Description	Image	Source
Entire En Gedi Trench	Entire En Gedi Trench - cleaned a bit more than on 16 Feb. 2023 Photo by Jefferson Williams 17 Feb. 2023	Jefferson Williams
Entire En Gedi Trench - closer in	Entire En Gedi Trench Photo by Jefferson Williams 16 Feb. 2023	Jefferson Williams
Entire East Section of En Gedi Trench	Entire East Section of En Gedi Trench Photo by Jefferson Williams 17 Feb. 2023	Jefferson Williams
Entire West Section of En Gedi Trench	Entire West Section of En Gedi Trench Photo by Jefferson Williams 17 Feb. 2023	Jefferson Williams
Entire Middle Section of En Gedi Trench	Entire Middle Section of En Gedi Trench Photo by Jefferson Williams 17 Feb. 2023	Jefferson Williams
Top of Middle Section of En Gedi Trench	Top of Middle Section of En Gedi Trench Photo by Jefferson Williams 17 Feb. 2023	Jefferson Williams
Middle 01 of Middle Section of En Gedi Trench	Middle 01 of Middle Section of En Gedi Trench Photo by Jefferson Williams 17 Feb. 2023	Jefferson Williams
Middle 02 of Middle Section of En Gedi Trench	Middle 02 of Middle Section of En Gedi Trench Photo by Jefferson Williams 17 Feb. 2023	Jefferson Williams
Bottom of Middle Section (Long shot) of En Gedi Trench	Bottom of Middle Section (Long shot) of En Gedi Trench Photo by Jefferson Williams 17 Feb. 2023	Jefferson Williams
Bottom of Middle Section (Medium shot) of En Gedi Trench	Bottom of Middle Section (Medium shot) of En Gedi Trench Photo by Jefferson Williams 17 Feb. 2023	Jefferson Williams
Bottom of Middle Section (closeup) of En Gedi Trench	Bottom of Middle Section (closeup) of En Gedi Trench Photo by Jefferson Williams 17 Feb. 2023	Jefferson Williams

• Click Scaniverse icon in lower left to open full screen lidar scan in a new tab

• from Kagan et al. (2011)

• from Migowski et al (2004)

• from Migowski (2001)

• from Neugebauer at al (2015)

• from Neugebauer at al (2015)

• from Neugebauer at al (2015)

• from Neugebauer at al (2015)

• from Williams et al. (2011:1-2)

The Dead Sea (31◦30 N, 35◦30 E) lies along the tectonically active Dead Sea Transform (DST), which separates the Arabian and Sinai plates (Garfunkel 1981). The DST is a mainly N–S-striking, left-lateral transform fault with normal faulting along its margins and at northwest bends and thrusting at northeast bends. A terminal lake, the Dead Sea, is situated in a pull-apart basin at the deepest location on land along the transform. Frequent seismic activity along the DST has been detected in the past century and recorded historically and archaeologically over the past 4000 years (Ben-Menahem 1991; Ambraseys et al. 1994; Salamon et al. 2003). Within the layered deposits of recent Dead Sea sediments lie subintervals which have been deformed, presumably due to earthquakes generated by fault movement along the DST (Marco and Agnon 1995; Enzel et al. 2000; Ken-Tor et al. 2001a; Migowski et al. 2004; Kagan et al. 2011).

In the fall of 1997, the GFZ German Research Centre for Geosciences in cooperation with the Geological Survey of Israel took three cores from the beach of the Ein Gedi Spa adjacent to the Dead Sea at a surface elevation of 415 m below sea level₁ (Figure 1). The cores sampled sediments that were originally deposited in a deep lacustrine environment (Migowski et al. 2004, 2006). A floating varve chronology was established (Migowski et al. 2004) after identifying and counting varves under the microscope and performing accelerator mass spectrometry (AMS) radiocarbon dating of wood fragments from the cores. Twenty-eight historically documented earthquakes were identified in a 1598-year interval between 140 BC and 1458 AD in Core Section A3 of the Ein Gedi core (Migowski et al. 2004).

While a previous study (Migowski et al. 2004) attempted to reconcile a varve-counted seismite date observed in the section to an earthquake date (33 AD) listed in the earthquake catalogues (e.g. Willis 1928; Amiran et al. 1994), this study makes no assumptions about the likely date of this early first-century seismite. A date was assigned to the seismite based on varve counting alone. Then, an attempt was made to determine the accuracy of that varve count and to compare this with an analysis of the historical sources which reveals a less well defined date assignment than the dominant 33 AD date that is present in most of the catalogues. By comparing the date from the varve count with the date range and date probabilities from the historical sources and conducting some geomechanical examination, we have come to some conclusions.

• from Williams et al. (2011:2-3)

The two fundamental assumptions that allow one to identify historically documented earthquakes in the Dead Sea sedimentary record of the Ein Gedi core are the following.

1. The sediments are varved (Heim et al. 1997; Migowski et al. 2004). Seasonal lamination patterns of white summertime precipitates (primarily aragonite, but also in a few cases gypsum and halite) and grey detritus from winter and spring time floods in the wadis (aka Nahal) can be counted as a varve, that is, 1 year of deposition.
2. Brecciated layers, also known as intraclast breccias (Agnon et al. 2006), mixed layers, or seismites, were created by deformation of the varved layers due to seismic shaking (Marco and Agnon, 1995, Marco et. al., 1996). An example containing annual varves and brecciated layers of Holocene Dead Sea sediments is shown in Figure 2.

Field observations of brecciated layers in cores and outcrops show that the upper contact is sharp whereas the basal contact can be sharp or gradual. Where the basal contact is gradual, folded and torn packets of laminae are abundant (Agnon et al. 2006).

There are several reasons that the brecciated layers are believed to have a seismic origin.

1. Field evidence. In Nahal Ze’elim, where lateral relations of the brecciated layers were readily observed, it was noted that the topography was flat at the time of deposition. Thus, there is no field evidence for gravitational slides. In addition, none of the aragonite fragments in the brecciated layers showed lateral grading, imbrications, or other transport indicators such as would be expected from lateral flows or turbidity currents (Ken-Tor et al. 2001a).
2. Similar layers are found elsewhere in the world. As noted by Ken-Tor et al. (2001a), similar soft sediment deformation structures have been documented in several other localities worldwide and interpreted as seismites (Sims 1973, 1975; Hempton and Dewey 1983; Allen 1986; Davenport and Ringrose 1987; Doig 1991).
3. Association with syndepositional faulting. In nearby Nahal Perazim, brecciated layers with similar lithology occur in association with syndepositional faults presumed to be caused by earthquakes (Marco and Agnon 1995; Marco et al. 1996).
4. Correlation to documented earthquakes. In the 1522-year varve-counting interval in the Ein Gedi core, brecciated layers were correlated to 28 historically documented earthquakes (Migowski et al. 2004). Seismites from other outcrops on the shores of the western Dead Sea have also been correlated to historically documented earthquakes (Enzel et al. 2000; Ken-Tor et al. 2001a, 2001b; Kagan et al. 2011).
5. Correlation to seismic intensity. In the Ein Gedi core, a relationship was discovered between estimated local intensity due to an earthquake and the probability that a brecciated layer would correlate to a seismic event. At historically estimated modified Mercalli intensities (MMI) greater than VII (i.e. VIII and higher), all well-documented earthquakes were correlated, whereas at intensities smaller than VI, none were matching (Migowski et al. 2004). In the Nahal Ze’elim outcrops, Williams (2004) independently estimated that the threshold intensity for seismic deformation is VIII. Williams was further able to develop a quantitative relationship between the historically estimated intensity of local ground shaking (expressed as peak horizontal ground acceleration) and the thickness of the brecciated layers themselves.

In the thinnest brecciated layers (e.g. less than 1 cm thick), the entire interval appears to have been fluidized, brecciated, suspended, and then redeposited after the seismic shaking ended (Migowski et al. 2004). In larger brecciated layers (several tens of centimetres thick and larger), the brecciated layers appear to have been deformed in situ, avoiding suspension and resettlement. We suspect that the thinnest brecciated layers originally (preseismically) formed a thin veneer of uncompacted and possibly not fully grain-supported sediment that was mechanically closer to a suspension of water and detritus than the underlying sediment. When relatively lower intensity ground shaking occurred, these thin layers were then fully suspended into the water column and resettled. The thicker layers appear to have been formed by longer lasting and more intense levels of ground shaking.

A strain softening type of liquefaction apparently played a significant role in the formation of the larger brecciated layers. This type of liquefaction has been observed in low-permeability sediments such as marine clays (e.g. see Vucetic and Dobry 1991) and is considered to have been operative in the Dead Sea sediments deposited from a deeper lacustrine environment such as was found in the Ein Gedi core. In strain softening liquefaction, increases in strain cause a reduction in the shear modulus of the soil, which reduces to such a point that the soil can no longer resist the deforming forces. More conventional pore pressure-induced liquefaction may be more operative in coarser grained, higher permeability Dead Sea sediments deposited in shore and near-shore environments (e.g. Enzel et al. 2000). Heifetz et al. (2005) proposed a model for soft sediment deformation that agrees well with field observations.

An important mechanical aspect to the deformation of Dead Sea sediments during earthquakes involves the sediment anisotropy. Cemented aragonite crusts provide lateral reinforcement. In fact, the sediments are significantly stronger in cyclic load tests when the load is applied laterally rather than vertically (Sam Frydman, personal communication 2000). In Holocene Dead Sea sediments, the aragonite crusts appear to undergo brittle rather than plastic failure during seismic shaking. Seismic loading away from the immediate epicentre of an earthquake is usually dominated by vertically propagating shear waves, which load the sediments horizontally. This appears to have been the case in most of the observed brecciated layers where the brittle aragonite crusts evidently fractured due to horizontal forces (Williams 2004).

It appears that in the fine-grained lacustrine sediments present in the Ein Gedi core, brecciated layers were formed by either suspension of a thin veneer of uncompacted water and detritus or a strain softening type of liquefaction coupled with brittle failure of aragonite crusts. Whatever the specific mechanism of failure, brecciated layer formation can be visualized in a simplified form as shown in Figure 3. The important point is that the brecciated layers apparently formed at the sediment–water interface, and the timing of each event is constrained by dating the first undisturbed layer overlying the disturbed sequence (Marco and Agnon 1995; Migowski et al. 2004).

