Morphotectonics inferred from the analysis of topographic lineaments auto-detected from DEMs: Application and validation for the Sinai Peninsula, Egypt

Tectonophysics 510 (2011) 291–308

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Tectonophysics j o u r n a l h o m e p a g e : w w w. e l s ev i e r. c o m / l o c a t e / t e c t o

Morphotectonics inferred from the analysis of topographic lineaments auto-detected from DEMs: Application and validation for the Sinai Peninsula, Egypt Alaa A. Masoud a,⁎, Katsuaki Koike b, 1 a b

Geology Department, Faculty of Science, Tanta University, Tanta, Egypt Graduate School of Science and Technology, Kumamoto University, Kumamoto, Japan

a r t i c l e

i n f o

Article history: Received 6 July 2010 Received in revised form 22 January 2011 Accepted 12 July 2011 Available online 28 July 2011 Keywords: Adaptive shading Segment tracing Lineament Fault type Sinai Peninsula Egypt

a b s t r a c t Morphotectonic lineaments observed on the Sinai Peninsula in Egypt were auto-detected from Shuttle Radar Topography Mission 90-m digital elevation model (DEM) and gravity grid data and then analyzed to characterize the tectonic trends that dominated the geologic evolution of this area. The approach employed consists of DEM shading, segment tracing, grouping, statistical analysis of the distribution and orientation of the lineaments, fault plane characterization, and smooth representation techniques. Statistical quantification of counts, mean lengths, densities, and orientations was used to infer the relative severity of the tectonic regimes, to unravel the prominent structural trends, and to demarcate the contribution of various faulting styles that prevailed through time. Restored to the present-day geographic position, prominent N50°–60°W, N20°–40°W, N50°–60°E, and N20°–30°E and less prominent N–S, E–W, and ENE trends were common. The prominence of these trends varied through time. The NW and NE trends showed relatively equal abundances in the Precambrian and the Cambrian whereas the prominence of the NW trends prevailed from the Carboniferous to the Holocene. Lineaments in all formations were near vertical and on average, about 65% showed as strike-slip, 22% as reverse, and 13% as normal faulting styles. Statistics from the detected linear features and the reference geological data reveal the relative severity of five dominant tectonic regimes: Precambrian compression followed by extension at its end, Cretaceous compression, Eocene compression, Miocene extension, and finally Holocene compression. Auto-detected lineaments and the severity of the characterized tectonic periods correspond well with reference data on the geologic structure, geodynamic framework, and the gravity anomaly. Furthermore, the recent significance of the broad structural zones was confirmed by the foci of the earthquake epicenters along and at the intra-plate intersections of the broad lineament zones and at the plate boundaries. The spatial distribution of trends with varying styles of faulting distinguished four main tectonic provinces of marked geodynamics variances. © 2011 Elsevier B.V. All rights reserved.

1. Introduction Crustal tectonic deformation and topography share a common origin in the large-scale dynamics of the underlying mantle and upper mantle thermal perturbations (Cloetingh et al., 2003, and references therein). These dynamics affect various mechanisms of observed large-scale crustal uplift (Gurnis et al., 2000; Lithgow-Bertelloni and Silver, 1998) and subsidence (Mitrovica et al., 1989; Stern and Holt, 1994). Stress fields driving lithospheric tectonic deformation are largely confined to plate boundaries, but can transmit to the plate interior where they interplay over the long term with climatic/ geomorphic processes and give rise to continental domains (e.g., Burbank and Anderson, 2001; Roessner and Strecker, 1997; Summerfield, 2000). The structures and orientations of these domains are largely ⁎ Corresponding author at: Geology Department, Faculty of Science, Tanta University, 31527 Tanta, Egypt. Tel.: + 20 40160377898, fax; + 20 403350804. E-mail address: [email protected] (A.A. Masoud). 1 Present: Graduate School of Engineering, Kyoto University, Kyoto, Japan. 0040-1951/$ – see front matter © 2011 Elsevier B.V. All rights reserved. doi:10.1016/j.tecto.2011.07.010

controlled by pre-existing crustal discontinuities and stress fields (Cloetingh et al., 2006, 2009, and references therein). Therefore, the orientation of landforms, drainage patterns, linear elements of rivers and valleys, ravines, escarpments, and lower edges of terrace slopes all have a bearing on mass transfer at the Earth's surface (Cloetingh et al., 2006) and can be used to constrain future earthquake ruptures (Gorshkov et al., 2000). Tectonic characterization of topographic features has recently become a major concern, motivated by the global availability of largescale digital topographic datasets. The morphometrics of fluvial channel networks and location of the modern knick points are important in the estimation of variations in rock uplift and hence faulting (e.g., Peakall et al., 2000; Snyder et al., 2003; Whipple, 1999, 2009). Numerical dating of geomorphic surfaces can provide rates for the tectonic processes (e.g., Barbero et al., 2010; Hetzel et al., 2002). The geomorphic process of preferential weathering commonly acts on topographic discontinuity zones and enhances their linear features to develop geomorphologic lineaments in the form of linear valleys and slope breaks or ridges (Jordan et al., 2005). Geomorphologic

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lineaments have been found to have common geologic origins due to compositional contrasts at geologic contacts or heterogeneity of rock strength as characterized by fault scarps, joints, and fold axes (Ramsay and Huber, 1987), which commonly extend to the subsurface (Sabins, 2000). Tectonics played a key role in establishing the present-day surface configuration of the Sinai sub- or micro-plate (Badawy et al., 2008; Ben-Menahem et al., 1976). The micro-plate exhibits complicated interactions at its active boundaries with Africa, Arabia, and Eurasia. Inter-plate movements and intra-plate adjustments between crust and mantle processes have dominated Sinai's geodynamic evolution. These processes are commonly associated with the rifting, uplift, and sometimes rotation that have controlled Sinai's morphotectonic structures and overprinted its geomorphology in the form of topographic lineaments. An approach that detects and relates such lineaments to dominant tectonics is therefore extremely desirable to better understand the regional tectonic evolution. Furthermore, such tectonically-controlled features are commonly identified as conduit zones for fluid migration; as such their accurate characterization is of paramount importance in the exploration for water, oil, gas, geothermal energy, and ore deposits (Krishnamurthy et al., 2000; Lattman and Parizek, 1964; Magowe and Carr, 1999). Despite the importance of Sinai tectonics, research efforts devoted to lineament detection and mapping on the peninsula are scarce and commonly rely on the visual inspection of either a set of filtered images of local geophysical data such as the total magnetic intensity (Rabeh and Miranda, 2008; Rabie and Ammar, 1990) or linear features traced from satellite image data (Kusky and El-Baz, 1998). These techniques are tedious, time-consuming, and likely to be biased by subjective interpretations. Moreover, the accuracy of lineament detection from satellite images is strongly conditioned by the sensor characteristics, illumination conditions, and spatio-spectral resolution (Smith and Wiseb, 2007). Digital-elevation-model (DEM) derived shaded relief images and terrain spatial parameters (e.g., gradient, aspect, and curvatures) have proven to be promising alternative data sources for lineament and fault extraction that avoid the limiting factors mentioned above (Ganas et al., 2005; Hooper et al., 2003; Jordan, 2003a, 2003b). Based on the above background, a set of integrated data-processing techniques has been employed to extract lineaments from the DEM and global gravity grid data of Sinai, which aims to unravel its crustal structures. Furthermore, spatial and statistical analyses of these features, with the incorporation of surface geology, are performed to characterize the tectonic regimes that were dominant through time to lead to better understanding of the geodynamic evolution of the Sinai Peninsula. 2. Geologic and tectonic setting of the Sinai Peninsula Widespread extensional processes alternating with short compressional processes (e.g., Be'eri-Shlevin et al., 2009; Guiraud and Bosworth, 1999) have interplayed to control Sinai's active boundaries, its present-day local geomorphology, and the patterns of rock distributions spanning from the Precambrian to Recent. Plate movements and adjustments between crust and mantle processes formed the basement complex (Be'eri-Shlevin et al., 2009; Blasband et al., 2000; Cochran and Karner, 2007; Mahmoud et al., 2005; Meert, 2003). Expansion and contraction of the ancient Tethys Ocean, mostly related to mantle thermal perturbations, controlled the northward dominance of the extensive veneer of Phanerozoic cover (Bumby and Guiraud, 2005; Guiraud and Bosworth, 1999). Details of the geology of Sinai, covering an area of ~ 61,000 km 2 are addressed by Said (1962, 1990) and references therein. Precambrian igneous and metamorphic basement rocks belonging to the Arabian– Nubian Shield reach an altitude of 2629 m (G. Saint Catherine) in the south and are overlain by a Phanerozoic sedimentary wedge in the north that gently dips and generally becomes more recent towards the Mediterranean coast (Fig. 1). The Phanerozoic sediments are

