing
models of the rift shoulder uplift (e.g., van Balen et al. 1995; van der Beek et al. 1995; Burov
& Cloetingh 1997).
Rift shoulder retreat—The rift shoulder retreat takes place during the erosional destruction
that follows the rift shoulder uplift, due to the fact that the lateral erosion is much faster than
the vertical one (e.g., van Balen et al. 1995; Burov & Cloetingh 1997).
Right-hand overstep (stepover)—Defined as an overstep, in which every following fault or
fold segment occurs to the right of the segment from which it is being viewed (EA Fig. 1;
Campbell 1958; Wilcox et al. 1973; Biddle & Christie-Blick 1985). In cross sections, the
direction from which the overstep is being viewed needs to be specified. Also referred to as
right-stepping overstep (stepover).
Right-lateral—A fault offset where the far side is apparently displaced to the right (EA Fig.
1) when compared to the near side (Biddle & Christie-Blick 1985).
Right separation—Refers to the separation of blocks in strike-slip faulting where the far side
of a fault is apparently displaced to the right when compared to the near side (Biddle &
Christie-Blick 1985).
Right slip—Slip component measured parallel to the strike of a fault, where the far side of
the fault is displaced to the right (EA Fig. 1) when compared to the near side (Biddle &
Christie-Blick 1985).
Right-stepping overstep (stepover)—See right-hand overstep (stepover).
Roll-over anticline—An anticline developed by the gravitational collapse of the hanging
wall above the listric normal fault, characterized by strata steepening towards the controlling
fault.
Rotation poles— Pole of rotation is a term often used in relation to plate reconstruction and
strike-slip faulting. In plate reconstruction, the motion between two rigid plates can be
described as rotation of the plates about a rotation axis. The rotation pole then refers to a
point where spreading (rotation) axis cuts the Earth’s surface (logically there are two rotation
poles on opposing sides of the earth for each rotation axis). This rotation pole is also called
the Euler pole. Pole of rotation and angular velocity are two parameters that need to be
defined when dealing with the relative motion of the plates (Kearey et al. 2009). It is said that
the pole of rotation generally doesn’t remain fixed (Harrison 1972).
Sag basin— Generally characterized as a sedimentary basin formed by the thermal
subsidence. There are several different types of sag basins interpreted by various authors,
including intra-cratonic sag basins (Middleton, 1989), sag basins in areas of failed rifts
(McKenzie, 1978; Sclater et al., 1980; Allen and Allen, 1990 and references therein) and
those that appear in rift zones before the final breakup (Karner et al., 2003; Moulin et al.,
2005; Huismans and Beaumont, 2008)). Sag basins in failed rifts are also called aulacogen
basins. They are formed by prolonged thermal relaxation after the thermal source causing the
rifting was completely withdrawn upon forming the mid-oceanic ridge in the other area. The
thermal relaxation causes a regional thermal subsidence, which results in a sag basin.
Aulacogen basins tend to be dominated by fluvial and lake facies. Sag basins present in rifts
prior to the final breakup are formed during quiescent period before the final breakup of the
lithospheric mantle and formation of first oceanic crust. Their sedimentary fill overlies the
typical rift-sequence and can be overlaid by salt.
Sea-floor spreading— Formation of a new oceanic crust by solidification of rising mantle
material, which occurs in mid-oceanic ridges with an active divergent motion of lithospheric
plates. According to various geological models, the first segments of oceanic crust are formed
when the asthenosphere thins down enough for it to concentrate thermal and mechanical
constraints, enabling the margin to enter a magmatic phase (Cannat et al. 2009). This
magmatic phase should be represented by a pulse in the magmatic activity, which triggers the
final lithospheric break-up, and it is supposed to be generally unrelated to the previous rifting
deformation phases (Péron-Pinvidic et al. 2013). Sea-floor spreading can be divided into
symmetric and asymmetric, depending on whether the oceanic crust on both sides of midoceanic ridge forms symmetrically. Generally, it tends to be modelled as symmetrical, but on
a regional scale, small- or large-scale asymmetries appear. The reason for asymmetric seafloor spreading is not well understood yet. It seems to appear mostly in areas overlying hotspots, which may indicate ridge propagation towards mantle plumes or minor ridge jumps
between ridges and plumes as the cause for asymmetric spreading (Müller et al. 1998).
Secondary synthetic fault (shear)—See P shear.
