Gulf of Aqaba

Tectonic and Bathymetric Maps

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Gulf of Aqaba in Google Earth

Gulf of Aqaba in Google Earth

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Tectonic setting of entire Gulf of Aqaba

from Matthieu et al. (2021)

Figure 1

(A) Tectonic setting of the sinistral strike-slip Dead Sea Fault (DSF). Seismicity from the ISC earthquake catalogue 1964 - 2015 (http://www.isc.ac.uk). The DSF connects to the North to the East Anatolian Fault System (EAFS) and to the South to the Red Sea ridge (modified from Le Béon et al., (2008)) GA: Gulf of Aqaba, ST: Strait of Tiran.

(B) Multibeam bathymetric map of GA and ST with the main active faults, combining R/V Thuwal (2018), F/S Meteor (1999) and Hall & Ben Avraham (1978) datasets. The main strike-slip faults are in red while normal faults are in black. Fault traces have been simplified for clarity. The grey focal mechanisms corresponding to the successive sub-events for the, M_w 7.3, 1995 earthquake, and location of the seismic swarms in 1983, 1990, 1993 and other focal mechanisms after Klinger et al., (1999). Grey background is Landsat 8 Imagery, courtesy of the U.S. Geological Survey (2018).

ArF: Arnona Fault
AF: Aragonese Fault
DF: Dakar Fault
EF: Eilat Fault
HF: Haql Fault
TF: Tiran Fault

Matthieu et al. (2021)

Bathymetric Maps

from Matthieu et al. (2021)

Figure 2

Bathymetric map of the Gulf of Aqaba combining R/V Thuwal (2018), F/S Meteor (1999) and Hall & Ben Avraham (1978) datasets
Shade bathymetry of the Gulf of Aqaba with an azimuth of 315N and a sun angle of 25°
Slope map of the Gulf of Aqaba from low slope angle (white: 0°) to high slope angle (black: >45°)

All maps are projected in WGS 84 - UTM 36N. On-land grey background from a Landsat-8 image, courtesy of the U.S. Geological Survey.

Matthieu et al. (2021)

Fault Map - North Gulf of Aqaba

from Matthieu et al. (2021)

Figure 3

Zoom-in of the northern part of the Gulf of Aqaba, along the morphological trace of the Haql fault (see location on Figure 2) with location of the cross sections shown in (B). The fault lines are more detailed than in Figure 1. Red lines represent the main strike-slip faults, black lines the main normal faults. Along the Eilat fault, a long-term displaced channel as well as the left-lateral displacement of a small hill confirm the strike-slip character of the Eilat fault.
Cross-sections along the longitudinal shape of the alluvial fans, North of the city of Haql. No vertical offsets are visible on these cross-sections, with the exception of a possible knickpoint along profile D-D’. The continuous convex shape of the fans suggests no recent activity of the Haql fault.
The trace of the Haql fault is buried by fans coming from the coastal plain, with no visible recent perturbations of the fans at this location. Nevertheless, the high relief shows the long-term normal or oblique character of the Haql fault. In few places, the shaded topography suggests that a small part of strike-slip motion is also accommodated along the Haql fault.
At the southern termination of the Haql fault, discontinuous small scarps across the fans suggest that this section of the fault might have been activated recently.

Matthieu et al. (2021)

Fault Map - Central Gulf of Aqaba

from Matthieu et al. (2021)

Figure 4

Detailed fault map of the sinistral strike-slip fault system in the central GA. Direct evidence of surface rupture associated to the main subevent (see Fig. 2) of the 1995 M_w = 7.3 Nuweiba earthquake are found in box B.
Sharp fault morphology suggesting very recent fault activation. Small changes of geometry along the Aragonese fault are responsible for small pull-apart (black squares) and counterslope scarp (white square).
Detail of the fault zone between Aragonese Deep and Arnona Deep resulting from a complexity in the geometry of the Arnona fault. The red line represents the main active strike slip fault.

Matthieu et al. (2021)

Fault Map - South Gulf of Aqaba

from Matthieu et al. (2021)

Figure 5

Southern part of the Gulf of Aqaba (see location on Figure 2). Dakar and Tiran Deeps are located between the sinistral strike-slip Arnona fault (red line) and the normal Dakar fault (bold black lines). The location of the main strike-slip fault is partly masked by diapiric foldings (black arrows) and secondary faulting (thin black and dashed black lines) associated with the destabilization of large salt deposits moving down from the Dahab plateau.
Cross-sections across the Dahab plateau showing the eastward sloping and the topographic drop from the Dahab plateau toward the Dakar and Tiran deeps.

Matthieu et al. (2021)

Fault Map - Strait of Tiran

from Matthieu et al. (2021)

Figure 6

Strait of Tiran (see location on Figure 2).

The sinistral strike-slip Tiran Fault is located between the Woodhouse and Jackson reefs. The sharp bathymetry to the North and to the South of the reef emphasizes the location of the fault. Red lines represent the main strike slip faults, black lines represent the main normal faults.
Slope map of the Strait of Tiran, from low slope angle (white: 0°) to high slope angle (black: >45°).

Matthieu et al. (2021)

R/V Mediterranean Explorer Cores (N Gulf of Aqaba)

Aerial Views

Gulf of Aqaba in Google Earth

Gulf of Aqaba in Google Earth

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Location Map

Fig. 1

(a) Regional tectonic map of the Dead Sea Transform and location of the Gulf of Aqaba/Elat

(b) Topographic image map of the southern Arava Valley showing location of the study area of the Elat Sabkha. Previously mapped faults in black lines (after Garfunkel, 1970; Garfunkel et al., 1981; Sneh et al., 1998). Previous study sites including Avrona Sabkha and Yovata Sabkha and locations of the paleoseismic trenches (in block circles)

QT = Qatar trench (Klinger et al., 2015)
AT = Avrona Trenches (Amit et al., 1999; Zilberman et al., 2005)
ST = Shehoret trenches (e.g. Amit et al., 2002)
GAE = Gulf of Aqaba/Elat

CMP shots discussed in this study from seismic lines SI-4047 and GI-2108 are plotted as light-blue dots and yellow dots, respectively. The blue rectangle marks the extent of the study area maps presented in Figs. 3 and 9. The pink line represents the location of the offshore high-resolution seismic profile by Hartman et al. (2014) detailed in Fig. 2b.

Kanari et al (2020)

Cores and other data

Article Excerpts

Ash-Mor et al. (2017)

Abstract

from Ash-Mor et al. (2017:36)

Submarine mass transport deposits (MTDs) are a well-known phenomenon in tectonically active regions. Evidence for such deposits is commonly found in the continental slope sedimentary records, as distinct units with coarser grain size compared to the usual and continuous pelagic sedimentation. The Gulf of Eilat/Aqaba is located between the southernmost end of the Dead Sea transform and the spreading center of the Red Sea, and is considered as an active tectonic region.

In this study, an innovative approach using symbiont-bearing Larger Benthic Foraminifera (LBF) to identify MTDs in the Gulf of Eilat/Aqaba (GEA) sedimentary record is presented. The abundance, size and preservation state of LBF shells were analyzed in two radiocarbon dated sediment cores collected at different deposition environments, at water depth of 532 m and 316 m.

The microfaunal and taphonomic results show that the coarse units are characterized by a generally higher numerical abundance of LBF, dominated by Operculina ammonoides, Amphistegina papillosa and Amphistegina bicirculata. These benthic assemblages are found in deeper depths than their original habitat, ranging between 50 and 120 m, in accordance with their symbionts light requirements. In the coarse units, LBF> 1 mm appear in high frequency, up to 161 specimens per g sediment, and poorly preserved shells are also abundant, containing up to 247 specimens per g sediment. In addition, these units also contain high numbers of yellowish and blackish colored LBF shells, as opposed to null in the non-disturbed units, and unlike their natural white color.

The large shell size indicates that high energy is involved in the displacement of the sediments. The poor state of preservation also suggests a turbulent flow during transportation, which requires a high-energy triggering mechanism. The color alteration is probably associated with a diagenetic process related to increasing burial time/depth, also supported by the stratigraphic older ages of the MTDs, suggesting a long burial before the sediments were displaced. In addition, according to the dating of the record, some units correlate with historical and pre-historical earthquakes, reinforcing LBF species as a reliable proxy for mass transport events.

Introduction

from Ash-Mor et al. (2017:36-37)

Submarine mass transport deposits (MTDs) are recognized as im portant sedimentary facies in the marine environment. These deposits exhibit distinct characteristics (Ducassou et al., 2013; Gao and Collins, 1994; Masson et al., 2006) and are used to infer transport processes in different geodynamic settings. The displacement process is known to be associated with sea level fluctuations, ice rifting, river mouths, high sedimentation rates, tropical storms, tsunami backwash, and particularly in tectonically active continental margins (Griggs, 2011; Hampton et al., 1996; Maslin et al., 2005; Polonia et al., 2015; Sugawara et al., 2009; Wright and Anderson, 1982; Yordanova and Hohenegger, 2002; Zabel and Schulz, 2001).

Mass transport deposits consist of recycled sediments initially deposited at the continental shelf and gravitationally transported down the continental slope to deeper water depths. The transport and deposition are strongly grain size selective, resulting in a distinctive texture of coarser sediments, often finning upwards, distinguishable from the finer pelagic continuous deposition (Ducassou et al., 2013; Gao and Collins, 1994; Masson et al., 2006). In some cases, increased organic carbon concentrations point to rapid burial and high preservation that also serve as indicators for MTDs (de Haas et al., 2002; Ducassou et al. 2013;Zabel and Schluz, 2001)

Benthic foraminifera species, which are generally restricted to a specific depth range due to their ecological adaptations, can also serve as indicators for MTDs. In undisturbed conditions, their assemblages vary depending on increasing water depth, substrate type, oxygen content and organic matter flux (Edelman-Furstenberg et al., 2001; de Stigter et al., 1998; Hohenegger, 2004; Jorissen et al., 1995; Murray, 2006). However, instantaneous mass movement events can transport benthic foraminifera along with the sediments and re-deposit them downslope, in a deeper environment compared to their natural habitat. Considering the depth ranges and the ecological requirements of the transported species, it is possible to infer the original deposition depth of the displaced sediments (e.g. Ducassou et al., 2013). Large symbiont-bearing benthic foraminifera (LBF), which are re stricted to the photic zone, are particularly good indicators for MTDs as their depth range is more limited than that of deep sea species (Hallock and Hansen, 1979; Hohenegger et al., 1999; Reiss and Hottinger, 1984). Therefore, sediments derived from a shallow water depth may be easier to recognize and their original depth of deposition can be determined accurately.

