Global and Planetary Change 103 (2013) 248–260
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Global and Planetary Change
journal homepage: www.elsevier.com/locate/gloplacha
Quaternary mass wasting on the western Black Sea margin, offshore of Amasra
Derman Dondurur ⁎, H. Mert Küçük, Günay Çifçi
Dokuz Eylül University, Institute of Marine Sciences and Technology, Bakü Street, No: 100, 35340 İnciraltı, İzmir, Turkey
a r t i c l e
i n f o
Article history:
Received 28 June 2011
Accepted 14 May 2012
Available online 26 May 2012
Keywords:
sliding
Amasra mass failure zone
debris flows
gas hydrate dissociation
a b s t r a c t
In recent years, the western Black Sea margin has become well-studied due to its potential for petroleum
plays in relatively deeper waters. In 2010, multi‐channel seismic, multibeam bathymetry and Chirp high
resolution seismic data were collected in order to define the existing geohazards along the margin, to identify
the seabed morphology and to determine mass movement types and their run‐out distances.
Seismic data indicate that the western Black Sea margin is an unstable region with sediment erosion. Particularly, an unstable area offshore of Amasra in the NW consisting of four slides and four buried debris lobes is
named the Amasra mass failure zone. Different types of sliding with varying sizes and different mechanisms
are observed. These include sliding in the steep slope zones where block‐type sliding occurs, smaller‐scale
slides on the canyon walls, and relatively larger slides in the Amasra mass failure zone. Block‐type sliding
is observed on the upper continental slope to the south as well as on the canyon walls. They are formed
along the rotational faults and occur due to the gravitational loading on the steep slope zones possibly
triggered by local seismic activity. In addition, seven large debris lobes identified in the northern toe of the
slope buried in the Quaternary sediments triggered by excess pore pressures due to high sediment input
and submarine fluid flow.
We suggest that earthquake activity may be an important agent for all kind of mass movements in the area. In
addition, we propose that the slides in the Amasra mass failure zone are triggered by excess pore pressures in
shallow sediments due to the submarine fluid flow possibly produced from gas hydrate dissociation. Warmer
Mediterranean seawater input during the rapid transgression period after the Last Glacial Maximum in the
Black Sea together with the rapid sedimentation resulted in destabilization of gas hydrates, which caused
excess pore pressures in shallow sediments leading to massive sediment failures. Small‐scale normal faults
around the scarps may be a secondary factor promoting the failures providing the suitable pathways for
the fluid flow as well as the suitable weak surfaces for the sliding.
© 2012 Elsevier B.V. All rights reserved.
1. Introduction
Continental slopes are the regions extending from shallower continental shelf to deep abyssal plain with relatively high bathymetric
gradient. Downslope sedimentary processes under the effects of gravitational loading, bottom currents, earthquake activity and existence
of seabed fluid flow and gas hydrates in the sediments are widely
observed along the slopes (Eschard, 2001; Casas et al., 2003; Dondurur
and Çifçi, 2007). These processes primarily affect the development,
form and morphology of the continental slopes.
The study of mass movements along the canyons and continental
slopes has important social and economic implications because they
constitute a potential geohazard for offshore engineering installations
such as submarine pipelines, cables and drilling platforms and they can
produce destructive tsunamis. These processes include gravity flows
and submarine landslides generally triggered by seismic activity,
⁎ Corresponding author. Tel.: + 90 232 2785565; fax: + 90 232 2785082.
E-mail address: [email protected] (D. Dondurur).
0921-8181/$ – see front matter © 2012 Elsevier B.V. All rights reserved.
doi:10.1016/j.gloplacha.2012.05.009
oversteepening and overloading of the slope as well as tides and
storm waves (Mulder et al., 2009). Sediment gravity flows are the
flow of sediments – or sediment/fluid mixtures – under the action of
gravity and are classified as turbidity flow, fluidized sediment flow,
grain flow and debris flow (Middleton and Hampton, 1973), which
are the effective processes in the evolution and morphology of the
margins.
Recent studies indicate that natural gas hydrates play an important
role on the dynamics of the seafloor (Sultan et al., 2004b; Mienert et
al., 2005; Gee et al., 2007; Owen et al., 2007). Variations in sub‐bottom
thermobaric conditions of gas hydrate accumulations are generally
due to the climate changes, e.g., an increase in the temperature and/
or a decrease in the ambient pressure due to sea level fall, that result
in dissociation of gas hydrates in the gas hydrate stability zone. The
dissociation produces huge amounts of methane in the uppermost
sediments and causes excess pore pressures in a weak layer close to
seafloor, along which the sediment failure may be triggered
(Mienert et al., 2005). The produced methane is released into the
water column and then into the atmosphere, which negatively affects
the climate conditions since methane is a greenhouse gas.