Sediment entry point— A place, where sediments enter a basin or other depositional
environment. It can be a breach in the mountain ridge, river channel or other environmental
path that allow sediments easy passage into their depositional area. Sediments in or close to
the entry point can still be present in the form of channel flow sediments, and are dominantly
clastic in nature (Cleary et al. 1977).
Separation—Defined as: (1) Apparent displacement between previously adjacent surfaces on
opposite sides of a fault, measured in any direction (Reid 1913; Crowell 1959, Biddle &
Christie-Blick 1985); (2) Distance between parallel, overlapping strike-slip faults, measured
perpendicularly to the fault direction (Rodgers 1980; Mann et al. 1983; Aydin & Nur 1985;
Biddle & Christie-Blick 1985); (3) Distance between parallel, overstepping strike-slip faults
(either overlapping or underlapping), which is measured in the sense parallel to the faults
(Segall & Pollard 1980; Biddle & Christie-Blick 1985). This definition of is almost
synonymous with overlap as defined by Rodgers (1980).
Shear—Characterized as a strain, which results from stresses that usually cause lateral
movement between parts of a body in a direction parallel to their plane of contact (Bates &
Jackson 1980; Biddle & Christie-Blick 1985).
Shear margin— Shear-slip margin (Rabinowitz & Labrecque 1979; Scrutton 1979), also
referred to as transform margin.
Sidewall rip-out—Asymmetric sidewall rip-out is a typical feature of strike-slip faults that
forms by a lateral jump of the active slip zone during adhesion along a section of the
principal displacement zone (Swanson 2005). Subsequently, the new irregular slip zone after
this jump creates an indenting asperity that plows through the host rock during continued
adhesion or is cut off by a renewed motion along the main section of the fault. A rip-out
translation during adhesion controls the structural asymmetry with trailing extensional and
leading contractional ends of the rip-out block.
Simple shear—Characterized as a homogeneous deformation that results in constant volume
internal rotation when under plane strain. In the deformed state, a single family of parallel
material planes remains undistorted and parallel to the same family of planes in the
undeformed state (Hobbs et al. 1976; Biddle & Christie-Blick 1985). Synonymous with term
rotational deformation (strain). In Wernicke (1985) a simple shear model for extensional
basin, the whole lithosphere undergoes a normal simple shear of uniform-sense. This model
is characterized by a lateral offset of the post-rift subsidence form the syn-rift subsidence.
Subsequent dynamic modeling revealed that this model is applicable to the post-orogenic
extensional collapse of the young and weak lithosphere (Braun & Beaumont 1989; Buck
1991; Govers 1993).
Simple strike-slip or wrench fault—Characterized as a strike-slip or wrench fault with
purely lateral displacement of adjacent blocks. No crustal shortening or extension occur in
this type of fault (Christie-Blick & Biddle 1985). Also referred to as simple parallel strikeslip or wrench fault and slip-parallel fault (Wilcox et al. 1973; Mann et al. 1983;
respectively).
Sinistral— Leftward in motion (sinistral slip is a left slip; EA Fig. 1), pertaining to the left
side (as defined by Biddle & Christie-Blick 1985).
Slickenside—Defined as polished or smoothly striated surface on both sides of a fault plane
that forms as a consequence of motion along the fault (Bates & Jackson 1980; Biddle &
Christie-Blick 1985).
Slip—Characterized as relative displacement of previously adjacent points on opposite sides
of a fault. It is measured along the fault surface (Reid 1913; Crowell 1959; Biddle & ChristieBlick 1985).
Slip-oblique fault – See oblique-slip fault.
Soft link—The situation where there is no fault, which is mappable at the scale of
observation, connecting the overlapping faults through their overlap zone, which can be
represented by a relay zone.
Splay—See fault splay. Terms are generally synonymous.
Spreading ridge— Mid-oceanic ridge (or mid-oceanic spreading center) is a narrow linear
zone in the central part of the oceanic floor, where new oceanic crust forms as a consequence
of the divergent movement of lithospheric plates.
Stepover—See overstep.
Strain hardening—The effect in which the controlling stress loading must be increased to
maintain a fixed strain rate.
Strain partitioning—Physical decomposition of strain into different components. For
example, a regional transpression can be partitioned into the areas undergoing pure shear and
areas undergoing simple shear.
Strain rate—The rate at which the strain accumulates.
Strain softening—The effect in which the controlling stress loading must be decreased to
maintain a fixed strain rate.