composed mainly of: (1) Paleozoic (Cambrian and Carboniferous) sediments, (2) Mesozoic limestone, marl, and chert, and (3) Tertiary (except Oligocene) and Quaternary sediments; these are interspersed with Phanerozoic volcanics. According to the Atlas of Israel (1985), the scattered occurrences of Oligocene sedimentary rocks, associated with Miocene sedimentary rocks, are limited to the Gulf of Suez side. Recent surfaces in the northern Sinai are dominated by dunes and sand plains made up of unconsolidated and aeolian sands, wadi alluvium, clays, conglomerates and gravel terraces, and uplifted beaches and coral reefs. Sandy dune ridges dominate in the northwest between Bardawil Lake and the Great Bitter Lake. In the southern part of the dune sheets, a few outliers of massive limestone, such as G. Yelleq, Maghara, and Halal, rise above the undulating surface. In the central Sinai the Tih and Egma Plateaux continue the dominant southward rise, culminating at the Tih escarpment with elevations as high as 1600 m. South of the escarpment, elevations drop where the drainage systems of Feiran (west) and Watir (east) flow between the plateau and the Precambrian complex and finally to the gulfs of Suez and Aqaba. Quaternary sediments largely made up of a series of Mesozoic and Tertiary rocks cover the Qaa Plain at the western coastal belt in the south. The Sinai microplate (Fig. 2) consists of a triangular continental crustal block locked between the major Arabia and Africa plates and the Anatolian–Aegean microplate (Joffe and Garfunkel, 1987; Le Pichon and Francheteau, 1987; McKenzie, 1970; McKenzie et al., 1970). The microplate is bounded eastward by the Dead Sea leftlateral transform and its submerged extension in the Gulf of Aqaba. The Dead Sea connects to the north with the left-lateral East Anatolian transform fault zone that is in continuity with the Cyprus active margin (Ben Avraham et al., 1995, and references therein) that demarcates the microplate's northern boundary. The Cyprean Arc is dominated by subduction and transcurrent processes along its western and eastern extents, respectively (Wdowinski et al., 2006). The Gulf of Suez is a confirmed segment of the microplate's western boundary. Recent evidence from shallow seismic activities (Badawy et al., 2008; Kebeasy, 1990) and marine geophysical data (Mascle et al., 2000) confirm the northward offshore extension of the Gulf of Suez as a boundary of the microplate. Plate motion along this boundary is still a matter of debate (Badawy et al., 2008). Minor transcurrent components have been recognized in the Gulf of Suez (Abdel-Gawad, 1969; Joffe and Garfunkel, 1987): their sense of motion favors a relative left-lateral strike-slip movement, which agrees with focal mechanisms of fault-plane solutions from the region (Abdel-Gawad, 1969; Badawy et al., 2008; Garfunkel and Bartov, 1977; Joffe and Garfunkel, 1987; McKenzie, 1977). Seismological and GPS constraints reveal the dominance of slab-pull rather than ridgepush forces for the motion of Sinai relative to Africa (Badawy et al., 2008). Badawy et al. (2008) show that the southern Gulf of Suez is characterized by extensional deformation with left-lateral strike slip that is consistent with the kinematic model of Badawy (2005), Badawy and Horváth (1999b), and Mahmoud et al. (2005). However, the northern segment is characterized by unexpected compressional deformation, inconsistent with earthquake focal mechanisms and regional tectonics. Earthquake focal mechanisms are dominated by normal faulting accompanied by a left-lateral strike slip component. The strike-slip component gradually increases northward (Badawy et al., 2008). The separation of the Sinai microplate from Africa in the Miocene, in the framework of the Red Sea opening, continued onwards (BenMenahem et al., 1976, Cochran, 2005, and references therein; McClay and Khalil, 1998; Steckler et al., 1988). The Suez Rift was faulted in the Miocene, jointly with the northern Red Sea (Bartov et al., 1980c; Garfunkel and Bartov, 1977). The rate of tectonic activity there has reduced considerably since the latest Miocene (Garfunkel and Bartov, 1977 and Steckler, 1985), concurrently with the opening of the Aqaba Rift. The Aqaba and the more regional Dead Sea transform fault (offset

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Fig. 1. Geologic and geomorphologic features of the Sinai Peninsula.

of 105–107 km) has been active since 15 Ma (Freund et al., 1970; Garfunkel et al., 1981; Mart and Rabinowitz, 1986; Quennell, 1958, 1984); however, most of its tectonic evolution has occurred since the Pliocene (Horowitz, 1987; Mart, 1991), with regional transtension responsible for rift depressions, which connected the extensional sea floor spreading in the Red Sea in the south with the Zagros–Taurus continental collision zone in the north. The motion of the Sinai microplate along the Dead Sea left-lateral transform fault is estimated at about 8–9 mm/year (Garfunkel et al., 1981), while slower extensional motion of about 1 mm/year (Bosworth and Taviani, 1996; Steckler et al., 1988) is registered along the Gulf of Suez. 3. Data and method Topography, represented by the Shuttle Radar Tomography Mission (SRTM) 90-m DEM, was used as the base source data for detecting lineaments. The sequence of processes followed for the autodetection and analysis of lineaments is: (1) DEM shading applying the Adaptive-Tilt Multi-Directional Shading (ATMDS) method of Masoud and Koike (2011), (2) tracing line segments using segment tracing

algorithm (STA) of Koike et al. (1995, 1998), (3) proximity and colinearity based grouping of line segments to lineaments (Koike et al., 1998), (4) smooth representation of lineament appraisal using of BSpline Curve technique (Masoud and Koike, 2009), 5) modeling faulting styles (normal, reverse, and strike-slip) and geometry (dipping criteria) (Masoud and Koike, 2011), and 6) statistical and orientation analyses. Due to the lack of detailed peninsula-wide coverage of proper structural data upon which to base the evaluation, a regional approach was needed to evaluate the detected lineaments and their trends. Line segments and lineaments were visually and statistically evaluated against structural features of 7277 main faults digitized from the geological map of Egypt (Klitzsch et al., 1987) and 1985 faults and dykes from the Sinai geological photomap (Bartov et al., 1980b), which are referred to through the remainder of this manuscript as CONOCO87 and BARTOV80, respectively. STA was applied to the 1-arc-minute gravityanomaly (mGal) grid by Sandwell and Smith (1997); prominent trends were also included in the evaluation process. This is because low gravity anomalies commonly represent deep structural zones that results from compositional density contrasts in the subsurface (Masoud and Koike,

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Fig. 2. Tectonic framework of the Sinai microplate and its surrounding area modified after Mascle et al. (2000). Movements in the Gulf of Suez area are quoted from Badawy et al. (2008).

2011; Smith and Sandwell, 1994), which provides a basis for joint interpretation with topography. Earthquake epicenters made available from the databases of the National Earthquake Information Center (NEIC) and the Egyptian National Seismograph Network (ENSN) were utilized to survey the recent (1900–2006) tectonic activity and significance of lineaments. Line segments and lineaments located within various geologic units were characterized to detect remarkable changes through time. Geology was compiled from five sheet maps covering the peninsula (EGSMA, 1995). Geologic contacts were first digitized from separate sheets, then patched together, and finally geo-referenced to UTM coordinate system zone 36. The GIS generalization technique implemented in GRASS GIS (GRASS Development Team, 2010) was applied to regionalize the twelve major geologic units out of the digitized eighty-seven rock varieties. A stand-alone software tool was encoded in visual basic for the digitization and the spatial and orientation analysis of the detected and digitized faults. The mathematical background and evaluation of the techniques applied have been addressed in detail by Masoud and Koike (2011). Therefore, only brief explanations of the core components of the techniques, e.g., the ATMDS, STA, and style of fault modeling, are given here. ATMDS sets the six northward illumination azimuths at 30° intervals, clockwise from west, and automatically sets the tilt angle from 0° to 45°, upward from the horizon. At each DEM grid node, shade intensities in the defined ranges are calculated and the

maximum value is selected: this process is repeated for all grid nodes. The reason for selecting these limited directions is because opposite southward illumination azimuths and tilt angles greater than 45° suppress shade intensities rather than enhance them. STA is a non-filtering approach to the accurate extraction of continuous straight lines. In STA, local variations of gray levels are examined using a local window (in this case, 11 by 11 pixels) centered sequentially on every grid node (Fig. 3). The direction which minimizes the variation is assumed to be a line segment. Judging whether the line segment represents a ridge or valley feature is performed using the curvature criteria from the DEM. Where segments along valleys are retained and those located at ridges are eliminated. A trace of a fracture or fault plane on the terrain composed of concatenated line segments can be used to calculate the fault plane geometry (i.e., strike and dip). The trace curvature on the surface depends on the dip angle of the interpreted plane; the style of faulting can be then judged based on the geometry of the terrain and the fault plane normals (Masoud and Koike, 2011). The directional relationships between the normals and both the terrain and fault planes are used to judge whether the obtained plane represents a normal or reverse fault feature Fig. 4. Three features were set: normal, reverse, and strike-slip types depending on whether the directions of the normals are the same (normal), opposite (reverse), and undetermined (strike-slip). Because strike-slip movements generally do not accompany downthrows on the fault plane, the fault type cannot be

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Fig. 3. Local window with array size of 11 by 11 pixels and scan lines at π/16 radian intervals used in STA. Numbers denote examples of the scan line directions for examining gray level variations.

specified by a simple geometrical relationship. Therefore, the undetermined type is assumed to have either pure strike-slip movement or to be combined with normal or reverse components. 4. Results 4.1. Trends and styles of faulting of line segments and lineaments Line segments and grouping-derived lineaments are shown in Figs. 5 and 6, respectively. These features are remarkably focused in areas of rolling topography — and absent in the flats of the central, northern, and Qaa plains, where there was no distinctive neighbor pixel-wise change in slope. Since their absence could be related to the limited spatial resolution of the DEM, a finer-scale DEM of ASTER 30 m data was tested; however, satisfactory results were not achieved. It seems that linear features in these areas would be strongly affected by water flow and sediment load, i.e., climatic rather than tectonic effects, which mean that there should be a resolution limit for local and regional topography to infer pure tectonic environments. Temporal evolution of the tectonic trends, fault type contributions, and dipping criteria for line segments and lineaments in various geologic formations are shown in Fig. 7 using rose diagrams and hemisphere projections. Tectonic trends for the whole peninsula are prominent along NW–SE and E–W. Geologic formations show varied trends judged from the lineaments and, when less prominent trends disappear, trends of the line segments are used instead. In general, prominent N40°–60°W, N50°–60°E, N20°–40°W, N10°– 30°E, and less-prominent N–S and E–W trends dominated from the Precambrian to the Holocene. Prominence of the ENE trend continued from Carboniferous to Jurassic, in the Paleocene, and from the

a

Miocene to Pleistocene. Jurassic and the Quaternary deposits do not show prominence for the NNE trend. Also, the NNW trend is not conspicuous in the Precambrian, Jurassic to Eocene, Pliocene, and Holocene. During the Paleocene the NW trend showed a 10° counterclockwise rotation. The disappearance of tectonic trends in the formations is likely related, either individually or in combination, to their limited spatial extent, relative strength, or to the distal position from the active structural zones during the development of these formations. This also can possibly be related to the spatial resolution of the DEM, in particular, formations covering extensive flat areas in wadis (Quaternary deposits), or for intensively weathered formations (Pliocene). The NE trends were more abundant in the Precambrian compared with the NW trends. Both trends showed similar abundance in the Cambrian, whereas the abundance of the NW trends dominated from the Carboniferous onwards. Faulting styles of lineaments are shown in Fig. 8. Lineaments in all formations were vertical to sub-vertical and on average, about 65% showed strike-slip, 22% reverse, and 13% normal faulting styles. Normal faulting was abundant in the Gulf of Suez rift area. The reverse faults were concentrated in the northern fold belt, and two zones were also distinctive in the Precambrian (see Fig. 8). Strike-slip styles of faulting dominated in the Precambrian, in particular, on the Gulf of Aqaba side. These styles, generally in conjunction with either normal or reverse faulting, emphasize the structural complexity of the northern fold belt. Normal faulting is most abundant in the Carboniferous, Triassic, and Jurassic. Reverse faults are most common in the Pliocene, Triassic, Paleocene, and Cretaceous. The highest lineament frequencies demarcated to have strike-slip components were recorded in Precambrian, Paleozoic, Eocene, and Miocene formations.

b

ti = (t1i , t 2 i , t3i ) T

li = (l1i , l 2i , l3i ) T

Average surface normal vectors

n

t

Lineament n Surface

Normal fault

n Reverse fault

Fig. 4. Diagram showing a) the definition of vectors used in estimating azimuth of “fracture” plane from lineaments, and b) schematic idea of judging fault type by the directional relationship between the average of surface normal vectors and the normal vector of the interpreted fracture plane.