D. Dondurur et al. / Global and Planetary Change 103 (2013) 248–260
This study is based on the interpretation of multi‐channel seismic
reflection (MCS), bathymetric and Chirp sub‐bottom profiler acoustic
datasets collected simultaneously along the western Black Sea Turkish
coastal waters (Fig. 1) including the Turkish continental slope and rise
during a survey aboard the R/V K. Piri Reis of Dokuz Eylül University,
Institute of Marine Sciences and Technology, İzmir in June 2010. The
aim of the study is (i) to identify the seabed morphology and recent
sedimentary processes in the western Black Sea Turkish margin,
(ii) to determine mass movement types and their run‐out distances,
and (iii) to describe the existing and potential geohazards associated
with the mass movements. We also discuss the distribution of the
mass wasting structures and their possible triggering mechanisms.
2. Regional setting
The Black Sea is a large intercontinental basin located on the western
flank of the active Arabia–Eurasia collision and north of the regional
right‐lateral strike–slip North Anatolian Fault–NAF (Finetti et al.,
1988; Robinson et al., 1996). It is a Mesosoic–Early Cenozoic marginal
back‐arc basin generated by the northwards subducting Tethys Ocean
(Okay et al., 1994; Nikishin et al., 2003). The Black Sea basin comprises
western and eastern Black Sea sub‐basins, which are separated by a
regional high, the Mid Black Sea Ridge. This ridge is subdivided into
the Andrusov and Archangelsky Ridges to the north and south, respectively (Fig. 1a). Although the Black Sea has an extensional origin, the
tectonic setting changed to a compressional system due to the collision
between Eurasia and Arabia in Eocene time, and the margins of the
Black Sea are currently characterized by compressive deformation
(Robinson et al., 1996; Spadini et al., 1996; Tarı et al., 2000).
The western and eastern sub‐basins have different kinematics and
separate origins with different rifting histories (Okay et al., 1994;
Spadini et al., 1996). It is proposed that the western Black Sea basin
was opened by the rifting of the western and central Pontides from
249
the Moesian Platform that initiated in the Aptian, when a part of the
Moesian Platform (now the western Pontides of Turkey) began to
rift and move away to the south‐east (Görür, 1988). However, the
eastern Black Sea basin is younger and opened by the separation of
the Shatsky Ridge and the Mid Black Sea Ridge by a rotation about a
pole west of Crimea during the Palaeocene to Eocene (Finetti et al.,
1988; Robinson et al., 1996; Spadini et al., 1996). The western Black
Sea basin is floored by oceanic crust and the thickness of the post‐rift
sediments since the Upper Cretaceous reaches approx. 13 km in the
center of the basin (Finetti et al., 1988; Robinson et al., 1996). Morphologically the Black Sea has two different margin types: The continental
shelf is not well developed along the eastern and southern margin,
and the continental slope is so steep that the water depth reaches to
approx. 1800 m depths only approx. 15 km beyond the shelf break;
whereas there is a wider shelf along the northern and western margins
with a gentle continental slope gradient.
The Black Sea and surroundings is defined as a low seismicity or
“silent” region (Tarı et al., 2000). Seismicity of the central parts of
the deep basin is negligible, and the most important seismic activity
along the southern margin is not related to the Back Sea itself but is
related to the NAF (Fig. 1a) which extends approx. 1500 km from
eastern Turkey to the Aegean Sea and separates the northern Turkey
province and the Black Sea regions from central Anatolian province.
According to recent seismicity data (Barka and Reilinger, 1997), compressional deformation in the Anatolia is still active in the western
sub‐basin, while the eastern margin is nearly completely aseismic. The
Bartın earthquake in 1968 (MS = 6.6) is the strongest instrumentally
recorded earthquake in the northwestern Turkey along the margin,
and the source mechanism indicates thrust faulting (Alptekin et al.,
1986).
The western Black Sea margin has been explored by Turkish Petroleum Co. by 2D MCS surveys and two commercial wells were drilled
close to study area (Fig. 1b). Akçakoca‐1 and Ayazlı‐1 wells penetrated
Fig. 1. (a) Major tectonic elements of the Black Sea and surroundings (modified from Finetti et al., 1988; Robinson et al., 1996; Spadini et al., 1996). Location and fault plane solution of
Bartın Earthquake are from Tarı et al. (2000). (b) Simplified bathymetric map of the study area and tracklines of the data collected along the western Black Sea margin. Bathymetric
contour interval is 200 m. Topographic map is from GeoMapApp (http://www.geomapapp.org). Thick parts of the solid lines correspond to the data illustrated in figures.