• from Williams et al. (2011:3-4)

Because of the ubiquity of sediment deformation due to the 31 BC earthquake throughout the Dead Sea (Reches and Hoexter 1981; Enzel et al. 2000; Ken-Tor et al. 2001a; Migowski et al. 2004; Kagan et al. 2011), once this event is identified in a given section, it can be treated as a chronological anchor (see Event B in Figure 2). Varve counting, for example, can proceed from the 31 BC event upward in the section towards more recent earthquake events. The primary historical source for the earthquake of the early spring of 31 BC is Josephus Flavius, who wrote in The Jewish War (Book 1, Chapter XIX, 370):

But as he [King Herod] was avenging himself on his enemies, there fell upon him another providential calamity; for in the seventh year of his reign, when the war about Actium₂ was at the height, at the beginning of the spring the earth was shaken, and destroyed an immense number of cattle, with thirty thousand men; but the army received no harm, because it lay in the open air. (Josephus et al. 1981)

This was evidently a powerful earthquake. Amiran et al. (1994) believe that local intensities were as high as X in several places, whereas Arieh (1993) assigned a maximum local intensity value of IX and ML = 7.0.₃ Ben-Menahem (1991) estimated ML = 6.7 and places the approximate epicentre ∼25 km north of where the Jordan River empties into the Dead Sea along the Jericho fault. Williams (2004) suggested that the fault break was most likely on the Jericho fault with a southern termination near Nahal Darga and a northern termination well up the Jordan Valley directly north of Gesher Adam (Jisr Damiya).

Ben-Menahem (1991) mentions damage in Jerusalem at the Second Temple, Masada, Qumran, and Jericho at Herod’s Winter Palace. Guidoboni et al. (1994) believes this earthquake is mentioned by Iohannes Malalas in Chronographia (Malalas et al. 1986) when he reported that a city in Palestine named Salamine (possibly present- day Lod, near Tel Aviv) was destroyed and rebuilt by Augustus and re-named Diospolis. Rahmani (1964) reports that Jason’s Tomb in Jerusalem was destroyed by this earthquake. Amiran et al. (1994) note that earthquake damage was severe in Galilee and Judea. Karcz (2004) and Ambraseys (2009), however, caution that there is limited textual evidence regarding the area affected by this earthquake and suggest that the extent of damage reported from archaeological sites in Israel (e.g. at Qumran, Masada, Jason’s tomb, and/or Jericho) due to a 31 BC earthquake may be overstated. Karcz (2004) estimates a magnitude of 6.0–6.5.

The 31 BC earthquake appears to have ruptured the ground surface near Jericho and shows up in trenches excavated by Reches and Hoexter (1981). A seismite from the 31 BC earthquake also appears to be present in outcrops at Nahal Darga as a thick deformed layer labelled Stratigraphic Unit 11 by Enzel et al. (2000). At Nahal Ze’elim, a brecciated layer labelled Event B was assigned to the 31 BC earthquake (Ken-Tor et al. 2001a). This brecciated layer is ∼17 cm thick and spatially continuous, appearing as shown in Figure 2. Kagan et al. (2011) assigned a 6 cm-thick brecciated layer in En Feshka to the 31 BC earthquake. Migowski et al. (2004) identified a 9 cm-thick brecciated layer with the 31 BC earthquake, and this layer is shown at the bottom of the thin-section log in Figures 4 and 6.

• from Williams et al. (2011:4-6)

Thin-section images along with interpretive tracks of the Ein Gedi core from the 31 BC earthquake to the early first- century earthquake are shown in Figures 5 and 6.

It should be noted that the dark cracks present in the thin-section slides were created during the epoxy impregnation process in creating the thin sections. Fresh sediment slices (10 × 2 cm) were impregnated with resin after freeze-drying (Brauer and Casanova 2001). The impregnated sample blocks were cut along the long axis so that the cutting plane could be used for large-scale thin-section preparation.

Excluding the depth track (in millimetres) on the far left, there are nine tracks in these logs. On the far left and far right are images of the thin sections themselves (slides A3-3-2 and A3-3-3). Adjacent to each microscope image is a varve quality track.

A varve quality index is defined below.

1. Discontinuous ambiguous clastic layer.
2. Clearly identifiable clastic layer but thickness estimate is not very accurate.
3. Well-preserved varve with good accurate estimate of thickness.

Every counted varve was assigned an index value of 1, 2, or 3. For the purpose of this study, where the goal is accurate chronological dating, a varve quality index value of 1 indicates that the varve is somewhat suspect and a varve quality index of 2 or higher indicates that the varve count is regarded as fairly certain.

In the 62 years counted from 31 BC to 31 AD, 36 (58%) had a varve quality rating of 1 and 26 (42%) had a varve quality rating of 2 or higher. In addition to the varve quality tracks, two more track types are present. They are the facies tracks and years tracks.

The facies tracks were developed for each individual slide (A3-3-2 or A3-3-3) and use simplified symbols geared towards the goal of counting varves and identifying earthquakes. Layers are defined, as shown in the facies legend of Figures 5 and 6, into one of three types of facies: clastic layers, evaporites, and brecciated layers. Evaporites halite can also be present.

A brecciated layer has undergone deformation due to ground motion (usually due to earthquakes) in the predominantly undisturbed and finely laminated lacustrine Ein Gedi sediments. A combination of one clastic layer (deposited primarily during winter) and one evaporite layer (deposited primarily during summer) is assumed to represent one varve or 1 year of deposition (Migowski et al. 2004). Varves inside of brecciated layers were counted based on observed discontinuous laminations and maintaining congruence in varve thickness with the average thickness of adjacent undeformed varves. Years tracks were constructed adjacent to the facies tracks based on this assumption. One years track was created for each individual thin section (A3-3-2 and A3-3-3), and a composite years track was created combining what was regarded as the most accurate varve counts between the two thin sections. Varve quality index values were used to decide which chronology to use where the thin sections overlapped (see Figures 5 and 6).

Palaeoclimate information may be contained in the sediments deposited immediately after the 31 BC earthquake. In his book Jewish Antiquities (Book XV, Chapter 9), Josephus reported a drought in Judea in 28 BC or possibly 25 BC₄:

Now on this very year, which was the thirteenth year of the reign of Herod, very great calamities came upon the country; whether they were derived from the anger of God, or whether this misery returns again naturally in certain periods of time (14) for, in the first place, there were perpetual droughts, and for that reason the ground was barren, and did not bring forth the same quantity of fruits that it used to produce. (Josephus 1930)

From 31 BC to 28 BC in Figure 4, we note that the aragonite layers are relatively thin and that there are a fairly large number of gypsum rhombs.₅ If these were years of drought, this would tend to support the thesis of Stein et al. (1997) and Barkan et al. (2001) that enhanced aragonite production requires a continuous supply of freshwater loaded with bicarbonate (Migowski et al. 2006), leading to the conclusion that thick aragonite layers were precipitated in the summers after years of heavier rainfall and abundant runoff into the Dead Sea, whereas thinner aragonite layers correspond to summer time precipitation following years of less rainfall and less runoff. In addition, the extra gypsum in these years may represent drier years, when the upper water mass of the Dead Sea was diminished due to lower water input and enhanced evaporation (Migowski et al. 2004).

Leroy et al. (2010) noted a decrease in cultivated pollen (Olea, Pistacia, Juglans, and Vitis) in the Nahal Ze’elim outcrop in the ∼5 years after the 31 BC earthquake, which could indicate drought conditions and/or a decrease in agricultural productivity due to destruction caused by the 31 BC earthquake.

• from Williams et al. (2011:7)

Migowski et al. (2004) assigned the brecciated layer in Figure 5 to an earthquake listed as occurring in 33 AD in the earthquake catalogues. Ken-Tor et al. (2001a), using the outcrops at Nahal Ze’elim, assigned a correlative seismite (labelled Event C) to the 33 AD earthquake. Kagan et al. (2011) also assigned 33 AD to an earthquake event identified in outcrops at En Feshka. All of these assignments could refer to an earthquake reported to have occurred immediately after the crucifixion of Jesus of Nazareth. The primary source document for the earthquake of the crucifixion is the 27th chapter of the Gospel of Matthew in the New Testament. It describes an earthquake occurring when Jesus of Nazareth died on the cross:

50But Jesus, again crying out in a loud voice, yielded up his spirit. 51 At that moment the curtain in the Temple was ripped in two from top to bottom; and there was an earthquake₆ with rocks splitting apart.

The curtain referred to comes from the Aramaic word parokhet, which was a 1 ft-thick piece of fabric covering the entrance to the holy of the holies in the Second Temple. The Gospels of Mark and Luke also mention the tearing of the temple curtain in the moments surrounding Jesus’ death, but do not cite an earthquake as the cause of destruction.₇ In Chapter 28, the Gospel of Matthew goes on to describe another earthquake roughly 36 hours after the one described above:

1After the Sabbath, toward dawn on Sunday, Mary of Magdala and the other Mary went to see the grave. 2Suddenly there was a violent earthquake, for an angel of God came down from heaven, rolled away the stone and sat on it.

In modern terms, this might be described as an aftershock event.

The day and date of the crucifixion are fairly well known. The year is not so well known. According to the four canonical gospel accounts (Matthew, Mark, Luke, and John), the crucifixion occurred on a Friday on either 14 or 15 Nisan, a month in the Jewish lunar calendar. The year, however, is not specified. One clue to the year is that the crucifixion occurred during the reign of Pontius Pilate who was the Procurator of Judea from 26 to 36 AD. This is agreed upon by all four gospels as well as Tacitus in Annals (Book XV, 44) (Tacitus et al. 1942).

Humphrey and Waddington (1983) tabulated the days between 26 and 36 AD when 14 or 15 Nisan fell on a Friday and came up with four possible years: 27, 30, 33, and 34 AD. Humphrey and Waddington (1983) further pointed out that 27 and 34 AD were unlikely dates when one tried to match the crucifixion with the time of Jesus’ ministry and the estimated date of Paul of Tarsus’ conversion on the road to Damascus.₈ Thus, they listed two dates as the most likely dates for the crucifixion: Friday 7 April 30 AD (14 Nisan) or Friday 3 April 33 AD (14 Nisan). They proposed that Friday 3 April 33 AD was the more probable of the two dates.

• from Agnon et al. (2006:197-198)

The following terms have been used in the literature to describe various types of deformed unconsolidated sedimentary layers associated with earthquakes:

• mixed layers (Marco and Agnon, 1995; Marco et al., 1996b): this term may cause confusion with a number of unrelated uses in the earth sciences;


• mixtites (Jackson and Bates, 1997): this term describes any clastic layer regardless of composition or origin; any flood deposit may fall in this category;


• homogenites (Kastens and Cita, 1981): the term stresses uniform composition of the deposit, yet does not account for systematic variations across the layer;


• seismites (Seilacher, 1969): this term is interpretative; and


• intraclast breccias (Marco and Agnon, 2005): we use this descriptive term to separate observations from interpretations: "intraclast" refers to the origin of the clasts being reworked from within the sedimentary section (Jackson and Bates, 1997), and "breccia" refers to the texture of the deposit.

• from Agnon et al. (2006:198)

Intraclast breccias are distinctive in sequences that are otherwise well-bedded or, better yet, laminated. Lamination is typical in the lacustrine facies of the Dead Sea deposits and makes recognition of intraclast breccias practical because of the conspicuous alternation between chemically precipitated white aragonite and darker detritus (Fig. 2) (Katz et al., 1977).