The state of shell preservation (taphonomy) can also be used to characterize mass transport processes. In a laboratory experiment, Beavington-Penney (2004) examined the effect of transport distances on shell breakage. Distinguishing between different preservation states of Palaeonummulites venosus, lead them to conclude that the most poorly preserved shells were transported under turbidity current conditions.

Shell coloration is also a taphonomic parameter that can be used to detect sediment mixing in transportation and resuspension processes. Yordanova and Hohenegger (2002) studied black and/or brown LBF shells at water depths of up to 100 m off the shore of western Okinawa, Japan, and suggested that the blackish color is the result of pyritisation and iron sulfides precipitation under anoxic conditions due to sediment burial. Furthermore, the yellowish-brown color is the outcome of limonitisation, a re-oxidation of the pyrite into ferric oxide, due to sediment mixing caused by tropical storms typical to the area.

Here, we focus on fossilized LBF assemblages as a biomarker for the identification and characterization of MTDs in the seismically active region of the northern Gulf of Eilat/Aqaba (GEA). The foraminiferal analysis of sediments in piston cores collected from the gulf enables to establish LBF as a reliable proxy for mass transport events.

Western slope core- MG10P22

from Ash-Mor et al. (2017:38-40)

The bulk of the sediments in this core is generally fine grained with less than 20% coarse size fraction greater than 63 μm. The core record is dissected by two distinct coarser sediment layers of MTDs occurring between 50–56 cm (P22A) and 170–180cm (P22E). These layers comprise of 75–90% coarse fraction. The sediment in these layers is composed of large biogenic (e.g. molluscs, corals, echinoids as well as LBF) and rock fragments. In addition, three layers of slightly coarser sediments are also identified at 65–68 cm (P22B), 79–82 cm (P22C) and 120–134 cm (P22D), containing 18–27% coarse fraction> 63 μm (Figs. 5, 6).

Core MG10P22 spans approximately the last 13 ka (Table 1; Fig. 5). Radiocarbon ages above the MTDs represent the chronological age of these events: unit P22A is dated to 4071 ± 55 cal years BP and unit P22E dates to 7416 ± 42. The former reveals no unconformity caused by the displacement event, while the latter reveals a ~ 5000-year hiatus (Table 1, Fig. 6). The sediments within these units are dated by the age of the displaced LBF: unit P22A is dated to 9364 ± 58 and 10,087 ± 79 cal years BP, and unit P22E to 11,074 ± 87 and 11,759 ± 142 cal years BP, with whitish shells slightly younger than the yellowish shells (Table 1; Fig. 5).

The most common LBF in the core material are Operculina ammonoides and several species of Amphistegina, mostly A. papillosa and A. bicirculata, and also A. aff. A. radiata and A. lessonii that occur in low numbers. The depth ranges of A. aff. A. radiata were similar to that of A. papillosa (Hottinger et al., 1993), and in some cases they were difficult to distinguish (especially juveniles and poorly preserved shells). Therefore, these two species were combined into a single group of A. papillosa & A. radiata. The abundance of the most common LBF species in the displaced sediment layers is much higher than in the pelagic sediments. Furthermore, the dominant species in this core are A. bicirculata and A. papillosa + A. radiata followed by O. ammonoides (Fig. 7).

The LBF occurred in the coarser units in higher numbers, with many shells greater than 1 mm, and with more poorly preserved shells frequently having a yellowish/blackish color. In contrast, the specimens in the fine pelagic sediments, if present, are mostly juvenile, and larger and colored shells are scarce (Fig. 7).

In the coarser units P22A and P22E, the total number of LBF shells larger than 1 mm amounts to 18.8 specimens per g sediment. In contrast, the pelagic sediments amount to 6.6 specimens per g sediment, as most of the specimens in these sections, if present, are smaller than 1 mm.

The number of broken shells (greater than 50%) of Amphistegina spp. in the coarse sediments amounts to 28.3 specimens per g sediment, as opposed to 3.7 specimens per g in the pelagic sediments. In addition, the number of broken shells (greater than 50%) of the less abundant O. ammonoides in the coarse sediments amounts to 3.3 specimens per g sediment, as opposed to 1.1 specimens per g in the pelagic sediments.

Total yellowish shells in the coarse units amounts to 18.1 specimens per g sediment, as opposed to 2.8 specimens per g in the pelagic sediments. Blackish shells are extremely rare in this core (Fig.7).

Submarine canyon core – MG10P27

from Ash-Mor et al. (2017:40)

MG10P27 spans a shorter time period compared to MG10P22, only the last 2300 years (Table 1, Fig. 5). The sediments in this core are generally fine grained as well, with less than 20% of the coarser greater than 63 µm size fraction. These sediments are intersected by five distinct units of coarse sediments occurring between 0 and 10 cm (P27A), 18 and 25 cm (P27B), 38 and 45 cm (P27C), 105 and 110 cm (P27E) and 112 and 145 cm (P27F) at the bottom of the core, with material > 63 µm comprising 70-98% of the entire sediment (Fig. 6). In addition, one more unit of slightly coarser sediments occur at 80-82 cm (P27D) containing 29% fraction > 63 µm. The composition of these units is similar to the material mentioned above in MG10P22.

Due to suspected mixing of the core top, the upper 25 cm were not dated or analyzed for LBF. The sediments below unit P27B were dated to 658 ± 34 cal years BP, while units P27C, P27D and P27E were dated to 1067 ± 42, 2093 ± 56 and 2261 ± 57 cal years BP, respectively. The sediments within these layers were dated between 3482 ± 41 and 5440 ± 49 and contained only whitish and blackish shells. No dating analysis was conducted between units P27E and P27F due to the lack of material for dating. Sediments within unit P27F were dated in a 10-cm resolution varying between 4409 ± 52 and 6523 ± 46, with white shells being the youngest and blackish shells the oldest. Yellowish shells occur only in this unit (Table 1, Fig. 5).

The same species that occur in MG10P22 also appear in MG10P27, although in this core O. ammonoides is the dominant species, followed by A. papillosa (Fig. 7). The overall abundance of these species in core MG10P27 is an order of magnitude higher than in MG10P22. A. bi-circulata is rare in this core. Other LBF species such as Sorites orbiculus, Peneroplis planatus and Heterostegina depressa also occur in both cores though in much lower numbers (Fig. 7).

In the MTDs of MG10P27, LBF larger than 1 mm consist of up to 161.2 specimens per g sediment. In contrast, the pelagic sediments in this core consist of up to 3 specimens per g. The no. of broken shells ( > 50%) of Amphistegina spp. in the coarse sediments amounts to 86.8 specimens per g sediment, as oppose to 7.1 specimens per g in the pelagic sediments. The number of broken shells ( > 50%) of the highly abundant O. ammonoides in the coarse sediments amounts to 190.5 specimens per g sediment, as oppose to 19 specimens per g in the pelagic sediments.

The total number of blackish shells in the coarse units amounts to 132 specimens per g sediment, as opposed to 13 specimens per g in the fine pelagic sediments. The total number of yellowish shells in the coarse units amounts to 12.4 specimens per g sediment, as opposed to 0.1 specimens per g in the fine pelagic sediments (Fig. 7).

Ages from within the mass transport deposits

from Ash-Mor et al. (2017:41)

The ages from within the MTDs are significantly older than the pelagic sediments above and below them, due to their recycling from a prior deposition site. Anomalous older age, unfitting the core stratigraphy, can serve to identify the occurrence of displaced sediments.

The difference between the age of the MTDs and the chronological age of the displacement event, dated above the MTDs, places the burial time of the sediments at the continental shelf before the mass transport event. In the two studied cores, the maximum age differences range from 2681 years (unit P27C) to 6016 years (unit P22A;Fig. 5), suggesting that the sediments were buried for ~2500 to ~6000 years on the continental shelf prior to their displacement.

In some cases, as occurs in unit P22E of the slope core, the age from within the MTD appears not to be anomalously old, and the sedimentary sequence may seem continuous (Fig. 5). However, the chronological age of this unit, dated to 7416 ± 42 ka BP, indicates ~4600 years of sediment removal by this event and an unconformity in the record (Figs. 5,6).

In the canyon core, all three intervals dated in unit P27F reveal similar dating results (Fig. 5), suggesting that this layer originated from the same sediment pack in one massive event. The shorter residence time of sediments on the northern shelf, feeding the canyon core record, together with the higher frequency of MTDs, points to a larger volume of sediments available for transport relative to the western slope.

Sediments availability and sedimentation rates

from Ash-Mor et al. (2017:41-42)

... Repetitive mass transport events may appear as a single event in the sedimentary record (Martín-Merino et al., 2014), which could serve as a possible explanation for the thick P27F unit in the canyon core. Nevertheless, this unit seems to be the result of a single massive event, as it presents a typical graded bedding accumulation pattern known to occur in turbidites (Mulder and Alexander, 2001). The LBF abundance supports this suggestion. As the sediments grow coarser towards the bottom, the numerical abundance of foraminifera gradually decreases(Figs. 5 and 7). Although the bottom of unit P27F was not recovered,these opposite trends of grain size and LBF occurrence, combined with the similar dating results from within this unit, strengthened the interpretation of a single massive event. In contrast, in the western shelf [e.g. slope core P22], where less sediments accumulate, a single high magnitude event may displace a large amount of sediments, and therefore reduce or even eliminate the volume of sediments available for transport in the following event, suggesting that the record is incomplete.

The diﬀerences in the sedimentary record of the two deposition environments show that, regardless of the small distance between them,the recorded events depend strongly on the accumulation rates and the sediment source site. While the pelagic material accumulates relatively equally throughout the water body, the MTDs in the canyon core comprise ~50% of the total sedimentary record, compared to ~11% in the slope core. This highlights the contribution of MTDs to the sedimentary record, and the importance of mass transport processes in the GEA. Moreover, it emphasizes the crucial understanding of the diﬀerent surroundings and bathymetric settings where a study is conducted.