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D. Dondurur et al. / Global and Planetary Change 103 (2013) 248–260
reverse‐faulted ramp anticlines in Eocene structures of the Pontides and
produced gas from Early Eocene turbidite reservoirs (Menlikli et al.,
2009). The lack of well information in deeper areas around the continental rise prevents an exact timing of the stratigraphy in deeper
waters. Robinson et al. (1996) correlated the Akçakoca‐1 stratigraphy
to the commercial MCS lines, and they suggested that their regional
MCS data indicates approx. 1.5 km thick Quaternary sequence in the
deep basin over thick (approx. 2.5 km) post‐rift sedimentary rocks of
Neogene age and thin (500 to 700 m thick) Late Eocene–Early Oligocene
sediments. MCS reflection data also indicate Cretaceous age syn‐rift
ridge formations, e.g., the Kozlu Ridge, located close to the western
part of the study area (Menlikli et al., 2009).
3. Data acquisition and processing
High resolution MCS, Chirp sub‐bottom profiler and multibeam
bathymetry data sets collected simultaneously along the margin. A
global DGPS system was utilized during the entire survey with an
integrated navigation system. A 216 channel, 1350 m‐long digital
streamer with 6.25 m group interval was used during MCS data
acquisition and a total of 1950 km of MCS data were recorded. The
record length and sampling interval were 6 s and 1 ms, respectively.
The seismic source was a 45 + 45 in. 3 Generator-Injector (GI) gun
fired every 25 m. GI guns suppress their own bubble pulse generating
a sharp seismic signal between 8 and 240 Hz frequency band, and
therefore, they are preferred for high resolution MCS exploration. The
MCS data were processed using Vista Seismic Processing Software
from Gedco and a conventional processing sequence was applied to
the MCS data as geometry definition, band‐pass filter (8 to 220 Hz),
trace editing, f–k dip filter, sort to 27‐fold CDP gathers, velocity analysis
(approx. every 1500 m along the lines), normal move‐out correction,
stacking, post‐stack time migration and gain recovery. All seismic line
interpretations are based on time-migrated data.
Bathymetric data were collected using a pole‐mounted Elac
SeaBeam 1050D multibeam system operating at 50 kHz. The system
utilizes 126 beams with 1.5° resolution providing a total swath coverage
of 153°. The multibeam data were processed using the Caraibes
software with the following processing steps: beam editing and de‐
spiking, correction of navigation errors, data interpolation, digital
terrain model (DTM) construction and gridding with 100 m grid
interval.
The shallow sedimentary structure was investigated using a pole‐
mounted Chirp sub‐bottom profiler utilizing a sweep signal between
2.75 and 6.75 kHz centered at 3.5 kHz. Delay‐time correction, gain
recovery, de‐chirping and amplitude envelope calculations were
applied to Chirp data.
4. Results
The Western Black Sea margin can be defined as a deep water
turbidite system with morphological elements of highly dissected
canyon systems with widespread overbank gullies, slides of various
sizes, stacked turbidity‐flow leveed channels and large buried debris
lobes in the distal area as well as sediment waves. The collected
acoustic data were analyzed by means of debris flow lobes, and slides
and slide scars. The main structural and morphological elements of the
study area, such as sedimentary ridges, buried debris lobes, sediment
wave fields and slides together with the bathymetry of the slope
deduced from our acoustic dataset, are shown on the generalized
map in Fig. 2. The map also shows the location of the Pontides thrust
belt to the south as well as the limit of extensional tectonics towards
the deep basin to the north revealed from the dataset.
The continental slope of the area between 120 and 1400 m depth
contour as well as some confined zones between the ridges have
extremely steep slopes named “steep slope zone‐SSZ”. SSZs are defined
as the anomalously steep areas with local seafloor slopes ranging from
10 to 27° indicated by blue polygons in Fig. 2. The MCS data has very
limited penetration on these oversteepened zones. The continental
rise, however, lying between 1400 and 2100 m water depths, is the
major depocenter where the analysis of erosional structures is concentrated. We identify a highly unstable area offshore of Amasra which we
name it “Amasra mass failure zone‐AMFZ” indicated in dashed pentagon in Fig. 2. This zone has four separate slides and four individual
buried large debris lobes.