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Fig. 5. DEM-derived line segments overlying shaded DEM.

4.2. Relative strength of tectonic regimes The prominent trends and statistics of detected line segments and lineaments were correlated with those derived from the statistical and orientation analyses of the segments digitized from reference geologic maps using the geologic scale as a basis for comparison (Table 1). Structural features digitized from the CONOCO87 and BARTOV80 data are shown in Figs. 9 and 10, respectively. Faults on CONOCO87 registered prominent NW, N–S, and NNW trends, whereas the long regional faults showed, in addition to the NNE trend, abundant ENE, NNW, and E–W trends. A rose diagram of fault orientations shows prominent NNE, N–S, and NW trends. Faults on BARTOV80 (Fig. 10) showed similar trends to the faults of CONOCO87 with more prominent ENE trends, whereas the dykes exhibited more abundant NW, E–W, and NNE trends. A rose diagram of the summed structural features indicates trends that are similar to those seen for the faults with N–S replacing the NNE trend. Reference data show similar trends to the DEM-derived segments and lineaments, but with varying prominence and abundances relative to the NE trends. In total, 101,409 DEM-derived line segments were identified on the peninsula; these were grouped into 7398 long concatenated

lineaments (Table 1). Segments had average lengths of 1.5 and 1.4 km, and maximum lengths of 15 and 25 km for the Precambrian and Cretaceous, respectively (see Table 1). Lineaments had average lengths of 5.2 and 4.3 km and maximum lengths of 38 and 30 km for the same geologic ages. This simply reflects the effective role of grouping, as a proximity- and co-linearity-parametric technique, to extrapolate lineaments in isolated flat areas located close to rolling topography — and thereby to enhance their continuity and connectivity at both local and broader scales, as will be shown later. Statistical parameters [particularly counts, length means (km) and densities (km/km 2)] of the DEM-derived line segments and lineaments and those derived from the reference data (Table 2) characterized the relative strength of the tectonic regimes dominant through time (Fig. 11). Counts show strong peaks in the Precambrian, Cretaceous, Eocene, and Holocene on all data. The length means and densities showed variable peaks with time. For the DEM-derived line segments and lineaments, length means showed peaks in the Cambrian, Triassic, Cretaceous, and Pliocene (line segments), and in the Precambrian, Cretaceous, and Pliocene (lineaments), whereas length densities peaked in the Cambrian, Jurassic, Miocene, and Holocene (line segments), and in the Precambrian, Cretaceous, and

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Fig. 6. DEM-derived lineaments overlying the geological map of Sinai.

Miocene (lineaments). Similar results were drawn from the statistics of the BARTOV80 and CONOCO87 data, in which length means showed peaks in the Precambrian, Carboniferous, Jurassic, Eocene,

and Holocene (BARTOV80), and in the Precambrian, Triassic, Paleocene, and Holocene (CONOCO87). The peaks of lineament density are common to both data sets for the Cambrian, Cretaceous,

Table 1 Descriptive statistics for DEM-derived line segments and lineaments. Geologic units

Holocene Pleistocene Pliocene Miocene Eocene Paleocene Cretaceous Jurassic Triassic Carboniferous Cambrian Precambrian

Area

Line segments

Lineaments

Length (km)

Count

(km2)

%

Mean

Max

Sum

20,859 9713 893 1661 8612 1639 12,980 475 98 267 1054 7768

31.6 14.71 1.35 2.52 13.05 2.48 19.66 0.72 0.15 0.4 1.6 11.77

1.33 1.42 1.49 1.43 1.37 1.37 1.43 1.38 1.45 1.45 1.49 1.5

16 18 13 29 13 12 25 9 7 8 14 15

36,155 12,663 1117 4597 18,057 2954 34,986 1786 367 882 3574 25,633

L/A = Length/Area, N/A = Number/Area

27,134 8886 747 3212 13,171 2160 24,401 1296 253 610 2391 17,143

Density

Length (km)

Count

L/A

N/A

Mean

Max

Sum

1.73 1.3 1.25 2.77 2.1 1.8 2.7 3.76 3.74 3.3 3.39 3.3

1.3 0.91 0.84 1.93 1.53 1.32 1.88 2.73 2.58 2.28 2.27 2.21

3.92 4.14 4.86 4.15 3.92 3.75 4.35 4.1 3.14 3.97 4.23 5.22

27 18 28 31 23 12 30 13 5 11 14 38

5847 1936 170 698 2660 165 6900 221 25 99 609 6791

1492 468 35 168 678 44 1586 54 8 25 144 1302

Density L/A

N/A

0.28 0.2 0.19 0.42 0.31 0.1 0.53 0.47 0.26 0.37 0.58 0.87

0.07 0.05 0.04 0.1 0.08 0.03 0.12 0.11 0.08 0.09 0.14 0.17

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and Miocene, in addition to the Triassic on the CONOCO87 data. From this, the most frequent peaks (3 out of the 4 data sets compared) of the length means were for the Precambrian and Cretaceous, whereas those for densities were recorded in the Cretaceous and Miocene. Therefore, it can be concluded that the tectonic regimes through the Precambrian, Cretaceous, Miocene, Eocene, and Holocene attained the highest strengths during the geodynamic evolution of Sinai. Lineament frequency was plotted on the hemisphere projections, where

the temporal variation of the relative strength of tectonic stress can be easily traced (see Fig. 7). Tectonic trends auto-detected from the DEM were congruent with the prominent trends of the BARTOV80 and CONOCO87 data. Gravity-derived line segments showed prominent NW, NE, NNE, E–W, and N–S trends in decreasing order of abundance. Dense line segments formed broad zones along the negative anomalies that correspond to the deepest crustal shears. These shear zones

Fig. 7. Rose diagrams and hemisphere projections of the DEM-derived line segments and lineaments and dominant stress fields (SF) with arrow length proportional to the strengths. Arrows indicate Sinai (black) plate motion (PM) relative to Eurasia (gray) since the Cretaceous (quoted from Guiraud and Bosworth, 1997 and Bunge et al., 2002); their lengths are proportional to the motion speed.

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299

Fig. 7 (continued).

are NW-oriented in the Gulf of Suez area, NNE-oriented in the Gulf of Aqaba area, NW-, N-S-, NNE-, and E-W-oriented in the basement, and NW-/NE-oriented beneath the plateaux; an almost conjugate set of NW- and NE-oriented zones underlie the northern Sinai fold belts (Fig. 12). Positive gravity features delimited by negative anomaly zones occur in the central plateaux area, the northern fold belts, and a roughly N–S-oriented granitic ridge that disturbs the continuity of the NW-oriented zones to the southeast in the basement. These trends agree well with those derived from the topography, which signifies the common deep origin of these structures (Fig. 11). Recent (1900–2006) large earthquakes (M N 5) have been concentrated along the gravity segments of the gulfs of Aqaba and Suez — particularly at the intersections of these segments with the intra-peninsula broad structural zones (Fig. 12). Although the epicenters within the peninsula are dispersed with low magnitudes (Badawy and Horváth, 1999a, 1999b, 1999c; Kebeasy, 1990), they are distributed at the intersections of the broad structural zones in low gravity anomaly zones. In the Gulf of Suez, the foci of strong

earthquake epicenters (M N 5) demarcate the intersection of the NW-oriented zones with N–S zone west of Qaa Plain comparable to the Durba and Ekma transfer faults of McClay and Khalil (1998), and with N–S and NE zones west of Ras Sidr. This intersection can explain the unexpected compressional deformation inconsistent with the earthquake focal mechanisms and regional tectonics in the Gulf of Suez raised by Badawy et al. (2008). Note that two distinctive reverse fault zones demarcated in the Precambrian (Fig. 8) also overlie two remarkable negative gravity anomalies trending roughly E–W and NW (Fig. 12). These two zones meet and intersect with the Gulf of Aqaba transform fault at the location of the strongest seismic activity (M N 8) recorded in the gulf. 5. Geodynamic evolution of tectonic trends The geodynamic evolution of NE Africa is well studied (e.g., Bentor, 1985; Be'eri-Shlevin et al., 2009; Blasband et al., 2000; Bumby and Guiraud, 2005; Cochran and Karner, 2007; Guiraud and Bellion, 1995; Guiraud and Bosworth, 1999; Meert, 2003; Stern, 2002; Stein and

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Fig. 8. Fault types modeled from the topographic lineaments. Note the two concentrated reverse fault zones (outlined in black).

Goldstein, 1996). Until the instigation of Red Sea rifting in the Miocene, the Sinai had a common geodynamic evolution with NE Africa. Since then, the Sinai has had its present-day configuration as a

sub- or micro-plate. A review is given here for the evolution of the dominant tectonic regimes from the Precambrian onwards. Their relevance to detected lineaments and tectonic trends is discussed.

Table 2 Descriptive statistics for the line segments derived from the reference data. Geologic units

Holocene Pleistocene Pliocene Miocene Eocene Paleocene Cretaceous Jurassic Triassic Carboniferous Cambrian Precambrian

CONOCO 87

BARTOV 80

Length (km)

Count

Mean

Max

Sum

4.09 3.85 3.54 3.38 3.62 3.74 3.61 3.1 3.95 2.92 3.82 4.32

47 18 18 14 32 33 29 7 7 9 15 24

3983 2361 301 929 4226 685 7889 229 83 172 853 6128

L/A = Length/Area, N/A = Number/Area.

974 613 85 275 1167 183 2186 74 21 59 223 1417

Density

Length (km)

Count

L/A

N/A

Mean

Max

Sum

0.19 0.24 0.34 0.56 0.49 0.42 0.61 0.48 0.85 0.64 0.81 0.79

0.05 0.06 0.1 0.17 0.14 0.11 0.17 0.16 0.21 0.22 0.21 0.18

3.8 2.9 2.2 3.5 3.8 2.7 3.3 4.4 4.2 4.3 3.9 4.2

17 9 3 9 16 7 31 15 6 9 11 22

842 570 16 286 698 144 2286 83 17 69 363 1612

223 195 7 82 182 54 699 19 4 16 92 385

Density L/A

N/A

0.04 0.06 0.02 0.17 0.08 0.09 0.18 0.18 0.17 0.26 0.34 0.21

0.01 0.02 0.01 0.05 0.02 0.03 0.05 0.04 0.04 0.06 0.09 0.05

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301

Fig. 9. Structures modified after Klitzsch et al. (1987) including faults (black) and major faults (red) and their generated rose diagrams.