In the present context, intraclast breccias within a laminated sequence can be distinguished by the following criteria:

1. The primary mineralogical composition of an intraclast breccia is identical to the underlying strata (and typically the overlying layers also). Thus, the intraclast breccia appears similar to the enclosing deposits. The absence of fine-scale lamination observed from a short distance helps to recognize the intraclast breccia in the field.
2. Fragments of laminae may vary in distributions of size. Tabular fragments of competent laminae (with the long dimension commonly 1-5 mm) float in a fine-grained matrix. Graded bedding is common, either fining or coarsening upward.
3. Intraclast breccia layers are typically several centimeters thick, but can be as thin as a few laminae (viewed under a microscope; cf. Fig. 2B).
4. The upper contact of an intraclast breccia is invariably sharp and is typically overlain by laminated beds.
5. Basal contacts can be gradual, but occasionally are sharp. In the former case, folded and torn packets of laminae are abundant (Figs. 2A, 2C).
6. The verified lateral extent of individual intraclast breccias is on the order of 100 m. Over lateral distances of several tens of meters, the layers vary little in thickness, except where they onlap local paleorelief that formed during earthquakes (Fig. 3).

Jones and Omoto (2000) suggest the following criteria for the identification of seismic triggering of soft sediment deformation:

1. geological setting,
2. extent of the deformed units,
3. absence of evidence indicating other potential trigger mechanisms, and
4. presence of evidence of other potential trigger mechanisms elsewhere in the stratigraphic section associated with undeformed sediment.

Intraclast breccia layers in the sediments studied satisfy all these criteria.

• from Agnon et al. (2006:198)

The Dead Sea intraclast breccias have many similarities to seismites described from the lacustrine environment as well as from glacial deposits and volcanic terrains. Breccia layers associated with microfaults and intraformational folds in glacial deposits were reported in Scotland (Davenport and Ringrose, 1987). Earthquake-induced soft sediment deformation in Late Pleistocene lacustrine beds associated with activity at the Narugo Volcano, Japan, has been reported by Jones and Omoto (2000), who suggest the above-mentioned criteria to identify seismic triggering agents.

Most of the paleoseismic studies focus on Pleistocene to Recent deposits, but seismites have been reported in significantly older rocks, including Silurian strata (Kahle, 2002) and laminated Neogene deposits in Spain (Rodriguez-Pascua et al., 2003). In the area of the Dead Sea Rift, intraclast breccia layers are also present in Senonian chert (Mishash Formation), which crops out near the Lisan Formation (Fig. 2B). Similar fabrics in the breccias in these formations prompted Kolodny et al. (2005) to suggest a similar mechanism of formation.

Interpreting the origin of ancient seismites often relies on intuition and on understanding models of the mechanism of their formation. Features interpreted as ancient seismites should resemble those formed by modern earthquake deformation. Soft sediment deformation associated with strong earthquakes is documented by observations of recent seismic events (Allen, 1974; Sims, 1973). Earthquakes have caused silting and resuspension of sediments in Canadian lakes in association with earthquakes and in turbid water observed in lakes <10 km from the epicenter of the 1935 Temiskaming, Canada, M 6.3 earthquake (Doig, 1990, 1991). Piston cores from the bottom of that lake recovered a 20-cm-thick chaotic layer composed of tabular fragments derived from a preexisting silt layer. Graded bedding has been suggested as a criterion for subaqueous liquefaction based on observations in Kobe, Japan, following the 1995 earthquake (Kitamura et al., 2002).

Earthquake-induced historical homogenites are reported in Lake Lucerne, Switzerland (Siegenthaler et al., 1987). Lake Le Bourget, France, has homogenites that correlate with the A.D. 1822 earthquake (local intensity VII-VIII), the strongest known historical earthquake of the French outer Alps (Chapron et al., 1999). Historical accounts of this earthquake report violent lake water oscillations, which were probably a seiche, and an earthquake-induced subaqueous slide may have formed the homogenite layer.

• from Agnon et al. (2006:198-201)

The formation of intraclast breccias involves five stages (Fig. 3). First, layered deposits at the lakebed (Fig. 3A) are disrupted and deformed by ground shaking, motion of the water column, and water escape from the underlying uncompacted sediment (Fig. 3B). During this stage, the pressure of pore fluids in the sediment exceeds the confining pressure of the overlying lake brine, resulting in liquefaction of the sediment. Subsequently, the top of the sedimentary succession becomes fluidized and suspended at the bottom of the water body; fault ruptures can create topographic steps at the lake bottom (Fig. 3C). Seismic waves can trigger mechanical instability in the sediment, expelling pore fluid into the overlying suspension (Hamiel, 1999; Heifetz et al., 2005). Long water waves that oscillate the entire lake (seiche) carry significant momentum at the bottom of the lake, keeping the sediment suspended. After the waves have dispersed and attenuated, an intraclast breccia is deposited from the suspension by grain settling and water escape (Fig. 3D). After settling, the intraclast breccia is capped by the continuing deposition of laminated sediments that gradually bury any fault-related topography (Fig. 3E).

The intraclast breccia's texture attests to the interplay between forces in the sediment; namely, the pressure of the pore fluids, the contact forces between solid particles, and gravity. The clasts were originally part of laminae, and rupture of the laminae is a precursor of a liquefied state, where pore pressure exceeds cohesive forces and drives cracks through the sediment. As long as the pore pressure exceeds the lake pressure head, fluids that escape the liquefied bed exert drag stress on particles in the top part of the layer. When this drag exceeds gravitational forces, particles are suspended and the sediment is fluidized.

The formation energy of seismites in general is supplied by seismic shaking, but gravitational energy contributes on slopes, where the disturbed bed slides downhill, expending potential energy. Gravitational energy can also contribute to the formation of seismites where the density profile of the undisturbed sediment is inverted (dense on top): overturning the sediment releases the gravitational potential for overcoming resistance. There is no evidence that gravitational energy was a factor in the formation of common Dead Sea intraclast breccias. The seismites were deposited on flat surfaces, and no evidence for density inversion was found. Some gravitational energy is involved when pore fluid is injected upward, but most of the formation energy is of seismic origin. Three agents of seismic energy for disruption can be considered: the shaking of the ground below, the motion of the water above, and the injection of pore water from below.

Fragments of laminae in upward-fining intraclast breccias indicate that the lake-bottom sediment was compacted and cohesive before the earthquake. During the event, laminae shattered, and fragments were suspended into the fluid. Liquefaction of sediments under earthquake shaking is well documented and is traditionally related to the passage of shear waves (e.g., Allen, 1982). Yet observations of liquefaction features from recent large earthquakes highlight the role of P-waves (Lin, 1997), and engineering design based on resistance to cyclic shear loading has occasionally failed (Hatanaka et al., 1997). Observations of intraclast breccias in the Dead Sea basin have stimulated new theoretical and experimental studies of liquefaction (Bachrach et al., 2001; Hamiel, 1999; Lioubashevski et al., 1996).

An alternative mechanism is the Kelvin-Helmholtz Instability (KHI) mechanism, in which stably stratified layers undergo a shear instability during relative sliding, which is set off by earthquake shaking (Heifetz et al., 2005). Analysis suggests a threshold for ground acceleration increasing with the thickness of the folded layers. The maximum thickness of folded layers, on the order of decimeters, corresponds to ground accelerations of up to 1 g. The application of the KHI model to earthquakes is based on a translation of the instrumentally measurable ground accelerations to pressure gradients. The KHI model is at a preliminary stage and does not provide precise correspondence between field observations and the actual driving ground accelerations. Moreover, it does not rule out alternative sources for pressure gradients, such as surface and internal waves in the depositing water body. Since water depth of Lake Lisan above the investigated area was several tens of meters (Bartov et al., 2002), ground acceleration waves might have dominated over water waves.

• from Agnon et al. (2006:201)

Several authors have cited criteria to distinguish between the seismic and nonseismic origin of soft sediment deformation features (for reviews, see Jones and Omoto, 2000; Obermeier, 1996). Marco and Agnon's (1995) studies of seismites were originally motivated by intraclast breccias juxtaposed to intraformational faults in the vicinity of Masada (Fig. 1), where a terrace capped by laminated sediments is present between the Dead Sea and the western fault escarpment (Agnon, 1983; Sagy et al., 2003). Similar exposures of fault zones juxtaposed to intraclast breccias are also present in the Lisan Peninsula. Analysis of the microstratigraphy at these sites shows simultaneity between two processes acting at the lake bottom; namely, faulting and homogenization of the lake bed (Marco and Agnon, 1995, 2005). The time interval between intraclast breccia formation and lake bottom faulting is shorter than the time it took to deposit a lamina, which is likely less than one year (see evidence below for varve-like lamination). This association is perhaps the strongest evidence for attributing datable sedimentary structures to earthquakes, hence naming them seismites. The geological observation that deformation occurred at the water-sediment interface makes intraclast breccias excellent markers to determine the times of past earthquakes, if the time of sedimentation can be determined.

• from Agnon et al. (2006:201)

The Masada fault zone has repeatedly ruptured the surface along several km (Fig. 1). In order to examine the subsurface continuity of the faults, Marco et al. (1996a) carried out ground penetrating radar (GPR) and high-resolution seismic reflection surveys to image the fault zone (Figs. 4-7).

The 275-m-long GPR survey shows fault planes extending several meters below the surface (Fig. 5). A parallel calibration profile runs at the top surface of the Lisan Formation above a buried fault whose location and dip are known from exposures (Fig. 4).

The 450-m-long seismic reflection line across the Masada fault zone overlaps the GPR profile but extends farther east and west (Fig. 6). In addition to conventional reflection data, we present diffraction data analyzed using the method proposed by Landa et al. (1987) and Kanasewich and Phadke (1988). The diffracted waves are sensitive to the discontinuities in beds due to faulting, providing an independent support for the interpretation of the reflection profile.

Three zones of discontinuous reflectors on the processed profile represent faults that extend from <0.05 s down to 0.25 s (two-way traveltime [TWTT]) at shot points 50-60, 80-95, and 125-135. The faults are also the sources of diffractions that are shown in the diffraction section from 0.05 s to 0.15 s (TWTT) at shot points ~60, 95, and 125-135 (Fig. 6).

The coincidence of the faults in the geological, radar, and seismic reflection sections shows that every outcropping fault can be traced down to ~250-300 m (Fig. 7). Eyal et al. (2002) reported similar results in another study of a fault zone in an alluvial fan. Several faults evident in the seismic section are not expressed in outcrops, which may be evidence for syndepositional faulting, or alternatively, faults that did not rupture the surface.