Fig. 2. The generalized map showing main structural and morphological elements of the study area revealed from our acoustic dataset. A detailed map of compressional tectonic
belt of Pontides and the limit of extensional tectonics of the deep Black Sea basin to the north is also provided. The dashed pentagon corresponds to a highly unstable area named
Amasra mass failure zone (AMFZ). Bathymetric contour interval is 100 m.
D. Dondurur et al. / Global and Planetary Change 103 (2013) 248–260
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4.1. Seabed morphology
4.2. Buried debris flow deposits
The morphology of the western Black sea continental margin has
the characteristics of a modern ocean margin with a shallow shelf, a
steep continental slope, a continental rise with gentle slopes and a
smooth abyssal plain. In our study area, the shelf break is located at
approx. 120 m depth contour. The continental shelf is very narrow
with a mean width of approx. 3 km. It is generally smooth and flat
but is deeply incised by submarine canyons immediately below the
shelf break. The continental rise of the area is considered as the
main depositional environment for the terrigenous sediment load
supplied via the continental slope. It is connected to the narrow
shelf with an extremely steep slope located between 120 and
1400 m water depths with a maximum inclination of 27°.
The canyons show tributary characteristics and broaden slightly as
traced downslope. Fig. 3a and b shows MCS and Chirp sub‐bottom
profiler lines respectively, which cross‐cut a number of gullies
between two small canyon banks (indicated as Bank A and B) on the
northern rim of a large canyon bank from the eastern part of the
area. The flanks of the gullies are of strong erosive nature and are identified by their distinct diffraction hyperbolas. The MCS line in Fig. 3a
shows that Plio‐Quaternary sediments overlying the Miocene unconformity are highly deformed by small‐offset normal fault activity,
most of which also affect the Miocene sediments. Between the seabed
and the deformed strata down to approx. 400 ms from seafloor, we
observe an unstable sediment zone consisting of semi‐chaotic reflections. The bottom of this zone is sub‐parallel to the seafloor and is
defined as the “primary glide surface” over which the sediment sliding
mainly along the rotational faults occurs resulting in a semi‐chaotic
reflection pattern (see also Fig. 6 and Section 4.3.1).
A number of buried large‐scale debris flow deposits or debris lobes
(DLs) in the deep water area (e.g., to the northernmost part of the
study area with relatively smooth bathymetric gradient) were observed
and they are indicated as DL1 to DL7 from west to east (Fig. 2). Flow
directions of all DLs are towards the deep basin from south to north.
They are generally lens‐shaped lobes with largest thickness at the
central part, and are easily recognized by their semi‐transparent to
transparent internal structure on the MCS data. They show almost no
internal reflections possibly due to an irregular deposition of the unconsolidated material during the failure. Their upper and lower surfaces are
characterized by well defined erosional surfaces.
Fig. 4 shows two example buried debris flow lobes (DL2 and DL3) on
the MCS data. The seismic line extends from the end of a steep slope
zone, crosscutting the Pontides thrust belt to the south, to deep abyssal
plain to the north. The eastern extension of Kozlu Ridge is located in the
central part of the section indicating a highly deformed ridge crest with
the small‐offset faults. The debris lobe DL3 is located at the northernmost part of the line buried in Plio-Quaternary sediments of the abyssal
plain (Fig. 4b). In addition to this type of larger debris lobes, indications
of small‐scale debris flows are also observed in shallower sediments
(not mapped here, see Fig. 4a for examples).
Table 1 shows some of the calculated geometrical properties of
observed DLs. The lobe DL3 has the largest run‐out distance of approx.
25 km, which has an estimated volume of 12 km3 and an area of approx.
263 km2. We classify debris lobes into two groups: the western (DL1 to
DL4) and eastern (DL5 to DL7) ones. The western DLs are located at
shallower sub-surface depths than those in the eastern area. Assuming
a 1600 m/s average velocity in the upperlying sediments for western
Fig. 3. (a) Part of MCS line showing gullies between two small canyon banks (Bank A and B). Dashed blue line represents the interface (named primary glide surface) between an
unstable sediment zone with semi‐chaotic reflections and rotational faults, and Plio‐Quaternary sediments highly deformed by small offset faults. (b) Chirp sub‐bottom profiler line
along the same profile in (a). Multibeam bathymetric map shows the location of the line. Vertical exaggeration is 6× and 9× for (a) and (b), respectively.
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Fig. 4. (a) MCS line extending from shallower SSZ to abyssal plain showing two buried large debris flow lobes DL2 and DL3, and (b) close‐up and interpretation of DL3 debris lobe.
