Identifying fault segments from 3D fault drag analysis (Vienna Basin, Austria)

Journal of Structural Geology 55 (2013) 182e195

Contents lists available at ScienceDirect

Journal of Structural Geology journal homepage: www.elsevier.com/locate/jsg

Identifying fault segments from 3D fault drag analysis (Vienna Basin, Austria) Darko Spahi
c a, *, Bernhard Grasemann a, Ulrike Exner b a b

Department for Geodynamics and Sedimentology, University of Vienna, Althanstrasse 14, A-1090 Vienna, Austria Naturhistorisches Museum Wien, A-1010 Vienna, Austria

a r t i c l e i n f o

a b s t r a c t

Article history: Received 16 February 2013 Received in revised form 24 July 2013 Accepted 25 July 2013 Available online 24 August 2013

The segmented growth of the Markgrafneusiedl normal fault in the late Miocene clastic sediments of the central Vienna Basin (Austria) was investigated by construction of a detailed three-dimensional (3D) structural model. Using high resolution 3D seismic data, the fault surface and marker horizons in the hanging wall and the footwall of the Markgrafneusiedl Fault were mapped and orientation, displacement and morphology of the fault surface were quantified. Individual, fault segments were identified by direct mapping of the deflection of the marker horizons close to the fault surface. Correlating the size of the identified segments with the magnitude of fault drag and displacement distribution showed that fault evolution progressed in several stages. The proposed method allows the detection of segments that are not recorded by the magnitude of displacement or fault morphology. Most importantly, detailed mapping of marker deflections in the hanging wall could help to constrain equivalent structures in the footwall, which may represent potential hydrocarbon traps. Ó 2013 Elsevier Ltd. All rights reserved.

Keywords: Fault growth Displacement gradient 3D seismic Fault curvature Fault drag Hydrocarbon trap

1. Introduction Extension of the Earth’s crust is a dynamic process that often results in deformations, which differ in scale and morphology. After a fault segment nucleates, this planar discontinuity propagates changing individual and immediate spatial configurations. Initially isolated fault segments may grow and link to larger fault systems under regional extension (Crider and Pollard, 1998). Consequently, a complex fault surface may represent the product of the cumulative displacement events accommodated by progressive enlargement and coalescence of the individual segments (e.g., Nicol et al., 1995; Willemse et al., 1996; Peacock, 2002; Walsh et al., 2002). Segmented normal faults have been described in rift systems (e.g., Coward et al., 1987; Davidson, 1994; McLeod et al., 2002), intracontinental basins (Lohr et al., 2008) and passive margins (e.g., Childs et al., 1996; Marchal et al., 2003) and occur over a wide range of scales (e.g., Kristensen et al., 2008; Peacock and Sanderson, 1991; Gawthorpe et al., 2003; Walsh et al., 2003).

* Corresponding author. Current address: Zoi i Aleksandra Kosmodemyanskikh 36A, 125130 Moscow, Russian Federation. Tel.: þ7 962 9209347. E-mail addresses: [email protected], [email protected] (D. Spahi
c). 0191-8141/$ e see front matter Ó 2013 Elsevier Ltd. All rights reserved. http://dx.doi.org/10.1016/j.jsg.2013.07.016

The individual segments were successfully identified by quantifying the displacement distribution across the faults in combination with studies of fault morphology (e.g., Watterson, 1986; Barnett et al., 1987; Walsh and Watterson, 1989; Cartwright et al., 1995; Nicol et al., 1996; Walsh et al., 2003; Marchal et al., 2003; Kim and Sanderson, 2005; Lohr et al., 2008). Displacement distributions along an isolated elliptical fault typically shows a decrease from the fault centre towards the tips (Walsh and Watterson, 1989). Such displacement gradients cause perturbation strain (Passchier et al., 2005; Kocher and Manctelow, 2005) in the host rocks resulting in a deflection of marker layers. This fault drag compensates the discontinuous displacement gradient along the fault by continuous deformation in the host rock (Barnett et al., 1987; Peacock and Sanderson, 1991; Reches and Eidelman, 1995; Kocher and Manctelow, 2006; Gupta and Scholz, 2000; Grasemann et al., 2005). Mechanical models suggest that fault drag, which forms a hanging wall anticline, corresponds with a syncline in the footwall of the fault (Resor and Pollard, 2012). Flanking structures (Passchier, 2001) or fault drag have been described from various metamorphic conditions, ranging from high-grade gneisses, marbles, to non-metamorphic sedimentary rocks (e.g., Wiesmayr and Grasemann, 2005). Rollover anticlines above listric faults, which can be also interpreted as fault drag (Barnett et al., 1987; Reches and Eidelman, 1995; Grasemann et al.,

D. Spahi
c et al. / Journal of Structural Geology 55 (2013) 182e195

2005) form important hydrocarbon traps in sedimentary basins and passive margins. However, the mechanical explanation for the rollover and fault drag model is very different. While the rollover concept positions displacement of a deformable hanging wall above a rigid footwall (e.g. McClay, 1990), fault drag represents a host rock deformation in the deformable hanging- and footwall induced by displacement gradient along the fault. Subsequently, the mechanical models predicting the formation of hydrocarbon traps play a critical role in the analysis of the main parameters of a petroleum system (Magoon and Dow, 1994) because the differences in the mechanics of faulting affect the time of source rock burial, fluid migration, trap and secondary faults formation. The fault drag model represents an attractive alternative solution for the formation of rollover anticlines where a connection to a low-angle detachment fault at depth is not evident (e.g. Spahi
c et al., 2011), replacing earlier interpretation of a rollover anticline in a hanging wall above a listric normal fault (e.g., McClay, 1990; for a discussion see also Brun and Mauduit, 2008). In this study, a large z16 km Markgrafneusiedl normal fault system located in the central Vienna Basin, Austria was investigated using a high-resolution industrial 3D seismic dataset (OMV Exploration and Production GmbH). A 3D structural model was created by detailed mapping of the fault plane and associated synand anticlines in the layered sediment host rocks. The depthconverted geometrical model enabled a spatial correlation between the displacement distributions along the fault with respect to the adjacent fault drag cluster to be demonstrated. We conclude that fault drag is a very useful additional criterion to constrain the development of linked normal fault systems originating from initial independent isolated segments.

a

183

2. Geological background The investigated fault, covering 5000 km2, is located in the central part of the Vienna Basin (Fig. 1a). The basin has been extensively studied for hydrocarbon exploration (e.g., Ladwein, 1988; Wessely, 1988; Strauss et al., 2006) and for seismic activity (Decker et al., 2005; Hinsch et al., 2005a; Beidinger and Decker, 2011; Beidinger et al., 2011). The system of Intra Carpathian Tertiary basins developed as a consequence of gravitational collapse of thick orogenic terrains (Ratschbacher et al., 1991; Corver et al., 2009). Subsequently a new system of thin-skinned basins became superimposed on the AlpineeCarpathian fold belt (Dolton, 2006). Internally, the system is divided into two main intramontane domains (Hamilton et al., 2000), the rhombohedral-shaped Vienna Basin (Royden, 1985) belonging to peripheral lowlands and the adjacent Great Pannonian Basin accommodated in the central Intra-Carpathian lowlands (Sclater et al., 1980). The two subsystems can be further separated by a prominent differentiation of the thermal gradients, whereby the present day Pannonian system reaches twice the average for continental systems (Sclater et al., 1980; Dolton, 2006). The low to average thermal gradients of the Vienna Basin positioned the main thermocatalytic oil window at a depth of almost 6000 m (Landwein, 1988). The respectively lower thermal gradients probably originated from the early pre-Miocene crustal thickening induced by the active NW-directed thrusting of the AlpineeCarpathian nappes (Sauer et al., 1992; Hölzel et al., 2010; Schmid et al., 2008). The Eggenburgian initiation of the early Vienna Basin was triggered by the rapid subsidence of the peripheral Pannonian Basin (Sclater et al., 1980). The sedimentation is

b

Fig. 1. (a) Tectonic sketch map of the Vienna Basin (modified after Strauss et al., 2006), (b) Position of the 3D seismic block Seyzamdue (OMV) near the Markgrafneusiedl Fault that was investigated in this study, and the two deep exploration boreholes Markgrafneusiedl 004 and Gänserdorf Übertief 003a. Fault heaves are from the structural map of the preNeogene basement (modified after Kröll and Wessely, 1993).