Travel distance and source area estimation

from Ash-Mor et al. (2017:42-43)

Based on the composition of the LBF assemblages, the displaced sediments in both regions originate from a water depth of approximately 50-120 m (Perelis-Grossowicz et al., 2008; Reiss and Hottinger, 1984; Fig. 1). However, the foraminiferal results show distinct differences between the two cores, reflecting the different expression of the same process in different environments. The abundance of LBF species per g dry sediment in the canyon core is ten times higher than in the slope core (Fig. 7). A suggested explanation is that the submarine canyon is transporting sediments that originate from a wider source area.

Based on GIS "watershed" analysis, the estimated source areas are 7.8 km² and 0.47 km² for the canyon and the slope core, respectively. Considering the 50-120 m depth range of the MTDs assemblage's habitat (Hottinger, 2008; Perelis-Grossowicz et al., 2008; Reiss and Hottinger, 1984), the sediment source areas are 3.5 km² and 0.25 km², respectively (Fig. 8), reinforcing this suggested explanation. The travel distance of sediments from the shelf edge at 120 m, is ~3.7 km to the location of the canyon core, and ~0.75 km. to the slope core. These distances are not as long as those known for turbidites in open ocean (Griggs, 2011; Hampton et al., 1996; Khripounoff et al., 2003; Locat and Lee, 2002; Mulder and Alexander, 2001; Tailing et al., 2007), yet four MTDs occur in the past 2500 years in the canyon core, while no such units appear in the slope core in this time period. The different MTDs occurrence in the two records is apparently related to the amount of available portable sediments, which is connected not only to the source area, but also to the bathymetric features of the continental shelf, providing the sediments accumulation space.

Shell size and mobilization

from Ash-Mor et al. (2017:44)

The coarse MTDs are characterized by LBF with a generally larger shell size (Fig. 7), with A. papillosa reaching a maximum diameter of 1.5 mm, A. bicirculata of 2 mm and O. ammonoides of ~4 mm. In contrast, the pelagic sediments contained only a few juvenile specimens, with a shell diameter of 150-250 µm, if any. The larger shell size represents adult specimens living and dying in their natural habitat prior to the abrupt event that triggered the displacement. Larger grains and shells require higher energy and current velocities in order to be moved as particles.

Yordanova and Hohenegger (2007) examined threshold friction and entrainment velocities and showed that A. bicirculata and A. papillosa with a shell diameter of 1.5 mm and O. ammonoides with a shell diameter of 3 mm, require velocities of ~18 cm/s for entrainment on a flat rough surface. The rare occurrence of LBF > 150 µm in the fine pelagic sediments suggests that sediments of this size are not transported from the shelf area to the deep sea-bed under natural conditions. Therefore, the larger shell size of the LBF in the MTDs is another indicator for transport from the outer continental shelf to a deeper depth by high velocity events.

Taphonomy

Degree of breakage

from Ash-Mor et al. (2017:44-45)

The coarse MTDs are also characterized by high abundance of broken LBF shells (Fig. 7). This indicates turbulent conditions during transport causing a high degree of shell abrasion and fragmentation, unlike the excellent preservation of planktonic and deep water benthic foraminifera found in the fine pelagic sediments. Beavington-Penney (2004) simulated the transport of Palaeonummulites venosus shells under laboratory conditions, and analyzed their fragmentation and abrasion features. According to this study, > 50% of shell fragmentation is related to transport distance > 70 km, predation by large bioeroders or transport within turbidity currents. Considering the relatively short distance of transport in the current study area (Figs. 2 and 8), it is believed that the turbulent flow associated with mass transport processes is the cause of the highly fragmented shells found in the MTDs.

Turbulent flow requires a high-energy triggering mechanism and steep bathymetry. The high abundance of > 1 mm and broken LBF shells in the MTDs, combined with the steep bathymetry of the GEA slope (Tibor et al., 2010), requires much higher current velocities than the velocities measured in the gulf (Biton and Gildor, 2011; Khripounoff et al., 2003; Wynn et al., 2000). Therefore, the GEA's regional tectonic activity is a potential trigger for these mass transport events.

Shell coloration

from Ash-Mor et al. (2017:45)

The displaced sediments in the MTDs are characterized by a relative abundance of colored LBF shells, corresponding to their larger shell size and poor preservation (Fig. 7). In the slope core, colored LBF shells found within the MTDs occurred with yellowish color, whereas in the canyon core shells appeared with both yellowish and blackish color (Fig. 7). The coloration of biogenic particles in the GEA has not been studied yet, although black shells of O. ammonoides were found to be present in surface sediments from the northern shelf (Perelis-Grossowicz et al., 2008).

LBF shell coloration is assumed to be associated with postmortem processes and burial depth (Maiklem, 1967; Yordanova and Hohenegger, 2002). The latter described a linear diagenetic process affecting foraminifera shells, starting with pyritisation due to anoxic conditions caused by sediment accumulation and burial, followed by limonitisation associated with re-ventilated conditions due to tropical storms. This led to the suggestion that colored shells may also serve as an indicator for identifying MTDs that consist of older recycled sediments.

In this study, the re-oxidation may be the outcome of the turbulent flow during the mass transport events. Sediments, which were long buried, were mixed and exposed once again to the oxygenic water column before their redeposition in the final deeper terminal accumulation area. Since the coloration is a diagenetic process developed over time, we expected an age difference with colored shells being older than the pristine white shells.

The dating results of LBF taken from the MTDs in the canyon core support the process described above, as the blackish and yellowish shells were found to be older than the white shells at a range of a few hundred up to 2060 and 1222 years, respectively. Yellowish shells were found only in unit P27F, yet their age was consistently younger than the blackish shells by 300 to 1400 years (Table 1. Fig. 5). In the slope core, no black shells occur, and the yellowish shells suggest that all pyrite containing shells are apparently oxidized to limonite upon their transport. However, the dating results of the yellowish shells from both units pre-date the pristine white shells by 700 years, suggesting a more complex process of diagenesis related to post-mortem secondary calcite precipitation. Moreover, the higher abundance of yellowish shells in unit P22A, rather than unit P22E (Fig. 7), suggests that only a part of the sediments from the source area were transported during the deposition of unit P22E. Therefore, the sediments of unit P22A were buried for a longer period on the continental shelf, enabling the diagenetic process to progress before being transported.

Foraminiferal proxies, both shell size and taphonomy, for MTDs also appear in units that cannot be distinguished based on grain size alone, as in units P22B - P22D of the slope core and unit P27D of the canyon core (Fig. 7). This reinforces the reliability of foraminifers as a proxy for the identification of small scale mass transport events, as well as large scale events.

Earthquakes as triggers for mass transport events

from Ash-Mor et al. (2017:45)

Mass transport events are known to be associated with tectonic activity (Griggs, 2011; Locat and Lee, 2002; Polonia et al., 2015). The northern GEA is a tectonically active zone (Ben-Avraham, 1985; Ehrhardt et al., 2005; Klinger et al., 1999; Shaked et al., 2011), and seismic activity is a possible trigger for mass transport events.

The chronology of MG10P27 covers the historical period, which is well documented in seismic catalogues and geological records (Ambraseys et al., 1994; Amit et al., 2002; Kagan et al., 2011; Ken-tor et al., 2001; Khair et al., 2000). According to Kanari (2016), unit P27C in the canyon core coincides, within the error range, with a ~7M_W earthquake which occurred in 948 years BP (1068 CE) and caused heavy destruction to Aqaba (Ambraseys et al., 1994; Ben-Menahem, 1991; Kagan et al., 2011). In addition, a major surface rupture of > 12 km in length documented north of Eilat, caused by a seismic event of at least 7M_W and dated between 900 and 1000 years BP (Zilberman et al., 2005), correlates to this event.

The chronological sequence of MG10P22 reveals a pre-historical period too old for documentation in seismic catalogues. Nevertheless, the two MTDs in this core, P22A and P22E, correlate well with two catastrophic events described by submerged fossilized coral reefs (Shaked et al., 2004, 2011). Unit P22A correlates well with an earthquake event suggested by Shaked et al. (2004, 2011) to have occurred ~4.7 ka BP. Unit P22E, dated to 7416 ± 66, correlates well with the initial growth of fossilized corals, dated to at least 7 ka BP, suggesting that this unit served as the substrate for the corals settlement. The occurrence of these two events documented in the coastal area of the gulf, in association with the slope core from the deep sea, reinforces the assumption of a physical barrier, as suggested above, preventing shallow water sediment and benthic fauna from being transported to the deep sea during these events. Evidence for the intensity and widespread influence of these two events was also identified at the northern extension of the Dead Sea Transform, in sedimentary cores from the shores of the Dead Sea (Kagan et al., 2011).

The correlation of the MTDs found in the studied cores with known and previously studied seismic events strengthens the hypothesis of seismic activity as the triggering mechanism in this study area. Furthermore, if the taphonomy of the LBF (% of poorly preserved shells), which is dictated by the mass transport intensity, is used as a proxy for the local intensity of the triggering event, it is possible to distinguish between small, intermediate and large-scale events vs. the pelagic sediments (Fig. 9). However, it should be noticed that the number of specimens is highly dependent on the depositional settings, and the MTDs of the western slope vs. the submarine canyon need to be distinguished.

Conclusions

from Ash-Mor et al. (2017:45-46)

The Gulf of Eilat/Aqaba (GEA) continental slope cores display coarse sediment units with distinct micropaleontological and taphonomic features, indicative of displaced sediments. These units are characterized by a sharp increase in the abundance of symbiont-bearing Larger Benthic Foraminifera (LBF) with large shell size and poor preservation, suggesting an abrupt and energetic triggering event and turbulent transport. Shell coloration appears to be associated with the large shell size and poor preservation, and probably indicates a long burial before the displacement and re-oxidation during an instantaneous transport event. Nevertheless, further geochemical analysis is required in order to understand the diagenetic processes involved.