Multibeam bathymetric map shows the location of the line. Vertical exaggeration is 13×.
DLs and 1800 m/s for eastern DLs, the burial depths of DLs change from
42 to 112 meters below the seafloor (mbsf) for the western DLs and 141
to 222 mbsf for eastern DLs. An example MCS line with an for eastern DL
(DL7) is also shown in Fig. 5.
Along with the DLs, we also observe sediment waves migrating
upslope possibly formed by turbidity currents along the continental
slope to the continental rise in the north where the slope of the seafloor is approx. 2 to 3° and water depth is around 2000 m (Fig. 2).
Sediment waves occupy a surficial area of approx. 1100 km 2 and are
located in Plio-Quaternary sediments (see Figs. 4a and 5a). General
geometric parameters describing the shape of the sediment waves
from high‐resolution seismic data are given in Fig. 5c. For our study
area, the wavelength of the sediment waves (defined as the horizontal
distance between two successive troughs or two peaks) generally
varies between 230 and 1350 m, and in places, it rarely reaches 2000 m
towards the west. Considering a 1500 m/s velocity both for water column
and upper sediments, their wave height (defined as the vertical distance
between successive troughs and peaks) ranges from 6 to 33 m and generally increases towards the seafloor. As an example, angle of climb for the
sediment waves in Fig. 5b ranges between 8.2 and 8.9°.
4.3. Sliding
The slides in the study area are subdivided into three groups based
on their morphology and formation environment, as
(1) block‐type sliding along the rotational faults in the SSZ
between 120 and 1400 m water depths and the steep areas
Table 1
Geometrical characteristics of buried debris lobes (DLs) calculated from MCS data.
Group
Western DLs
Debris lobe
Surficial area (km2)
Total volume (km3)
Water depth (m)a
Thickness (m)b
Depth (mbsf)c
Run‐out distance (km)
DL1
116.3
2.3
1812
22
112
20.44
a
DL2
30.79
0.9
1846
26
80
9.57
Eastern DLs
DL3
263.50
12.2
2014
36
64
24.82
DL4
90.91
3.8
1906
28
42
18.25
DL5
108.2
2.2
2150
15
141
12.67
DL6
23.76
0.4
2180
20
230
7.4
DL7
40.16
1.1
2084
19
222
11.1
Water depth at the head zone (commonly southernmost point).
Maximum thickness of the lobe at the center assuming a 1600 m/s average velocity
for western DLs and 1800 m/s for eastern DLs.
c
Burial depth at the head zone assuming a 1600 m/s average velocity for the
upperlying sediments for western DLs and 1800 m/s for eastern DLs.
b
between the ridges (the areas limited by blue polygons in
Fig. 2),
(2) smaller‐scale slides on the canyon walls in the continental rise
to the north, and
(3) relatively larger slides in the AMFZ.
4.3.1. Slides in the SSZ
In the SSZs of the study area, generally block‐type sliding is
observed. This type of sliding occurs above a primary gliding surface
which is either sub‐parallel to the seabed or concave upwards
(Figs. 3a and 6b). There are well defined rotational faults between
the blocks, which allow the rotational sliding along the fault planes
which produce secondary glide surfaces. The back‐rotated blocks
create positive topographic relief on the seafloor resulting in a
hilly surface, and their internal structure has generally chaotic
reflection pattern. Scar faces located at the upslope side of the
blocks can easily be recognized on the MCS data, indicating that
the sliding along the rotational faults on the upper slope is a recent
process.
Fig. 6a shows an example MCS line extending from shallower
upper slope to the end of continental rise. The southern part of the
line is defined as SSZ and successive block‐type mass failures exist
in this anomalously steep area. To the north, the line shows a bottom
simulating reflector (BSR) with a possible free gas accumulation
imaged below as an acoustic turbidity zone. The line also crosscuts
the distal part of a large canyon which is 8 km wide across the channel
banks and its depth to thalweg is approx. 460 m. The bathymetric
gradient of the seafloor along these canyon walls is as much as 10°.
The close‐up in Fig. 6b shows the acoustic structure of the canyon
walls in detail. The primary glide surface is easily distinguishable
approx. 200 mbsf. Above this surface there are rotational faults
concave to the channel banks at both sides of the canyon. The average
distance between the rotational faults is 250 m. They separate the
back‐rotated slide blocks with hummocky crests on the seabed
forming small‐scale scarps at the upslope sides. The depth extend
of the rotational faults can be traced down to the primary glide
surface.