Strength—The amount of stress a rock can sustain before it fails or yields. See Brittle
deformational environment and Ductile deformational environment for further details.
Stretching—Stretching, s, equals 1 + e, where e is the elongation.
Stretching stage—It is the initial stage of rifting, during which the continental crust starts
stretching under extensional forces. It is characterized by the formation of distributed listric
normal faults, which cut through the brittle upper crust and sole out at mid-crustal levels. The
faults bound the rift basins up to 4 km deep and 30 km wide, fault offsets are less than 10 km.
During this stage, the total extension of the area is limited (low to moderate) and both
hanging wall and footwall are subsiding (Lavier & Manatschal 2006). During this stage, the
crust or lithosphere does not undergo any major thinning.
Strike-slip bridge—Synonymous with relay ramp, overstep and stepover (Gamond 1987;
Ramsay & Huber 1987).
Strike-slip duplex—A duplex formed in the strike-slip fault zone.
Strike-slip margin—Strike-slip margin (Nagel et al. 1986) is synonymous for transform
margin.
Strong old stable lithosphere—A strong old stable lithosphere can be characterized as a
lithosphere, which was stabilized in pre-Cambrian times and represents an old cratonic
lithosphere (Christensen & Mooney 1995; Huismans & Beaumont 2005). Sometimes the
strong old stable lithosphere does not contain a weak lower crustal layer. This results in its
integrated yield strength profile indicating a brittle strength from its upper surface to the base
of the upper mantle and ductile strength controlling only the deformation of the lower mantle
(e.g., Brun 1999; Huismans & Beaumont 2005). Rifting in the strong cratonic lithosphere is
prone to narrow rift development. It is characterized by rift asymmetry and coupled upper
crust with upper mantle.
Subsidence curve of the orthogonal rift basin—It contains a relatively long-lasting
moderately steep subsidence curve segment representing the rift event, followed by a
shallower-dipping one representing the thermal subsidence event.
Subsidence curve of the pull-apart basin—It contains a fairly short and steep subsidence
curve segment representing the pull-apart opening event, followed by a fairly shallowlydipping one representing the post-pull-apart regime.
Syn-rift sediments—Strata thickening towards the controlling rifting-related faults and
terminated by the post-rift unconformity.
Syn-tectonic sediments—Strata displaying the thickness changes controlled by syndepositional activity of the fault that controls their depocenter.
Synthetic fault (shear)—See Riedel shear.
Taper break – The nearest point to the coast where the crustal thickness is reduced to ≤
10 km. It is located at the outer end of the necking domain (Osmundsen & Redfield 2011).
Tear fault—Refers to a strike-slip or oblique-slip fault, which bounds or is located within an
allochthon. Tear faults form as a result of regional extension or regional shortening, and
accommodate displacement within the allochthon or between the allochthon and adjacent
structural units, depending on their position within the allochthon (Biddle & Christie-Blick
1985).
Tectonic subsidence—Characterized as the part of subsidence caused by tectonics. Its
component in the whole subsidence of the area can be calculated by a back-stripping
technique, which allows us to remove sedimentary sequences, compaction, water-depth
changes and balancing isostasy, and their effect on the subsidence of the area (Biddle &
Christie-Blick 1985 and authors within).
Tension—Characterized as a system of stresses that usually results in the lengthening or
increasing the volume of a material. Two main types of tension are uniaxial tension, which
has one principal tensile stress of non-zero value; and general tension with two principal
tensile stresses. Tensile principal stress can also occur with a compressive principal
stress/stresses (Means 1976; Biddle & Christie-Blick 1985).
Tension (T) fracture—See extension fracture.
Thermal expansion-controlled uplift—Heat transfer into the transform margin from the
adjacent oceanic crust and the spreading ridge during the active continent-oceanic stage
causes temporary thermal expansion that causes an uplift (Todd & Keen 1989). In
comparison to other uplift-driving factors, this one is relatively small (Nemčok et al. 2015c).
Thermal subsidence— Characterized as the component of overall tectonic subsidence
caused by thermal contraction (Sleep 1971; Parsons & Sclater 1977; Biddle & Christie-Blick
1985).