Larger symbiont-bearing benthic foraminifera were found to be a useful tool to identify mass transport deposits (MTDs). According to the LBF assemblage found in the MTDs at the GEA, these deposits originate from the deeper shelf area, at a water depth of 50-120 m. The dating results of the displaced LBF are anomalously older than the pelagic sediments above them, suggesting that sediments accumulated at the deep shelf, a few thousand years before the transport.

Although both cores present similar LBF characteristics, their different deposition environment also dictates differences in the MTDs record. The canyon core, fed by a wider and moderate shelf area, presents a higher frequency of events and a larger volume of transported sediments, with a chronologically younger age of the accumulating MTDs. The slope core shows a lower frequency of events transporting a smaller sediment volume. In addition, considering that mass transport events are not necessarily expressed by anomalous ages, as seen in unit P22E, it is concluded that in the study of MTDs, age anomalies should be used only to support other proxies such as grain size, organic carbon content and displaced benthic fauna.

The correlation between the young MTDs and known earthquakes reinforces the hypothesis that seismic events are the triggering mechanism. We conclude that LBF serve as a useful and reliable proxy for the identification and investigation of mass transport events in general, and those triggered by earthquakes in particular.

Kanari et al. (2015)

Abstract

from Kanari et al. (2015:240)

Located at the Northern tip of the Gulf of Aqaba-Elat, the on-land continuation of the submarine Avrona Fault underlies the Hotels District of Elat, where seismic deformation was documented after the 1995 Nuweiba (Sinai) earthquake (7.2 M_W). This active segment of the Dead Sea Fault is the transition between the deep marine basin of the Gulf and the shallow continental basin of the Arava Valley. Paleoseismic trenching revealed the fault, based on surface rupture and liquefaction features. Radiocarbon dating of the offset strata and liquefaction suggest that it ruptured in the historical earthquakes of 1068 and 1458 AD, yielding a vertical slip rate of ~1.1 mm/yr. Independent dating of anomalous coarse grain events in core sediments from offshore nearby suggests these earthquakes triggered marine sediment mass-flow. Using this pattern, we analyze anomalous coarse grain events in several cores to compile a paleoseismic record dating back to the late Pleistocene.

Introduction

from Kanari et al. (2015:240-241)

At the north tip of The Gulf of Aqaba-Elat (the northeast extension of the Red Sea; Fig. 1), reside the cities of Elat (Israel) and Aqaba (Jordan): major economic, cultural, and recreational centers of southern Israel and Jordan, and vital aerial and naval ports. It so happens that they are both also built on active faults, which have ruptured in the past. Aqaba was completely destroyed in the 1068 AD earthquake (Ambraseys et al., 1994; Avner, 1993), and significant damage to structures in both Elat and Aqaba was inflicted by the Nuweiba (Sinai) earthquake (22.11.1995; M_W 7.2) even though the epicenter was located 70 km to the south (Klinger et al., 1999). The estimation of seismic hazard to these neighboring cities is therefore vital. The peaceful hotels and beaches of Aqaba and Elat are located on a tectonic plate boundary, which is also a transition zone between two crustal realms of the Dead Sea Fault system (DSF): the deep en echelon submarine basins of the Red Sea (Ben-Avraham, 1985) and the shallow continental basins of the Arava (Frieslander, 2000), localizing into a single fault strand heading northward.

Previous studies of the submarine structure of the Northern Gulf of Aqaba-Elat (NGAE) suggest slip on the east and west boundary faults is predominantly normal and recently active (Ben-Avraham, 1985; Ben-Avraham et al., 1979; Ben-Avraham and Tibor, 1993). However, recent high-resolution seismic and bathymetric data (Tibor et al., 2010; Hartman, 2012; Hartman, 2015) revealed a complex fault system across the shelf of the NGAE with varying degrees of recent seismic activity. Hartman et al. (2015) conclude that during the Holocene, the submarine Avrona Fault (Evrona Fault in some papers) accommodates most of the strike-slip faulting in this transform plate boundary, between the Sinai sub-plate and the Arabian plate, with an average sinistral slip-rate of 0.7±0.3 mm/yr through the Late Pleistocene and 2.3 3.5 mm/yr during the Holocene. (Fig. 2), and a Holocene vertical slip rate of 1.0 ±0.2 mm/yr, suggesting that its seismic activity has increased through recent time.

On-shore, several works estimated the location of the Avrona Fault at the border of the Elat Sabkha (Garfunkel et al., 1981) and in the vicinity of the Elat hotel district (Wachs and Zilberman, 1994). Using seismic imaging, Rotstein et al. (1994) suggested a vertical deformation band of several hundred meters wide below the eastern part of the Elat Hotel District. Further seismic data was used by Frieslander (2000) to suggest a distinct sub vertical discontinuity in the sediments in the same area in Elat. Active surface faulting was observed following the Nuweiba (Sinai) earthquake in 1995 (epicenter 70 km south to Elat), when an offset street was reported in the same hotels area (Wust, 1997). Some 15 km farther north, Paleoseismic trenching in the Avrona Playa revealed late Pleistocene earthquake ruptures displaced 1-1.5m with estimated magnitudes M6.7-M7, and Holocene earthquakes displacing 0.2-1.3m with estimated magnitudes M5.9-M6.7 (Amit et al., 2002). Zilberman et al. (2005) had extensively detailed the surface rupture of the fault in the Avrona Playa, relating observed surface rupture to the two historical earthquakes affecting the southern Arava valley and the ancient city of Aila: the 1068 AD and the 1212 AD earthquakes. They suggest an earthquake recurrence interval of 1.2±0.3 ka for this fault zone. However, the location and the paleoseismic record of the on-land continuation of the marine Avrona Fault, as it emerges from submarine to terrestrial domain, was not known, and surface rupture from the 1068 AD earthquake south of the Avrona playa was not observed so far. Zilberman et al. (2005) report that there was no way to determine the length of the surface rupture in the Avrona Playa due to obscuring by erosion, younger deposits and incision of alluvial fans.

Results and Discussion

from Kanari et al. (2015:241-242)

... In an independent analysis of the submarine core P27 (Fig. 4; see Fig. 2 for core location) - several anomalous coarse grain (>2mm, up to several cm maximum) events were observed, while most of the core is of typical pelagic deposition of less than 250 um in grain size. Radiocarbon dating of the anomalous events in the core resulted in a good match between the estimated ages of two anomalous events from the top of the core and the 1068 and 1458 AD earthquakes (Fig. 4). We therefore suggest that the anomalous events in the submarine core P27 correspond to the earthquakes of 1068 AD and 1458 AD, which were also observed independently in T1 and T3 trenches on-land, just several km away to the north.

Following this similar pattern of dating anomalous events in core P27 (validated by historical and on-land observations), several other piston cores were analyzed, and their coarse grain anomalous events ages were determined using radiocarbon dating of foraminifera, gastropod and bivalves: P12, P17, P22 and P29 (460, 540, 320 and 280 mbsl). For some events, more than one anomalous events appear to coincide in time in different cores. We suggest that where anomalous events in different cores coincide in their age constraints – it is most likely evidence for mass flow triggered by earthquake events, driving coarse material from the shallower shelf edge into the deep basin (as opposed to sporadic slumping, or mass flow triggered by flashfloods). These anomalous events, observed in several cores from across the NGAE (Fig. 5), serve as basis for the compilation of an earthquake record dating back to late Pleistocene. We discriminate between events validated in more than one core (high confidence level) and events that appear in one core (low level of confidence). In total, we count seven earthquake events (excluding the 1068 AD and the 1458 AD historically validated core events) of which four are of high confidence level; one event is dated to ca 40ka, but could be of less confidence to to the limitations of the 14C dating method. Zilberman et al. (2005) suggest that 5 earthquakes ruptured the Avrona Playa between 14.2±0.3 and 3.7±0.3 ka, which conform with our marine core sediment dated events, as we identify an event ca 2.5 ka, and event ca 40 ka, and five events in a similar time range.

To conclude, we suggest that by correlating on-land and offshore paleoseismic observations, we have evidence for past earthquakes of the late Pleistocene and Holocene around 2.5, 3-3.3, 4.0-4.2, 5.8-6.3, 7.5, 14-14.5 and possibly an event around 40 ka BP. Some of these events may support evidence for past earthquakes suggested by previous authors.

Canyon Core P27

Grain Size Distribution Logs

Ash-Mor et al. (2017)

Figure 5

3D grain size distribution up to 2 mm (left) and radiocarbon dating results (right) along the canyon core MG10P27. Color bar represent % of grain size differential distribution by volume. Black dots represent the chronological age of the pelagic sediments, whereas diamonds represent the different color groups of LBF shells from within the MTDs.

Ash-Mor et al (2017)

Kanari et al. (2015)

Figure 4

Grain size distribution (downcore spectrum of % volume per grain diameter) and ¹⁴C age determinations (cal BC/AD) of core P27 from the northern Gulf of Aqaba Elat. ¹⁴C age calibrated using Calib 7.0 (Stuiver and Reimer, 1993) and Marine13 calibration curve (Reimer et al, 2013).

Kanari et al (2015)

Sedimentary Characteristic Log

Figure 6

Sedimentary characteristics of the studied core MG10P27

Core image and scheme describing the MTDs
% sediment > 63 μm
sedimentation rates calculated according to the chronological dating results of the pelagic sediments

Ash-Mor et al (2017)

Foraminifera Log

Figure 7

The distribution of the most common LBF in the canyon core MG10P27

Total count of A. papillosa + A. radiata, A. bicirculata andO. ammonoides
number of LBF smaller and larger than 1 mm
number of pristine, moderately and poorly preserved A. ammonoides
number of pristine, moderately and poorlypreserved A. papillosa + A. radiata
number of white, yellowish and blackish LBF shells

All categories are normalized to 1 g of dry sediment.

Higher values occur in all categories at 18–25 cm (P27B), 38–45 cm (P27C), 82–84 (P27D), 105–110 cm (P27E) and from 112 cm down to the bottom of the canyon core (P27F).

Ash-Mor et al (2017)

Slope Core P22

Grain Size Distribution Log

Ash-Mor et al. (2017)

Figure 5

3D grain size distribution up to 2 mm (left) and radiocarbon dating results (right) along the slope core MG10P22. Color bar represent % of grain size differential distribution by volume. Black dots represent the chronological age of the pelagic sediments, whereas diamonds represent the different color groups of LBF shells from within the MTDs.