4.3.2. Smaller‐scale slides on canyon walls
There are also small‐scale slides from canyon banks to the canyon
axes located on the walls of the distal parts of the major canyons
towards the north (Fig. 7). They tend to appear as multiple slides
D. Dondurur et al. / Global and Planetary Change 103 (2013) 248–260
253
Fig. 5. (a) Part of MCS line showing buried debris lobe DL7 and sediment waves. Vertical exaggeration is 13×. Thick dashed lines correspond to the peaks of the sediment waves. See
Fig. 1b for location. (b) Close‐up of sediment waves with peaks marked with dashed lines, and (c) general geometric parameters describing the shape of the sediment waves from
high‐resolution seismic sections (modified from Bøe et al., 2004).
with small slide scars. Their major difference from block‐type sliding
is that they do not occur along the rotational faults. Instead, the
sediment failures are formed along well defined slide surfaces and
they produce distinct erosional scarps behind the head zones. Their
internal facies is generally preserved possibly due to the smaller
run‐out distances which are generally less than 400 m.
4.3.3. Slides in the AMFZ
MCS and bathymetric data indicate that there is a submarine slide
zone offshore of Amasra where a number of relatively larger slides
exist. Together with four buried large debris lobes in the proximity,
four individual slides are identified in the AMFZ and are indicated as
ASL1 to ASL4 in Fig. 2.
Table 2 shows some of the calculated geometrical properties of the
slides in the AMFZ. The headwalls of the slides are located at water
depths between 1383 and 1733 m and their run‐out distances change
from 10.9 to 15.9 km. The largest slide is ASL1 with 244 m maximum
headwall scarp height and it affects an area of approx. 109 km 2
(Table 2). The slides cover a total surficial area of 280 km 2.
The MCS line in Fig. 8 crosscuts ASL1, ASL2 and ASL3 slides in the
AMFZ. Failure directions of ASL1 and ASL3 slides are perpendicular to
the seismic line. Below the canyon banks at both sides of the slides,
several small offset vertical faults around the headwalls are observed,
which are possibly formed by differential compaction. A chaotic facies
of slide debris is observed over the glide planes. A somewhat intact
coherent facies overlies this slide debris. Assuming a 1600 m/s
Fig. 6. (a) MCS line extending from shallower SSZ to the end of continental rise. Vertical exaggeration is 6×. (b) Close‐up of the large canyon showing block‐type sliding along the
rotational faults over a primary glide surface. Multibeam bathymetric map shows the location of the line.
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Fig. 7. Two MCS data examples for smaller‐scale sliding across the canyon walls. This type of sliding occurs along a well defined slide surface on the seismic data and produces smaller
scarps at the head zone. Multibeam bathymetric map shows the location of the lines. Vertical exaggeration is 6× and 13× for (a) and (b), respectively.
average velocity, it is approx. 100 m thick for ASL1 and ASL4 while it
is approx. 50 m for ASL2 and ASL3 (Table 2).
The coherent facies has irregular seafloor topography and its internal
structure can be correlated with that in the intact area at both sides
beyond the sidewalls. Fig. 8c shows Chirp sub‐bottom profiler data
from ASL2 and ASL3 slides, in which the internal structure of the coherent facies undergoes some internal deformation by small‐offset sub‐
parallel faults. These faults, however, can only be observed on the
Chirp data and are far beyond the resolution limit of MCS data. In
some places, a thin drape of post-slide sediments (generally 7.5 to
12 m thick) overlying the coherent facies can be distinguished on the
Chirp sub‐bottom profiler data.
MCS data also indicates widespread BSR occurrences approx.
180 mbsf especially below the slide zones. These BSRs are inferred
to be the base of gas hydrate accumulation zone. Possible gas accumulations and gas chimneys are also identified especially beneath
the headwalls of the slides (Fig. 8a and b).
The sidewall and headwall erosions of the slides as well as the
glide planes are well defined on the MCS data. A BSR reflection with
an acoustic turbidity zone behind the SE sidewall of ASL1 and another
BSR directly below the ASL2 are observed (Fig. 9). The sidewalls of the
slides can be identified by erosional truncation of the sediments at
both sides. The coherent facies has some minor deformation with
small‐offset normal faults penetrating down to the slide debris facies,
which are possibly formed during the sliding event.
Table 2
Geometrical characteristics of the slides in the AMFZ calculated from MCS and bathymetric
data.