5.1. Precambrian trends Precambrian lineaments show prominent N50°–60°W, N50°–60°E, N20°–30°E and less prominent N–S and E–W trends. Vertical to subvertical NW and NE trends showed nearly equal abundances, which were dominated by strike-slip styles of faulting (59%) followed by reverse (27%) and normal types (14%). The NW-oriented lineaments were judged to have mixed styles of faulting, with dominant strike-slip and subordinate normal and reverse types. The N50°–60°E trend clarified principal southeasterly dipping strike-slip and normal types and northwesterly dipping strike-slip and reverse types. Strike-slip and reverse types are common on the N20°–30°E trend. Strike-slip and reverse types characterized the N–S and the E–W trends, respectively. The NW-oriented spatial distribution (80% of the basement) and orientation of the metamorphic core complexes, their mesoscopic and macroscopic structures as well as the fold axes, the associated NE-oriented normal faults, and the NE-trending basic and felsic dyke swarms (latest Neoproterozoic to Early Cambrian) are commonly aligned along or parallel to the detected Precambrian tectonic trends (see, e.g., Be'eri-Shlevin et al., 2009; El-Shafei and Kusky, 2003; Fowler and Hassan, 2008). These belts were formed during extensional

tectonism (Blasband et al., 2000; Fowler and Hassan, 2008). Postorogenic A-type granites, molasse type sedimentary basins, and mafic and felsic dykes are examples of the geologic features developed in the extending crust slightly after the core complexes (El-Sayed, 2006; Greiling et al., 1994; Stern et al., 1984). Ion-probe U–Pb dating of zircon in the basement of the Sinai (Be'eri-Shlevin et al., 2009) clarified that Late Proterozoic NW–SE extension was synchronous with the transition to alkaline A-type magmatism commencing at c. 600 Ma, which followed older postcollisional calc-alkaline (c. 635–590 Ma) and within-plate alkaline magmatism (c. 608–580 Ma). Fowler and Hassan (2008) proposed an extensional regime for the gneissosity-forming and migmatization event of the Feiran core complex. They mapped steep NE to NNE trending dextral faults and NW-trending sinistral faults that followed the granite intrusion and were crosscut by the younger NE-trending basaltic and felsite dyke swarms. Blasband et al. (2000) suggested that extension in the Kid core complex was active until c. 580 Ma based on the reported Rb–Sr ages of intruding A-type granites of this age (Moghazi et al., 1998). They reported two deformation phases in the Wadi Kid metamorphic belt: an older arcaccretion with WNW- to NW-oriented compression (720–650 Ma)

302

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Fig. 10. Structures after Bartov et al. (1980a) including faults (black) and Miocene dykes (red) and their generated rose diagrams.

followed by NW–SE extension synchronous with the emplacement of the mafic and felsic dykes. The NNE trend in the Precambrian basement is related to the later left-lateral movement distinguished into two phases of offset: 62 km in the Miocene and 45 km in the Pliocene to Recent associated with the Dead Sea Transform fault initiated since the Late Miocene (Quennell, 1958). The c. 24–21 Ma rift initiation dyke swarms of Sinai are offset by ~107 km, the same amount as basement structures (Eyal et al., 1981; Quennell, 1958). The N–S and E–W trends form subordinate sets likely resulting from older N–S-directed compressional regimes (Youssef, 1968). 5.2. Paleozoic trends The Cambrian is characterized by the appearance of a remarkable NNW trend that continued into the Carboniferous. Paleozoic NW trends showed a gradual increase in abundance, whereas NE trends decrease in concentration as compared with their relatively widespread abundance in the Precambrian. Dominant faulting styles in the Cambrian trends are similar to the Precambrian, whereas in the

Carboniferous normal styles dominate over reverse types. The NW trends mainly showed southwesterly dips on strike-slip and normal faults. Paleozoic trends are associated with extension that continued through successive extensive erosional phases associated with gentle tectonic deformation that caused regional uplift (Fabre, 1988) and consequently developed systematic hiatuses and slight unconformities (Issawi, 1996; Wennekers et al., 1996). These had limited exposure (2%) in the sedimentary units close to the northern boundary of the peneplained basement. Paleozoic tectonic activities likely reactivated the Precambrian trends as evidenced by the similarity of the trends and the abundance of NW trends relative to NE trends that continues to the Recent. This abundance suggests a strong link with the early evolution of the Neotethys that initiated at the end of the Paleozoic (Bumby and Guiraud, 2005; Guiraud and Bosworth, 1999). 5.3. Mesozoic trends In general, NW trends dominate over NE trends throughout the Mesozoic; NW, NNE, and E–W trends are common. The NNW trend

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a

b

4

6 Density (length/area)

Density (length/area)

3.5 3

303

Mean

5

Mean Counts/1,000

Counts/10,000

4

2.5 2

3

1.5

2

1 1

0.5 0

0

Line segments

c

Lineaments

d

4.5 4

7 6

Density (length/area) Mean

2.5

Mean Counts/100

3.5 3

Density (length/area)

Counts/1,000

2

5 4 3

1.5

2

1 1

0.5

0

0

CONOCO87

BARTOV80

Fig. 11. Statistical characteristics (counts, length mean and density) of a) line segments, b) lineaments, c) segments of CONOCO87, and d) segments of BARTOV80. Short vertical lines are used to show the position of weak peaks.

showed prominence in the Triassic, less prominence in the Jurassic, and disappeared in the Cretaceous. The N–S trend showed prominence only in the Cretaceous. Interestingly also, a 10° counterclockwise rotation (discussed below) deviates from its common trend in the Paleozoic. The NW-trends showed common southwesterlydipping normal and strike-slip faulting in both the Triassic and Jurassic and mixed styles of faulting during the Cretaceous. The abundance of normal faulting over reverse styles characterized the Jurassic, whereas increased Cretaceous reverse faulting was more frequent than normal faulting. Mesozoic extensional trends are associated with the end of Paleozoic to Late Cretaceous rifting. These trends include successive episodes of Neotethys rifting, the far-field effects of extension between Madagascar and Africa, and Early through Late Cretaceous N–S-oriented extensional to NE–SW-aligned tensional or transtensional tectonic regimes related to the opening of the South and Equatorial Atlantic Oceans (Guiraud and Bosworth, 1999). Tectonic instability, underscored by uplifting, block tilting, and sometimes slight folding, affected the passive margin initiated in and developed since the Triassic and Jurassic. This instability resulted from local transpression (Guiraud, 1998) that is synchronous with the Eastern Mediterranean basin that developed in the Triassic then underwent continental crustal thinning supported by deep-seated magmatic intrusions (El Toukhy et al., 1998; Garfunkel and Derin, 1988). These deformations represent the distal effects of the Eo-Cimmerian

orogenic event in the Black Sea area (Nikishin et al., 1999). Intraplate continental tholeiitic basalts erupted in the Early Triassic in west Central Sinai (Meneisy, 1986) and alkali basaltic sills intruded into southwest Sinai (Ibrahim et al., 2000). Their paleomagnetic characteristics are inconsistent with their counterparts in Africa because of local tectonic rotations and/or geomagnetic secular variation (Ibrahim et al., 2000; Wassif, 1991), which were facilitated by the normal faulting dissecting these areas (Meshref, 1990). This rotation was apparent on the Jurassic N60°–70°W trend that was deviated ~ 10° counterclockwise from its common N50°–60°W orientation. Frequent unconformities and gaps in the sedimentary series marked the ends of Triassic and Jurassic synchronous with gentle tectonic activity in the region (Alsharhan and Nairn, 1997; Guiraud, 1998). Alkaline dykes and plugs were emplaced in the Early Cretaceous (140 Ma) on the western side of the Gulf of Suez, parallel to G. Hammam Faroun in Sinai, associated with the initial rifting of the South Atlantic and the corresponding Africa–South America compression (Meneisy, 1986). A strong compressional phase pulsed in Late Cretaceous (Rosenbaum et al., 2005) accompanying the N–S-directed convergence of Africa– Arabia with Eurasia in a counterclockwise rotational mode (Guiraud and Bosworth, 1999). These pulses resulted in transpressional deformation and development of the ‘Syrian Arc’ structures in northern Sinai along the Mediterranean margin (e.g., Guiraud et al., 2005; Moustafa and Khalil, 1990).

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5.4. Cenozoic trends NW trends continued to dominate over NE trends in the Cenozoic. Prominence of the N–S trend recommenced in the Cretaceous then gradually decreased in abundance to the Holocene, with the exception of a marked prominence in Eocene lineaments. The 10° rotation was apparent and common on all Cenozoic trends. Reverse faulting became secondary in abundance after strike-slip, with normal faulting being the most rare. Strike-slip styles were common along NW, N–S, NNE, and ENE trends. Reverse faulting characterized the E–W trend, whereas strike-slip and normal faulting were often mixed and dominated the NNW trends. These trends are associated with the convergence that has continued since the Paleocene at lower rates. The Paleocene trends resemble those of the disconformably underlying Cretaceous and the overlying Eocene, because there were no major changes in the direction of the continued motion of Africa relative to Europe during these times (Guiraud and Bosworth, 1997). The only marked difference is seen in the prominence of the ENE trend and the more than 10° counterclockwise rotation of the NW trend, likely related to this motion. The recommencement of convergence during the Late Eocene (Guiraud and Bosworth, 1999; Rosenbaum et al., 2005), caused northern Sinai and the ‘Syrian Arc’ structures to be reactivated, again in transpression (Guiraud et al., 2001, 2005; Moustafa and Khalil, 1994). Also, many E–W-striking extensional faults were reactivated as reverse faults and local thrusts (e.g., at G. Maghara, Halal, Yeleg) (Aal et al., 1992; El Toukhy et al., 1998; Guiraud and Bellion, 1995). Strike-slip movement accompanied by drag folds occurred along E–W-trending, dextral transcurrent fault zones in Central Sinai (Themed fault zone; Moustafa and Khalil, 1994). The major shortening direction in the Late Eocene was oriented nearly NW–SE. The precise age of this event can be deduced from very frequent and well-dated angular unconformities (Said, 1990) that correspond to the Africa–Arabia collision with the Eurasian plate. The continued counterclockwise rotation of Arabia relative to Africa in the Early Miocene (25–21 Ma) led to opening of the Red Sea and its northern extension, the Gulf of Suez (Bosworth and McClay, 2001). The main phase of rift shoulder uplift began then (e.g., Omar and Steckler, 1995) and continued to the Early Mid-Miocene associated with faulting. The geometry of this phase was controlled mainly by the reactivation of Precambrian basement fabrics (Khalil and McClay, 2001). Rift-synchronized extensional NW-oriented basaltic dykes from the lithospheric mantle diffusely intruded (20 Ma) parallel to the long axis of the rift (Baldridge et al., 1991; Eyal et al., 1981; Ibrahim et al., 2000). These dykes extend for more that 100 km (See Fig. 8) in length and 100 m in width crossing Sinai from W. Kid near the Gulf of Aqaba coast to Abu Zeniema on the Gulf of Suez and cutting across the granite ridge of G. Saint Catherine (Ibrahim et al., 2000). Paleomagnetic studies showed that Sinai was at about 7° south of its present position at 20 Ma at the time of intrusion of these dykes (Ibrahim et al., 2000). Burke (1996) suggested that the onset of the Oligocene–Miocene rifting, with its extraordinary volcanic signature, was triggered by the cessation of movement of the African plate with respect to the underlying asthenosphere. This facilitated penetration of the lithospheric mantle and crust by plumes, and thus changed the character of extension within the crust. Arabia and Anatolia collision-related rift coalescence and failing in the Gulf of Suez initiated the Gulf of Aqaba rifting, ongoing since the Middle Miocene (e.g., Bosworth and McClay, 2001). This was underpinned with the major change in the direction of motion of Africa relative to Eurasia from NNE in the Miocene to NNW since the late Miocene to present (Guiraud and Bosworth, 1997; Mazzoli and Helman, 1994). This change enhanced the prominence of the NNW and the ENE trends relative to the Eocene trends (Fig. 7). Miocene extension accommodated 30 km at 0.2 cm/yr in the southern part of the Gulf of Suez and 62 km at 0.7 cm/yr in the form of sinistral movement along the Dead Sea transform fault (Bosworth et al., 2005,