Ash-Mor et al (2017)

Kanari et al. (2015)

Figure 5

Grain size distribution (downcore spectrum of % volume per grain diameter) of core P22 (316 mbsl) from the Northern Gulf of Aqaba-Elat; see Fig. 2 for core locations

Kanari et al (2015)

Sedimentary Characteristic Log

Figure 6

Sedimentary characteristics of the studied core MG10P22

Core image and scheme describing the MTDs
% sediment > 63 μm
sedimentation rates calculated according to the chronological dating results of the pelagic sediments

Ash-Mor et al (2017)

Foraminifera Log

Figure 7

The distribution of the most common LBF in the slope core MG10P22

Total count of A. papillosa + A. radiata, A. bicirculata andO. ammonoides
number of LBF smaller and larger than 1 mm
number of pristine, moderately and poorly preserved A. ammonoides
number of pristine, moderately and poorlypreserved A. papillosa + A. radiata
number of white, yellowish and blackish LBF shells

All categories are normalized to 1 g of dry sediment.

Higher values occur in all categories at50–56 cm (P22A), 65–68 cm (P22B), 79–82 cm (P22C), 124–125 cm (P22D) and 170–180 cm (P22E) in the slope core

Ash-Mor et al (2017)

Core P17

Grain Size Distribution Log

Figure 5

Grain size distribution (downcore spectrum of % volume per grain diameter) of P17 (540 mbsl) from the Northern Gulf of Aqaba-Elat; see Fig. 2 for core locations

Kanari et al (2015)

Core P29

Grain Size Distribution Log

Figure 5

Grain size distribution (downcore spectrum of % volume per grain diameter) of P29 (282 mbsl) from the Northern Gulf of Aqaba-Elat; see Fig. 2 for core locations

Kanari et al (2015)

Fence of Grain Size Distribution Logs for P17, P22, and P29

Figure 5

Grain size distribution (downcore spectrum of % volume per grain diameter) of cores P17 (540 mbsl), P22 (316 mbsl) and P29 (282 mbsl) from the Northern Gulf of Aqaba-Elat; see Fig. 2 for core locations

Kanari et al (2015)

Bathymetric distribution of common larger symbiont-bearing foraminifera (LBF)

Figure 1

The bathymetric distribution of common larger symbiont-bearing foraminifera (LBF) from the Gulf of Eilat/Aqaba (GEA). The darker color represents the maximumabundance depth interval. Based on Haunold et al. (1997), Hottinger et al. (1993), Oronet al. (2014), Perelis-Grossowicz et al. (2008) and Reiss and Hottinger (1984).

Ash-Mor et al (2017)

Larger symbiont-bearing foraminifera (LBF) as a proxy for Mass Transport Deposits (MTDs)

Figure 9

LBF as proxy for MTDs in diﬀerent deposition environments. The % of poorly preserved shells distinguishes between the pelagic sediments and small, intermediate and large-scale events. The number of specimens/g sediment distinguishes between the western slope as opposed to the submarine canyon.

Ash-Mor et al (2017)

Sediment source areas for slope core P22 and canyon core P27

Figure 8

Sediment source areas estimated for the slope core MG10P22 (light blue) and the canyon core MG10P27 (pink) based on GIS “watershed” analysis. The opaque color represents the 50–120 m water depth of the deeper shelf.

Ash-Mor et al (2017)

R/V Thuwal Cores (N, S, and Central Gulf of Aqaba)

Aerial Views

Gulf of Aqaba in Google Earth

Gulf of Aqaba in Google Earth

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Fig. 1 - Location Map

Normal Size

from Bektaş et al. (2024)

Fig. 1

Sediment coring locations in the Gulf of Aqaba.

a) Tectonic setting in the eastern Mediterranean and the northern Red Sea.

b) The bathymetry of the Gulf of Aqaba and active faults (Ribot et al., 2021; Le B´eon et al., 2012) with black dots showing the coring locations. The main strike-slip faults are in red

TF: Tiran Fault
ArF: Arnona Fault
AF: Aragonese Fault
EF: Eilat Fault
WAF: Wadi Araba

whereas normal faults are in black. Paleoseismic trenching sites are labelled as

x (Klinger et al., 2015)
y (Amit et al., 1999; Zilberman et al., 2005)
z (Amit et al., 2002)
t (Kanari et al., 2020)

The top right inset shows the lateral extent of historical earthquakes on the WAF (Klinger et al., 2015).

c-g) Close-up views of the coring locations in the Eilat, Aragonese, Dakar, Tiran and Hume deeps.

Bektaş et al. (2024)

Magnified

from Bektaş et al. (2024)

Fig. 1

Sediment coring locations in the Gulf of Aqaba.

a) Tectonic setting in the eastern Mediterranean and the northern Red Sea.

b) The bathymetry of the Gulf of Aqaba and active faults (Ribot et al., 2021; Le B´eon et al., 2012) with black dots showing the coring locations. The main strike-slip faults are in red

TF: Tiran Fault
ArF: Arnona Fault
AF: Aragonese Fault
EF: Eilat Fault
WAF: Wadi Araba

whereas normal faults are in black. Paleoseismic trenching sites are labelled as

x (Klinger et al., 2015)
y (Amit et al., 1999; Zilberman et al., 2005)
z (Amit et al., 2002)
t (Kanari et al., 2020)

The top right inset shows the lateral extent of historical earthquakes on the WAF (Klinger et al., 2015).

c-g) Close-up views of the coring locations in the Eilat, Aragonese, Dakar, Tiran and Hume deeps.

Bektaş et al. (2024)

Table 1 - Core Location Coordinates

from Bektaş et al. (2024)

Table 1

List of location coordinates (UTM Zone 36) and lengths of collected sediment cores.

Bektaş et al. (2024)

Cores

Article Excerpts on Description and Interpretation

Turbidites and sedimentary events

from Bektaş et al. (2024:4-9)

Radiographic images of the cores reveal that the background sediments in the Gulf of Aqaba do not exhibit lamination due to intense bioturbation (Figs. 2 and 3). Since the shell density of planktonic foraminifera (1.4 – 1.5 g/cm3; Fok-Pun and Komar, 1983) is generally lower than that of the sediments, they are seen in the radiographic images as light-colored spots. Within the intensely bioturbated background sediments having high biogenic content (predominantly planktonic foraminifera), we identified numerous sedimentary anomalies that are significantly different from these complex background sediments in our cores. These sedimentary anomalies, which produce signals that differ significantly from the background sediments in proxies and radiographic images, are labeled alphabetically from the top to the bottom for each core (Figs. 2, 3, and 4). They appear in the radiographic images as darker intercalations, implying that they have higher density compared to the background sediments.

The examples of the sedimentary events presented in Fig. 2 are the most prominent and well-preserved turbidites in the studied cores. Due to their instantaneous deposition and thickness mostly over 4 cm, bioturbation is limited to only the topmost parts of these turbidites, resulting in well-preserved internal structures. Parallel-to-subparallel laminations (Fig. 2: 2-E, 15-B, 3-A, 4-A, 15-C, 16-B, 17-G, and 15-A), and even cross laminations in some cases (3-E, 4-D, 14-B, and 16-A; Fig. 2), just above the sharp bottom boundaries of these events, can be attributed to multiple coarse sediment pulses that are due to multiple successive mass wasting events along basin slopes, likely caused by an earthquake (Shiki et al., 2000; Nakajima and Kanai, 2000; Goldfinger et al., 2007, 2008; Goldfinger, 2011; Van Daele et al., 2014; 2017; Wils et al., 2021). These distinct laminations are mostly overlapped by more homogeneous and probably finer-grained sediments that are clearly lacking carbonaceous biogenic content, i.e. mostly planktonic foraminifera seen as whitish spots (e.g., 13-I in Fig. 2). Although boundaries at the bottom of turbidites are sharp and distinct, the boundary between the top of the turbidite and the above background sediments is difficult to determine precisely since it is gradational and bioturbated (Goldfinger et al., 2008; Goldfinger, 2011). In addition to their well-preserved internal structures, the lack of biogenic content in these sediments confirms their quasi-instantaneous deposition, which must have been too fast for the biogenic carbonates precipitating from the water column to be included into the sediments. In this study, we classify all of the sedimentary events that appear like the ones presented in Fig. 2 as “Type I: Turbidites”.

Some hazy, but still relatively darker levels, can be seen in the radiographic images (e.g., 9-B, 17-B, 2-C, and 18-B in Fig. 3), although they neither include multiple laminations nor foraminifera-free homogenous parts. We also classified these levels as sedimentary events, since their darker radiographic view implies a sudden influx of coarser hemipelagic sediment arrival at the coring location. We interpret them as either thin and singular turbidites or flood deposits that were dispersed within the background sediments by intense bioturbation. Hence, we classify them as “Type II: Turbidite or Flooding” events. Type III events are characterized by their thickness, darker appearance in radiographic images (e.g., 17-C, 17-F, and 18-D in Fig. 3), and coarser grain-size compared to the background sediments, which is evident from their Sand (%) values (9-F, 11-F, 11-I, 11-M, 17-C, 17-F, and 18-D in Fig. 5). Unlike Type I events, Type III events are bioturbated, lacking any lamination, and containing biogenic remains. We consider three possible explanations regarding their origin. Firstly, it is possible that Type III events are turbidites that underwent complete bioturbation after deposition. However, considering that coarser sediments typically exhibit less vertical penetration of bioturbation (Wheatcroft, 1992), and given that Type III events are noticeably coarser than the background sediments, it is unlikely that these events under went extensive bioturbation after deposition. Another possibility is that Type III events originated from hemipelagic sediments highly rich in biogenic content, hence their turbidites contain abundant biogenic remains as well (e.g., Van Daele et al., 2017; Polonia et al., 2023). However, if this was the case, we would expect to observe some evidence of multiple laminations or fining-upward grain-size trends in Type III events, which is not observed. The most plausible explanation for these events is that, during their deposition, there was sufficient time for organisms to dig and burrow, and for biogenic remains from the water column to be included into the sediments. Consequently, the deposition of Type III events was likely slower than that of Type I and Type II events. These events likely represent a series of successive flooding events that occurred over a period of several years or decades. We classify these events as “Type III: Thick flooding sequence”. For a more detailed classification of sedimentary events, please refer to the Supplementary Material (E-SUPP 1).