Slide name
ASL1
ASL2
ASL3
ASL4
Surficial area (km2)
Water depth (m)a
Height of scarp (m)b
Run‐out distance (km)
Mean drape thickness (m)c
Failure direction
109.2
1440
244
14.7
100
NE to SW
91.5
1636
85
15.9
56
SE to NW
43.4
1733
132
10.9
48
SW to NE
35.5
1383
182
11.8
96
SW to NE
a
Water depth at the headwall zone.
Maximum height of the headwall.
c
Mean thickness of the sedimentary drape overlying the slide debris assuming a
1600 m/s average velocity.
b
5. Discussion
5.1. Sedimentation and source area
The characteristics of source area, such as its dimensions and
cementation of the sediments it contains, widely affect the sedimentation and the nature of the sediments in the deposition (or sink)
area. The climate changes in the source area are, however, the most
important factor controlling the sediment supply. In addition, the
greater the topographic relief and uplift in the source area, the greater
the sediment supply. The catchment area supplying the sediments is
the western Pontides region, while the main depositional environment is the continental rise of the study area. The sedimentation
rate in the continental rise is relatively high. Ross (1977) suggested
a sedimentation rate of >30 cm/1000 years. Duman (1994) proposed
100 to 200 cm/1000 years for the same region, and showed that
turbidites are very common towards the basin reaching >50 cm
thickness during the last 200 years. We do not have more detailed
information on the composition, grade, amount and contents of the
sediments deposited in the area.
There are a number of large sedimentary basins on land close to
the study area filled with Cretaceous to Tertiary sedimentary and
volcanic rocks (Tüysüz, 1999). The closest basin on land is Zonguldak
Basin extending from Ereğli–Zonguldak to Amasra parallel to the
Black Sea coast line. We suggest that the terrigenous sediments are
sourced from this large sedimentary basin and they are re‐mobilized
by river erosion and then transported to the coastal area by two large
rivers, Filyos and Bartın Rivers (see Figs. 1 and 2 for locations). This
sediment input from land is later transported from the continental
shelf to the deep basin most probably by turbidity current activity
along the canyon systems.
5.2. Debris flows
We interpret the buried large debris flow lobes as gravity flows of
unconsolidated sediments and as slurry in the areas of low slope
gradient. All of the DLs are located in the Quaternary sediments. However, the timings of the western and eastern DLs are clearly different;
the eastern DLs are located in deeper sediments and hence they
are older (Table 1). The existence of smaller‐scale debris lobes in
D. Dondurur et al. / Global and Planetary Change 103 (2013) 248–260
255
Fig. 8. (a) Part of MCS line crosscutting three major slides in the AMFZ (ASL1, ASL2 and ASL3), and (b) its line drawing interpretation. Failure directions of ASL1 and ASL3 are
perpendicular to the seismic line while it is in‐line with the seismic section for ASL2 slide. BSRs, gas fronts and gas chimneys exist below the slides. (c) Chirp sub‐bottom profiler
data showing ASL2 and ASL3 slides. Multibeam map shows the location of the seismic line and the slides in the AMFZ. Vertical exaggeration is 13× and 9× for (a) and (c),
respectively.
shallower sediments on the upper slope indicates that mass wasting
as debris flows is an ongoing process along the margin.
We do not know the composition and internal facies of the debris
flow deposits. Their base reflector has an erosional character and the
upper surface has generally rough and hummocky pattern indicating
that the flows erode the upper sedimentary layer and carry significant
amount of slurry. Similar, but larger (4200 km 2 total area, 56 km 3
total volume and up to 100 km run‐out distances), debris flow lobes
were also reported along the East Greenland continental margin
(Wilken and Mienert, 2006).
Theoretical investigations on the frontal dynamics of a debris flow
indicate high mobility of the frontal part due to the hydroplaning
which is the intrusion of a lubricating water layer underneath the
front reducing the basal friction (Ilstad et al., 2004). This results in a
flow with large run‐out distances even on the low gradient slope
such as our study area. We suggest that higher sedimentation rates
in the area can generate high pore pressures in a thin sediment
layer and result in the initiation of various‐size debris flows in the
toe of the continental slope, as proposed by Imbo et al. (2003) for
the Trinity Peninsula Margin, Antarctica. In addition, submarine
fluid flow could also be an important contributing factor for the
high pore pressures in the shallow sediments along the continental
rise.
5.3. Slides
In the SSZs, the mass movements occur as block‐type sliding along
the rotational faults. Similar observations were also reported from
eastern Black Sea margin by Dondurur and Çifçi (2007) where they
observed rotational faults in the areas defined as open continental
slope or canyon banks. Using high resolution sub‐bottom profiler
data, they showed stacked slide blocks as relict slides between the
rotational faults and discussed a model for their periodic occurrence.