Thermal transients at transform margins— Thermal transients at transform margins are
thermal perturbations developed during the active continental-oceanic stage of transform
development. Their spatial and temporal distribution is controlled by (Nemčok et al. 2015c):
(a) the regional cooling of the elevated thermal regime developed by rifting in neighbor pull-
apart terrains that culminated at break-up, (b) the short-term local extreme heating driven by
the passing-by spreading center, (c) the long-term regional heating by the passing-by
spreading center, (d) the long-term heating by the newly accreted oceanic crust on both sides
of the spreading center, and (e) the effect of the background pre-rift heat flow on the cooling
rate of the above-mentioned thermal events.
Thick-skin pull-apart basin—A pull-apart basin whose detachment fault reaches deeper
into the lithosphere, resulting in a thicker slab of crust involved in the pull-apart motion
(Royden, 1985). Theoretically the entire lithosphere may get involved in the case of a thick
skin pull‐apart basin. Example of a thick-skin pull-apart comes from the Gulf of Paria in
Venezuela (Sims et al., 1999).
Thin-skin pull-apart basin—A pull-apart basin whose detachment fault does not reach
deeper parts of the lithosphere, resulting in a thinner slab of crust involved in the pull-apart
motion (Royden, 1985). Example of a thin-skin pull-apart comes from the Death Valley in
California (Sims et al., 1999).
Thinning stage—The second stage of rifting, following after the stretching phase. This stage
occurs in the necking zone and is said to be the reason for its formation (Péron-Pinvidic et al.
2013). During this stage, the continental crust thins down from ± 30 km to less than 10 km. It
contains large-scale conjugate detachment structures, also called thinning faults, which affect
the upper crust and the lower crust/upper mantle. These fault are decoupled at the level of
shear zone in the weak middle crust. The thinning stage affects the future distal margin in a
narrow zone, which tends to be buried under thick sediments. Little to no evidence of upper
crustal deformation accompanying the thinning has been evidenced (Lavier & Manatschal
2006; Péron-Pinvidic et al. 2013).
Thrust fault—A fault that results in the shortening of material, commonly (but not
necessarily) bedding. It is a contraction fault of a map-scale. The term applies to faults of any
dip, a special distinction going to faults in an active slip, which usually dip at less than 30°
(McClay 1981; Biddle & Christie-Blick 1985)
Transcurrent fault—(l) Defined as a strike-slip fault, which commonly cuts through
supracrustal sediments into an igneous and metamorphic basement. Typically subvertical at
depth, almost synonymous with the term wrench fault (Moody & Hill 1956; Freund 1974;
Biddle & Christie-Blick 1985); (2) A long, subvertical strike-slip fault, which tends to cut
strata in a direction more-or-less perpendicular to strike (Geikie 1905; Dennis 1967; Biddle &
Christie-Blick 1985).
Transfer strike-slip faults—Transfer faults serve as accommodation structures linking rift
units or rift zones undergoing different amounts of extension or two different extensional
provinces (Gibbs 1984; Rosendahl 1987; Milani & Davison 1988; Versfelt & Rosendahl
1989; Tari et al. 1992). There are numerous examples of transfer faults that were developed
by the reactivation of pre-existing crustal weakness, e.g., examples from the Tanganyika and
Malawi rift zones (Versfelt & Rosendahl 1989) and the Gulf of Suez-Red Sea rift (Younes &
McClay 2002). Obtuse angles between rift-bounding faults and reactivated shear zones
promote local extension, acute ones local contraction (Younes & McClay 2002).