Fig. 4 shows radiographic images of all the 18 cores side-by-side along the gulf, together with the depths of ¹⁴C and ²¹⁰Pb measurements and their results (Table in Fig. 4). The stratigraphic order of the dated samples confirms that our samples were collected in regular sedimentation sections and not in any anomalous sedimentary event. Radiocarbon results of 16 samples show that most of the cores include the sedimentary record for at least the last 1000 years.

We applied the “Constant Flux Constant Sedimentation Rate” model (Goldberg, 1963) on the ²¹⁰Pb_ex (excess lead) from the cores 3, 7, and 17, which yielded sedimentation rates of 0.25, 0.19, and 0.38 mm/yr for the top parts of these cores, respectively. It should be noted that event 17-A was excluded from the depth scale for the sediment rate calculation of core 17. Although the number of radionuclide samples collected from core 7 is insufficient to achieve a statistically meaningful sedimentation rate, the ²¹⁰Pb_ex results from cores 3 and 17 can be compared with the radiocarbon ages. For core 3, ¹⁴C at the bottom of the core yields a bulk sediment rate of 0.44 mm/yr, which is inconsistent with the rate obtained by ²¹⁰Pb_ex for the same core (0.25 mm/yr). However, ¹⁴C-based sediment-rate calculation by using composite depths, which are obtained by excluding the sedimentary events, yields a composite sediment rate of 0.26 mm/yr that is consistent with the ²¹⁰Pb_ex sediment rate. Similarly, ¹⁴C-based bulk rates for cores 14 and 15 (0.56 and 0.50 mm/yr, respectively) in Eilat Deep are much higher than the ²¹⁰Pb_ex-based sediment rate obtained for core 17 (0.38 mm/yr) from the same basin. However, once corrected by removing the event layers, the ¹⁴C-based composite sediment rates for cores 14 and 15 (0.34 and 0.33 mm/yr, respectively) are rather consistent with the 0.38 mm/yr derived from the ²¹⁰Pb_ex measurements. Thus, these two comparisons confirm the importance of determining both the bottom and top boundaries of sedimentary events to properly exclude them from the sequences and hence to construct reliable sediment chronology.

In order to cross-check the existence and extent of the sedimentary events in the cores, we compared our visual inspections of the radiographic images to magnetic susceptibility, grain-size (only sand content) and µ-XRF measurements (Fig. 5). For these analyses, at least one core per basin was selected. During floods or mass wasting events, coarser sediments originating from the shallower parts of the basin or from drainages onshore are expected to reach bottom of the basins. Thus, sand content in the event deposits is expected to increase compared to the background sedimentation. Deep marine sediments are normally a mixture of terrigenous clastics (mainly aluminosilicate minerals) and bio/chemical carbonates produced in the water column. However, there is almost no bio/chemical carbonate input during the almost instantaneous deposition of turbidite or flood deposits, which makes them richer in terrigenous clastics compared to the background sediments. Since aluminosilicates have higher magnetic susceptibility values than carbonates (Nowaczyk, 2001), turbidite and flood deposits should show as anomalies along the magnetic susceptibility profiles of the cores. Similarly, Zr/Sr profiles, where Zr and Sr represent aluminosilicates and foraminiferal calcite, respectively (Rothwell et al., 2006; Croudace and Rothwell, 2015), should show anomalies at turbidite and flood levels.

Among the 67 events shown in Fig. 5, significant magnetic susceptibility anomalies are observed for 51 of them. While 11 events (1-C, 3- D, 7-A, 9-C, 10-B, 10-C, 13-I, 15-C, 17-E, 17-G, and 18-B) show almost no magnetic susceptibility anomalies, only three events (3-A, 15-A, and 17-A) have clearly lower magnetic susceptibility values than the background sediments. Sand fraction profiles show distinct anomalies for 57 events out of 67. On the other hand, events 7-B, 15-A, 15-B, and 17-A have lower sand content compared to the background sediments. No sand anomalies are observed for four events (1-A, 1-B, 13-I, and 17-G). Since no µ-XRF scanning was done for core 18, 62 events can be tested for Zr/Sr anomalies. Among these, 49 events show higher and only two events (3-A and 17-A) lower Zr/Sr anomalies. Nine events (3-C, 3-D, 7-A, 9-C, 10-B, 11-A, 11-B, 13-D, and 13-I) show no Zr/Sr anomalies. Accordingly, from the data presented in Fig. 5, we note that magnetic susceptibility, sand content, and Zr/Sr profiles are successful in detecting sedimentary events observed on the radiographic images at rates of 81%, 91%, and 79%, respectively. Although Type I and Type III events are already evident in the radiographic images, some Type II events are unclear in the images (e.g., 7-A in Fig. 3). Magnetic susceptibility, sand content, and Zr/Sr profiles, together with the radiographic images, were therefore particularly useful to confirm/detect the thickness of Type II events, so that all sedimentary events were successfully excluded from the sequences covered by the cores to achieve reliable stratigraphical correlations and sediment chronology.

Chemostratigraphical correlation and sediment chronology

from Bektaş et al. (2024:9-10)

Coevality of turbidites at different locations and even in different basins should be tested to achieve successful submarine paleoseismological records (e.g., Goldfinger, 2011), which can be achieved by careful high-resolution stratigraphical correlations. In Fig. 4, some of the sedimentary events can be visually correlated between cores collected from the same basin according to their stratigraphical order, e.g., between the cores 2, 3, and 4 in Tiran basin, cores 9 and 10 in Dakar basin, and cores 14, 15, 16, 17 and 18 in Eilat basin. However, visual observations are not reliable enough for inter-basin correlations between the cores as one cannot assume that the number of turbidites in different basins is the same. Stratigraphical correlations between cores can be achieved by using data reflecting geophysical and geochemical properties of sediments, which may include magnetic susceptibility, bulk density, grain-size distribution, computed tomography (CT) image analysis, μ-XRF data and paleomagnetic secular variation (PSV) records (Patton et al., 2013; Drab et al., 2015; Ikehara et al., 2016; Goldfinger et al., 2017; Usami et al., 2018). In our study, geophysical and geochemical properties of sediments were evaluated by magnetic susceptibility and grain-size measurements (Sand percent), and μ-XRF μ scanning, of which resolutions were 5 mm, 10 mm and 0.5 mm, respectively. μ-XRF data, which has significantly higher resolution than the other proxies, was preferred for core correlation. Reliable chemostratigraphical correlations over large distances, like in our case where the cores are distributed along the ~180-km-long gulf, can only be successfully achieved by using a sedimentary geochemical proxy recording regional environmental conditions effective for the entire gulf. It is known that Sr/Ca ratios of planktonic foraminifera have a strong positive correlation with sea-surface temperature (e.g., Clerouxet al., 2008). Since the sediments of the Gulf of Aqaba are rich in planktonic foraminifera, and by assuming that the surface seawater temperatures would synchronously change over the entire gulf, we used Sr/Ca ratio profiles to correlate the cores in this study.

We intentionally collected core 11, the longest core in this study, from a ridge just to the south of the Aragonese Deep, rather than from the depocenter of the basin (Fig. 1d), so that it would dominantly reflect the background sedimentation recording the climatic conditions rather than being dominated by turbidites. To achieve reliable stratigraphical correlations, sedimentary events should be excluded from the sequences, and correlations should be done on composite profiles representing background sedimentation (Arnaud et al., 2002; Schwab et al., 2009; Avs¸ar et al., 2015; Moernaut et al., 2017). After excluding the sedimentary events, we correlated the composite Sr/Ca profiles of all the cores to the composite Sr/Ca profile of core 11 (Fig. 6). As the sedimentation rates are different between basins, this resulted in squeezing or stretching of the core records. This calibration was done using 5 to 8 characteristic reference levels that can be recognized in all the cores, which we used as tie-points to ensure consistency between the cores through the calibration process. (Gray lines in Fig. 6). Detailed explanation of the procedure for the removal of the sedimentary events and chemo-stratigraphical correlation can be found in Supplementary Material (E-SUPP 2). The original and modified depths of these tie-points are also presented as bi-plots next to each correlation plot in Fig. 6. Almost linear and smooth appearances of these bi-plots confirm that depth modifications did not result in abnormal sedimentation rates. For five cores (1, 6, 8, 12, and 13), the chemostratigraphical correlations result in almost perfect overlaps with core 11. For most of the other cores, although exact overlaps are not achieved, similarities between general trends and fluctuations along the cores are still within the range of uncertainties and thus are deemed acceptable. The discrepancies are probably due to the semi-quantitative nature of ITRAX µ-XRF scanning data or local differences in sedimentation. In addition, cores 11 and 13 were collected by a piston corer, which sometimes causes a disturbance and sediment loss close to the water/sediment interface. However, a good correlation between cores 11 and 12, which was taken by a multicorer ensuring undisturbed recovery of water/sediment interface, confirms that there was no sediment loss at the top of core 11. On the other hand, correlation between cores 11 and 13 shows that approximately 15 cm-thick sediment was lost at the top of core 13 during coring operation. A close-up view of the correlation between cores 13 and 11 is also provided in Fig. 6. Correlation between cores 17 and 18 is based on the stratigraphical order of the events since no µ-XRF data is available for core 18.

The raw ¹⁴C dates listed in Fig. 4 were included in the OxCal P_Sequence code (E-SUPP 3) with respect to their modified depths obtained by the chemostratigraphical correlations, as if all the radiocarbon samples came from core 11. The k value for the P_Sequence function was selected as 3, which is generally used for deep sea environments where hemipelagic sedimentation can be assumed rather constant (e.g., Polonia et al., 2023). Although the dates show a reasonable trend along core 11 (Fig. 7), two dates, the one from the bottom of core 7 (7–38.70) and the youngest date from core 9 (9–33.04), are clearly older than the general trend. Hence, we interpret them as reworked material and did not use them in the P_Sequence code. The resulting age-depth model, which is presented with 68% and 95% confidence intervals in Fig. 7, yields approximate sedimentation rates of 0.35, 0.16 and 0.22 mm/yr for the intervals of 0–32 cm, 32–55 cm and 55–80.5 cm, respectively. All of the sedimentary events projected on core 11 are also plotted with respect to depth in Fig.7, and are included in the P_Sequence code (E-SUPP 3), so that probability density functions (PDFs) for each event detected in Gulf of Aqaba can be calculated.