and references therein). The convergence significantly decreased and has remained low from the late Miocene (19 Ma) to the present (Rosenbaum et al., 2005). The low rate of convergence since the late Miocene strongly supports the importance of sinistral movements along the Dead Sea transform fault that offset 62 km in the Miocene (Quennell, 1958). During the Pliocene, synchronous with the change in the direction of motion of Africa relative to Europe, extension continued but at a lower rate (0.1 cm/yr) for the last 5 km in the Gulf of Suez and higher rate at 0.9 cm/yr for the 45 km sinistral movements along the Dead Sea transform fault zone (Bosworth et al., 2005, and references therein). This northward motion, accompanied by counterclockwise rotation, strongly enhanced the prominence of the NNE and the NW trends, perpendicular to the NE trend, whereas the N–S and E–W trends showed less prominence compared with those that preceded in the Miocene. The continued rotational motion resulted in local NE- to NNEoriented Late Pleistocene upwarping (Horowitz, 1975). In turn, this led to the diversion of the flow direction of the old W. Al-Arish upstream into the Dead Sea (Garfunkel and Horowitz, 1966; Ginat et al., 2002) (e.g., Wadi Girafi of Pliocene exposures) and also the change of its downstream flow direction from westward and northward in the Pliocene to the present northeastward flow (Kusky and El-Baz, 2000). These tectonic activities enhanced some of the Pliocene and Pleistocene trends compared with those that preceded them. The prominence of the NW/NE trends of the Pliocene and Pleistocene, Pliocene NNE trend, and the Pleistocene NNW trend (see Fig. 7) was relatively enhanced. The effect of this motion on the E–W trends is subtle, whereas the N–S trend gradually faded from their reappearance in the Cretaceous to the Holocene. The NW and the E–W trends dominated in the Holocene with subordinate contributions from the NNE and NNW trends. These trends favor a compressional regime with a horizontal field minimum stress oriented roughly N–S with a mixture of normal and sinistral strike-slip faulting predominant in both the Gulf of Suez rift area (Bosworth and Taviani, 1996) and the alluvium deposits along the western margin of the Gulf of Aqaba (Guiraud and Bosworth, 1997). Active Quaternary fault zones include: (1) the NW to NNW fault swarms delimiting tilted blocks and uplifted shoulders along the Gulf of Suez, sometimes very active since the Late Pleistocene (Bosworth and Taviani, 1996), and (2) the transtensive NNE Gulf of Aqaba Dead Sea fault zone, and associated NE or NW transfer or linking faults. The same conclusion was drawn from the analysis of recent (1900–2006) seismic activities in that the foci of the strongest epicenters (M N 5) are concentrated along the structural segments of the Gulf of Aqaba and the Gulf of Suez, in particular, at the intersections of these segments with the intra-peninsula broad structural zones. The low and dispersed intra-plate seismic activities showed marked foci along and at the intersections of the detected structural zones. The disappearance of the NNW and the ENE trends that dominated in the Pleistocene is likely related to the formation of the sand dunes that are now dominant in the area between the Bardawil and Great Bitter Lakes. 5.5. Temporal change of stress fields The detected structural features and their trends correspond well with those derived from the reference data. This correspondence can be supported further by the intimate association of these features with regional deep-crustal shear zones derived from gravity data (Fig. 12). Recent seismic activity has been focused along these shear zones and at their intersections, within and at the boundaries of the plate. Most of the events that played key roles in developing the configuration of NE Africa and the Sinai are associated with worldwide plate reorganizations (Bunge et al., 2002; Scotese et al.,

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1988). The principal extensional events affecting the region occurred during (1) the Late Proterozoic from c. 600 Ma (Be'eri-Shlevin et al., 2009; Blasband et al., 2000), (2) from the end of the Paleozoic to the Late Cretaceous, and (3) in the Oligocene–Miocene (Guiraud and Bosworth, 1999). The strongest sharp compressive events affecting broad areas of Africa – including the Sinai – occurred during the Precambrian (pre-600 Ma), Late Cretaceous, and Late Eocene, and demonstrate the role of far field stresses in altering intra-plate tectonic stability (Blasband et al., 2000; Guiraud and Bellion, 1995; Guiraud and Bosworth, 1997). An older late-Precambrian compressional phase was followed by extension from c. 600 Ma (e.g., Be'eriShlevin et al., 2009; Blasband et al., 2000). Extensional tectonics synchronized with rifting characterizes the Paleozoic to the late Cretaceous. Strong compression initiated in the Late Cretaceous with convergence between Africa and Eurasia and recommenced in the Late Eocene (e.g., Rosenbaum et al., 2005). Convergence has been accommodated by subduction processes, continental collision, and lithospheric deformation (Dewey et al., 1973, 1989) and has continued to the present at variable rates as indicated by the alternation between short compressive or transpressive events and longer mild or even distensive periods (Guiraud and Bellion, 1995; Ziegler, 1990). The convergence slowed due to the reversed direction of motion in the Campanian (c. 74 Ma) and from the Late Miocene to present (c. 10–0 Ma) and accelerated in the Eocene-Oligocene (c. 43– 25 Ma) with the faster movement of Africa towards Eurasia as inferred from the mantle-circulation models of Bunge et al. (2002). The Miocene was locally extensional, resulting principally from interplate compression. This compression was the result of rotation and localized intra-plate uplift caused by mantle thermal perturbations as evidenced by the emplacement of Miocene dykes (Eyal et al., 1981). Temporal changes in the direction of previously reviewed stress fields and plate motions (e.g., Bunge et al., 2002; Guiraud and Bosworth, 1997) are shown on Fig. 7. Evolutionary N–S directed compression dominated most of the Precambrian followed by extension at its end (Be'eri-Shlevin et al., 2009) that extended to the Late Cretaceous with counterclockwise rotation. This rotation accompanied the Late Cretaceous northward convergence of the African plate with Eurasia to the present with variable rates owing to changes in the motion directions (see Fig. 7) of the two plates (Rosenbaum et al., 2005).

305

In the northern province, the style of deformation becomes complex where large uplifted blocks of Mesozoic domes and asymmetrical anticlines oriented N65°–85°E, as part of the Syrian Arc System, dictate the distribution and character of the Quaternary alluvial plains. The southerly limbs of the anticlines are markedly steeper and crossed by normal and reverse faults that commonly trend at large angles and less commonly are aligned parallel to the fold axis (Moustafa and Khalil, 1995). The belt was formed in response to successive compressional Cretaceous episodes (Bartov et al., 1980c; Jenkins, 1990) and well developed by the Tertiary (Guiraud and Bosworth, 1999, and references therein). According to microtectonic analyses (Eyal and Reches, 1983; Letouzey and Tremolieres, 1980), the shortening during the development of the Cretaceous/Paleocene structures resulted from E–W to WNW–ESE horizontal compression where this was generally NNW–SSE directed in the Western Desert, shifting progressively to NW in Sinai and nearly E–W in neighboring regions to the east (Sehim, 1993). This eastward increase in the shortening along the Tethyan margins (Guiraud and Bosworth, 1997) is synchronous with counterclockwise rotational northward drift of the African–Arabian plate and its increased collisional coupling with the Eurasian plate (Le Pichon et al., 1988; Ziegler, 1990). The Gulf of Suez rift province extends from the Bitter Lakes southwards to Ras Mohammed. This province is characterized by the wide-spread presence of NW-oriented normal faults of varying lengths and displacements in the Phanerozoic cover in the north (El Shazly et al., 1974). Strike-slip faulting prevails in the basement with subordinate normal and reverse faulting. The Gulf of Aqaba province exhibits the basement stretching from Ras Mohamed to Taba and the unconformably overlying Paleozoic and Cretaceous sediments in the north with very limited Holocene coastal deposits. Focal mechanism solutions and maximum shear analysis show that extensional strains and normal faulting are predominant in the Gulf of Suez province and compressional strains with left lateral strike slip styles prevail in the Gulf of Aqaba province (Badawy and Horváth, 1999a, 1999b, 1999c; El-Fiky, 2005; Mahmoud, 2003). The conspicuously dispersed occurrence of the Cretaceous in the basement of W. Dahab and Eocene deposits juxtaposing the Cretaceous close to the basement in W. Watir strongly support ~ 45 km of offset from their northern outcrops through strike-slip movements mostly related to broad fracture zones parallel to the sinistral Dead Sea transform fault.