• from Agnon et al. (2006:201,203)

Based on correlations with historical earthquakes, Migowski et al. (2004) concluded that local intensity is a critical factor in the formation of intraclast breccia layers. We review this work below and define the conditions required to form intraclast breccias.

The Dead Sea region has not experienced strong earthquakes during the instrumentally recorded twentieth century (Fig. 1). Therefore, to assess the intensity required to brecciate laminated sediments, we must rely on historical records of earthquakes. The Dead Sea region has a long, rich record of historical earthquakes, which consists of information on shaking in settlements nearby (Ambraseys et al., 1994; Amiran et al., 1994; Guidoboni, 1994). Recent retreat of the Dead Sea shorelines has exposed sediments deposited in the past millennia, where the effects of known historical earthquakes can be seen in the sediments. Exposures revealing historical deposits are limited to where the lake level was higher than the level during the past decade. Even for these intervals, the sediments may lack fine bedding or may have been eroded during times of low water levels. Drill cores in lacustrine laminated facies overcome these limitations.

Another analysis, which considers both the thickness of the breccia beds and the lithology of beds directly overlying them, is applied in order to identify the stronger (M > 7) earthquakes within the record recovered from the Lisan Formation (Begin et al., 2005). The analysis is based on the occurrence of gypsum immediately above 11 breccia layers between 54 and 16 ka, a coincidence that is explained by the triggering of a strong seiche, which mixed the stratified waters of Lake Lisan. Mixing of the sulfate-rich upper water layer with the calcium-rich lower water layer could trigger the deposition of gypsum (Stein et al., 1997a). The resulting time series of earthquake recurrence interval is similar to the M ≥ 7.2 recurrence interval in the Dead Sea basin, as extrapolated from present seismicity; therefore, Begin et al. (2005) suggest that the present seismic regime in the Dead Sea basin, as reflected in its magnitude-frequency relationship, has been stationary for the past ~40 k.y.

• from Agnon et al. (2006:203-205)

Ken-Tor et al. (2001a) studied Holocene outcrops of Dead Sea sediments in Ze'elim fan, east of Masada (Fig. 1), where eight late Holocene seismites are exposed due to the accelerated recess of the Dead Sea during the past decades (Fig. 8). Six intraclast breccia layers (A-F) are identified in the lacustrine laminated facies. The uppermost part of the section exposes only the near-shore sandy facies, showing two liquefied sand units (G and H). Some intraclast breccias in the lower section grade laterally into liquefied beds showing flame structures as these beds change facies into beach sands. Twenty-four radiocarbon ages of plant debris are largely consistent with the stratigraphic order in the section sampled. Ken-Tor et al. (2001a) were able to fit a model of moderately varying deposition rates (3-9 mm/yr) between hiatuses by assuming that all intraclast breccias were formed during historically recorded earthquakes (Fig. 8). Significant uncertainty remained only regarding event D (~4 m above the base of the section, Fig. 8) that could correspond to either of two historical earthquakes: 363 or 419 A.D. Ken-Tor et al. (2001a) have considered both events (compare their Figures 3 and 4).

By combining radiocarbon dating with the precise dates of historical earthquakes and with field evidence suggesting subaerial exposures, Ken-Tor et al. (2001a) determined that two unconformities exist in the section between 400 (or 420) and 1200 A.D., and between 1300 and 1750 A.D. (Fig. 9). This interpretation is remarkable in that the eight major historical earthquakes that are missing from the Ze'elim fan section have all happened during the hiatuses thus dated (551, 559, 749, 1033, 1068, 1160, 1456, 1546 A.D.). This suggests that the Ze'elim earthquake record is complete for the periods of deposition, with the exception of either 363 or 419 A.D., for which only a single intraclast breccia is found. We return to this dilemma after reanalyzing the outcrop data, and again after presenting the results from drill cores.

Subsequently, Ken-Tor et al. (2001b) used the historical earthquakes to refine calibration of their radiocarbon dates and to infer that the time between death of the plant and burial in the sediment is 50 yr or less.

Figure 8 offers a revised correlation scheme between the Ze'elim stratigraphic record and the historical earthquake record that satisfies two conditions:

1. All model dates match event horizons with historical earthquakes.
2. A uniform sedimentation rate between successive event horizons and a slowly varying sedimentation rate between hiatuses. This allows interpolation between event horizons and determination of ages based on historical dates exceeding the precision of radiocarbon.

We used 24 calibrated radiocarbon ages (Bookman et al., 2004; Ken-Tor et al., 2001a) to guide the matching of the sediment height-age model. Nineteen ages are compatible with the model in that the higher bound on the age is older than the model age. For most cases, the model line goes through the calibrated age range. Five calibrated ages are younger than the model by several decades. Some incongruence can be resolved by considering low-probability ranges in the calibrated date distributions. Discrepancy may also result from sampling of roots debris that might have deteriorated in situ. Alternatively, our uniform deposition rate may be an oversimplification for this arid climate featuring irregular flash floods.

Our model indicates a strikingly uniform mean sedimentation rate during the three periods of continuous sedimentation: 5-6 mm/yr. Where data is ample, we note that the deposition rate may fluctuate by 50% around that mean rate: at the short period before the Christian era (between the events of 140 and 31 B.C.), we see an anomalously high rate of deposition (7 mm/yr). Subsequently, between 31 B.C. and 33 A.D., the rate declines to 3 mm/yr, maintaining an average of ~5 mm/yr.

This uniform deposition rate model results in breccia layer A in the section of Ken-Tor et al. (2001a) correlating with the 140 B.C. earthquake. The 64 B.C. earthquake, which was originally correlated with breccia layer A, cannot be distinguished from the 31 B.C. breccia layer. We prefer this correlation to the correlation of breccia layer A with the 64 B.C. earthquake because the latter correlation implies an excessive deposition rate of 24 mm/yr.

Our model correlates breccia layer D with the 419 A.D. earthquake, noting that this results in the 363 A.D. earthquake being uncorrelated. Our reconstruction agrees with lamina counting data (Migowski et al., 2004) reviewed in the next section.

• from Agnon et al. (2006:205-206)

Continuous cores from three sites along the Dead Sea shore were drilled during the fall of 1997 (Migowski et al., 2004), including cores from Ze'elim fan and Ein Gedi Spa (Fig. 1). The staggered-pair drilling technique recovered a continuous record of the subsurface sediments. We chose the Ze'elim fan site to correlate subsurface strata with the outcrops (Ken-Tor et al., 2001a, 2001b) and verify that both surface and drill core methods agree. The two other boreholes were drilled very close to the contemporary shoreline to avoid hiatuses due to lake level drops beyond the current level (Bookman et al., 2004) and to obtain data from very recent sediments. The 20-m-long Ein-Gedi core provides a continuous sedimentation record that spans the past 10,000 yr (Migowski et al., 2004). Above a 10 ka salt layer, the core contains two alternating principal facies: laminated fine-grained chalk (laminites) and bedded to massive silt. The laminated chalk contains aragonite laminae, resembling the lacustrine facies of Lisan Formation (Bartov et al., this volume).

The Ein Gedi core has penetrated 53 deformed intervals (Migowski et al., 2004, their Table 2). Many of the 53 deformed intervals in the core resemble the intraclast breccia beds of Marco and Agnon (1995). Migowski et al. (2004) focused on a 2.25-m-long interval of aragonite-rich laminites for a detailed inspection under an optical microscope. They counted couplets of detritus and chemically precipitated laminae as single depositional cycles. In some cycles, laminae of gypsum added to form triplets. They identified 1500 deposition cycles and suggested that a cycle represented one year of sedimentation. One way to test the annual depositional cycle hypothesis is to evaluate the lamina chronology with intraclast breccia events to see whether the time intervals match the historical record. Within that microscopically analyzed interval, Migowski et al. (2004) found 22 intraclast breccia layers and developed a chronological model for the sequence in which each cycle represents one year.

Migowski et al. (2004) constrained their chronological model to minimize the number of breccia beds for which no historical earthquake is known and found only one model that matched as many as 20 out of 22 breccia layers with historical earthquakes since ca. 150 B.C. (e.g., Fig. 10). They found that four additional earthquakes correspond to periods in which the record was destroyed by brecciation associated with subsequent earthquakes, postdating the missing events by several years (and, in a single case, by 33 yr). Figure 10 shows a unique situation where the contact between two breccia layers is preserved, so two earthquakes separated by 10 yr may be resolved.

The chosen matching, leaving out two subcentimeter breccia layers at 90 A.D. and 175 A.D., is significantly better than any other chronology model. The chosen model also minimizes the number of historical earthquakes for which no disturbance is shown in the sediment: six historical earthquakes from the entire region are missing. Clearly, some of these missing historical earthquakes were too distant or too weak to generate significant shaking in Ein Gedi.

While direct historic information on local shaking intensity at Ein Gedi is rare, empirical formulas can evaluate the historic data and estimate the magnitude of a given earthquake, as well as its location (Ambraseys, 1988). In a recent compilation of Middle East intensity data, Ambraseys and Jackson (1998) develop a formula that relates mean earthquake magnitudes to the intensity of shaking at given source distances:

M = -1.74 + 0.66I + 0.0015R + 2.26 log R,   (1)

where R is the distance from the earthquake focus (assumed at 7.4 km depth), I is the local intensity, and M is the magnitude. Formulas of this form, known as attenuation relations, estimate the local intensity at a site if both the earthquake magnitude and its location are known. A sufficiently close earthquake source with sufficiently large magnitude should generate intraclast breccia layers; sources that are too weak or too distant would not generate intraclast breccia layers. Figure 11 shows a compilation of all historical and instrumental earthquake sources in terms of magnitude and distance from Ein Gedi. The diagonal bold curves separate three domains: strong and close sources on the lower right, all matched with breccia layers; weak and far sources on the upper left, all unmatched; a median domain where about half of the events are matched by intraclast breccia layers. These lines can be fitted to equations of the form of equation (1). The straight line is given by:

M = 1.9 log R + 2.8.   (2)

The dotted and dashed lines in Figure 11 mark intensities V and VI, respectively. Note that for distances R > 50 km and magnitudes M > 5.5, we can consider the isoseismal I = V as a domain boundary. At lower distances and magnitudes, a higher intensity seems to be required for generating breccia layers.

As shown in Figure 11, all historical earthquakes with calculated local intensity at Ein Gedi I > V are matched with intraclast breccia. The earthquake of 363 A.D. may constitute an exception to that rule.

The unique match of the two independent records, namely the historical and the one derived from the core, supports three assumptions used to develop the chronological model:

1. breccia layers form by seismic shaking,
2. strong shaking results in breccia layers, and
3. the lamination is seasonal with a detectable annual cycle.