184

D. Spahi
c et al. / Journal of Structural Geology 55 (2013) 182e195

characterized as an early piggyback basin system (ca. 23e17 Ma). This early thin-skin deformation-controlled sedimentation produced the sequences distributed exclusively within today’s central basin part (Strauss et al., 2006). During the second deformation stage (Middle and Upper Miocene, ca. 16e7 Ma, Decker and Peresson, 1996), the basin evolved into a pull-apart basin along the sinistral Vienna Basin fault system (Fodor, 1995; Decker and Peresson, 1996; Peresson and Decker, 1997). Unlike the general transgression in the Pannonian area (Sclater et al., 1980), here a combination of sea level fluctuations and main fault activity produced moderate subsidence of the main depocentres accommodating in total z5.5 km of sediments (Lankreijer et al., 1995; Hölzel et al., 2008) during 9 Ma. The investigated 3D seismic dataset imaged the Matzene Schönkirchen gas and oil field (Fig. 1) (Brix and Schultz, 1993). The initial formations are discordantly overlying the pre-Neogene basement (e.g., Sauer et al., 1992). The oldest succession associated with the central part of the Basin is characterized by a limnice fluvial succession comprising conglomerates, sandstones and interbedded pelites of Upper Karpatian age (ca. 16.5 Ma). These lacustrine, terrestrial and clastic depositional sequences are further characterized as the Backfloess (Rogl et al., 2002), Gaensendorf (Kreutzer, 1992b) and Aderkalaa Formations (Weissenbäck, 1996). The sequences were interrupted by a few main sea-level regressions (Wagreich and Schmid, 2002). A new marine transgression in the early Badenian is characterized by distal and proximal deltaic clastics and carbonates (Strauss et al., 2006). After a period of sea-level drop, dating the Badenian/Sarmatian boundary, a new transgression cycle in the Sarmatian resulted in the deposition of marls and clays. In the subsequent Early to Middle Pannonian period, another transgression covered most of the Sarmatian succession with a clayey and sandy lacustrine sequence (Harzhauser et al., 2004; Strauss et al., 2006). The clastic sequence of the “Middle Pannonian Formation Top” is the uppermost chronostratigraphically constrained section in the seismic data set. The normal faults accommodated in the vicinity of the central part of the Vienna Basin are kinematically linked to strikeeslip faults, forming extensional flower structures along the releasing bends of the Vienna Basin Transfer Fault. The most prominent growth strata are between the Lower Sarmatian (13 Ma) and Lower Pannonian (z11.5 Ma; e.g., Decker et al., 2005; Hinsch et al., 2005b; Beidinger and Decker, 2011). Nevertheless, the faults accommodated around the main depocentres record no Sarmatian (13 Ma) sediment growth (see also Fig. 3, the Matzen Fault system in Paulissen et al., 2011). The investigated 3D seismic data indicated the possibility of the formation of a thin hanging wall sedimentary wedge, exclusively associated with the end of the Middle Pannonian (Fig. 2). However this late Middle Pannonian effect has not been recorded in key regional sequenceestratigraphic works (e.g., Strauss et al., 2006). Therefore, in order to omit the interpretation uncertainties associated to eventual late Middle Pannonian thickening, we isolate investigations to a lower fault section, which is characterized by a population of prominent near-fault deflections. 3. Construction and quantification of the 3D structural model The database used for this study was derived from a 3D seismic survey of the central part of the Vienna Basin (Fig. 1). 3D seismic data were provided by the OMV Exploration and Production GmbH (Austria). The time-migrated 3D seismic reflection dataset covers z64 km2 with a recording time of 4000 ms (corresponding to about 7 km in depth) and has a line spacing of 25 m with a ca. 30 m vertical resolution. The geometry of the horizons and the faults were constrained by detailed choosing of the horizons in the numerous inlines (Fig. 2a),

crosslines and lines orientated roughly perpendicular to the fault strike. The seismic interpretation software SeisVision (LMKR Geographix) was used for seismic mapping. The integrated subsurface chronostratigraphic constraints were provided by the OMV. The time-migrated seismic data provided a traceable fault plane almost down to the basement domain. Unfortunately, the reflections in a vicinity of the lower fault tip exhibited unclear seismic patterns. Thus the final 3D model contains a section of Markgrafneusiedl normal fault from z400 ms down to z1230 ms. This section, adjacent to the investigated normal fault, clearly records anti- and syn-forms (Fig. 2a and b). In order to analyse the geometry of the offset and deformed sediments, we mapped continuously traceable sub-horizontal seismic amplitudes were mapped continuously: h1 (Lower Sarmatian), h2 (Lower Pannonian), h3 (Middle Pannonian), h4 (Middle Pannonian) and h5 (Middle Pannonian). The chosen amplitudes represent the continuous seismic response that captured a general stratal geometry. Subsequently, the investigated fault plane and the five horizons were constructed by triangulation using the mapped data points and equilateral triangle algorithm in Gocad (Paradigm). In order to compute morphoekinematic attributes, in particular the finite displacement and three-dimensional near-fault configuration, the elements of the resulting 3D structural model were converted to depth. The depth-conversion was based on the exponential increase in the seismic velocity with depth (Hinsch et al., 2005b). The idea was that the equation should simplify, but still reflect the interchange of clastic sequences with respect to the undisturbed compaction trend of the main Tertiary clastic sequences. The investigated seismic record gave neither indications of overpressure effects nor of high-pressure zones within the Neogene clastic rocks in the Vienna Basin (Gier et al., 2008), evidencing a non-disturbed decrease in porosity with depth. Note that sharp velocity structure change is exclusively associated with the deep-settled basement sediment interface, way below the investigated fault section (>3 s) (Hinsch et al., 2005b). The 3D depth migrated and georeferenced structural model allowed the calculation of the fault azimuth, fault dip, fault curvature, horizon dip and fault displacement. The set of attributes was calculated for the fault section ca. at 360 m and 1250 m (Fig. 3a). Calculation of Gaussian curvature (kG) across the entire 3D fault surface allowed visualization of the fault topography. The Gaussian curvature is the product of the minimum and maximum curvature at a point. Cylindrical structures or flat planes have a value of kG ¼ 0, while saddles have kG < 0 and domes or basins kG > 0 (e.g., Mynatt et al., 2007; Lisle and Toimil, 2007). Areas with kG s 0 represent deviations from a cylindrical or flat fault surface, probably associated with older linked fault segments (e.g., Walsh et al., 1999; Mansfield and Cartwright, 2000; Marchal et al., 2003). Accordingly, areas with positive curvature are convex upwards, while areas with negative curvature are concave upwards. A change from positive to negative curvature might represent the linkage of earlier segments (Lohr et al., 2008). The fault displacement was calculated from the 5 traced horizons. Displacements below h1 and above h5 record sudden decreases that are actually artefacts. These artefacts were induced by the absence of traceable horizons and are thus excluded from the interpretation. By defining the upper and lower boundary of the horizons (cut-off lines), the Gocad software calculates 3D displacement using the fault throw and a centre line between the two cut-off lines. The resulting 3D fault displacement is contoured and colour-coded onto the mapped fault plane. Clearly, the accuracy of the final structural model and the analysed attributes are dependent on the resolution of the seismic data and the mapping of the horizons. In order to avoid correlation errors, only a few, but continuously traceable, reflectors were

D. Spahi
c et al. / Journal of Structural Geology 55 (2013) 182e195

185

a

h4 & h5- Middle Pannonian

h3- Middle Pannonian h2- Lower Pannonian

h1- Lower Sarmatian

b

Fig. 2. (a) The seismic inline 344-1. The reverse drag markers are clearly visible in the hanging wall, but also gently developed in the footwall of the Markgrafneusiedl Fault. The five most prominent seismic amplitudes as stratigraphic markers were seismically constrained throughout the entire 3D cube volume: h1 (Lower Sarmatian), h2 (Lower Pannonian), h3 (Middle Pannonian – 20), h4 (Middle Pannonian – 12) and h5 (Middle Pannonian – 5). Note that the positions of the intersected markers lines are displayed in accordance with the intersection angle between the inline and horizon map, thereby not perfectly illustrating the thicknesses of the markers in the exhibited inline view. (b) The mapped horizon h3 (Middle Pannonian) within the 3D seismic block indicated as the rectangle in Fig. 1b. Map of the horizon h3 is clearly separated by the Markgrafneusiedl Fault into two seismic domains: hanging wall (SE) and footwall (NW).

mapped. Additionally, only structures of amplitude at least twice the seismic resolution (i.e., 30 m) were mapped. The validation of the depth-converted model and the existence of the mapped dome structures were confirmed by a previous hydrocarbon exploratory drilling (P. Strauss e OMV Exploration and Production GmbH, personal communications).

