• from Lu et al. (2020:2)

• from Lu et al. (2020:2)

Similar to the structures preserved in the Dead Sea margin outcrops, the folded layers in the drilling core from the Dead Sea depocenter also appear as folded aragonite-detritus laminae in various forms of (i) linear waves (Fig. 2, A and B), (ii) asymmetric billows (Fig. 2, C to F), and (iii) coherent vortices (Fig. 2, G to J) (18). These delicate aragonite laminae are well preserved and can be traced in the strata, indicating that the layers are deformed in situ and have not undergone any notable transportation. That is, notable transportation would disaggregate and destroy the delicate submillimeter-thick aragonite laminae. Layer-parallel displacements characterize these in situ folded layers. The ICDP Core 5017-1 is positioned in the center of the Dead Sea abyssal plain. This makes improbable postdepositional causes for layer-parallel shears such as sloping substrates or downhill water flow above the sediments. In total, we identify 367 in situ folded layers in the ICDP Core 5017-1 (table S2). Figure S2 shows more examples of in situ folded layers in the drilling.

• from Lu et al. (2020:2-3)

Similar to the structures preserved in the Dead Sea margin outcrops, this type of layer from the Dead Sea depocenter consists of mixed aragonite-detritus laminae fragments (Fig. 2, K to M). The in situ deformation process of this type of layer is recognized by (i) the lack of erosion processes at the base of the layer and (ii) some remaining parts of in situ folded layer can be observed directly in the base of the intraclast breccia layer (Fig. 2, K to M, and fig. S3). Together, these features provide evidence for the evolution of the intraclast breccia texture under seismic shaking and differentiate it from any other detrital layers with laminae fragments that formed by secondary sedimentary processes such as mass transport. In total, we identify 46 intraclast breccia layers in the ICDP Core 5017-1 (table S2). Figure S3 shows more examples of intraclast breccia layers in the core.

Our observations reveal that some in situ deformed layers have been re-deformed by a latter deformation within a short core section. Using geophysical and chemical datasets and careful sedimentary structure analysis, we recognize these redeformed seismites. The potential uncertainties in thickness measurements (for redeformed seismites) that are induced by the redeformation range from a few millimeters to several centimeters and therefore have no notable effects on shaking intensity estimation. In addition, regarding the latter in situ deformations, the shape and thickness of total folded sediments constrain the intensity of seismic shaking regardless of the type and thickness of the redeformed seismites. In total, we identify 413 independent seismic shaking markers from the Dead Sea center.

• from Lu et al. (2020:3)

We update the previous computational fluid dynamics modeling of Wetzler et al. (18) by extending the upper limit of layer thickness and ground acceleration from 0.5 m and 0.6g to 1.0 m and 1.0g, respectively (Fig. 3; Materials and Methods). We run a series of two-dimensional numerical simulations using the physical properties of the soft sediments at the bottom of the Dead Sea based on the Kelvin-Helmholtz instability mechanism. According to the numerical simulations, formation of a layer of (i) linear waves, (ii) asymmetric billows, (iii) coherent vortices, and (iv) intraclast breccia requires a minimum acceleration of 0.13, 0.18, 0.34, and 0.50g, respectively (Fig. 3). Also, our numerical simulations indicate that the thickness of the deformed layer scales with acceleration (Fig. 3 and fig. S4). Because we could not directly determine the Reynolds number of an individual unstable layer at the time of seismic shaking, we constrain only the lower boundary of acceleration needed to initiate deformation of a layer with a certain thickness. By considering both the shape and thickness of the deformed layers, we identified 18, 67, 139, 141, 240, and 413 events with acceleration of ≥0.65g, ≥0.50g, ≥0.34g, ≥0.26g, ≥0.18g, and ≥0.13g, respectively (table S2).

• from Lu et al. (2020:3-4)

Because of the intrinsic geometrical spreading, dissipation, and possibly dispersion of the seismic energy, ground motion effects of a moderate earthquake nearby generate similar shaking intensities to those generated by a large earthquake farther away. Therefore, determining the location of an earthquake along the length of the Dead Sea Fault or on other nearby faults is the key for the magnitude constraint. It follows that magnitude estimation based on a single station is difficult and dependent on the location of possible source faults. For this sinistral boundary between the African and Arabian plates, we model the potential source region of earthquakes as having a fixed width and a length of a few hundred kilometers (Fig. 1). The nearest faults are ~5 km from the Core 5017-1 drilling site.