• from Agnon et al. (2006:212)

The systematic approach to accumulating observations on intraclast breccia layers permits their analysis as recorders of paleo-earthquakes. Such breccia layers, previously called "mixed layers," are abundant in sedimentary sections of Quaternary lakes from the Dead Sea basin. The finding of intraclast breccia layers juxtaposed against surface faults has driven a wide range of studies focused on these long-term seismic recorders. The fault scarps form a fault zone, traced to the subsurface by high-resolution geophysical surveys. This fault zone is a subsidiary structure to the master fault bounding the Dead Sea basin from the west and extending north to the transform plate boundary. While the record of displacement of the fault zone is limited to particular slip events that activate this secondary structure, the record of breccia layers may be complete for events that rupture the plate boundary.

We define field criteria for the identification of intraclast breccias, focusing on features that can indicate a seismic origin. The wealth of data for such earthquake indicators collected in natural outcrops within the Dead Sea basin offers insights into the phenomenology and systematics of earthquakes on time scales that are not obtainable elsewhere.

Cores from the receding shore of the Dead Sea contain continuous sedimentary records of the past 10 k.y., undisturbed by lowstands. The Ein Gedi core features 3 m of alternating seasonal laminae. The two independent earthquake records—historical and sedimentary—offer a simultaneous test regarding two hypotheses: the earthquake origin of breccia and the annual cycle of laminae. The likelihood of matching historical earthquakes with an arbitrary time series that corresponds to the breccia layers in the core is negligible. The observation that the breccia layers match earthquakes from the historical catalogues that are strong and/or close to the coring site supports both hypotheses (Fig. 11). Changes in the rate of recurrence of earthquakes in the Dead Sea record during the historical period seem to correlate with changes in the Anatolian Fault system. If the rates of recurrence could be taken as indicators of activity of the plate boundaries, then these plate boundaries might be coupled on the time scale of 500 yr.

The brecciation from an earthquake that succeeds another strong earthquake might obliterate the breccia layer of the predecessor. This hampers the potential that lies in the laminated sediment to resolve pairs of earthquakes. The resolution of an interseismic interval is no better than the ratio of the thickness of a breccia layer to the rate of deposition. But the resolution limit of individual earthquakes does not affect the observation of clustering in the record, which is evident in the long periods of quiescence alternating with periods of recurrence of earthquakes. During most of the period recorded, we find that the apparent recurrence interval is significantly longer than the resolution limit. During a cluster of earthquakes ca. 52 ka, the interseismic interval becomes shorter than 200 yr, which is close to the resolution limit for the Lisan outcrops.

• from Kagan et al. (2011:3)

[10] The sediments of the Holocene Dead Sea comprise the Ze’elim Fm. of the Dead Sea Group. The sediments represent various depositional environments: fluvial, fan deltas, shores, and lacustrine (see detailed description given by Bookman (Ken‐Tor) et al. [2004]). The current (2009) lake level is 422 m below sea level (mbsl), reflecting mainly anthropogenic diversion of freshwater inflow, while during the Holocene the natural (climate related) level varied from ∼370 mbsl (e.g., ∼6000 years ago) to lower than ∼430 mbsl around 8000 years ago [Bookman (Ken‐Tor) et al., 2004; Bookman et al., 2006; Migowski et al., 2006; Bartov et al., 2007]. The drop of the current lake level (12 m from 1980 to 2000 [Bookman (Ken‐Tor) et al., 2004]) has caused the formation of deep gullies along the retreating shores. These gullies provided an excellent opportunity to study the late Holocene sedimentary sections in detail. The paleoseismic data in the current study was derived from the outcrops of the Ze’elim Gully and the Ein Feshkha Gully. Another site used for comparison, the Ein Gedi core site, was studied by Migowski et al. [2004].

[11] The study site of Ein Feshkha Gully (at the Ein Feshkha Nature Reserve) is located at the northwestern shore of the Dead Sea, 60 km north of the Ze’elim Gully (Figure 1). Ein Feshkha is an oasis of brackish streams and pools. Nearly exclusively lacustrine sediments are exposed in the Ein Feshkha site by a ∼6.5 m deep gully (as of 2008). The site is close to the Jordan Valley segment of the DST and may be located on the WNW continuation of the Kalia transverse fault mentioned in section 2.

[12] The Ze’elim gullies (site ZA) are dissected into the Ze’elim Plain east of the ancient fortress of Massada (Figure 1). Currently (as of 2009) the ZA Gully is ∼11 m deep (approximately at lake level, to slightly above). The gully exposes a stratigraphic sequence of lacustrine, shore environment, and fluvial sediments. The ZA site is closer to the Arava segment of the DST, about 50 km away, than to the Jericho fault. The active eastern normal boundary fault of the DST is at a similar distance to all sites, less than 5 km away (Figure 1).

• from Kagan et al. (2011:3-4)

[13] El‐Isa and Mustafa [1986] used intraformational folds to generate an earthquake archive. Subsequent studies presented more complete archives recognizing that folds might present the weakest events. “Mixed layers” were renamed “intraclast breccia layers” [Agnon et al., 2006].

[14] Agnon et al. [2006] define field criteria for the recognition of intraclast breccias, focusing on features diagnostic of a seismic origin. The field criteria reflect the mechanisms of breccia formation, which include ground acceleration, shearing, liquefaction, water escape, fluidization, and resuspension of the originally laminated mud.

[15] In the current study we recognize deformed structures such as intraclast breccias, liquefied sands, folded laminae, and small faults (centimeter scale) (Table 1). Figure 2 displays photographs and photo tracings of seismites from the study sites. In addition we recognize another type of deformation termed microbreccia or homogenite. This type of mid‐gray‐color sedimentary layer ranges in thickness from a few mm to 1–2 cm and appears homogenous in the field. Thin‐section investigation under a polarizing microscope shows that these are actually brecciated laminae, and include a mixture of detritus, aragonite, and in places gypsum fragments.

[16] In the more fluvial Ze’elim section there are instances of seismites with a combination of lacustrine breccia and sand liquefaction. For example (see Figure 2c), ZA seismite III is the product of the deformation of a lower sandy layer and an upper laminated marl layer, resulting in brecciated marl laminae (near top of Figure 2c) with injection of sand fingers (near bottom of Figure 2c) from below.

• from Kagan et al. (2011:5)

[17] Fieldwork included the detailed description and sampling of subvertical to vertical outcrops in the Ein Feshkha and Ze’elim gullies. Columnar sections were prepared with emphasis on measurement and description of the deformations. Adjacent outcrops were examined in order to describe spatial variations in lithology and seismites. Sediment blocks (∼10 × 10 × 10 cm in size) were collected for further analysis in the lab. At Ein Feshkha, 58 continuous blocks of wet sediment were retrieved from the gully wall at the columnar section site, from the surface plain down to 40 cm below the autumn 2005 water level of the spring outflow (see our previous paper, Neumann et al. [2007]). At the Ze’elim Gully sediment blocks were retrieved from the various lithological units. In addition, organic debris (typically short‐lived leaves or twigs), found in the two outcrops, were sampled for radiocarbon dating.

• from Kagan et al. (2011:7-9)

[20] The chronologies of the Ein Feshkha and Ze’elim sections were constructed by radiocarbon dating of terrestrial organic debris (mainly small pieces of wood and twigs). All the recovered wood in Dead Sea sections can be considered driftwood, however their transport time is relatively short. We made an effort, where possible, to date only short‐lived organic debris. Nine samples from EFE and twelve samples from ZA2 were prepared for radiocarbon dating at the Radiocarbon laboratory, Weizmann Institute, Rehovot, Israel. The samples were then measured by accelerator mass spectrometry (AMS) or liquid scintillation counting (LSC) at the NSF radiocarbon facility in Arizona. Eight additional organic debris samples from EFE were taken from a core drilled a mere few meters away, on the cliff bounding the gully, and analyzed at the AMS facility in Kiel. The core was correlated with the outcrop by Marcus Schwab at GFZ‐Potsdam. Radiocarbon ages are reported (Table 2) in conventional radiocarbon years (BP = before present; present defined as 1950 A.D.) in accordance with international convention [Stuiver and Polach, 1977]. Calibrated ages (=cal BP) were calculated by applying the INTCAL04 calibration scheme of Reimer et al. [2004] by means of the OxCal v4.1 program of Bronk Ramsey [1995, 2001, 2008]. Age‐depth models (Figures 3 and 4) and seismite model ages (see Table 3) were also created with OxCal (v4.1) [Bronk Ramsey, 1995, 2001, 2008].

[21] Radiocarbon data are listed in Table 2. Table 2 presents the measured ages, calibrated ages, and deposition model ages applying the Bayesian statistics of the OxCal v4.1 program. The depositional model ages were used to establish an age‐depth chronological model for the seismites. The fundamental assumption in Bayesian modeling of stratigraphic sequences is that age increases with depth. This requires use of a function usually termed “Boundaries” in OxCal. The boundaries separate different sedimentary units that may have different sedimentation rates, grain sizes, and facies. They are also placed on the top and bottom of the entire series to constrain the model to a specific time interval. With no other information, this would be treated by what is usually termed the “Sequence” model by OxCal. A uniform sedimentation rate would be treated with the “U_Sequence” type model. Depth and other dating information can be included in a less rigid way using Poisson distribution priors, termed “P_Sequence” models, where the time gap between deposition of grains varies, and the events are basically random but deposition is given approximate proportionality to depth. This requires the estimation of the uniformity of the deposition (given as the k parameter), which signifies the increment size (conceptually the grain size, or size of deposition events) and indicates the relation between the events and the stratigraphic process [Bronk Ramsey, 2008].

[22] In this study a P_sequence (Poisson distribution) Bayesian depositional model was used, with a k factor value of 1 (see Bronk Ramsey [2008] and Kagan et al. [2010] for details of Bayesian factors used). In the work of Kagan et al. [2010] the main objective was to test the Bayesian model with and without historical earthquake anchor points. The conclusion of the work was that the “known‐earthquake‐anchors” do not significantly improve the age model. For that reason, and due to the complexity in choosing definite historical anchors, in this study no anchors are used and the models are based solely on radiocarbon data, stratigraphic data, and the P_sequence and k factor constraints discussed in this section and by Kagan et al. [2010].

• from Kagan et al. (2011:9-10)

[23] For the EFE section the chronological model is based on the treatment of seventeen radiocarbon ages of which five were excluded as outliers (Table 2). In the last 2500 years, the period with historic earthquake correlations and implications, there were only two outliers, both of which were too old and probably represent long‐lived organic debris from the shores. One of these two outlier samples also appeared in the work of Neumann et al. [2007] (169 cm depth) and was considered an outlier. One interval, from 230 to 390 cm, is slightly anomalous: the sediment is much darker than the rest of the section and has less aragonite layers. Within this interval, between 230 and 330 cm, we did not recognize any deformed layers (Table 3). No organic debris was found from 220 to 410 cm depth (Table 2 and Figure 3).