Transform fault—Characterized as a strike-slip fault, which acts as a lithospheric plate
boundary, terminating at both ends against major tectonic feature that also represents a plate
boundary. Such features can involve oceanic ridges, subduction zones, other transform faults,
etc. (Wilson 1965; Freund 1974). The development of a transform fault contains an active
continental stage, active continental-oceanic stage and a passive margin stage, described
under respective terms. It is important to note that the entire fault is not active at a given time
and different segments of the transform fault can be coevally in different stages of its
development. There are many examples of transform faults that were initiated by the
reactivation of the pre-existing crustal weakness, e.g., the South Tasman transform
reactivating structural grain of the Paleozoic Avoca fold belt (Gibson et al. 2013), the De
Geer transform reactivating structural grain of the Caledonian Inuitian orogen (Harland 1965;
Faleide et al. 1993; Doré et al. 2015), the Coromondal transform reactivating structural grain
of the Proterozoic orogen (Nemčok et al. 2012c; Sinha et al. 2015), some segments of the
Romanche transform margin reactivating structural grain of the Pan-African orogen
(Antobreh et al. 2009), Pernambuco-Ngaoundere and Patos transforms reactivating shear
zones of the Pan-African orogen (e.g., Chang et al. 1988). There are six main transform
settings where the transform gets into contact with the oceanic crust (Morley et al. 2013). The
first setting is an intra-continental rifting that develops two rift zones separated by a transfer
zone, such as in the Tanganyika-Rukwa-Malawi case (see Rosendahl 1987; Versfelt &
Rosendahl 1989) that eventually develops into the continental break-up, represented by the
Equatorial Atlantic analog (e.g., Scotese 1998; Golonka 2000). The second setting as the
subduction zone with a reentrant characterized by the opposite subduction polarity, the sides
of which are linked with its back-arc basin that reached sea-floor spreading with two
transforms like in the Scotia arc case (Dalziel et al. 2013). The third setting is the transform
linking the rift system that reached sea-floor spreading with collisional orogen, such as the
Dead Sea transform-Red Sea rift system-Taurus Mountains case (Morley et al. 2013). The
fourth setting is the advance of the orogenic arc overriding the oceanic lithosphere at the
subduction zone characterized by slab roll-back, along the continental plate, such as the
Caribbean example (Garciacaro et al. 2011; Sanchez et al. 2015). The fifth setting is the
continent overriding an obliquely trending oceanic ridge-transform system, represented by
the San Andreas case (Wallace 1990). The sixth setting is an obliquely convergent margin
with a segment of highly oblique plate convergence, represented by the Andaman Sea
example (Curray 2005; Morley 2015).
Transform margin—Characterized as a plate margin dominated by strike-slip deformation,
which is formed by one or a system of transform faults (Biddle & Christie-Blick 1985) that
reached the passive margin stage everywhere along its strike during their development
(compare Wilson 1965 with Mascle 1976, Lonsdale 1985, and Mascle & Blarez 1987).
Transform margins are present at various scales. While the entire Romanche margin is
understood as the transform margin (e.g., Mascle & Blarez 1987; Basile 1990; Pontoise et al.
1990; de Caprona 1992; Basile et al. 1993, 1998; Mascle et al. 1995, 1996, 1998; Benkhelil
et al. 1998a, b; Clift & Lorenzo 1999; Bigot-Cormier et al. 2005), it can be further
subdivided into three smaller-scale transform margin segments and two smaller-scale pullapart segments (Antobreh et al. 2009). While some of the transform margins, such as the
Coromondal transform margin (Nemčok et al. 2012c, Sinha et al. 2015) are neighbors to rift
margins, others can be neighbors to pull-apart margins (e.g., several transform margins
segments of the Sergipe-Alagoas margins and North Gabon margins – Rosendahl et al. 2005;
Nemčok et al. 2012a, St Paul and Romanche transform margin segments – Nemčok et al.
2012b, 2015b). The ends of the transform margin are divided into the inner and outer corners
(Basile 2015). The inner corner represents the landward end of the transform, represented by
the horse-tail structure that is not affected by the transform itself. The outer corner,
representing the oceanward end of the transform contains the horse-tail structure
subsequently cut by the transform itself as it kept developing through all the stages of its
development. The transform margin is characterized by a relatively narrow shelf and slope
(Towle et al. 2012a, b; Addis et al. 2013a, b). Transform margins are usually characterized
by the sediment entry point system remaining relatively stable after the onset of the passive
margin stage of the transform margin development. Examples come from the Romanche and
Guyana transform margins. The same applies to the shelf break location that does not
undergo distinct shifts (e.g., Singh 1999 – Konkan segment of the West Indian transform
margins; Nemčok et al. 2015a). All these sedimentary characteristics are also related to a lack
of the post-breakup uplift of this margin.
Transform margin provenance—The provenance on a transform margin may change
through time as the juxtaposition changes as a result of the lateral motion during its first two
development stages (see Active continental and Active continental-oceanic stages of the
transform margin development), in addition to any changes that result from the progressive
exhumation of the provenance areas.
Transpression—Characterized as a system of stresses that results in both lateral movement
(sinistral or dextral) and crustal shortening (EA Fig. 1). It occurs in areas of oblique
shortening (Harland 1971; Sylvester & Smith 1976; Biddle & Christie-Blick 1985).
Transpressional strike-slip fault zone—A strike-slip fault zone that contains a component
of contraction perpendicular to the zone (Naylor et al. 1986). It usually contains different
types of shears in the map view and a positive flower structure in the cross section.