4. Morphoekinematical analysis of the fault plane and marker horizons Examination of the surface of the composite Markgrafneusiedl Fault spans the present day 3D morphology (Fig. 3aec) and the 3D finite displacement analysis (Fig. 3e and f). Displacementedistance

186

D. Spahi
c et al. / Journal of Structural Geology 55 (2013) 182e195 shown in Figs 6 & 7

a

b

c

d

e

f

Fig. 3. (a) Oblique-frontal view to the 3D modelled surface of the investigated part of the Markgrafneusiedl Fault, including the position of the section analysed in detail in (bef). Fault attribute maps of (b) dip direction (azimuth), (c) dip, (d) Gaussian curvature, (e) displacement and (f) combined plot of dip and displacement. The position of the displacementedistance andedip profiles A and B (Fig. 5) are indicated in (e).

diagrams (Fig. 4) and vertical displacement plots (Fig. 5) were also constructed, analysing the combined morphoekinematical attributes with respect to the near-fault marker deflections (Figs. 6 and 7). 4.1. Fault azimuth and dip The calculated values of the fault dip direction range from 140 to 165 (Fig. 3b). Mean azimuth variations divided the fault plane into the sub-areas labelled as Fault 1 (Fig. 3b, mainly blue coloured area), Fault 2 (Fig. 3b, mainly green coloured area) and Fault 3

(Fig. 3b, mainly blue and red coloured area). F3 is comprised of the values indicating that the fault strikes more towards the SE in this particular section. The mean azimuth (Fig. 3b, blueered regions) is in accordance with the mean dip distribution (Fig. 3c, blueedarkblue regions) and correspond to areas with the maximum dip values, ranging z47 e57 (Fig. 3c). Note that the dip values of the investigated fault surface range between 28 and 76 (Fig. 4) having a mean of 51. The boundary zone, separating F1 from F2 is characterized by abrupt changes in the azimuth and dip values (Fig. 3b and c). In fact, F3 represents a curved bend in the fault plane having dip values >60 (Fig. 3c, redepinkish regions).

D. Spahi
c et al. / Journal of Structural Geology 55 (2013) 182e195

187

the SW part (especially in the upper half of the fault) record lower displacement values (zone tapering towards a tip). The transition between low and high displacement zones roughly occurs at the centre of the investigated fault section (F3).

Fig. 4. The dipedisplacement point-based diagram illustrates the relation between maximum displacement and the dip angle. Fault sections between 30 and 40 have a maximum displacement of z280 m, whereas the sections of (ii) 40 e45 have maxima of 400 m and finally the sections of (iii) 45e50 are characterized by maximal values of 480 m.

4.2. Fault curvature The values of calculated Gaussian curvature are highly variable across the fault surface (Fig. 3d). Areas having a zero Gaussian curvature form an irregular network on the fault plane and represent either cylindrical or planar zones. Between these cylindrical/ planar corrugations, isolated patches of various sizes have negative or positive Gaussian curvatures. Given the regional observation in the Vienna Basin (Hinsch et al., 2005; Beidinger and Decker, 2011) that normal faults have a convex upwards shape, the patches with a negative Gaussian curvature possibly represent previous isolated normal fault segments. Patches with a positive Gaussian curvature may indicate ancient relay zones, where the segments linked. Furthermore, there is also a cluster of fairly round-shaped patches having very large negative values (0.002). Associated with the maximum deviations of the azimuth and dip attributes (Fig. 3b and c) these fields are relief-prominent features and are distinctive in respect to the observed segmentation pattern. Thus, these morphological peculiarities may represent imaged remnants of ancient overprinted dead splays. Geometrical irregularities qualifying dead splays can also be observed along the entire fault trace (Fig. 2b and Fig. 3a), Comparison between the Gaussian curvatures contoured on the fault plane with the dip contours exhibit a good correlation between the zones of negative Gaussian curvature and dip values of z50 . In contrast, contours of positive Gaussian curvature are correlated with fault dip values of z40 e50 .

4.3.1. Relationship between fault dip and finite displacement The overlapped dip and displacement contoureattribute map (Fig. 3f) highlights the areas with a steep displacement gradient (narrow dark blue areas between higher and lower displacements, Fig. 3e). The areas with steep displacement gradients are correlative with local dip maxima of z60 . In order to analyse the relationship between the calculated fault dip and maximum displacement, the extracted dip values were plotted against the maximum displacement of the corresponding area. The results illustrated a relationship between the fault dip and the fault displacement (Fig. 4), associating the fault sections of 30e 40 dip with a maximum displacement of z280 m, whereas the sections of 40e44 have significantly higher maxima reaching 400 m. The fault sections having the steepest dip angles of 45e50 are characterized by maximum displacement values of up to 480 m. 4.3.2. Vertical displacement component The kinematic attributes contoured onto the 3D fault plane in Fig. 3e were further investigated for vertical segmentation (Section A and B, Fig. 5). The vertical displacement value is actually a 3D vector value measured down-dip between two separated points along the intersection between of the marker in the hanging wall and footwall. The profiles A and B exhibit three local maxima, additionally indicating a general increase in the finite displacement progressing towards deeper structural levels (Fig. 5a and c). The fault dipedistance plots (Fig. 5b and d) show areas of similar dip values, whereby sudden variations in the dip angle could indicate to the existence of an initial proto-segment boundary. Still, the general segmentation pattern is rather unclear.

a

profile A

b

c

profile B

d

4.3. Fault displacement The fault displacement map shows an irregular displacement pattern with horizontally aligned displacement contours (Fig. 3e). As mentioned earlier, the zones of zero displacement at the top and the bottom are artefacts produced by the section boundaries. The map yields the two prominent areas of displacement maxima (z400 m and 480 m). The zones are interrupted by several local minima having 150e200 m of displacement. These transition zones between the maxima and the local minima are marked by an abrupt change in the displacement gradient. Generally, the NE part of the investigated fault section displays considerably larger areas characterized by higher displacements (central fault zone), while

Fig. 5. Displacementedistance plots (a, c) of two sections (A and B, marked in Fig. 3e) recorded along the dip direction of the Markgrafneusiedl Fault, and the corresponding dipedisplacement profiles (b, d), measured from the contoured displacement and dip diagrams in Fig. 3c and e, respectively. The maximum displacement along both sections varies between 200 and 425 m. The displacement maxima are separated by several minima, but a general increase in displacement downwards can be observed. Changes in dip angles commonly correlate with changes in displacement along a fault.