[24] Several different models were run: (1) No internal boundaries from 0 to 500 m depth; 500 cm to base modeled separately. (2) Two internal boundaries in the 0–537 cm interval, at 230 and at 500 cm depth, which allow, but do not force, the model to have sediment rate changes. (3) The 0–230 cm and 390–500 cm deep segments run separately. (4) Various other options with different boundaries and various outliers.

[25] We choose option 2 from the above list (Figure 3). This curve yields the best “agreement indexes” for the Bayesian model, with one index value under 60% (at 17%) while the other models have lower agreement indexes. Alternative models give seismite ages (from at least the 5th century B.C. and on) that are very similar to the chosen model and do not change the paleoseismic conclusions (for an example, see option 1 in Figure S1).1 The slight facies change at 230 cm depth is allowed a degree of freedom to coincide with sedimentation rate change, but in the resulting model shows no significant rate change (see Figure 3).

[26] The chronology of the top 537 cm of the section is Bayesian‐modeled as one space with two internal boundaries at 230 cm and 500 cm, implying continuous sedimentation and allowing, but not forcing, sedimentation rate change at these boundaries. Agreement values are found to be well above 60% at most depths of the model. The resulting model ages of the section indicate a maximum range of 1261 B.C. to 1383 A.D., but more likely from ∼1100 B.C. to 1312 A.D. (Table 2 and Figure 3). The top unit, from 0 cm (surface) to 500 cm, shows ages that range from the 5th century B.C. to the 14th century A.D., with a 0.27 ± 0.03 cm/yr sedimentation rate (based on 2s age ranges). The age range of the lower unit (500 to 537 cm) is from approximately 11th–5th century B.C. (0.07 ± 0.03 cm/yr sedimentation). The base of the seismite‐bearing investigated section is at 590 cm, however in the bottom 53 cm no organic matter was found and therefore the age was not modeled. The sedimentation rate of the top 500 cm calculated here (0.27 cm/yr) is approximately constant, in comparison to that stated for the same section by Neumann et al. [2007] (0.14, 0.51, and 0.11 cm/yr for three stratigraphic units within the same depth interval). The rates presented here, based on the new Bayesian model, are more similar to published Holocene rates (e.g., Migowski et al. [2004]: ∼0.15 cm/yr for the entire Holocene Ein Gedi core) and more congruous with homogeneous pollen concentrations [Neumann et al., 2007], which are normally closely linked to sedimentation rate [Horowitz, 1992].

[27] The truncation of the last six centuries from the studied EFE section eliminates recording the key instrumental earthquake M6.2, 7 November 1927, the source zone of which spans the site (Figure 1). Macroseismic evidence for the 1927 A.D. instrumentally recorded earthquake was reported along the Jordan River [Hough and Avni, 2011]. Niemi and Ben‐Avraham [1994] interpreted large submarine slumps in the northern Dead Sea basin to have been caused by this earthquake. For the purpose of the discussion (section 5), this event will be considered recorded in the northern Dead Sea basin.

• from Kagan et al. (2011:10)

[28] Twelve organic debris samples from the 10.7 m deep Ze’elim (ZA2) outcrop were measured. Their calibrated radiocarbon ages range from 1056 to 1276 A.D. to 4843–4583 B.C. (95% probability). A deposition model is calculated for the top 8 m of this section. Model ages of samples are given in Table 2. The more western ZA1 section (∼100 m away) was dated by Ken‐Tor et al. [2001a, 2001b] and revised by Agnon et al. [2006]. In the Ze’elim Gully previous studies infer the sedimentation rate to range between 0.28 to up to ∼1.3 cm/yr [Ken‐Tor et al., 2001a; Agnon et al., 2006; Neumann et al., 2007] reflecting the additional detrital‐clastic sediments that are more abundant in the fan delta environment. The lower sedimentation rate (0.3 ± 0.03 cm/yr) at the ZA2 section of Ze’elim (current study) reflects the proximity of this section to the lacustrine environment. The ZA2 outcrop is interpreted to show continuous deposition according to the age‐depth model (Figure 4), as opposed to the numerous unconformities at the more landward ZA1 outcrop. However, at ZA2 there is the possibility of short hiatuses compensated by additional sediments at sandy facies which therefore are not manifested in the age‐depth diagram.

• from Kagan et al. (2011:10-11)

[29] Ages of seismites (Table 3) are interpolated from the radiocarbon age‐depth data using Bayesian stratigraphic constraints. The ages and their uncertainties are interpolated using the OxCal program and take into consideration the asymmetrical probability distribution of radiocarbon ages. Each seismite is assigned a probability distribution histogram with a 68% (∼1s) and 95% (∼2s) probability age range (Figures 3 and 4). Model ages are presented (Table 3) with a nominal precision of a single year, however due to the Bayesian statistical modeling each model run produces slightly different age ranges and therefore ages could be rounded off by 10 years. Although the annual dates are shown, they are dealt with as if rounded off; for example, when giving the historical “fit” in Table 3, the age ranges are considered in decades.

[30] Seismite ages have been compared to historical catalogs as a major component in the assessment of the validity of the interpretation of the breccia layers as seismites [e.g., Ken‐Tor et al., 2001a; Migowski et al., 2004]. At the same time, seismites can be used for the corroboration of individual earthquakes in the historical record. Ken‐Tor et al. [2001a, 2001b] used the radiocarbon ages of the individual breccia layers or liquefied sands for direct comparison with the historical records and noted that notorious historic earthquakes unrepresented in the geological record lie within sedimentary hiatuses in the western Ze’elim Gully section (termed here ZA1). Migowski et al. [2004] positively identified these “missing” earthquakes in the continuous lacustrine section of the Ein Gedi core, supporting the hiatus‐hypothesis. Moreover, by counting the laminae in the intervals between seismites they were able to correlate almost the entire historical and Ein Gedi core records.

[31] Table A1 presents the historic earthquakes in the region with information regarding damage, casualties, sources of historical data, and, in the footnotes, selected archeological and paleoseismic data. Table A1 is based on earthquake catalogs, whereas the information in the catalogs is derived from historical sources. Table A1 is reliable mostly during the past two millennia (from the Roman period and onward), but less information is available for the time interval 750–1100 A.D. (when the Muslim empire center moved from Damascus to Baghdad). The historical accounts in the pre‐Christian era are rare and if they do exist tend to be vague [Karcz, 2004]. A mid‐8th century event and its paleoseismological and historical implications are discussed in detail in Appendix C. Local source moderate earthquakes are probably missing in the historical catalogs. For seismite ages where only very distant correlative historic earthquakes exist, we propose small local source events as possible sources of seismite genesis (marked LS on Table 3). A map of historical locations is given in Appendix B (Figure B1).

[32] For the past two millennia we correlated almost all of the seismites in the Ze’elim and Ein Feshkha records to historic earthquakes (details in Tables 3 and 4). All historic earthquake dates that correspond to the 95% probability range of each seismite age are given in Table 3 (right column). Those that correspond to the 68% probability range are in bold.

[33] The protocol for assigning a particular historic earthquake to a seismite in the sedimentary section is the following: (1) We consider all known earthquakes within a time segment of the age‐depth model pertaining to the seismite depth (segment = within 1–2s uncertainty of the radiocarbon model age); this step is given in column titled “all possible events” in Table 3. (2) Among the earthquakes within this time segment, we select the one that is most consistent with age‐depth models of Figures 3 and 4 (preserving the sedimentation rate); see the correlation in Figure 5. We also considered the local intensity for the earthquakes estimated for the study area when deliberating, in certain cases, between the various earthquakes.

[34] Table 4 and Figure 5 present the results of the correlation of the paleoseismic evidence (Table 3) with the historical record (Appendix A) and the comparison of these results from four sections: EFE, EG, ZA1, and ZA2. In Figure 5 the historical dates of seismites are superimposed on the age‐depth models to display the matching of the two models; the deposition model and the historical correlation model.

[35] There are two possible sources of errors in a comparison between two archives, such as the historic earthquakes and the radiocarbon dated seismites. As noted in Table 3 the uncertainty in the age‐depth model is variable but typically less than 100 years (2s). This reflects the errors derived from the age‐depth Bayesian model. The uncertainty in the “historic ages” of specific seismites reflects the spread of all historic earthquakes that lie within the 2s model age range of the specific seismite depth (the right‐hand column of Table 3). Thus, the errors on the Bayesian curve are the reasonable estimate of errors in the historical ages–seismite comparison. In other words, we say that the maximum error in our comparison is less than 100 years and, as Table 3 shows, typically lower than 50 years.

[36] A special case is the couplet of earthquakes at 1202 and 1212 A.D. that, with the typical temporal resolution in Dead Sea sediments, are not resolvable. We chose to present them as a pair of events as 1202/1212 A.D. The seismite at 28 cm depth at EFE has a 1s model age of 1199–1240 A.D. Both the 1202 and 1212 A.D. events are large M >7 earthquakes that ruptured far from the Dead Sea (north of the Sea of Galilee to Lebanon, minimum 130 km [Marco et al., 2005] and south of the Arava, minimum 250 km, respectively). Agnon et al. [2006] show two adjacent seismites at this time in the EG core record and interpret these to represent both the 1202 and 1212 A.D. events. Both are candidates for this EFE seismite.

[37] The Ein Gedi core was dated by 20 radiocarbon ages and by laminae‐counting of ∼1500 years, from 200 B.C. to 1300 A.D. [Migowski, 2001; Migowski et al., 2004]. The laminae‐counted floating chronology of the seismites was matched with the historic earthquake catalog. The best‐fit history of Migowski et al. [2004] gave ages younger than their radiocarbon ages by 50–200 years, consistent with reworking of organic debris (e.g., wood) in the nearshore environment before settling to the bottom of the dense saline lake. In our analysis, the chronologies of the Ze’elim and Ein Feshkha section indicate no long reworking time of the organic debris before settling in the sediment. When referring to the seismites from the Ein Gedi (EG) core only we use the shifted laminae‐counted chronology of Migowski et al. [2004] for the EG section.

[38] Note in Table 3 that the type B seismites, “homogenites,” clearly correlate with important historic earthquakes, which supports their interpretation as seismites.

[39] The recording of earthquakes by seismites, as well as by historical documents, requires intensity above respective thresholds. In this study our data suggest that these two thresholds are similar. Quiescence intervals are more robust than specific earthquakes, because they are less sensitive to individual date correlation. Specifically there is a quiescence interval in the seismite archive of the three sites from the end of the 2nd to the beginning of the 4th century A.D. (Figure 7). This correlates to an historical earthquake quiescence period noted without a single historically documented earthquake in the region from 127 to 306 A.D. (Appendix A).

• from Kagan et al. (2011:11-13)

[40] Here we summarize the occurrence of seismites at the three sites presented in section 5.1 and in Table 4: Ein Feshkha (EFE), this study; Ein Gedi (EG core; after Migowski et al. [2004]); and the two Ze’elim Gully subsites: ZA1 (western, landward; after Ken‐Tor et al. [2001a] and Agnon et al. [2006]) and ZA2 (eastern, lakeward; this study), considered henceforth as one site. We limit the comparison to the historical period starting at the 2nd century B.C.