Discussion

from Bektaş et al. (2024:10-13)

PDFs of Type I and Type II events in the cores through the gulf are plotted with respect to calendar dates in Fig. 8, which were obtained according to the age-depth relation presented in Fig. 7. The probability density values for Type II events were multiplied by 0.5 since they may also be floods. Type III events, which are most probably not related to earthquakes, were not included in this plot. In Fig. 8, the plot of summed PDFs is also presented in order to statistically express the spatial extent and coevality of turbidites in the gulf; i.e., multiple coeval turbidites should be seen as distinct anomalies on the summed PDFs profile. Several sedimentary events appear to be coherent for several cores, including cores located in different basins. Hence, they are seen as anomalies in the summed PDFs profile (e.g., around late 10th, 16th and 20th centuries CE), and we interpret these events as the signature of past earthquakes that triggered turbidites (Type I or II) in the Gulf of Aqaba.

The turbidites dated to the late 20th century are unambiguously seen as Type I events at the tops of cores in Aragonese and Eilat Deeps (Cores 11, 12, 14, 15, 16, 17, and 18) and as Type II event in Dakar Deep (Core 10). They constitute a perfect benchmark of our sedimentary system as they are almost certainly the sedimentary traces of the 1995 Nuweiba (M_W 7.2) earthquake, the most recent major earthquake in the Gulf of Aqaba. Given that the Aragonese Fault, and probably partially the Eilat Fault as well, were the source faults for the 1995 Nuweiba earthquake and that the earthquake rupture propagated northward (Baer et al., 2008; Hofstetter, 2003; Klinger et al., 1999; Shamir et al., 2003; Ribot et al., 2021), it is not surprising that turbidites (Type I events) from this earthquake are not found in cores from the southern part of the gulf (i.e., cores 1 to 10). Despite the proximity to the epicenter of the 1995 earthquake, no turbidite associated with this event is visible in core 13 due to sediment loss at the top of that core (Fig. 6).

Among older turbidites, two events stand out and are recognized almost in every core, one in about the early 12th century CE and one in about the late 16th century CE. Starting from the oldest turbidites, one series of turbidites are clearly visible as Type I events in all cores (except core 11), which are long enough to cover at least the last millennium (Fig. 8). Since core 11 is not a turbidite-targeting core, and it was collected from a ridge rather than from a depocenter, it can be expected to see only limited evidence for turbidites in this core. The significant anomaly on the summed PDFs profile, near the beginning of the 12th century CE is most probably associated with the only major earthquake known in the region during that period: the CE 1068 earthquake. The temporal discrepancies observed for dates of turbidites in some cores are most probably related to the difficulties of the inter-core chemostratigraphical correlations due to intense bioturbation in the sediments of the Gulf of Aqaba that can lead to small time shifts when it comes to estimate exact age of specific core sections. The 18 March 1068 Aqaba-Hijaz earthquake was reported to have devastating effects in many locations from the city of Aila (Eilat) to Medina and Cairo (Ambraseys, 2009). Presence of seismo-turbidites all along the gulf in addition to evidence found at the onshore paleoseismic sites north of the Gulf (Amit et al., 1999, 2002; Zilberman et al., 2005; Klinger et al., 2015; Kanari et al., 2020) indicate that the CE 1068 earthquake was a major earthquake in the region that ruptured the Eilat, Aragonese, Arnona and probably Tiran faults together, in addition to the southernmost part of the Wadi Arabah Fault near the gulf. Hence, the total rupture length of this earthquake could have been at least ~200 km.

Similar as for the CE 1068 event, numerous coeval turbidites along the Gulf of Aqaba are represented by the anomaly on the summed PDFs profile about late 16th century CE, which is consistent with the well-documented earthquake of 4 January 1588 (Ambraseys, 2009). Majority of them are of Type I events, except the ones in core 2 in Tiran Deep, and in cores 9 and 10 in Dakar Deep, which are of Type II. The absence of turbidites in core 6 can be explained by the fact that this core was collected from a small and isolated basin (Fig. 1e). Indeed, this basin is isolated from turbidity flows that would come from the main slopes of the gulf, and its relatively smaller slopes may not be sensitive to earthquake shaking as much as the larger main slopes of the gulf. An absence of a turbidite in core 14 is more difficult to explain. One possible explanation could be the sediment clearance on the slopes and the banks of the submarine channels during the preceding earthquake, i.e., in CE 1068, leaving nothing to be wasted during the CE 1588 earthquake. Furthermore, another possibility is that the turbidity flows due to 1588 earthquake might have bypassed the location of core 14 (Goldfinger et al., 2017). Apart from turbidite absence in these three cores, it seems that the CE 1588 earthquake triggered seismo-turbidites along the entire gulf. This earthquake is also known as a devastating event in the historical records, affecting many places from the cities of Eilat and Aqaba at the northern end of the gulf to Cairo in Egypt (Ambraseys, 2009). A second event is also reported in the historical chronicles that happened on 7 April 1588, which was felt in Cairo and in the northern Red Sea. Although it cannot be ruled out that this second event was completely independent of the event in January 1588, it could be an aftershock of the former event (Ambraseys, 2009), triggering additional turbidites for example along the Tiran fault section. Klinger et al. (2015) report no surface rupturing evidence for the CE 1588 earthquake at the Qatar trench site (x in Fig. 1b). On the other hand, Kanari et al. (2020) report both rupturing and paleoliquefaction evidences that could be related to this earthquake in the Eilat Sabkha (t in Fig. 1b). Hence, based on our seismo-turbidite observations, it appears that the CE 1588 earthquake likely also ruptured the entire fault system in the Gulf of Aqaba. Unlike in CE 1068, however, the 1588 rupture does not seem to have propagated inland beyond the northern end of the gulf. In the south, based on turbidites we can trace the rupture to the south of the Tiran Strait and it may have ruptured even further to the south.

Some additional sequences of turbidites are found in our record, although they are not as extensive as the sequences associated respectively to the CE 1068 and CE 1588 earthquakes. According to the historical records, the CE 1212 earthquake caused widespread damage in an extensive area from Al-Shaubak and Al-Karak in the north (ca. 150 km north of Eilat and Aqaba cities) to the St. Catherine Monastery in the south (ca. 50 km east of Dahab) (Ambraseys, 2009). Although Klinger et al. (2015) reported evidence of surface rupturing along the southern Wadi Arabah Fault (Qatar Site, x in Fig.1b) that could be related to the CE 1212 earthquake, the trenching studies closer to the gulf (Zilberman et al., 2005; Kanari et al., 2020; y and t in Fig. 1b, respectively) claim that the CE 1212 earthquake was likely generated by a secondary fault on the eastern edge of the Eilat depression, rather than by the Wadi Arabah Fault. The probable turbidites of the CE 1212 earthquake are only seen in cores 17 and 18 from the northernmost part of the gulf (Fig. 8). Thus, it appears as a minor anomaly in the summed PDFs profile. The turbidite in core 18 is thicker and better preserved with its laminated internal structure. Furthermore, there are no turbidites in cores 14 and 15 around 1200s. Given the absence of consistently coeval turbidites through the gulf that are dated to the beginning of the 13th century (except the one in core 13), and only two turbidites in the northernmost cores (17 and 18), we conclude that the CE 1212 earthquake was likely significantly smaller than the CE 1068 and CE 1588 events. The exact location of this event north of the gulf remains uncertain, but it might have caused a discontinuous rupture on fault segments of the Wadi Arabah Fault system.

Prominent turbidites with well-preserved internal structures are seen around the mid-19th century CE in cores 2, 3 and 4. No turbidites were found around this period in the neighboring cores, except Type II events in cores 1, 7, and 8. The summed PDFs profile has a minor anomaly due to these turbidites that are limited to the most southern part of the gulf. These turbidites are probably related to the CE 1839 earthquake, which is described in the historical records as causing minor damage on the walls of St. Catherine Monastery (Ambraseys, 2009). Furthermore, Purkis et al. (2022) reported an incipient submarine landslide on the southeastern slopes of Tiran Deep that failed within the last 500 years, plausibly triggered by the CE 1839 earthquake. Given that the turbidites in cores 2, 3, and 4 are limited only in Tiran Deep, it is likely that the CE 1839 earthquake only partially ruptured either the Tiran or Arnona fault, or activated one of the secondary faults in the southernmost part of the gulf. Except for the ones related to the CE 1839 earthquake, the absence of distinct coeval turbidites in cores 1 to 10 (southern half of the gulf) since CE 1588 implies that most of the Tiran and Arnona faults has not ruptured since then.

While the cores provide a detailed record for the past millennium, only a few cores provide information for earlier periods. Around the middle 5th century CE, there are implications of coeval turbidites in cores 9, 10, and 11 in Dakar and Aragonese basins. Klinger et al. (2015) report evidence of surface rupturing at the Qatar trenching site on the Wadi Arabah Fault between 9 BCE and CE 492 (Fig. 8), which they attribute to the CE 363 earthquake that affected significantly the southern part of the Dead Sea Fault (Thomas et al., 2007; Ambraseys, 2009). However, this event could not be found north of the Qatar site (Lefevre et al., 2018). Hence, the turbidites around the middle 5th century CE could be due to a large earthquake rupturing several faults in the gulf, in addition to a 50 km long on-shore fault section. In core 11, three more Type I events, which can be other earthquakes in the gulf, are dated to ca. 250, 850 and 1350 BCE. The one around 250 BCE temporally correlates well with the event between 338 and 213 BCE, detected by Klinger et al. (2015) (Fig. 8). Although Klinger et al. (2015) do not report any event around 850 BCE, they report two events occurred during the period 2797–1245 BCE, which also coincide with the date of the turbidite ca. in 1350 BCE.