D. Spahi
c et al. / Journal of Structural Geology 55 (2013) 182e195

188

4.4. Distribution of the near-fault saddle-dome structures The five mapped horizons represent the internal stratal architecture of several large near-fault folds embedded in the adjacent markers (Fig. 6aed). In a 3D perspective view, each of the depth colour code attribute maps illustrates a vertical section, depicting the geometry of the three distinctive high-amplitude-wavelength anticlines. Considering the investigated fault section as a complex finite length discontinuity (Fig. 2b) with a record of steep displacement gradient

e

d

(Fig. 3e), it is assumed here that a population of a faulteadjacent saddle and dome features could be considered as a system of normal and reverse drag features. The most prominent anticlineshaped hanging wall structures or a reverse drag population are labelled as R1, R2 and R3 (Fig. 6aed, yellow-coloured sections of the markers h2, h3, h4 and h5). The individual reverse drags are separated by saddle shape forms or normal drag (N1 and N2). These near-fault structures are also designated in the footwall, showing moderate amplitudes (Fig. 6e). Still, the footwall deformations are visible and mathematically measurable. In the following, the nearfault structures R1, N1, R2, N2, R3, and N3 are analysed in detail with reference to a particular horizon. The notation Rx*hx means a section through the reverse drag Rx is recorded on horizon hx. Rx þ y indicates that drag Rx and Ry are linked into a single structure. The relationship between fault morphology and adjacent matrix geometry is analysed by the juxtaposition of a hanging wall fault drag cluster (shape, i.e., sense and magnitude) with fault morphoe kinematic attributes. Starting from h2, the horizon depicts a population of three reverse drags (R1*h2, R2*h2 and R3*h2) separated by normal drag synclines (N1*h2, N2*h2, N3*h2). Approximately 100 m above h2, the horizon h3 reveals the same structures (reverse drag R1*h3, R2*h3 and R3*h3) having lower amplitudes with respect to the horizon h2 (Fig. 6b). The structures R2*h3 and R3*h3 are now in contact, and are almost linked into a single reverse drag (R2 þ 3*h3). The horizon h4 exhibits an appended amplitude decrease. Now, R1*h4 and R2*h4 are linked into a single structure and R1 þ 2*h4 is separated by N4 from R3*h4. The drag amplitudes are further decreased on the level of h5. Nevertheless, the structures R1*h5, N1*h5, R2*h5, N2*h5, R3*h5 and N3*h5 can be clearly discriminated.

c 4.5. Correlation between fault drag and fault attributes

b

a

Fig. 6. (a) Oblique view on the 3D modelled reverse and normal drag developed within the footwall and hanging wall of the Markgrafneusiedl Fault. The sections of the 3D model illustrated the local amplitude change of reverse and normal drags distributed within the hanging wall tracked from the bottom towards higher fault domains: (b) marker h2, (c) marker h3, (d) marker h4, (e) marker h5.

In order to analyse the correlation between fault drag and fault parameters such as, curvature, displacement or fault dip, attention was focused on R1*h3, R2 þ 3*h3, N1*h3 and N3*h3 (Fig. 7aec). The R1*h3 structure is accommodated against the negative Gaussian curvature corrugation (Fig. 7a). This section is also characterized by a local displacement maximum with a finite displacement of z400 m (Fig. 7b) and having fault dip values of z50 (Fig. 7c). The good correlations between both the fault morphology and kinematics against the reverse marker deflection indicate a mechanical interaction between the fault slip and the propagating fault drag. The normal drag N1*h3 separates the reverse drags, R1*h3 and R2 þ 3*h3. N1*h3 developed where the fault has a positive Gaussian curvature, i.e., convex upwards and has high dip values of about 60 . Additionally, this feature is juxtaposed to a local decrease in the finite displacement of z230 m. The shape of R2 þ 3*h3 indicates a possible linkage of the two adjacent fault drag structures. This complex near-fault feature is characterized by three mathematically measurable amplitude maxima (R2 þ 3a*h3, z42 m; R2 þ 3b*h3, z21 m and R2 þ 3c*h3, z31 m). R2 þ 3*h3 is accommodated against a fault morphology transition zone, characterized by the changeover from a positive to a negative curvature area (Fig. 7a). The area is also characterized by moderate to high fault dip values, ranging from 45 to 60 (Fig. 6f). This highest amplitude drag feature, reaching an offset of z400 m, is found in the centre of a local displacement maximum (Fig. 7b). Of the adjacent feature, R2 þ 3b*h3, is of lower amplitude, while R2 þ 3c*h3 can be correlated with a decrease in the finite displacement and a zone of negative curvature. The amplitude of R2 þ 3*h3 further decreases, eventually transitioning into a zone of a high magnitude syncline or normal drag (N3*h3). N3*h3 is

D. Spahi
c et al. / Journal of Structural Geology 55 (2013) 182e195

189

a

b

c

Fig. 7. (a)e(c) Enlarged oblique view of the Markgrafneusiedl Fault. Contact between the marker h3 and the fault section. The spatial relationship between the fault attributes (dip, curvature and displacement) and the fault drag geometry: (a) an excellent example of an almost perfect juxtaposition of the reverse drag (e.g., R1) and a positive corrugation (green patch), potentially presenting an ancient fault segment; (b) the white line additionally displays the most gentle, large scale reverse and normal drag superposed onto smaller scale features.

located adjacent to the area of F3 and is characterized by a displacement maximum (Fig. 7b). 5. Discussion The 3D structural model of the surface of the Markgrafneusiedl Fault exhibits a complex fault morphology combined with heterogeneous displacement patterns (Figs. 6 and 7). In the following, how the geometric parameters are related to the mechanical linkages of the fault segments will be discussed and the growth history

of the Markgrafneusiedl Fault surface. Furthermore, a prediction of the footwall deformation, which is less well constrained by the 3D seismic data than the deformation in the hanging wall, will be included. 5.1. Fault displacement and morphology as criteria for segmented faults As strain increases, a fault population of initially consisting of a large number of planar, low displacement faults will eventually

D. Spahi
c et al. / Journal of Structural Geology 55 (2013) 182e195

190

evolve into a composite large fault system. Moreover, the geometric features of a fault are considered to be helpful when attempting to identify individual fault segments that might have linked during fault growth (Walsh et al., 2003). However, in highly complex large fault zones where fault tips are reprinted being not morphologically distinctable, simple displacement e distance analyses may result in rather ambiguous or unclear results. For example, the along-strike displacement e distance profile of the segments of the Markgrafneusiedl Fault (Fig. 8) show variations in the displacements which cannot be directly related to the fault segments. The displacement profile (Fig. 8c) does not exhibit a common displacement distance pattern, i.e., a large displacement is not associated with the centre of a fault, but it is rather asymmetric and peculiar. In fact, contradictorily, it seems the low displacement zones are juxtaposed to the fault central parts (Fig. 8c). Most

probably, as a consequence of the initial proto segment linkage, transient displacements on the isolated proto faults are reprinted by displacements of the hard linked fault. Similar asymmetric slip distributions and/or multiple slip maxima on normal faults may result from the linkage of individual fault segments or may reflect the mechanical interaction between intersecting faults (Peacock and Sanderson, 1996; Maerten et al., 1999). Numerical elastic models demonstrate that multiple slip maxima forced by intersecting faults are located neither along the intersections, nor at the fault centres. Therefore, the use of displacement distance plots alone is of limited practical use for the detection of fault segments. In order to increase the limited information on slip distributions along faults, other studies combined an analysis of the 3D geometry of segmented faults with displacement asymmetry measurements (e.g., Childs et al., 2003; Lohr et al., 2008). Analysing a

a

b

c

d

Fig. 8. (a) Schematic map (top view) and (b) the Gaussian curvature 3D map of the fault plane (oblique view), where two large, slightly overlapping fault segments (F1 and F2) are connected by a relay fault (F3). The presumably older segments F1 and F2 show a reverse drag of the adjacent markers on a large scale (see Fig. 6a), while the relay fault F3 is related to markers with a normal drag. (c) Correlation of the 3D displacement attribute map with the along-strike throwedistance plot. (d) Schematic sketch of the relay zone between F1 and F2 breached by the relay fault F3.

D. Spahi
c et al. / Journal of Structural Geology 55 (2013) 182e195

synsedimentary fault, Childs et al. (2003) constructed varying throw contour patterns on the strike fault projections, thereby demonstrating that the locations of fault maxima and minima indicated local growth directions. Lohr et al. (2008) combined 3D fault morphology data with displacement distance graphs, suggesting that the fault segmentation is reflected by triangular to half elliptical shaped real displacement profiles superimposed on the 3D fault segmentation pattern. By using true displacements from along slip normal movements, the authors demonstrated significant differences from the real, vertical and horizontal displacements identifying the former fault segments. The authors caution the applicability of throw values, which lead to a smoothing of the real displacement curves. Displacement analyses along the Markgrafneusiedl Fault demonstrated that the results of both techniques, the displacement distance profiles (Fig. 8c) and the 3D vertical displacement distance plot (Fig. 5), show a series of local displacement maxima and minima that are poorly correlative with the morphological and kinematic parameters computed from the mapped fault surface. The mechanical linking of differently sized segments could result in a significant change in the inherited displacement pattern (Zee and Urai, 2005) and therefore exclusive analysis of the alonge strike plot or 3D displacement attribute map of the Markgrafneusiedl Fault provided no clear discrimination between initially isolated and linked fault segments.