Transtension—Characterized as a system of stresses that results in both lateral movement
(sinistral or dextral) and crustal extension (EA Fig. 1). It occurs in areas of oblique extension
(Harland 1971; Biddle & Blick 1985).
Transtensional strike-slip fault zone—A strike-slip fault zone that contains a component of
extension perpendicular to the zone (Naylor et al. 1986). It usually contains either throughgoing fault zone or several slightly en echelon Riedel shears in a map view, and a negative
flower structure in the cross section.
Underlapping faults—Characterized as sub-parallel faults, which overstep without
overlapping (Pollard & Aydin 1984; Biddle & Christie-Blick 1985).
Uplift—The uplift is a mechanism that causes the ascent of the reference surface with respect
to geoid, causing an overall cooling of the uplifting area (England & Molnar 1990).
Vertical hydrocarbon migration— Capability of oil and gas to transport in the vertical
direction, along faults or similar structures that allow passage in the upwards-downwards
direction (EA Fig. 7). Most often the vertical migration is achieved by transport along fault
structures or possibly in a more-or-less straight up-down motion (Wang et al. 2014).
Hydrocarbon migration (vertical and lateral) is often caused by two factors – buoyancy and
gravity. Buoyancy-driven migration tends to result from existing lateral fluid density gradient
(Nemčok et al. 2005). Gravity-driven migration includes migration resulting from
compaction, when the rock porosity decreases and the hydrocarbon is being pushed out (Shi
& Wang 1986). Compaction- and gravity-driven migration is further constrained by
topography (faults and presence, and the dip of pervious sedimentary layers).
Volcanic margin—See magma-rich margin.
Weak young thickened lithosphere—The weak young thickened lithosphere can be
characterized as a lithosphere, which did not have a sufficiently long time for its thermal
equilibration after the last orogenic event and finds itself in the stage of thermal thinning (see
Zoetemeijer 1993). The integrated yield strength profile of this lithospheric type indicates a
four-layer strength distribution characterized by a brittle upper crustal and upper mantle
layers separated by ductile lower crust and lower mantle layers (see Brun et al. 1994; Brun
1999; Corti et al. 2003), which, contrary from the intermediately strong young stable
lithosphere, is characterized by strong decoupling along the ductile lower crust (Huismans &
Beaumont 2005). The rifting of the weak young thickened lithosphere further highlights the
loss of both rift asymmetry and its narrow character with weaker lithospheric strength. The
presence of a highly viscous lower crust essentially decouples the upper crust from the mantle
lithosphere.
Wedge graben—See fault-wedge basin.
Wrench fault—See transcurrent fault. Terms are nearly synonymous.
Y-shear—Characterized as a fault, which forms in response to simple shear. With ongoing
deformation it gradually accommodates the majority of the movement along the principal
displacement zone (Bartlett et al. 1981; Biddle & Christie-Blick 1985).
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Figure captions - Glossary
EA Fig. 1. Sketches illustrating structural patterns in map view, including (a) en echelon
pattern, (b) overstep pattern, (c) left stepping pattern, (d) right stepping pattern, (e) horsetail
splay, (f) restraining (convergent) bend, dextral strike-slip fault, right lateral movement, right
slip and transpression, and (g) releasing (divergent) bend, sinistral strike-slip fault, left lateral
movement, left slip and transtension (modified from Biddle & Christie-Blick, 1985).
EA Fig. 2. 3D block diagrams of (a) negative (normal) and (b) positive (reverse) flower
structures (modified from Woodcock & Fischer, 1986).
EA Fig. 3. Idealized cross-section across a magma-poor rifted margin, luustrating terms such
as oceanic crust, ocean-continent transition, serpentinized mantle, hyper-extended zone
(domain), distal margin, necking zone and proximal margin (modified from Mohn et al.,
2012).
EA Fig. 4. Types of stratal terminations (modified from Catuneanu, 2002).
EA Fig. 5. Block diagram showing basic types of faulting along with terms such as hanging
wall block, footwall (block) and dip (modified from http://www.britannica.com/science/faultgeology).
EA Fig. 6. (a) Grounds for formation of gravitational instability (Park, 1997). (b) Result of
gravitational instability (Park, 1997).
EA Fig. 7. Sketch showing vertical and minor horizontal hydrocarbon migrations (Wang et
al., 2014).