6. Tectonic provinces of the Sinai 7. Conclusions The tectonic trends and the styles of faulting reveal four main structural provinces located in the central and northern Sinai, and the gulfs of Suez and Aqaba. In the central province, the most prominent structural feature is the Ragabet El-Naam-Themed E–W-oriented wrench fault cutting through the peninsula from the Gulf of Suez to its eastern border with a variable displacement up to 2.5 km (Abd El Samie and Sadek, 2001, and references therein; Corchete et al., 2007). This shear zone was detected here as two segments (Fig. 8), the western Ragabet El-Naam and the eastern Themed, with the latter showing a minimum displacement of 300–750 m (Moustafa and Khalil, 1994, and references therein). Southward downthrows differentiate the Ragabet El-Naam segment from the Themed segment where downthrows are to the north (Abd El Samie and Sadek, 2001, and references therein). This fault marks the southernmost border of the Early Mesozoic passive continental margin and the front of the Syrian Arc structures in the north. This has been rejuvenated in association with the continued closure of the Eastern Mediterranean basin that proceeded by pure wrenching in the post-Middle Eocene to pre-Early Miocene (Moustafa and Khalil, 1994). Also, sub-horizontal Mesozoic and Tertiary sediments, forming the plateaux of El Tih and Egma, are affected by uplift-related normal faulting (unfolded central Sinai stable foreland) distinguished into N–S to NNE–SSW, NW–SE and E–W trends.

The methods employed in this study provided a complete framework for characterizing the tectonics and geodynamics of the Sinai Peninsula. Spatial and statistical analyses of lineaments extracted from SRTM 90-m data successfully identified the prominent tectonic trends and the relative strength of dominant tectonic regimes. Prominent trends and faulting styles provided important clues for the timing of their development as well as a structural inheritance model. The pattern of the lineaments on the surface represents not only the overprints of the deep tectonic structures, but also the results of the interaction between inter- and intra-plate forces dominant throughout its geodynamic evolution. Phanerozoic intraplate deformation and the related processes of erosion and sedimentation were generally controlled by structural trends that were frequently reactivated along existing fault systems. The grouping technique enhanced the line segment connectivity and hence the significance of the tectonic trends. Fault type and geometry modeling revealed the styles of faulting that dominated either individually or in combination along major trend orientations. Tectonic trends auto-detected from the DEM were congruent with the prominent structural trends identified in the BARTOV80 and CONOCO87 data. The detected lineaments can further update the structures of the reference data.

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Fig. 12. Gravity-derived segments with rose diagram and earthquake epicenters (M N 2) overlying the gravity anomaly map. Note the foci of epicenters where the two gravity anomaly zones (dotted in pink) meet and intersect with the Gulf of Aqaba.

Noteworthy linear features were used to constrain the relative severity of five periods with a new interpretation of the dominant tectonic regimes: (1) Precambrian compression followed by extensional at its end, (2) compressional Cretaceous, (3) compressional Eocene, (4) extensional Miocene, and (5) compressional Holocene. The severity of these tectonic periods controlled the structural make up that accounts for the present configuration of the Sinai. Furthermore, the recent significance of the broad structural zones identified from the linear features was confirmed by the locations of earthquake epicenters along and at the intersection of the sub-vertical to vertical topographic and gravity anomaly zones in the gulfs of Suez and Aqaba. There is a high probability that the detected lineaments represent crustal discontinuities, which can be evidenced by the gravity-derived shear zones in the basement rocks and the overlying sedimentary formations. These structures could have pre-existed the basement rock, been initiated and/or reactivated later, or be newly developed. These tectonically-controlled lineaments may be favorable conduits for the accumulation of natural resources such as water, hydrocar-

bons, geothermal energy and minerals. Their accurate mapping, therefore, integrating modern digital technologies combined with the reference geological and structural data can improve the exploration process, in terms of cost, accuracy, and time.

References Aal, A.A., Day, R.A., Lelek, J.J., 1992. Structural evolution and styles of the northern Sinai, Egypt. 11th EGPC Exploration Seminar, Egypt, 1, pp. 546–563. Abd El Samie, S.G., Sadek, M.A., 2001. Groundwater recharge and flow in the Lower Cretaceous Nubian Sandstone aquifer in the Sinai Peninsula, using isotopic techniques and hydrochemistry. Hydrogeol. J. 9 (4), 378–389. Abdel-Gawad, M., 1969. New evidence of transcurrent movements in the Red Sea areas and petroleum implications. Am. Assoc. Pet. Geol. Bull. 53, 1466–1479. Alsharhan, A.S., Nairn, A.E.M., 1997. Sedimentary Basins and Petroleum Geology of the Middle East, 1st Edn. Elsevier Science, UK, p. 843. Atlas of Israel, 1985. Cartography, physical and human geography. Survey of Israel. Macmillan Publishing. Badawy, A., 2005. Present-day, seismicity, stress field and crustal deformation of Egypt. J. Seismol. 9, 267–276. Badawy, A., Horváth, F., 1999a. Recent stress field of the Sinai subplate region. Tectonophysics 304, 385–403.

A.A. Masoud, K. Koike / Tectonophysics 510 (2011) 291–308 Badawy, A., Horváth, F., 1999b. Seismicity of the Sinai sub-plate region: kinematic implications. J. Geodyn. 27, 451–468. Badawy, A., Horváth, F., 1999c. The Sinai sub-plate and tectonic evolution of the northern Red Sea region. Geodynamics 27, 433–450. Badawy, A., Mohamed, A.M.S., Abu-Ali, N., 2008. Seismological and GPS constraints on Sinai sub-plate motion along the Suez rift. Stud. Geophys. Geod. 52 (3), 397–412. Baldridge, W.S., Eyal, Y., Bartov, Y., Steinitz, G., Eyal, M., 1991. Miocene magmatism of Sinai related to the opening of the Red Sea. Tectonophysics 197, 181–201. Barbero, L., Jabaloy, A., Gómez-Ortiz, D., Pérez-Peña, J.V., Rodríguez-Peces, M.J., Tejero, R., Estupiñán, J., Azdimousa, A., Vázquez, M., Asebriy, L., 2010. Evidence for surface uplift of the Atlas Mountains and the surrounding peripheral plateaux: combining apatite fission-track results and geomorphic indicators in the Western Moroccan Meseta (coastal Variscan Paleozoic basement). Tectonophysics. doi:10.1016/j. tecto.2010.01.005. Bartov, Y., Eyal, M., Shimron, A.E., Bentor, Y.K., 1980a. Sinai—Geological photomap, scale 1:500,000., Survey of Israel. Bartov, Y., Levy, Z., Steinitz, G., Zak, 1980b. Mesozoic and Tertiary stratigraphy, paleogeography and structural history of the Gebel Areif en Naqa area, eastern Sinai 1. Isr. J. Earth Sci. 29, 114–139. Bartov, Y., Steinitz, G., Eyal, M., Eyal, Y., 1980c. Sinistral movement along the Gulf of Aqaba—its age and relation to the opening of the Red Sea. Nature 285, 220–222. Be'eri-Shlevin, Y., Katzir, Y., Whitehouse, M., 2009. Post-collisional tectonomagmatic evolution in the northern Arabian–Nubian Shield: time constraints from ion-probe U–Pb dating of zircon. J. Geol. Soc. 166, 71–85. Ben Avraham, Z., Tibor, G., Limonov, A.F., Leybov, M.B., Ivanov, M.K., Tokarev, M.Y., Woodside, J.M., 1995. Structure and tectonics of the eastern Cyprean arc. Mar. Pet. Geol. 12, 263–271. Ben-Menahem, A., Amos, N., Moshe, V., 1976. Tectonics, seismicity and structure of the Afro-Eurasian Junction: the breaking of an incoherent plate. Phys. Earth Planet. Inter. 12 (1), 1–50. Bentor, Y.K., 1985. The crustal evolution of the Arabian–Nubian Massif with special reference to the Sinai Peninsula. Precambrian Res. 28, 1–74. Blasband, B., White, S., Brooijmans, P., De Boorder, H., Visser, W., 2000. Late Proterozoic extensional collapse in the Arabian–Nubian Shield. J. Geol. Soc. Lond. 157, 615–628. Bosworth, W., Taviani, M., 1996. Late quaternary reorientation of stress field and extension direction in the southern Gulf of Suez, Egypt: evidence from uplifted coral terraces, mesoscopic fault arrays, and borehole breakouts. Tectonics 15 (4), 791–802. Bosworth, W., McClay, K.R., 2001. Structural and stratigraphic evolution of the Gulf of Suez rift, Egypt: A synthesis. In: Zeigler, P.A., Cavazza, W., Robertson, A.H.F.R., & Crasquin-Soleau, S. (Eds.), Peri-Tethys Memoir 6: ‘Peritethyan Rift/Wrench Basins and Passive Margins’. Memoires du Museum National d’Historie Naturelle de Paris: v. 186, pp. 567–606. Bosworth, W., Huchon, P., McClay, K., 2005. The Red Sea and Gulf of Aden Basins. J Afr. Earth Sci. 43 (1–3), 334–378. Bumby, A.J., Guiraud, R., 2005. The geodynamic setting of the Phanerozoic basins of Africa. J. Afr. Earth Sci. 43, 1–12. Bunge, H.-P., Richards, M.A., Baumgardner, J.R., 2002. Mantle circulation models with sequential data assimilation: inferring present-day mantle structure from platemotion histories. Philos. Trans. R. Soc. Ser. A 360, 2545–2567. Burbank, D.W., Anderson, R.S., 2001. Tectonic Geomorphology. Blackwell Sciences, Inc, Malden. 274 pp. Burke, K., 1996. The African plate. S Afr. J. Geol. 99 (4), 339–409. Cloetingh, S., Ziegler, P.A., Cornu, T., ENTEC Working Group, 2003. Investigating environmental tectonics in the Northern Alpine Foreland of Europe. EOS Trans. AGU 84 (349), 356–357. Cloetingh, S., Cornu, T., Ziegler, P.A., Beekman, F., Environmental Tectonics (ENTEC) Working Group, 2006. Neotectonics and intraplate continental topography of the northern Alpine Foreland. Earth Sci. Rev. 74 (3–4), 127–196. Cloetingh, Thybo, H., Faccenna, C., 2009. TOPO-EUROPE: studying continental topography and Deep Earth—surface processes in 4D. Tectonophysics 474 (1–2), 4–32. Cochran, J.R., 2005. Northern Red Sea: Nucleation of an oceanic spreading center within a continental rift. Geochemistry Geophysics Geosystems 6, Q03006. doi:10.1029/ 2004GC000826. Cochran, J.R., Karner, G.D., 2007. Constraints on the deformation and rupturing of continental lithosphere of the Red Sea: the transition from rifting to drifting. Geol. Soc. Lond. Spec. Publ. 282, 265–289. Corchete, V., Chourak, M., Hussein, H.M., 2007. Shear wave velocity structure of the Sinai Peninsula from Rayleigh wave analysis. Surv. Geophys. 28, 299–324. Dewey, J.F., Pitman, W.C., Ryan, W.B.F., Bonnin, J., 1973. Plate Tectonics and the evolution of the Alpine system. Geol. Soc. Am. Bull. 84, 3137–3180. Dewey, J.F., Helman, M.L., Turco, E., Hutton, D.H.W., Knott, S.D., 1989. Kinematics of the western Mediterranean. In: Coward, M.P., Dietrich, D., Park, R.G. (Eds.), Alpine Tectonics: Geological Society, London, Special Publication, 45, pp. 265–283. EGSMA. 1995. Geological Map of the Sinai (Five sheets), Arab Republic of Egypt. 1:250, 000. Egyptian Geological Survey and Mining Authority. El Shazly, E.M., Abdel Hady, M.A., El Kassas, I.A., El Shazly, M.M., 1974. Geology of Sinai Peninsula from ERTS-2 satellite images. Rem. Sens. Center, Cairo. 10 pp. El Toukhy, M., Charmy, H., Studer, M., 1998. Structural evolution of the eastern part of the western Desert and its implications for hydrocarbon exploration. Proceedings of 14th Petroleum Conference October Exploration Egyptian General Petroleum Corporation, Cairo, 1, pp. 156–177. El-Fiky, G., 2005. GPS-derived velocity and crustal strain field in the Suez-Sinai Area, Egypt. Bull. Earthquake Res. Inst. Univ. Tokyo 80, 73–86. El-sayed, M., 2006. Geochemistry and petrogenesis of the post-orogenic bimodal dyke swarms in NW Sinai, Egypt: constraints on the magmatic-tectonic processes during the late Precambrian. Chem. Erde 66 (2), 129–141.