191

a

b

c

5.2. Fault drag as an additional criterion to identify fault segments As the segmentation and fault linkage clearly influence the displacement pattern along a fault (Maerten et al., 1999; Walsh et al., 2002) and displacement gradients induce fault drag, quantification of fault drag may aid the detection of linked proto-fault segments. As a result of a fault slip, heterogeneous stress and displacement fields develop in the surrounding rock (Pollard and Segall, 1987). The elastic theory predicts and some natural faults demonstrate that the displacement field around an isolated fault is elliptical, reaching a maximum at the centre of the fault and dropping to zero at the fault tips (Rippon, 1985; Barnett et al., 1987; Pollard and Segall, 1987; Walsh and Watterson, 1989; Martel, 1997; Cowie and Shipton, 1998; Bürgmann et al., 1994). Besides being dependant on the material properties, the wavelength and the amplitude of fault drag are mainly a function of the size of the fault, the aspect ratio of the fault, the orientation of the fault with respect to the regional stress field and finite displacement (Grasemann et al., 2005; Exner and Dabrowski, 2010; Reber et al., 2012). Larger amplitude drags are defined by constant wavelengths (e.g., laterally confined faults) that record a higher maximum displacement versus length ratio than laterally unconfined propagating faults (Grasemann et al., 2011 and references cited therein). During fault segment linkage, different wavelengths and amplitudes of fault drag may be superimposed on inherently smaller wavelength and amplitude drag geometries (Fig. 9) and even the sense of the drag (e.g., normal drag) may be inherited into a linked larger structure (e.g., reverse drag). As mentioned earlier, fault drag has been described and analysed in natural examples on various scales, as well as in analogue and numerical models (e.g., Barnett et al., 1987; Passchier, 2001; Grasemann and Stüwe, 2001; Grasemann et al., 2005; Wiesmayr and Grasemann, 2005; Resor, 2008, 2012; Spahi
c et al., 2011). The modelling results of drag around a single finite fault plane predict different drag senses on a central marker line (reverse or normal) as a function of the orientation of the marker line with the fault. In order to maintain strain compatibility, the magnitude, but also the sense of drag must change on marker horizons that do not meet the

Fig. 9. Normalized fault (segment) length e fault drag amplitude diagram illustrating the mechanism of progressive fault drag amplitude development (top view). (a) Initial growth of an isolated non-restricted fault segment with an associated symmetric drag amplitude (stage 1). (b) Schematic growth model and coalescence of two different laterally juxtaposed fault segments with similar amplitudes. After the segments linked, a normal drag at the linkage location is preserved. (c) Growth model of a linked fault surface associated with an overall reverse fault drag but preservation of local drags inherited from previously isolated segments.

centre of the fault (Grasemann et al., 2005). Consequently, a smaller wavelength/amplitude drag may change its sense during fault propagation and/or linkage if the relative position of the marker with respect to the centre of the fault changes. An excellent example of inherited drag is illustrated by the horizons h2eh5 developed along the Markgrafneusiedl Fault (Figs. 6 and 7). Here the association of reverse and normal drag (R1eR3, N1eN5) on the different horizons are interpreted to be the result of fault drag, which developed along initially isolated segments, linking up to form fault F1. By considering not exclusively the displacement and morphological analysis of the fault alone, a novel interpretation of the growth history of the normal Markgrafneusiedl Fault (Fig. 10) is suggested in which the mapped drag features embedded in the marker horizons are additionally included. 5.3. Evolution of the Markgrafneusiedl Fault The investigated section of the Markgrafneusiedl Fault (dating from the Lower Sarmatian up to almost the Middle Pannonian) consists of the two large fault segments F1 and F2, linked by a relay fault F3. The combined morphoekinematical model (fault

192

D. Spahi
c et al. / Journal of Structural Geology 55 (2013) 182e195

Fig. 10. A schematic illustration of the evolution of the Markgrafneusiedl Fault (left, oblique 3D view of a schematic reconstruction; right, schematic cross-section): (a) isolated small fault segments produced reverse fault drag in the adjacent marker horizons; (b) propagation and linkage of initially localized segments towards the larger fault segments F1 and F2, which are also associated with a larger reverse drag; (c) the relay fault F3 breaches the overlap between the large segments and generates normal drag geometries in the adjacent marker horizons.

morphology, displacement distribution and fault drag) suggests that F1 and F2 formed through the linkage of initially smaller segments. The nucleation of the initial proto fault segments was most likely induced for several reasons. After the post-Lower Sarmatian interruption (Wagreich and Schmid, 2002; Strauss et al., 2006), the activation of underlying precursor faults embedded in a preNeogene basement (Kröll and Wessely, 1993) most likely renewed

(pull-apart phase) the intrabasinal perturbations. The investigations indicate that this period was rather characterized by slow subsidence further influencing a gentle sedimentation. Sedimentation and compaction may influence the finite displacement on growth faults but not the sense of drag. Comparison of imposed tectonic displacements with the preserved displacements indicates that although the thicknesses of sand/shale sequences may decrease by up to c. 55% due to compaction, associated losses in

D. Spahi
c et al. / Journal of Structural Geology 55 (2013) 182e195

displacement on growth faults are typically much less than 20% because they displace partially compacted sediments (Taylor et al., 2008). Reverse drag or rollover above listric faults mechanisms are responsible for the hanging wall anticline and/or footwall syncline and not the sedimentation and compaction. Both models work with and without syndepositional sedimentation and compaction. Locally, isolated fault planes embedded in a pre-Tertiary and Sarmatian matrix slowly propagated generating initial fault drag deflections (Fig. 10a). After the vertically elongated faults of average size ranging from z200 to 400 m had linked (Fig. 10b), the ongoing extension phase induced of a new, larger generation of isolated segments (segment size z1500 to over 2000 m), as in the segments F1 and F2. Finally, the progressive stress magnitude increase resulted in the linkage of F1 and F2 thereby encompassing the fault F3 or the relay zone (Fig. 10c). After the relay zone had been breached, the new fault F3 impeded further lateral propagation of F1 and F2, enabling a cumulative slip equilibration along the linked faults (Ackermann et al., 2001). Thus, the largest finite displacement values are associated with the today’s central fault section that is located at the segment F3. This peculiar maximum offset area may indicate that the displacement variations along this normal fault were developed after the segments had been linked into a single large-scale fault, and indeed showed full displacement equilibration (Ackermann et al., 2001; McLeod et al., 2002) as the large segments grew and coalesce. The linkage of F1, F2 and F3 can be additionally authenticated by the along-strike “completely breached relay zones” geometry (Crider and Pollard, 1998; Marchal et al., 2003). Moreover, the magnitude of the N3 syncline or normal drag records a progressive decrease in amplitude going towards the up-sections (from h2 towards h5). The youngest mapped horizon being almost unaffected by the below normal drag may indicate that the latest evolution stage is rather characterized by a coalesced principal fault surface with no record of displacement gradient. Thus, the activity pattern of the most upper part of the investigated fault section could be associated with a minor sin-sedimentary wedge (e.g. Fig. 2). Zones of hard and soft linkage are characterized by a distinctive morphology of the branch domains (branch lines) (e.g., Ferrill et al., 1999; Walsh et al., 1999). Kinematically, the overlapping fault zones can be characterized by a steepening and the positioning of the maximum slope near or inside relay zones (e.g., Peacock and Sanderson, 1991; Marchal et al., 2003). The F3 zone is characterized by a curved fault morphology, and changes in the local azimuth, dip and curvature (Fig. 3bed). This morphological pattern in addition to displacement equilibration is typical for linkage zones of overlapping faults (e.g. Walsh et al., 1999; Rykkelid and Fossen, 2002). 5.4. Fault drag geometry in the footwall Deformation or fault drag in a footwall close to normal faults is frequently observed in outcrops and seismic data (Kasahara, 1981; McConnell et al., 1997; Mansfield and Cartwright, 2000). However, besides the numerous studies of hanging wall deformation (e.g., Tearpock and Bischke, 2003 and references cited therein), the mechanism of footwall deformations has received less attention (Resor and Pollard, 2012) for the following reasons. Firstly, the slip ratio or displacements due to a drag effect on a footwall matrix can be less intensive than those in a hanging wall, especially when inclined faults interact with the Earth’s surface (Grasemann et al., 2005). Although host rock deformation in the footwall of a normal fault branching to the surface is considerably smaller in magnitude than that of a hanging wall dome, a record of fault drag in the footwall around a normal fault still provides