We apply three empirical attenuation relations (28–31) developed for the Dead Sea region to constrain the lower magnitude limit of paleoseismic events that the ICDP Core 5017-1 records (Materials and Methods). Among them, two attenuation relations (28–30) are described by macroseismic intensity, magnitude, and epicentral distance (D), and one relation (31) is described by peak ground acceleration (PGA), magnitude, and epicentral distance. To apply the two attenuation relations that are described by seismic intensity, we convert accelerations into seismic intensity via the linear relationships between PGA and the modified Mercalli intensity scale (MMI) that Wald et al. (32) proposed (table S2). According to the three regional empirical attenuation relations, PGA ≥ 0.13g or MMI ≥ VI½ corresponds to Mw ≥ 5.5, 5.3, and 5.7, by taking Dmin = 5 km (table S3). Therefore, we interpret the lower-bound magnitude of the recorded PGA ≥ 0.13g (MMI ≥ VI½) events as Mw ≥ 5.3. Considering that the incompleteness of moderate earthquakes in the paleoseismic record, we infer that the mean recurrence of Mw ≥ 5.3 earthquake on the central Dead Sea Fault Zone is <530 ± 40 years. This value is much shorter than the previously obtained mean recurrence of ~1600 years for the same magnitude, based on the paleoseismic record between 60 and 14 ka ago (4).

• from Lu et al. (2020:4-5)

The ICDP Core 5017-1 records 139 strong seismic shaking events with PGA ≥ 0.34g (MMI ≥ VIII) (table S2). According to the regional empirical attenuation relations (28–31), one felt intensity, for example, MMI ≥ VIII (PGA ≥ 0.34g) at the drill site, will require a moderate earthquake (6.0 < Mw < 7.0) with a D between 5 and 30 km, an Mw ≥ 7.0 earthquake with a D of 30 km, or an Mw ≥ 8.0 earthquake with a D of up to 150 km. We adopt a maximum magnitude in the region of Mw 8.0, which will require a rupture of ~300 km along a major fault. Therefore, we consider only large earthquakes with D ≤ 150 km (the central Dead Sea Fault Zone) as triggers of the strong seismic shaking events (MMI ≥ VIII) that Core 5017-1 records. On this slow-slipping plate boundary, the Dead Sea Fault (zone) is the only real contributing fault because most of the transform margin slip rate is on the Dead Sea Fault, and the other faults were either far away or had low slip rates. Within the Dead Sea region, no other major faults have sufficient length and slip rate to materially contribute to the rate of Mw ≥ 7 earthquakes. The historic earthquake catalog from the region shows that the 1822, 1712, 1408, 1170, 1139/1140, and 859/860 CE Mw ≥ 7.0 earthquakes occurred with distance to the drill site ≥300 km north of the Dead Sea (the northern Dead Sea Fault Zone). However, none of them have an expression in the Dead Sea Core 5017-1 record.

The spatial distribution of instrumental and historic moderate and large earthquakes on the central Dead Sea Fault Zone during the past 2 ka do supply additional clues for magnitude constraint for these strong seismic shaking events. The instrumental (33) and historical (5, 34–36) earthquake catalogs reveal that during the past 2 ka, all major earthquakes (Mw ≥ 6.0) occurred with D ≥ 30 km from the drilling site (Fig. 1B). By taking the past 2 ka earthquake scenario as an analogy for the paleoseismic record, we assume that most Mw ≥ 6.0 earthquakes occurred with D ≥ 30 km from the drilling site. Under this basic assumption and the three regional empirical attenuation relations, (i) an intensity of MMI ≥ VIII (PGA ≥ 0.34g) requires an earthquake with Mw ≥ 7.0, ≥7.0, and ≥7.3; (ii) an intensity of MMI ≥ VIII½ (PGA ≥ 0.50g) requires an earthquake with Mw ≥ 7.4, ≥7.3, and ≥7.6; and (iii) an intensity of MMI ≥ IX (PGA ≥ 0.65g) requires an earthquake with Mw ≥ 7.8, ≥7.6, and ≥7.8 (Fig. 4E and table S3). Therefore, we interpret the corresponding lower-bound magnitudes of strong seismic shaking events in the ICDP Core 5017-1 to be Mw ≥ 7.0, 7.3, and 7.6, respectively.

We test our magnitude conversion versus known historic earthquakes on the central Dead Sea Fault Zone. Six seismites in Core 5017-1 dated at −2 ± 44 years before the present (yr B.P.), 42 ± 44 yr B.P., 148 ± 44 yr B.P., 1248 ± 44 yr B.P., 1555 ± 47 yr B.P., and 1626 ± 47 yr B.P. correspond to the 1956 CE (Mw 5.5; D, ~5 km), 1927 CE (Mw 6.25; D, ~30 km), 1834 CE (Mw ~ 6; D, ~60 km), middle 8th century (Mw > 7; D, ~100 km), 419 CE (Mw ~ 6; D, ~40 km), and 363 CE (Mw ~ 6.8; D, ~70 km) earthquakes, respectively (table S4) (36). According to the regional empirical attenuation relations, we constrain the magnitudes of the six paleoearthquakes (seismites) with intensities of VII (0.18g), VI½ (0.13g), VI (0.09g), VII (0.18g), VI (0.09g), and VII (0.18g) as Mw 5.6, Mw 6.1, Mw 6.2, Mw 7.1, Mw 6.0, and Mw 6.9, respectively, which are in line with recorded historic magnitudes (table S4). This test supports our magnitude conversion based on the regional empirical ground motion attenuation relations.