[41] 1. Seismites that appear in all three sites (termed here intrabasin seismites (IBS)): Mid‐2nd century and 31 B.C. and 33, 419, 551, 749, 1202/1212, 1293, and 1927 A.D.

[42] 2. Seismites that appear only in Ein Gedi: 76, 90, 112, 500/502, 1042, 1546, 1588, 1656, 1712, 1759, and 1822 A.D.

[43] 3. Seismites that appear only in Ein Feshkha: 64 B.C., 349, 363, 634, 847, 859, 956, 1063, 1170, and 1312 A.D., and numerous older prehistoric seismites.

[44] 4. Seismites that appear in Ein Gedi and Ein Feshkha but not in Ze’elim: 92 B.C. and 660, 757, 873, 991*, 1033*, 1114/1117*, and 1068* A.D. Stars indicate dates at which time there is no archive at Ze’elim.

[45] 5. There is one quiescence interval at ZA and EG ∼500–150 B.C. and another at all three sites from the end of the 2nd to the middle of the 4th century A.D.

[46] The new chronologies of the seismites in the Ze’elim (ZA) and Ein Feshkha (EFE) sedimentary sections are integrated with the high‐resolution seismite chronology of the Ein Gedi (EG) core to produce a comprehensive archive of late Holocene paleoseismic earthquakes from the entire Dead Sea basin. The paleoseismic archives also provide an opportunity to reevaluate a number of earthquake histories with timing and patterns that were not well established (e.g., single or several episodes).

• from Kagan et al. (2011:13-16)

[47] The chronologies that were established for the Ein Feshkha and Ze’elim sections combined with that of the Ein Gedi core [Migowski et al., 2004] allow us to compare the recurrence time of the seismites in these sites and to produce an integrated picture for the appearance of seismites in the northern Dead Sea basin (Table 4 and Figure 6). The number of seismites in the Ze’elim Gully sections is significantly smaller than at Ein Feshkha and Ein Gedi for the same time interval. Ken‐Tor et al. [2001a] and Agnon et al. [2006] recognized that the missing seismites at ZA1 (explained in section 1) relate to sedimentary hiatuses in the section. The new section we described at ZA2 yielded an apparently continuous age‐depth profile, and the hiatuses in the ZA1 section can be correlated with clastic‐sandy sequences in the ZA2 section. One of the missing (sedimentary hiatus) earthquakes (in the landward ZA1 section) does appear in the continuous ZA2 section as liquefaction in a sandy unit (correlative to the historical earthquake of 551 A.D.). In two instances the situation is reversed, where two seismites, correlated to 1293 and 1212 A.D. appear in the more landward ZA1 outcrop, and do not appear in the more lakeward ZA2 section. This specific period is characterized by a sandy facies at ZA2 (Figure 4) and detailed detection is also inhibited by difficult access at this part of the section.

[48] The EFE section has 52 seismites, while for the same time period the EG section shows ∼30 seismites. A quiescence period at EFE at around mid‐1st to 3rd century A.D. is concurrent to a period in EG with microscopic seismites (Type III of Migowski et al. [2004]). This could reflect the higher detection resolution of the Ein Gedi study. Despite this resolution difference, the situation is reversed in the pre‐2nd century B.C. period where EFE has 25 seismites (<1 cm to >9 cm) compared to 7 at EG. The recurrence of earthquakes in each one of these sections is illustrated as a cumulative function in Figure 7.

[49] A quiescence at ZA and EG during a period of enhanced seismicity in the north (EFE) at ∼500–150 B.C. (Figure 7) may suggest a period of moderate earthquakes concentrated north of the Dead Sea (i.e., Kalia fault). Additionally, there is a quiescence interval in the seismite archive of the three sites from the end of the 2nd to the middle of the 4th century A.D., which correlates to an historical earthquake quiescence period 127–306 A.D. (Appendix A). This is in line with the low‐seismicity interval during this period along the DST, the high‐seismicity period on the North Anatolian Fault, and the mechanical coupling and alternation of activity of the two faults suggested by Migowski et al. [2004] and Agnon et al. [2006].

[50] The comparison of EFE versus both EG and ZA clearly suggests higher activity in EFE. This can be explained by a difference in sensitivity between the sites, or the proximity of EFE to the Kalia transverse fault bounding the Dead Sea basin from the north (Figure 1). The EFE site is located on the continuation of this fault to the WNW, and has likely recorded local earthquakes of magnitude ∼5.5 that were too far to affect EG and ZA. Also, several seismites (during the time interval of the historical charts) were recorded only at the northern site of Ein Feshkha (e.g., 64 B.C. and 349, 363, 634, 847, 859, 956, 1063, 1170, and 1312 A.D.). Most of these events have destruction documented mainly in the northern Holy Land or further north (Antioch, Tyre, Turkey; see Appendix A), 1312 A.D. being the main exception. Since the work of Russell [1980], the 363 A.D. earthquake is often considered as one that ruptured from the north to the Arava. We suggest that this interpretation congeals two earthquakes, one northern and another southern (see Appendix A). The lack of documentation of earthquakes in the south can reflect bias due to population density, the south being more arid. However, the excess of recorded earthquakes at Ein Feshkha may corroborate higher seismic activity in the north. First let us consider the local setting of the Ein Feshkha Nature Reserve site: it is positioned at the edge of both the Jericho fault and the Kalia transverse fault (Figure 1). Ze’elim Gully, on the other hand, is several tens of kilometers from both Jericho and Arava faults, the likely sources of M > 6.5 events. Therefore, earthquakes rupturing the northern part of the Jericho segment will record at Ein Feshkha but not at the southern sites. Likewise, magnitudes 5.5–6 from the Kalia fault may be recorded locally but not at the southern sites.

[51] Our sites are located on the western shore of the lake, close to the western strand of the transform duplex. This observation may suggest an alternative explanation to the excess of earthquakes in the northern site EFE: The site is close to the Jordan (aka Jericho) fault that might act as a waveguide, a property documented for the plate boundary south and north of the Dead Sea [Haberland et al., 2003; Shtivelman et al., 2005]. Guided earthquake waves have been invoked to explain anomalous accelerations and damage in instrumentally recorded Dead Sea events [Wust‐Bloch, 2002]. The seismite sites in the south (EG, ZA) are farther from the Jordan fault, and disconnected from the Araba/Arava fault. This explanation can be tested by a similar research on the eastern shore: it would predict that the southern sites there will show more frequent events.

• from Kagan et al. (2011:16-19)

[52] In this section we discuss the temporal distribution of seismites that are recorded at all of our sites (intrabasin seismites (IBS)). Eight seismic events are recorded in all three sections, north, center and south. In addition we add to this list the 1927 A.D. instrumentally recorded event that formed seismites at the EG and ZA1 sites for which macroseismic evidence was found along the Jordan River [Avni, 1999] and caused slumping under the Dead Sea waters (interpreted by Niemi and Ben‐Avraham [1994]) near the EFE site. The 1927 A.D. event also produced the most pronounced sedimentary structure (in the ZA Gully) with sand liquefaction reaching >1 m in thickness (Figure 8). In addition, two seismites that were recovered from the Ze’elim and Ein Gedi sections and correlated to the 1458 and 1834/1837 A.D. historical events are not represented in Ein Feshkha since this part of the section is missing. However, we predict that processing of the upper part of the section preserved east of our EFE study site will recover these events. Note that the age of the seismite at ZA2 correlated to 1458 A.D. is above the dated and modeled part of the section and its age is extrapolated from the deposition model (see Figures 4 and 5). Part of this group of IBS seismites (mid‐2nd century and 31 B.C. and 33, 419, 1212, and 1293 A.D.) appears in sedimentary sequences of the lacustrine facies indicating clearly offshore conditions of at least 10–20 m of water above the sediment. Other IBS seismites (551, 749, and 1927 A.D.) were at nearshore conditions (hiatus at ZA1, sand and lacustrine sediments at ZA2, lacustrine sediments at EFE and EG). Thus, we see no clear correlation between lacustrine conditions and the three‐site seismite appearance. This observation is corroborated by the lack of seismites in intervals of the lacustrine section at Ze’elim while they appear in Ein Feshkha and Ein Gedi (e.g., between 830 and 1200 A.D.; see Table 4). The conclusion that we can draw from these observations is that the temporal and spatial appearance of the seismites does not depend strongly on the limnological‐sedimentological conditions. Seismites appear in both sandy facies and clay‐evaporite (marly) sequences. The Ze’elim sections contain prominent sand layers that were clearly affected by earthquakes, producing liquefied structures. Significant earthquakes, such as 1927, do appear in all lithological units. This does not imply that low‐magnitude or remote earthquakes have no effect on sandy layers. The topic clearly requires more investigation. If sediments were deposited in the lake they are affected by the earthquakes whether they comprise sands or marls. Figure 10 indicates that soil liquefaction and lacustrine brecciation have apparently similar thresholds. Hence, the archives we documented provide a reasonable picture of the earthquake activity and its effects in the lake basin, not filtered by the lacustrine environment. This conclusion opens the way for using the seismite spatial and temporal distribution to evaluate basin effects and recurrence patterns.