193

evidence that the fault terminates at a depth (except if the footwall is mechanically linked to another fault at a depth; e.g. Spahi
c et al., 2011). Secondly, the resolution of 2D or 3D seismic records above a fault plane may frequently result in a much higher resolution than in a footwall (Tearpock and Bischke, 2003). Footwall visualization on a basin scale largely depends on many factors, whereby the depth is one of the most critical. In dipeslip noninverted faults, a footwall is commonly positioned below the fault surface, deeper than a hanging wall, and therefore requiring better seismic preparation affecting the exploration costs. Moreover, it was shown that flattening normal faults disappear from seismic data with increasing depth (e.g., Tearpock and Bischke, 2003). To summarize, the quality of seismic records depends of the fault plane inclination, eventual fluid presence in or below the discontinuity, sedimentation rate, etc. Listric fault models (e.g. Shelton, 1984; McClay, 1990; Krészek et al., 2009) have been frequently used to define a structural model for hydrocarbon exploration near large normal faults (e.g., Matenco and Radivojevi
c, 2012). Being that many of these models are based on the assumption that the footwall below a fault surface behaves as a rigid body (e.g., Yamada and McClay, 2003), petroleum exploration targets have been mainly confined to hanging wall anticlines. As a contrast to the listric model, and the frequently observed hanging wall anticline above a normal fault, the drag concept could contribute an additional trap-based play-fairway. An additional exploration target corresponds to footwall drag (e.g., Porras et al., 2003). The employment of the conceptual model around normal faults allows indeed a better understanding of the behaviour of fault fluid with respect to lateral and top sealing. If a fault, for example, behaves as a conduit, there is a high probability that fluid migration pathways would end in a footwall syncline. Inherently, the hydrocarbon carrier has the same sealing horizon as the superimposed anticline prospect, therefore the concept further facilitates prediction of a top seal behaviour. According to the elastic analytical solutions, both reverse and normal fault drag develops around planar fault segments of finite length (Pollard and Segall, 1987; Grasemann et al., 2005). Modelling results suggest that the maximum amplitudes of the hanging wall reverse drag are developed around the central fault sections. As the displacement gradually decreases away from the central fault section towards the tips, contemporaneously, the reverse drag amplitude should decrease towards the fault tips. Eventually in the vicinity of fault tips, the reverse drag is juxtaposed to a normal drag (e.g., Fig. 10c). As a consequence, the fault drag magnitude lessens whereby the sense rotates along the slip direction. Therefore, the variations in amplitude may indicate the position of fault tips, which might be very useful for discriminating inherited tips among linked segments (Wiesmayr and Grasemann, 2005). A detailed structural investigation of reverse and normal drag in the hanging wall of a normal fault could help to detect footwall drag that is not or is poorly recorded by seismic data. In hydrocarbon exploration, footwall synclines are rather poorly explored because these structures are often regarded as features without potential for hydrocarbon reservoirs (because of the undesired U-shape geometry of the layers). 6. Conclusions The combination of 3D morphological and kinematic subsurface observations demonstrated a combination of methods that may enable the identification of deformation zones downscaling to an initial proto fault segment. The nature of the composite, large-scale Markgrafneusiedl Fault, located in the central section of the Vienna pull-apart system, was investigated. Employing a novel approach combining well-known analytical methods, the complex fault

194

D. Spahi
c et al. / Journal of Structural Geology 55 (2013) 182e195

configuration was analysed over a population of near-fault deformations embedded in marker horizons. The following could be concluded: (1) Early Middle Pannonian slip along the Markgrafneusiedl high angle normal fault system induced a complex heterogeneous displacement fields. (2) The Markgrafneusiedl normal fault system developed from several generations of linked fault segments (3) Linkage of initial isolated fault segments overprinted the previous individual displacement patterns. (4) Progressive bulk deformation and linkage of individual segments result in mature composite single faults, which loose the inherited displacement pattern from the initially isolated segments (i.e. displacement equilibration); (5) Convex upwards-shaped high angle faults with a hanging wall reverse drag (or rollover anticlines) do not have listric fault geometries and are not necessarily mechanically linked with a horizontal detachment at depth. (6) The fault drag related to displacement gradients is not restricted to the small-scale or mesoscopic flanking structures, but also occurs on basin scale fault systems.

Acknowledgement This study was funded by the Austrian Science Fund (FWFProject P20092-N10). We thank OMV EP Austria for providing the 3D seismic block along with datasets from neighbouring boreholes, and we appreciate discussion on the specifics of this fault with P. Strauss, M. Frehner and A. Beidinger. We thank the editor Ian Alsop and two anonymous reviewers for important comments, which improved the content of the manuscript.

References Ackermann, R.V., Schlishe, R.W., Withjack, M.O., 2001. The geometrical and statistical evolution of normal fault systems: an experimental study of the effects of mechanical layer thickness on scaling laws. J. Struct. Geol. 23, 1803e1819. Barnett, J.A.M., Mortimer, J., Rippon, J.H., Walsh, J.J., Watterson, J., 1987. Displacement geometry in the volume containing a single normal fault. AAPG Bull. B71, 925e937. Beidinger, A., Decker, K., 2011. 3D geometry and kinematics of the Lassee flower structure: implications for segmentation and seismotectonics of the Vienna Basin strike-slip fault, Austria. Tectonophysics 499, 22e40. Beidinger, A., Decker, K., Roch, K., 2011. The Lassee segment of the Vienna Basin fault system as a potential source of the earthquake of Carnuntum in the fourth century A.D. Int. J. Earth Sci. 100, 1315e1329. Brix, F., Schultz, O., 1993. Erdöl und Erdgas in Össtereich, p. 688. Brun, J.-P., Mauduit, T.P.O., 2008. Rollovers in salt tectonics: the inadequacy of the listric fault model. Tectonophysics 457, 1e11. Bürgmann, R., Pollard, D.D., Martel, S.J., 1994. Slip distributions on faults: effects of stress gradients, inelastic deformation, heterogeneous host-rock stiffness, and fault interaction. J. Struct. Geol. 16, 1675e1690. Cartwright, J.A., Trudgill, B.D., Mansfield, C.S., 1995. Fault growth by segments linkage: an explanation for scatter in maximum displacement and trace length from the Canyonlands Grabens of SE Utah. J. Struct. Geol. 17, 1319e1329. Childs, C., Nicol, A., Walsch, J.J., Watterson, J., 1996. Growth of vertically segmented normal faults. J. Struct. Geol. 18, 1389e1397. Childs, C., Nicol, A., Walsch, J.J., Watterson, J., 2003. The growth and propagation of synsedimentary faults. J. Struct. Geol. 25, 633e648. Coward, M.P., Dewey, J.F., Hancock, P.L. (Eds.), 1987. Continental extensional tectonics. Geol. Soc. Spec. Publ., 28, p. 637. Corver, M., Doust, H., Diederik van Wees, J., Bada, G., Cloetingh, S., 2009. Classification of rifted sedimentary basins of the Pannonian Basin System according to the structural genesis, evolutionary history and hydrocarbon maturation zones. Marine Pet. Geol. 25, 1452e1464. Cowie, P.A., Shipton, Z.K., 1998. Fault tip displacement gradients and process zone dimensions. J. Struct. Geol. 20, 983e997. Crider, J.G., Pollard, D., 1998. Fault linkage: three-dimensional mechanical interaction between echelon normal faults. J. Geophys. Res. 103, 24,373e24,391. Davidson, I., 1994. Linked fault systems extensional, strike-slip and contractional. In: Hancock, P.L. (Ed.), Continental Deformation, pp. 121e142.