307

El-Shafei, M., Kusky, T.M., 2003. Structural and tectonic evolution of the Neoproterozoic Feiran–Solaf metamorphic belt, Sinai Peninsula: implication for the closure of Mozambique Ocean. Precambrian Res. 123, 269–293. Eyal, Y., Reches, Z., 1983. Tectonic analysis of the Dead Sea Rift region since the Late Cretaceous based on mesostructures. Tectonics 2, 167–185. Eyal, M., Eyal, Y., Bartov, Y., Steinitz, G., 1981. The tectonic development of the western margin of the Gulf of Elat (Aqaba) Rift. Tectonophysics 80, 39–66. Fabre, J., 1988. Les séries Paléozoïques d'Afrique: une approche. J. Afr. Earth Sci. 7 (1), 1–40. Fowler, A., Hassan, I., 2008. Extensional tectonic origin of gneissosity and related structures of the Feiran–Solaf metamorphic belt, Sinai, Egypt. Precambrian Res. 164 (3–4), 119–136. Freund, D., Garfunkel, Z., Zak, I., Goldberg, M., Weissbrod, T., Berin, B., 1970. The shear along the Dead Sea rift. Philos. Trans. R. Soc. Lond Ser. A267, 107–130. Ganas, A., Pavlides, S., Karastathis, V., 2005. DEM-based morphometry of range-front escarpments in Attica, central Greece, and its relation to fault slip rates. Geomorphology 65 (3–4), 301–319. Garfunkel, Z., Bartov, Y., 1977. The tectonics of the Suez rift. Isr. Geol. Surv. Bull. 71, 1–44. Garfunkel, Z., Derin, B., 1988. Reevaluation of the Latest Jurassic–Early Cretaceous history of the Negev and the role of magmatic activity. Isr. J. Earth Sci. 37, 43–52. Garfunkel, Z., Horowitz, A., 1966. The Upper Tertiary and Quaternary morphology of the Negev, Israel. Isr. J. EarthSci. 15, 101–117. Garfunkel, Z., Zak, I., Freund, R., 1981. Active faulting in the Dead Sea rift. Tectonophysics 80, 1–26. Ginat, H., Zilberman, E., Amit, R., 2002. Red sedimentary units as indicators of early Pleistocene tectonic activity in the southern Negev Desert, Israel. Geomorphology 45, 127–146. Gorshkov, A.I., Kuznetsov, I.V., Panza, G.F., Sololiev, A.A., 2000. Identification of future earthquake sources in the Carpatho-Balkan orogenic belt using morphostructural criteria. Pure Appl. Geophys. 157, 79–95. GRASS Development Team, 2010. Geographic Resources Analysis Support System (GRASS) Software, Version 6.4.0. Open Source Geospatial Foundation. http://grass. osgeo.org. Greiling, R.O., Abdeen, M.M., Dardir, A.A., El Akhal, H., El-Ramly, M.F., Kamal El-Din, G.M., Osman, A.F., Rashwan, A.A., Rice, A.H.N., Sadek, M.F., 1994. A structural synthesis of the Proterozoic Arabian–Nubian Shield in Egypt. Geol. Rundsch. 83, 484–501. Guiraud, R., 1998. Mesozoic rifting and basin inversion along the northern African Tethyan margin: an overview. In: Macgregor, D.S., Moody, R.T., Clarck-Lowes, D.D. (Eds.), Petroleum Geology of North Africa: Geological Society, London, Special Publication, 132, pp. 217–229. Guiraud, R., Bellion, Y., 1995. Late carboniferous to recent geodynamic evolution of the West Gondwanian cratonic Tethyan margins. In: Nairn, A., Dercourt, J., Vrielynck, B. (Eds.), The Ocean Basins and Margins, 8: The Tethys Ocean. Plenum Press, New York, NY, pp. 101–124. Guiraud, R., Bosworth, W., 1997. Senonian basin inversion and rejuvenation of rifting in Africa and Arabia: synthesis and implications to plate-scale tectonics. Tectonophysics 282, 39–82. Guiraud, R., Bosworth, W., 1999. Phanerozoic geodynamic evolution of northeastern Africa and the northwestern Arabian platform. Tectonophysics 315, 73–108. Guiraud, R., Bosworth, W., Thierry, J., Delplanque, A., 2005. Phanerozoic geological evolution of Northern and Central Africa: an overview. J. Afr. Earth Sci. 43 (1–3), 83–143. Gurnis, M., Mitrovica, J.X., Ritsema, J., van Heijst, H.J., 2000. Constraining mantle structure using geological evidence of surface uplift rates: the case of the African Superplume. Geochem. Geophys. Geosyst. 1 (7). dx.doi.org/10.1029/ 1999GC000035. Hetzel, R., Niedermann, S., Ivy-Ochs, S., Kubik, P.W., Tao, M.X., Gao, B., 2002. Ne-21 versus Be-10 and Al-26 exposure ages of fluvial terraces: the influence of crustal Ne in quartz. Earth Planet. Sci. Lett. 201, 575–591. Hooper, D.M., Bursik, M.I., Webb, F.H., 2003. Application of high-resolution, interferrometric DEMs to geomorphic studies of fault scarps, Fish Lake Valley, NevadaCalifornia, USA. Remote. Sens. Environ. 84 (2), 255–267. Horowitz, A., 1975. The Quaternary Stratigraphy and Paleogeography of Israel. Paléorient 3, 47–100. Horowitz, A., 1987. Palynological evidence for the age and the rate of sedimentation along the Dead Sea Rift, and structural implications. Tectonophysics 141, 107–115. Ibrahim, E.H., Odah, H.H., El Agami, N.L., Abu El Enen, M., 2000. Paleomagnetic and geological investigation into southern Sinai volcanic rock and the rifting of the Gulf of Suez. Tectonophysics 321, 343–358. Issawi, B., 1996. Tectono-sedimentary synthesis of the Paleozoic basins in Egypt. In: Proceeding of 13th Exploration Conference, Egyptian General Petroleum Corporation, Cairo, v. 1, pp. 1–23. Jenkins, D.A., 1990. North and Central Sinai. In: Said, R. (Ed.), Geology of Egypt. A.A Balkema Publisher, Netherlands, pp. 361–389. Joffe, S., Garfunkel, K., 1987. Plate kinematics of the circum Red Sea, a re-evaluation. Tectonophysics 141, 5–22. Jordan, G., 2003a. Application of digital terrain modeling and GIS methods for the morphotectonic investigation of the Kali Basin, Hungary. Z. Geomorphol. 47 (2), 145–169. Jordan, G., 2003b. Morphometric analysis and tectonic interpretation of digital terrain data: a case study. Earth Surf. Processes Landforms 28 (8), 807–822. Jordan, G., Meijninger, B.M.L., van Hinsbergen, D.J.J., Meulenkamp, J.E., van Dijk, P.M., 2005. Extraction of. morphotectonic features from DEMs: development and applications for study areas in Hungary and NW. Greece. Int. J. Appl. Earth Obs. Geoinf. 7, 163–182.