[53] All seismites in the Dead Sea basin are marked on an epicentral distance versus magnitude diagram along with the field of instrumental earthquake data (Figure 9). This diagram highlights domains of intensity, which is a function of magnitude and distance of epicenter from the recording site. In Figure 9 the intensity lines are plotted according to the equation proposed by Ambraseys and Jackson [1998] (here termed A&J):

M_s = 1.54 + 0.65 I_i + 0.0029 R_i + 2.14 log R_i + 0.32p     (1)

where

R_i = (r_i² + 9.7²)^0.5

r_i, in kilometers, is the mean isoseismal radius of intensity I, and p is zero for mean values and one for 84 percentile values. This attenuation relationship is based on 123 instrumentally recorded shallow (depth <26 km) earthquakes from the eastern Mediterranean from a period of 85 years and ∼9000 intensity points. Different coefficients may be more appropriate for the magnitude‐distance field of the earthquakes associated specifically with Dead Sea Rift seismicity. The earthquakes plotted are mainly after the similar diagram by Migowski et al. [2004] and Agnon et al. [2006], where modified input data are explained below and in Appendix A. Each symbol represents a historical earthquake documented in the region, most matched to seismites (open squares), and some matched to seismites at all three sites in the study, the intrabasin seismites (IBS, solid squares). Distances are from the Ein Gedi site, for consistency with previous publications. A field corresponding to earthquakes not matched to seismites is demarcated by the thick gray curves (solid gray curve: earthquakes from historical catalogs; dashed gray curve: instrumentally recorded events). The magnitude‐distance data for each historic earthquake has significant uncertainties (for examples, see Figure 10); however this type of diagram has been shown to be useful [Migowski et al., 2004] for portraying a pattern in the presence of a large sample, barring any systematic bias. Figure 10 depicts only the IBS with estimated uncertainties. Each earthquake shows as a rectangle. We were especially cautious when estimating the upper left corner for each IBS rectangle. This corner, minimum magnitude and maximum distance, corresponds to the minimum intensity at the seismite site, which may be a threshold for intrabasin seismites. The considerations we applied when defining the IBS positions and uncertainties in Figure 10 are given here:

[54] Mid‐2nd century B.C.: Guidoboni et al. [1994] cite one event or more recorded at Antioch (for a summary of historic earthquakes in the region, see Appendix A; for locations of historical cities and towns, see Appendix B). The only traceable historical record for an earthquake comes from the cultural and political center at Antioch, where buildings were reportedly damaged, and Sbeinati et al. [2005] assign local intensity I = VII. For comparison, the 1202 A.D. event was only felt in Antioch, no damage reported [Ambraseys and Melville, 1988; Ambraseys, 2009; Guidoboni and Comastri, 2005]. Therefore if the magnitude of the mid‐2nd century B.C. event is smaller than M7.5 assigned for 1202 A.D., then the source was closer to Antioch and farther from the Dead Sea. Hence for the mid‐2nd century B.C. event we assign an uncertainty rectangle constrained by a bottom left corner coinciding with the 1202 A.D. position. The rectangle represents a range of local intensities spanning V–VII at Antioch, where the distance is calculated to the closer end of the respective rupture (consistent with the magnitude) along the DST. For this specific earthquake we cannot, at present, constrain the top left corner.

[55] 31 B.C.: The magnitudes of 31 B.C. and 749 A.D. are set at 7.2 assuming similarity in rupture length, both reported to have ruptured 110‐km‐long Jordan Valley segment [Reches and Hoexter, 1981]; the sites of damage attributed to the 31 B.C. event demarcate that segment. Ambraseys [2009] points out that 3.5 m dip‐slip displacement reported by Reches and Hoexter [1981] would correspond to an earthquake too large comparing with the historical reports. However, the displacement is measured locally on unconsolidated sediments. Reches and Hoexter [1981] explicitly avoid rejecting the possibility that a part of the slip occurred during several centuries following the event. Moreover they are aware of local complications in the strike of the fault that amplify dip‐slip. Hence we tentatively adopt the identification of the surface rupture with the 31 B.C. event. Gardosh et al. [1990] reevaluated the trench data in light of a newer geomorphic surface faulting study in the Dead Sea area. They conclude that slip accumulation reaches 1.2 m for two events in the past 2000 years on the trench strand. The uncertainty range of the magnitude of this event (Figure 10) is projected from a minimum given by Karcz [2004] and a maximum given by the rupture length discussed here.

[56] 419 A.D.: Damage from this event was reported for Jerusalem and “many cities and towns” and “all great cities” (sources in the works of Russell [1985] and Guidoboni et al. [1994]). Archaeological damage from Antipatris (central Holy Land) has been attributed to this earthquake [Karcz and Kafri, 1978] suggesting a Jordan Valley rupture. We think that it is feasible that the source of this event was similar to that of 1927 A.D. earthquake (see below). We assume a 6 ≤ M ≤ 6.5, with a maximum distance of 50 km.

[57] 551 A.D.: The event was updated to a larger magnitude offshore Lebanon earthquake, as is more widely accepted in the literature (Appendix A). Magnitude estimation is based on sonar imaging of seafloor morphology [Elias et al., 2007] and historical account compilation [Sbeinati et al., 2005].

[58] 749 A.D.: The historical sources are consistent with a rupture event or two in the Jordan Valley (between the Dead Sea and Sea of Galilee). The range of magnitude (M6.6–7.7) in Figure 10 reflects either a single event or a double event with a cumulative rupture of that 110‐km‐long segment (calculated using the results of Wells and Coppersmith [1994], Marco et al. [2003], and Karcz [2004]; see Appendix A).

[59] 1202/1212 A.D.: A single event brecciated the sediments in the EFE section in the early 13th century. Two events are recorded in EG. ZA recorded one or two events. Therefore only one of them is an IBS and the dating cannot rule which. For the 1202 event we use M7.4–7.6 based on historical analysis of Ambraseys and Melville [1988] and Ellenblum et al. [1998]. Paleoseismic and archaeoseismic trenching corroborate these assessments [Ellenblum et al., 1998; Marco et al., 2005; Daeron et al., 2007; Nemer et al., 2008]. The distance of the rupture edge from the farthest seismite site is 165 ± 10 km, based on rupture uncovered in trenching at the northern shore of the Sea of Galilee [Marco et al., 2005]. For the 1212 event, Ambraseys et al. [1994] suggest a rupture south of the Dead Sea or in the Gulf of Eilat (Red Sea). In severity of damage and aftershock occurrence it is seemingly similar to the 1995 modern event [Hofstetter, 2003], or could have been closer to the Dead Sea, according to the high level of damage at Aila and Karak. This similarity prompts us to give a best estimate of 7.2 magnitude and 300 km distance.

[60] 1293 A.D.: Based on evidence at an archeological site built on the Arava segment of the DST, the northern Arava did not rupture during this event [Haynes et al., 2006]. We consider the 12‐km‐long Amatzyahu fault (Figure 1) as the source for this event. This rupture length is consistent with a 6.2–6.7 magnitude earthquake. The maximum intensity recorded for this event was recorded at Karak (eastern Dead Sea), 45 km from the Amatzaya fault [Ambraseys et al., 1994; Guidoboni and Comastri, 2005], consistent with a magnitude of 6.7 according to the A&J equation (equation (1)). Taking into account poor construction and site effects this intensity could be achieved at a somewhat lower magnitude.

[61] 1927 A.D.: This event was recorded instrumentally [Shapira et al., 1993] (M6.2) and its distance uncertainty range is based on the distance from the ZA site to the Kalia transverse fault in the northern Dead Sea (Figure 1). It is also a possible scenario that the main fault of the DST ruptured along a limited length causing the 1927 earthquake.

[62] We have excluded the 33 A.D. IBS event from Figure 10 for lack of reliable historical evidence [see Ambraseys, 2009].

[63] In addition to the IBS magnitude‐distance discussion in this section (above), other modifications (Figure 9) made to the published magnitude‐distance diagrams are explained here. Regarding the 363 A.D. event, our review of the evidence indicates two or more separate earthquakes from ∼362 and 363 A.D., with damage in geographically disparate regions (see Appendix A). Also, symbols were added (in Figure 9) for 331 and 199 B.C. and 835 and 847 A.D. historic earthquakes, which are matched to seismites in this study, but not in previous studies at the Dead Sea basin. For the 331 B.C. event, Sbeinati et al. [2005] give intensity VI in the general region of “Syria.” For this ancient and not well‐covered event only a rough calculation is possible. An isoseismal distance of 70 km is consistent with a M6.5 earthquake using the attenuation relation of A&J. This is a relatively ancient event, population density was low, and a distance of ∼70 km from seismic source to historic source is reasonable. For the 199 B.C. event, assuming the intensities documented are from the same event (VII and VIII in “Syria,” probably Damascus, and Sidon, respectively [Sbeinati et al., 2005]), the magnitude is estimated in the same way to Ms6.8. For 847 A.D. the magnitude is taken from the analysis of Sbeinati et al. [2005]. The 873 and 956 A.D. events [Ambraseys et al., 1994; Guidoboni et al., 1994], matched to seismites in this study, are not on the distance‐magnitude diagram for lack of sufficient information.

[64] Second earthquakes were added at ∼mid‐2nd century B.C. and at 362/363 and 747/749 A.D. The location and magnitude of these added events are not known, each appearing in the diagram as a small circle on the symbol of the previously published single event.

[65] The intrabasin seismites that were recorded in all three sites (EFE, EG, ZA) define a well‐constrained field in the magnitude‐distance diagram, which cuts the A&J intensity lines plotted (Figures 9 and 10). Of the earthquakes matched to seismites on this diagram, 60% occupy the field of intensities larger or equal to IV. Eighty‐nine percent of the IBS seismites occupy the field of intensities larger or equal to V (or 100% if 1202 is chosen over 1212 A.D.; see above discussion), as opposed to 46% of all seismites.

[66] Figures 9 and 10 suggest that farther and stronger earthquakes require lower local intensities for being recorded in the entire basin (IBS). If we accept that 1212 A.D. is the IBS (as opposed to 1202 A.D.) at the beginning of the 13th century then it is the farthest (300 km) with M7 and I = IV. Otherwise the 551 A.D. earthquake and the mid‐2nd century B.C. earthquake are the farthest. The intensity threshold for magnitude 6.2 seems to be VII (419 and 1927 A.D.). A possible explanation for this observation is sensitivity to long‐period waves. The frequency content of the wave train is biased to long periods in earthquakes from large and remote sources. A Ms6 earthquake shows a corner frequency fc ∼ 0.1 Hz (period ∼10 s), whereas Ms7 shows fc ∼ 0.04 (period 25 s) [e.g., Geller, 1976]. Attenuation of the wave during travel, where the waves are damped according to the number of cycles between the source and the site in question, results in further bias toward lower frequencies.

[67] The sensitivity to low frequency may indicate that the critical condition for brecciation may depend on ground velocity rather than ground acceleration, where the frequency equals the ratio of the latter to the former. Heifetz et al. [2005] and Wetzler et al. [2010] suggest a Kelvin‐Helmholtz instability mechanism for the disturbances in the sediments (intraformational folding leading ultimately to brecciation). In this scenario the sediment bed is considered to have a gradient in the horizontal velocity (due to a density decrement). If the duration of the wave cycle is sufficiently long (or the frequency sufficiently low), a disturbance can be sustained: the growth rate of a disturbance must be larger than the driving frequency.

[68] The thick gray curve on Figure 10 represents the farthest epicentral distance of liquefaction of soil caused by modern earthquakes in the Aegean region [Papathanassiou et al., 2005]. Note that if 1202 A.D. is the date of the early 13th century IBS (as opposed to 1212 A.D.) then the threshold for intrabasin seismite genesis is very similar to this soil liquefaction curve.

[69] The average recurrence time of IBS is ∼200 years, which is significantly longer than the ∼50–95 years based on all seismites in the Ein Gedi core during the past 1600 years [Migowski et al., 2004] or ∼50 years at EFE since 525 B.C. The possibility to establish a high‐resolution comparison between distinct sedimentary sections located in different sites of the Dead Sea basin opens the way to further explore the response of the lacustrine system to various sources of seismic activity and thus extends the paleoseismic study to older sections such as those of the last glacial Lake Lisan. Such a comparison is currently under investigation.