Decker, K., Peresson, H., 1996. Tertiary kinematics in the Alpine-CarpathianPannonian system: link between thrusting, transform faulting and crustal extension. In: Wessely, G., Liebl, W. (Eds.), Oil and Gas in Alpidic Thrustbelts and Basins of Central and Eastern Europe, pp. 69e77. Decker, K., Peresson, H., Hinsch, R., 2005. Active tectonics and Quaternary basin formation along the Vienna Basin Transform fault. Quat. Sci. Rev. 24, 307e322. Dolton, G., 2006. Pannonian basin province, central Europe (Province 4808) e petroleum geology, total petroleum systems, and petroleum resource assessment. USGS Bull. 2204B, 47. Exner, U., Dabrowski, M., 2010. Monoclinic and triclinic 3D flanking structures around elliptical cracks. J. Struct. Geol. 32, 2009e2021. Ferrill, D.A., Stamatakos, A.J, Sims, D., 1999. Normal fault corrugation: implications for growth and seismicity of active normal faults. J. Struct. Geol. 21, 1027e1038. Fodor, L., 1995. From transpression to transtension: Oligocene-Miocene structural evolution of the Vienna basin and the East Alpine-Western Carpathian junction. Tectonophysics 242, 151e182. Gawthorpe, R.L., Jackson, A.-L.C., Young, M.J., Sharp, I.R., Moustafa, A.R., Leppard, C.W., 2003. Normal fault growth, displacement localisation and evolution of normal fault populations: the Hammam Faraun fault block, Suez rift, Egypt. J. Struct. Geol. 25, 883e895. Gier, S., Worden, R.H., Johns, W.D., Kurzweil, H., 2008. Diagenesis and reservoir quality of Miocene sandstones in the Vienna Basin, Austria. Mar. Petrol. Geol. 25, 681e695. Grasemann, B., Stüwe, K., 2001. The development of flanking folds during simple shear and their use as kinematic indicators. J. Struct. Geol. 23, 715e724. Grasemann, B., Martel, S., Passchier, C., 2005. Reverse and normal drag along a fault. J. Struct.Geol. 27, 999e1010. Grasemann, B., Exner, U., Tschegg, C., 2011. Displacementelength scaling of brittle faults in ductile shear. J. Struct. Geol. 33, 1650e1661. Gupta, A., Scholz, C.H., 2000. A model of normal fault interaction based on observations and theory. J. Struct. Geol. 22, 865e879. Hamilton, W., Wagner, L., Wessely, G., 2000. Oil and gas in Austria. Mitteilungen der Geologischen Gesellschaft 92, 235e262. Harzhauser, M., Daxner-Höck, G., Piller, W.E., 2004. An integrated stratigraphy of the Pannonian (Late Miocene) in the Vienna Basin. Austrian J. Earth Sci. 95/96, 6e19. Hinsch, R., Decker, K., Wagreich, M., 2005a. 3-D mapping of segmented active faults in the southern Vienna Basin. Quat. Sci. Rev. 24, 321e336. Hinsch, R., Decker, K., Peresson, H., 2005b. 3D seismic interpretation and structural modelling in the Vienna Basin: implications for Miocene to recent kinematics. Austrian J. Earth Sci. 97. Hölzel, M., Wagreich, M., Faber, R., Strauss, P., 2008. Regional subsidence analysis in the Vienna Basin (Austria). Austrian J. Earth Sci. 101, 88e98. Hölzel, M., Decker, K., Zámolyi, A., Strauss, P., Wagreich, M., 2010. Lower Miocene structural evolution of central Vienna Basin. Mar. Petrol. Geol. 27, 666e681. Kasahara, K., 1981. Earthquake Mechanics. Cambr. Unive. Press, Cambridge. Kim, Y.-S., Sanderson, D.J., 2005. The relationship between displacement and length of faults: a review. Earth-sci. Rev. 68, 317e334. Kocher, T., Manctelow, N., 2005. Dynamic reverse modelling of flanking structures: a source of quantitative kinematic information. J. Struct. Geol. 27, 1346e1354. Kocher, T., Manctelow, N., 2006. Flanking structure development in anisotropic viscous rock. J. Struct. Geol. 28, 1139e1145. Krészek, C., Adam, J., Grujic, D., 2009. Mechanics of fault and expulsion rollover systems developed on passive margins detached on salt: insights from analogue modelling and optical strain monitoring. In: Jolley, S.J., Barr, D., Walsh, J.J., Knipe, R.J. (Eds.), Structurally Complex Reservoirs, Geol Soc Lond Spec Publ, vol. 292, pp. 103e121. Kreutzer, N., 1992. Das Neogen des Wiener Beckens. In: Brix, F., Schultz, O. (Eds.), Erdöl und Erdgas Österreich, pp. 403e424. Kristensen, M.B., Childs, C.D., Korstgård, J.A., 2008. The 3D geometry of small-scale relay zone between normal faults in soft sediments. J. Struct.Geol. 30, 257e272. Kröll, A., Wessely, G., 1993. Wiener Becken und angrenzende Gebiete. In: Geologische Themenkarte der Rep. Österreich 1:200.000. Geolog. Bundesanst, Wien. Ladwein, H.W., 1988. Organic geochemistry of Vienna Basin: model for hydrocarbon generation in overthrust belts. AAPG Bull. 72, 586e599. Landwein, H.W., 1988. Organic geochemistry of Vienna Basin: model for hydrocarbon generation in overthrust belts. AAPG Bull. 72 (5), 586e599.  ák, P., Hlôska, M., Biermann, C., 1995. Lankreijer, A., Ková c, M., Cloetingh, S., Piton Quantitative subsidence analysis and forward modelling of the Vienna and Danube basins: thin-skinned versus thick-skinned extension. Tectonophysics 252, 433e451. Lisle, R.J., Toimil, N., 2007. Defining folds on three-dimensional surfaces. Geol. Soc. Am. 35, 519e522. Lohr, T., Krawczyk, C.M., Oncken, O., Tanner, D.C., 2008. Evolution of a fault surface from 3D attribute analysis and displacement measurements. J. Struct. Geol. 30, 690e700. Maerten, L., Willemse, E.J.M., Pollard, D.D., Rawnsley, K., 1999. Slip distributions on intersecting normal faults. J. Struct. Geol. 21, 259e271. Magoon, L.B., Dow, W.G., 1994. The Petroleum System, From Source to Trap, AAPG Memoir, vol. 60, p. 655. Mansfield, C.S., Cartwright, J.A., 2000. Stratal fold patterns adjacent to normal faults: observations from the Gulf of Mexico. In: Cosgrove, J.W., Ameen, M.S. (Eds.), Forced Folds and Fractures, Geol. Soc. Spec. Pub, vol. 169, pp. 115e128.