308

A.A. Masoud, K. Koike / Tectonophysics 510 (2011) 291–308

Kebeasy, R., 1990. Seismicity. In: Said, R. (Ed.), The Geology of Egypt. Balkema, Rotterdam, pp. 51–59. Khalil, S., McClay, N., 2001. Tectonic evolution of the Northwestern Red Sea — Gulf of Suez Rift system. In: Wilson, R.C.L., Whitmarsh, R.B., Taylor, B., Froitzheim (Eds.), Non-Volcanic Rifting of Continental Margins: A Comparison of Evidence from Land and Sea: Geological Society of London, Special Publication, 187, pp. 453–473. Klitzsch, E., List, F., PÖhlmann, G., 1987. Geological Map of Egypt, scale 1:500,000, Sheets NG 36 NE Quseir; NH 36 NW Cairo; NH 36 NE North Sinai; NH 36 SW Beni Suef; and NH 36 SE South Sinai, Conoco Coral and Egyptian General Petroleum Corporation (EGPC), Cairo. Koike, K., Nagano, S., Ohmi, M., 1995. Lineament analysis of satellite images using a Segment Tracing Algorithm (STA). Comput. Geosci. 21, 1091–1104. Koike, K., Nagano, S., Kawaba, K., 1998. Construction and analysis of interpreted fracture planes through combination of satellite-image derived lineaments and digital elevation model data. Comput. Geosci. 24, 573–583. Krishnamurthy, J., Mani, A., Jayaraman, V., Manivel, M., 2000. Groundwater resources development in hard rock terrain — an approach using remote sensing and GIS techniques. Int. J Appl. Earth Obs. Geoinf. 2 (3–4), 204–215. Kusky, T., El-Baz, F., 1998. Structural and tectonic evolution of the Sinai Peninsula, using Landsat data: Implications for ground water exploration. Egypt. J Remote Sens. Space Sci. 1, 69–100. Kusky, T.M., El-Baz, F., 2000. Neotectonics and fluvial geomorphology of the northern Sinai Peninsula. J. Afr. Earth Sci. 31 (2), 213–235. Lattman, L.H., Parizek, R.R., 1964. Relationship between fracture traces and the occurrence of ground water in carbonate rocks. J. Hydrol. 2 (2), 73–96. Le Pichon, X., Francheteau, J., 1987. A plate tectonic analysis of the Red Sea Gulf of Aden area. Tectonophysics 46, 369–406. Le Pichon, X., Bergerat, E., Roulet, M.-J., 1988. Plate kinematics and tectonics leading to the Alpine belt formation; a new analysis. Geol. Soc. Am. Spec. Pap. 218, 111–131. Letouzey, J., Tremolieres, P., 1980. Paleo-Stress Fields around the Mediterranean since the Mesozoic from Microtectonics. Comparison with Plate Tectonic Data. Rock Mech. Supply 9, 173–192. Lithgow-Bertelloni, C., Silver, P.G., 1998. Dynamic topography, plate driving forces and the African superswell. Nature 395, 269–272. Magowe, M., Carr, J., 1999. Relationship between lineaments and ground water occurrence in Western Botswana. Ground Water 37, 282–286. Mahmoud, S., 2003. Seismicity and GPS-derived crustal deformation in Egypt. Geodynamics 35, 333–352. Mahmoud, S., Reilinger, R., McClusky, S., Vernant, P., Tealeb, A., 2005. GPS evidence for northward motion of the Sinai block: implications for E. Mediterranean tectonics. Earth Planet. Sci. Lett. 238, 217–227. Mart, Y., 1991. The Dead Sea Rift: from continental rift to incipient ocean. Tectonophysics 197, 155–179. Mart, Y., Rabinowitz, P.D., 1986. The northern Red Sea and the Dead Sea Rift. Tectonophysics 124, 85–113. Mascle, J., Benkhelil, J., Bellaiche, G., Zitter, T., Woodside, J., Loncke, L., 2000. Prismed II Scientific Party, marine geologic evidence for a Levantine–Sinai plate, a new piece of the Mediterranean puzzle. Geology 28, 779–782. Masoud, A., Koike, K., 2009. Use of Bézier and B-spline curves for smart representation of lineaments auto-detected from DEM. Proceedings of the International Association of Mathematical Geology (IAMG09), August 23–28, 2009, Stanford University, CA, USA. Masoud, A.A., Koike, K., 2011. Auto-detection and Integration of Tectonically Significant Lineaments from SRTM DEM and Remotely-sensed Geophysical Data. ISPRS Journal of Photogrammetry and Remote Sensing, doi:10.1016/j.isprsjprs.2011.08.003. Mazzoli, S., Helman, M., 1994. Neogene patterns of relative plate motions for Africa– Europe: some implications for recent central Mediterranean tectonics. Geol. Rundsch. 83, 464–468. McClay, K., Khalil, S., 1998. Extensional hard linkages, eastern Gulf of Suez, Egypt. Geology 26 (6), 563–566. McKenzie, D.P., 1970. Plate tectonics of the Mediterranean region. Nature 226, 239–243. McKenzie, D.P., 1977. Surface deformation, gravity anomalies and convection. GJRAS 48, 211–238. McKenzie, D.P., Davies, D., Molnar, P., 1970. Plate tectonics of the Red Sea and East Africa. Nature 226, 243–248. Meert, J.G., 2003. A synopsis of events related to the assembly of eastern Gondwana. Tectonophysics 362, 1–40. Meneisy, M.Y., 1986. Mesozoic igneous activity in Egypt. Qatar Univ. Sci. Bull. 6, 317–328. Meshref, W.M., 1990. Tectonic framework. In: Said, R. (Ed.), The Geology of Egypt, Rotterdam, Netherlands. A. A. Balkema Publishers, pp. 113–156.Robertson, A.H.F., 1998. Mesozoic-tertiary tectonic evolution of the easternmost Mediterranean area: integration of marine & land evidence. Proceedings of the ODP, V. 160, pp. 723–782. Mitrovica, J.X., Beaumont, C., Jarvis, G.T., 1989. Tilting of continental interiors by the dynamical effects of subduction. Tectonics 8, 1079–1094. Moghazi, A.M., Anderson, T., Oweiss, G.A., El Bouselly, A.M., 1998. Geochemical and Sr– Nd–Pb isotopic data bearing on the origin of Pan-African granitoids in the Kid area, southeast Sinai, Egypt. J. Geol. Soc. Lond. 155, 697–710. Moustafa, A.R., Khalil, M.H., 1990. Structural characteristics and tectonic evolution of north Sinai fold belts. In: Said, R. (Ed.), The Geology of Egypt. A. A. Balkema, Rotterdam, pp. 381–389.

Moustafa, A.R., Khalil, M.H., 1994. Rejuvenation of the eastern Mediterranean passive continental margin in northern and central Sinai: new data from the Themed Fault. Geol. Mag. 131, 435–448. Moustafa, A.R., Khalil, M.H., 1995. Superposed deformation in the Northern Suez rift, Egypt: relevance to hydrocarbons exploration. J. Pet. Geol. 18 (3), 245–266. NEIC: National Earthquake Information Center, U. S. Geological Survey, http://neic.usgs. gov. Nikishin, A.M., Ziegler, P.A., Stephenson, R.A., Ustinova, M.A., 1999. Santonian to Paleocene tectonics of the East-European Craton and adjacent areas. Bull. Inst. Royal Sci. Nat. Belg. Sci. Terre 69, 147–159 Suppl. A. Omar, G.I., Steckler, M.S., 1995. Fission-track evidence on the initial opening of the Red Sea: two pulses, no propagation. Science 270, 1341–1344. Peakall, J., Leeder, M., Best, J., Ashworth, P., 2000. River response to lateral ground tilting: a synthesis and some implications for the modeling of alluvial architecture in extensional basins. Basin Res. 12, 413–424. Quennell, A.M., 1958. The structural and geomorphic evolution of the Dead Sea rift. Q. J. Geol. Soc. Lond. 114, 1–24. Quennell, A.M., 1984. The Western Arabia rift system. In: Dixon, J.E., Robertson, A.H.F. (Eds.), The Geological Evolution of the Eastern Mediterranean: Geol. Soc. Lond., Spec. Publ., 17, pp. 375–402. Rabeh, T., Miranda, M., 2008. A tectonic model of the Sinai Peninsula based on magnetic data. J. Geophys. Eng. 5, 469–479. Rabie, S.I., Ammar, A.A., 1990. Pattern of the main tectonic trends from remote geophysics, geological structures and satellite imagery, Central Eastern Desert, Egypt. Int. J. Rem. Sens. 11 (4), 669–683. Ramsay, J., Huber, M., 1987. The techniques of modern structural geology. Volume 2: folds and fractures. . 679 p. Roessner, S., Strecker, M., 1997. Late Cenozoic tectonics and denudation in the Central Kenya Rift: quantification of long term denudation rates. Tectonophysics 278, 83–94. Rosenbaum, G., Regenauer-Lieb, K., Weinberg, R., 2005. Continental extension: from core complexes to rigid block faulting. Geology 33, 609–612. Sabins, F.F., 2000. Remote Sensing - Principles and Interpretation: New York. W.H, Freeman and Company. 494 p. Said, R., 1962. The Geology of Egypt. Elsevier, Netherlands. Said, R., 1990. The Geology of Egypt. A. A. Balkema Publishers, Rotterdam, Netherlands. 734 p. Sandwell, D.T., Smith, W.H.F., 1997. Marine gravity anomaly from Geosat and ERS-1 satellite altimetry. J. Geophys. Res. 102, 10039–10050 available at http://topex. ucsd.edu/cgi-bin/get_data.cgi. Scotese, R.C., Gahagan, L., Larson, R.L., 1988. Plate tectonic reconstructions of the Cretaceous and Cenozoic ocean basins. Tectonophysics 155, 27–48. Sehim, A., 1993. Cretaceous tectonics in Egypt. Egypt. J. Geol. 37, 335–372. Smith, W.H.F., Sandwell, D.T., 1994. Bathymetric prediction from dense satellite altimetry and sparse shipboard bathymetry. J. Geophys. Res. 99B, pp. 21803–21824. Smith, M.J., Wiseb, S.M., 2007. Problems of bias in mapping linear landforms from satellite imagery. Int. J. App. Earth Obs. Geoinf. 9 (1), 65–78. Snyder, N.P., Whipple, K.X., Tucker, G.E., Merritts, D.J., 2003. Channel response to tectonic forcing; field analysis of stream morphology and hydrology in the Mendocino triple junction region, Northern California. Geomorphology 53 (1–2), 97–127. Steckler, M., 1985. Uplift and extension at the Gulf of Suez—indications of induced mantle convection. Nature 317, 135–139. Steckler, M.S., Berthelot, F., Lyberis, N., Le Pichon, X., 1988. Subsidence in the Gulf of Suez: implications from rifting and plate kinematics. Tectonophysics 153, 249–270. Stein, M., Goldstein, S.L., 1996. From plume head to continental lithosphere in the Arabian-Nubian shield. Nature 382, 773–778. Stern, R.J., 2002. Crustal evolution in the East African Orogen: a neodymium isotopic perspective. J. Afr. Earth Sci. 34, 109–117. Stern, T.A., Holt, W.E., 1994. Platform subsidence behind an active subduction zone. Nature 368, 233–236. Stern, R.J., Gottfried, D., Hedge, C.E., 1984. Late Precambrian rifting and crustal evolution in the northeast Desert of Egypt. Geology 12, 168–172. Summerfield, M.A. (Ed.), 2000. Geomorphology and Global Tectonics. John Wiley and Sons, Indianapolis, E.E.U.U. Wassif, N.A., 1991. Paleomagnetism and opaque mineral oxides of some basalt from west central Sinai, Egypt. Geophys. J. Int. 104, 319–330. Wennekers, J.H.N., Wallace, F.K., Abugares, Y.I., 1996. The geology and hydrocarbons oft he Sirt Basin: a synopsis. In: Salem, M.J., Mouzughi, A.J., and Hammuda, O.S. (Eds.), The geology of the Sirt Basin, First Symposium on the sedimentary basins of Libya, Tripoli, 10-13 October 1993. Elsevier, Amestrdam; New York, v. 1, pp. 3–58. Wdowinski, S., Ben-Avraham, Z., Arvidsson, R., Ekström, G., 2006. Seismotectonics of the Cyprean arc. Geophys. J. Int. 164, 176–181. Whipple, K.X., 1999. Geomorphic limits to climate-induced increases in topographic relief. Nature 401, 39–43. Whipple, K.X., 2009. The influence of climate on the tectonic evolution of mountain belts. Nat. Geosci. 2, 97–104. Youssef, M.I., 1968. Structural pattern of Egypt and its interpretation. Am. Assoc. Petrol. Geol. Bull. 52 (4), 601–614. Ziegler, P.A., 1990. Geological Atlas of Western and Central Europe, Shell International Petroleum Mij. B.V., distributed by Geological Society, London, 2nd. Ed. Publishing House, Bath. 239 p.