From the oldest turbidite to the youngest in 1995, the recurrence intervals for major earthquakes in the gulf seem to vary between 400 and 700 years (Fig. 8), with a mean value of 560 years. According to the paleoseismic trenching studies conducted in the north of the Gulf of Aqaba region, the Dead Sea Fault is characterized by earthquake clusters lasting 100–200 years, followed by seismic quiescence periods of 350–400 years (Klinger et al., 2015; Lefevre et al., 2018). On the other hand, Lefevre et al. (2018) have proposed that recurrence intervals for events should be longer due to partial stress partitioning between offshore and on-land fault segments, which is consistent with the 560 years recurrence interval found in the Gulf of Aqaba. In addition, the left-lateral relative plate motion between the Sinai and Arabian plates at this latitude is almost 5 mm/year (Castro-Perdomo et al., 2022; Viltres et al., 2022), which yields about 2 to 3.5 m fault slip deficit in a 400–700 year period, corresponding to an expected surface rupture length of 100–180 km (Wells and Coppersmith, 1994), which would be consistent with multi-segment rupturing events in the gulf. However, recent studies based on InSAR and GNSS observations (Li et al., 2021; Castro-Perdomo et al., 2022) report fault-locking depths that become shallower towards the south and possibility of partial fault creep along the southernmost fault strands of the gulf. Still, our results indicate that the entire Gulf of Aqaba fault system was activated in the 1068 and 1588 earthquakes and probably during the previous major earthquakes. This implies that the southern gulf can be regarded as being close to the end of the earthquake cycle as it has not ruptured in a major earthquake for more than 400 years.

Conclusions

from Bektaş et al. (2024:13)

Based on ITRAX µ-XRF scanning, radiographic imaging, magnetic susceptibility measurements and grain-size measurements, we detected a total of 86 sedimentary events in 18 sediment cores collected from the Gulf of Aqaba. Of these events, 46 were classified as distinct turbidites, 9 events described as thick flooding sequences due to high precipitation periods lasting probably several decades, and the remaining 31 less distinct events as turbidites or floods. Careful chemostratigraphical inter-core correlations and radiometric dating of these events provide a robust submarine paleoseismic record for the last millennium. The results show that the historical earthquakes in 1068 and CE 1588 were major characteristic earthquakes in the gulf that probably ruptured all the main faults (Tiran, Arnona, Aragonese and Eilat faults) in the gulf. On the other hand, the historical earthquakes of 1839 and CE 1212 were smaller and triggered only local turbidites in the southernmost and northernmost parts of the gulf, respectively. Information on older major events, together with the 1068 and CE 1588 earthquakes, suggests a recurrence interval of 400–700 years (average = 560 years), indicating that the southern gulf is a ripe for a major earthquake.

Fig. 2 - Radiographic Images of Type I events (prominent and well-preserved turbidites)

Normal Size

from Bektaş et al. (2024)

Fig. 2

U-channel radiographic images of prominent and well-preserved turbidites (Type I; red bars) in the sediments of the Gulf of Aqaba. Note the bioturbation and biogenic content (e.g., carbonaceous shells seen as whitish spots) in the background sedimentation. Event labeling is given in the lower left of each image.

Bektaş et al. (2024)

Magnified

from Bektaş et al. (2024)

Fig. 2

U-channel radiographic images of prominent and well-preserved turbidites (Type I; red bars) in the sediments of the Gulf of Aqaba. Note the bioturbation and biogenic content (e.g., carbonaceous shells seen as whitish spots) in the background sedimentation. Event labeling is given in the lower left of each image.

Bektaş et al. (2024)

Fig. 3 - Radiographic Images of Events of Types I (turbidites), II (turbidites or flood deposits), and III (flood deposits)

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from Bektaş et al. (2024)

Fig. 3

U-channel radiographic images of different types of sedimentary events.

Type I: Turbidites (red bars)
Type II: Turbidite or Flooding (gray bars)
Type III: Thick Flooding Sequence (yellow bars)

Bektaş et al. (2024)

Magnified

from Bektaş et al. (2024)

Fig. 3

U-channel radiographic images of different types of sedimentary events.

Type I: Turbidites (red bars)
Type II: Turbidite or Flooding (gray bars)
Type III: Thick Flooding Sequence (yellow bars)

Bektaş et al. (2024)

Fig. 4 - Sedimentary Events detected in all cores

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from Bektaş et al. (2024)

Fig. 4

Sedimentary events detected in the radiographic images of the Gulf of Aqaba cores, labelled by letters and indicated by red, gray, and yellow vertical bars. Depths of ²¹⁰Pb and ¹⁴C measurements are also shown by green and blue rectangles next to the images, respectively. In the table, raw and calibrated ¹⁴C results are listed. The raw and composite depths, and the depths corresponding on core 11 (after stratigraphical correlation) for each ¹⁴C sample are also given. In the lower right, results of ²¹⁰Pb_ex measurements on cores 3, 7, and 17 and the corresponding sedimentation rates (SR) are presented.

Bektaş et al. (2024)

Magnified

from Bektaş et al. (2024)

Fig. 4

Sedimentary events detected in the radiographic images of the Gulf of Aqaba cores, labelled by letters and indicated by red, gray, and yellow vertical bars. Depths of ²¹⁰Pb and ¹⁴C measurements are also shown by green and blue rectangles next to the images, respectively. In the table, raw and calibrated ¹⁴C results are listed. The raw and composite depths, and the depths corresponding on core 11 (after stratigraphical correlation) for each ¹⁴C sample are also given. In the lower right, results of ²¹⁰Pb_ex measurements on cores 3, 7, and 17 and the corresponding sedimentation rates (SR) are presented.

Bektaş et al. (2024)

Fig. 5 - Core Logs

Normal size

from Bektaş et al. (2024)

Fig. 5

Magnetic susceptibility (MS), sand content, and Zr/Sr profiles produced along selected cores (at least one core per basin) through the Gulf of Aqaba. Sedimentary events shown in Fig. 4 are also shown as horizontal bars in this figure. Event labeling and color code for different types of events are the same as in Fig. 4. See the Supplementary Material (E-SUPP 1) for detailed descriptions of sedimentary events for all of the cores.

Bektaş et al. (2024)

Magnified

from Bektaş et al. (2024)

Fig. 5

Magnetic susceptibility (MS), sand content, and Zr/Sr profiles produced along selected cores (at least one core per basin) through the Gulf of Aqaba. Sedimentary events shown in Fig. 4 are also shown as horizontal bars in this figure. Event labeling and color code for different types of events are the same as in Fig. 4. See the Supplementary Material (E-SUPP 1) for detailed descriptions of sedimentary events for all of the cores.

Bektaş et al. (2024)

Fig. 6 - Chemostratigraphical correlations

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from Bektaş et al. (2024)

Fig. 6

Chemostratigraphical correlation of composite Sr/Ca ratio profiles (event-free) of each core (red curves) to core 11 (black curves). Depths of sedimentary events and radiocarbon dates are shown as dots and blue rectangles, respectively. Bi-plots next to each graph show the original composite depths (y-axes) versus the modified depths on core 11 (x-axes) of the tie-lines (gray lines). Details of the chemostratigraphical correlation procedure are presented in the Supplementary Material (E-SUPP 2).

Bektaş et al. (2024)

Magnified

from Bektaş et al. (2024)

Fig. 6

Chemostratigraphical correlation of composite Sr/Ca ratio profiles (event-free) of each core (red curves) to core 11 (black curves). Depths of sedimentary events and radiocarbon dates are shown as dots and blue rectangles, respectively. Bi-plots next to each graph show the original composite depths (y-axes) versus the modified depths on core 11 (x-axes) of the tie-lines (gray lines). Details of the chemostratigraphical correlation procedure are presented in the Supplementary Material (E-SUPP 2).

Bektaş et al. (2024)

Fig. 7 - Age-Depth Model

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from Bektaş et al. (2024)

Fig. 7

Calibrated radiocarbon dates, and the age-depth model determined by OxCal P_Sequence function. Two reworked samples (9–33.04 and 7–38.70) were not included in the P_Sequence code, which can be found in the Supplementary Material (E-SUPP 3). The list and depths of all sedimentary events are also shown.

Bektaş et al. (2024)

Magnified

from Bektaş et al. (2024)

Fig. 7

Calibrated radiocarbon dates, and the age-depth model determined by OxCal P_Sequence function. Two reworked samples (9–33.04 and 7–38.70) were not included in the P_Sequence code, which can be found in the Supplementary Material (E-SUPP 3). The list and depths of all sedimentary events are also shown.

Bektaş et al. (2024)

Fig. 8 - Probability Distribution Functions (PDFs) of Type I and Type II events along the Gulf of Aqaba

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from Bektaş et al. (2024)

Fig. 8

Plot of Probability Distribution Functions (PDFs) of Type I and Type II events along the Gulf of Aqaba. Horizontal black lines mark the dates of historical earthquakes in the region. Major earthquakes that triggered extensive turbidites all along the gulf are seen as major anomalies on the summed PDFs profile. The scale of x-axis changes around CE 800. Time windows for the prehistorical surface rupturing events at the Qatar trenching site (x in Fig. 1, Klinger et al., 2015) are also shown. Note the recurrence intervals varying between 400 and 700 years and also the absence of extensive coeval turbidites in the southern half of the gulf since CE 1588.

Bektaş et al. (2024)

Magnified

from Bektaş et al. (2024)

Fig. 8

Plot of Probability Distribution Functions (PDFs) of Type I and Type II events along the Gulf of Aqaba. Horizontal black lines mark the dates of historical earthquakes in the region. Major earthquakes that triggered extensive turbidites all along the gulf are seen as major anomalies on the summed PDFs profile. The scale of x-axis changes around CE 800. Time windows for the prehistorical surface rupturing events at the Qatar trenching site (x in Fig. 1, Klinger et al., 2015) are also shown. Note the recurrence intervals varying between 400 and 700 years and also the absence of extensive coeval turbidites in the southern half of the gulf since CE 1588.

Bektaş et al. (2024)

Individual Core Logs

Core 1

from Bektaş et al. (2024) Supplemental