D. Spahi
c et al. / Journal of Structural Geology 55 (2013) 182e195 Marchal, D., Guiraud, M., Rives, T., 2003. Geometric and morphologic evolution of normal fault planes and traces from 2D to 4D. J. Struct.Geol. 25, 135e158. Martel, S.J., 1997. Effects of cohesive zones on small faults and implications for secondary fracturing and fault trace geometry. J. Struct. Geol. 19, 835e847. Matenco, L., Radivojevi
c, D., 2012. On the formation and evolution of the Pannonian Basin: constraints derived from the structure of the junction area between the Carpathians and Dinarides. Tectonics 31, TC6007. http://dx.doi.org/10.1029/ 2012TC003206. McClay, K.R., 1990. Extensional fault systems in sedimentary basins: a review of analogue model studies. Mar. Petrol. Geol. 7, 206e233. McConnell, D.A., Kattenhorn, S.A., Benner, L.M., 1997. Distribution of fault slip in outcrop-scale fault-related folds, Appalachian Mountains. J. Struct. Geol. 19, 257e267. McLeod, A.E., Dawers, N.E., Underhill, J.R., June 2002. The influence of fault Array evolution on Synrift sedimentation patterns: controls on deposition in the StrathspeyeBrenteStatfjord half Graben, northern North sea. AAPG Bull. 86, 1061e1093. Mynatt, I., Bergbauer, S., Pollard, D.D., 2007. Using differential geometry to describe 3-D folds. J. Struct. Geol. 29, 1256e1266. Nicol, A., Walsh, J.J., Watterson, J., Breatn, P.G., 1995. Three-dimensional geometry and growth of conjugate normal faults. J. Struct.Geol. 17, 847e862. Nicol, A., Watterson, J., Walsh, J.J., Childs, C., 1996. The shapes, major axis orientations and displacement patterns of fault surfaces. J. Struct.Geol. 18, 235e248. Passchier, C.W., 2001. Flanking structures. J. Struct.Geol. 23, 951e962. Passchier, C.W., Mancktelow, N.S., Grasemann, B., 2005. Flow perturbations: a tool to study and characterize heterogeneous deformation. J. Struct. Geol. 27, 1011e1026. 
Paulissen, W., Luthi, S., Grunert, P., Cori c, S., Harzhauser, M., 2011. Integrated highresolution stratigraphy of a Middle to Late Miocene sedimentary sequence in the central part of the Vienna Basin. Geologica Carpathica 62, 155e169. Peacock, D.C.P., 2002. Propagation, interaction and linkage in normal fault systems. Earth-sci. Rev. 58, 121e142. Peacock, D.C.P., Sanderson, D.J., 1991. Displacements, segment linkage and relay ramps in normal fault zones. J. Struct.Geol. 13, 721e733. Peacock, D.C.P., Sanderson, D.J., 1996. Effects of propagation rate on displacement variations along faults. J. Struct.Geol. 18, 311e320. Peresson, H., Decker, K., 1997. The Tertiary dynamics of the northern Alps (Austria): changing palaeostresses in a collisional plate boundary. Tectonophysics 272, 125e157. Pollard, D.D., Segall, P., 1987. Theoretical displacements and stresses near fractures in rocks. In: Atkinson, B.K. (Ed.), Fracture Mechanics of Rock. Academic Press, London, pp. 277e349. Porras, J.S., Vallejo, E.L., Marchal, D., Selva, C., 2003. Extensional folding in the Eastern Venezuela Basin: examples from fields of Oritupano-Leona Block. Search Discov. p. Article #50003:7. Ratschbacher, L., Frisch, W., Linzer, H.-G., Marle, O., 1991. Lateral extrusion in the Eastern Alps, part 1. Structural analysis. Tectonics 10, 257e271. Reber, J.E., Dabrowski, M., Schmid, D.W., 2012. Sheath fold formation around slip surfaces. Terra Nova 24, 417e421. Reches, Z., Eidelman, A., 1995. Drag along faults. Tectonophysics 247, 145e156. Resor, P.G., 2008. Deformation associated with a continental normal fault system, western Grand Canyon, Arizona. GSA Bull. 120, 414e430. Resor, P.G., Pollard, D.D., 2012. Reverse drag revisited: why footwall deformation may be the key to inferring listric fault geometry. J. Struct. Geol. 4, 98e109. Rippon, J.H., 1985. Contoured patterns of throw and hade of normal faults in the coal measures (Westphalian) of northeast Derbyshire. Proc. Yorkshire Geol. Soc. 45, 147e161.

195



Rögl, F., Spezzaferi, S., Cori c, S., 2002. Micropaleontology and biostratigraphy of the Karpatrian-Badenian transition (Early-Middle Miocene boundary) in Austria (Central Paratetis). Cour. Forschungen. Senckenberg 237, 47e67. Royden, L.H., 1985. The Vienna Basin: a thin-skinned pull-apart basin. In: Biddle, K.T., Christie-Blick, N. (Eds.), Strike Slip Deformation, Basin Formation and Sedimentation: Society of Economic Paleontologists and Mineralogists, Spec. Pub., vol. 37, pp. 319e338. Rykkelid, E., Fossen, H., 2002. Layer rotation around vertical fault overlap zones: observation from seismic data, field examples, and physical experiments. Mar. Petro. Geol. 19, 181e192. Sauer, R., Seifert, P., Wessely, G., 1992. Guidebook to excursions in the Vienna Basin and the adjacent AlpineeCarpathian thrustbelt in Austria. Mitteil. der Geolog. Gesell. 85, 1e264. Schmid, S.M., Bernoulli, D., Fügenschuh, B., Matenco, L., Schefer, S., Schuster, R., Tischler, M., Ustaszewski, K., 2008. The AlpineeCarpathianeDinaridic orogenic system: correlation and evolution of tectonic units. Swiss J. Geosci.101 (1),139e183. Sclater, J.G., Royden, L., Horvath, F., Burchfiel, B.C., Semken, S., Stegena, L., 1980. The formation of the intra-carpathian basins as determined from subsidence data. Earth. Planet. Sci. Lett. 51, 139e162. Shelton, W., 1984. Listric normal faults: an illustrated summary. Am. Assoc. Pet. Geol. Bull. 68, 801e815. Spahi
c, D., Exner, U., Behm, M., Grasemann, B., Haring, A., Pretsch, H., 2011. Listric versus planar normal fault geometry: an example from the Eisenstadt-Sopron Basin (E Austria). Int. J. Earth Sci. 100, 1685e1695. Strauss, P., Harzhauser, M., Hinsch, R., Wagreich, M., 2006. Sequence stratigraphy in a classic pullapart basin (Neogene, Vienna Basin). A 3D seismic based integrated approach. Geol. Carp. 57, 185e197. Taylor, S.K., Nicol, A., Walsh, J.J., 2008. Displacement loss on growth faults due to sediment compaction. J. Struct. Geol. 30, 394e405. Tearpock, D.J., Bischke, R.E., 2003. Applied Subsurface Geological Mapping, p. 821. Wagreich, M., Schmid, H.P., 2002. Backstripping dip-slip fault histories: apparent slip rates for the Miocene of the Vienna Basin. Terra Nova 14, 163e168. Walsh, J.J., Watterson, J., 1989. Displacement gradients on fault surfaces. J. Struct. Geol. 11, 307e316. Walsh, J.J., Watterson, J., Bailey, W.R., Childs, C., 1999. Fault relays, bends and branch lines. J. Struct. Geol. 21, 1019e1026. Walsh, J.J., Nicol, A., Childs, C., 2002. An alternative model for the growth of faults. J. Struct. Geol. 24, 1669e1675. Walsh, J.J., Bailey, W.R., Childs, C., Nicol, A., Bonson, C.G., 2003. Formation of segmented normal faults: a 3-D perspective. J. Struct. Geol. 25, 1251e1262. Watterson, J., 1986. Fault dimensions, displacements and growth. Pure Appl. Geophys. 124, 365e373. Weissenbäck, M., 1996. Lower to Middle Miocene sedimentation model of the central Vienna Basin. In: Wesssely, G., Liebl, W. (Eds.), Oil and Gas in Alpidic Thrustbelts and Basins of the Central and Eastern Europe, EAGE. Spec. Publ., vol. 5, pp. 355e363. Wessely, G., 1988. Structure and development of the Vienna basin in Austria. In: Royden, L.H., Horváth, F. (Eds.), AAPG Bull., vol. 45, pp. 333e346. Wiesmayr, G., Grasemann, B., 2005. Sense and non-sense of shear in flanking structures with layer-parallel shortening: implications for fault related folds. J. Struct.Geol. 27, 249e264. Willemse, E.J.M., Pollard, D.D., Aydin, A., 1996. Three-dimensional analyses of slip distributions on normal fault arrays with consequences for fault scaling. J. Struct. Geol. 18, 295e309. Yamada, Y., McClay, K., 2003. Application of geometric models to inverted listric fault systems in sandbox experiments. Paper 1: 2D hanging wall deformation and section restoration. J. Struct. Geol. 25, 1551e1560. Zee, W.v.-d., Urai, J.L., 2005. Processes of normal fault evolution in a siliciclastic sequence: a case study from Miri, Sarawak, Malaysia. J. Struct.Geol. 27, 2281e2300.