We observe similar blocky structures and stepped seafloor topography
between the rotational faults over a primary glide surface. The relict
slides, however, cannot be distinguished by our MCS dataset indicating
that they do not exist or they are beyond the resolution limit of the MCS
data. It is assumed that block‐type sliding is formed by gravitational
loading where the slope gradient exceeds the critical limit for the
shear strength of the material.
The coherent facies over the debris material of AMFZ slides can
be closely correlated with the intact sediments at both sides of
the slides. The irregular seafloor topography and the small‐scale
deformation inside of the coherent facies indicate that it does not
represent a post‐rift drape. The upwards inclination of layers in the
coherent facies at the sidewalls also supports this interpretation
(Figs. 8a and 9a). We therefore speculate that the slide material
consists of two different sedimentary facies which were deformed
differently during the sliding event. The weaker facies at the bottom
is deformed widely and can be recognized as slide debris on the
MCS data, while the upperlying coherent facies remains somewhat
intact with respect to the underlying slide debris facies. Exact timing
of the slides is not known, yet we tentatively suggest that they
are recent slides possibly formed during the rapid transgression
period occurred after Younger Dryas in the Black Sea (see also
Section 5.4.5).
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D. Dondurur et al. / Global and Planetary Change 103 (2013) 248–260
Fig. 9. Parts of MCS lines (a) perpendicular to the ASL1 slide to illustrate the sidewall erosion, and (b) parallel to the ASL2 slide to illustrate the headwall erosion after sliding. SM
represents slide material. Multibeam map shows the locations of the seismic lines. Vertical exaggeration is 6× and 8× for (a) and (b), respectively.
5.4. Triggering factors for AMFZ slides
A number of factors that control the slope stability along the margins
have been proposed, which include (1) oversteepening of the slope,
(2) excess pore pressure due to high sedimentation, (3) shallow gas
accumulation, gas seepage and gas hydrate dissociation, (4) earthquakes,
(5) sea‐level variation and wave loading, (6) seafloor erosion, (7) glacial
loading, (8) existence of faults, and (9) volcanic processes (Imbo et al.,
2003; Sultan et al., 2004a; Owen et al., 2007; Cauchon‐Voyer et al.,
2008; Mulder et al., 2009). Here, these factors are discussed to explain
possible triggering mechanism(s) for the slides in AMFZ.
5.4.1. Oversteepening
The MCS data show clear evidences of the Pontides thrust belt
along the Turkish Black Sea margin, and the extensional tectonics of
the deep basin to the north. The northern boundary of the structural
ridges roughly determines the limit of this extensional tectonics to
the north (Fig. 2). We suggest that Pontides thrust belt has a primary
effect on the narrowed shelf and oversteepening of the continental
slope. A similar case is also suggested by Dondurur and Çifçi (2007)
for the eastern Black Sea Turkish margin.
There is, however, no indication of large slides in these SSZs and
generally small‐scale sliding and block‐type small mass failures are
observed in the SSZs. When comparing to the continental slope, the
gradient in the AMFZ is considered to be low (it is approx. 5° around
the headwalls), indicating that the oversteepening is not a factor
promoting the large slides in the AMFZ. We tentatively suggest that
the oversteepened slope could principally cause the block‐type sliding
along the continental slope, where the sediments become unstable
under the effect of gravitational loading.
5.4.2. Existence of faults
Several small offset near‐vertical normal faults are observed in the
uppermost unit beneath the slides. We interpret that they are originated either due to uplifting of the ridges (Fig. 4a), or because of
the differential compaction (Fig. 8a). These faults are located very
close to the headwalls of the slides ensuring an appropriate surface
for sliding along the sidewalls. In addition, they may act as potential
conduits for the submarine fluids to migrate from deeper sediments
to shallower subsurface depths. Because of these factors, we propose
that the faulting close to the headwalls may be a secondary factor
which promotes the sliding.
5.4.3. Excess pore pressure due to high sedimentation
During the consolidation in normal hydrostatic pressure conditions,
the pore water can easily escape. If the layer is over‐pressurized, however, pore water circulation is restricted and loading from continuous
sedimentation is directly responded by pore water resulting in an
under‐consolidation state (Sultan et al., 2004a). When excess pore pressure in fine‐grained sediments exceeds the confining pressure, the
sediment will fail resulting in massive submarine slides. The high sedimentation rates together with the observed submarine fluid flow in the
D. Dondurur et al. / Global and Planetary Change 103 (2013) 248–260
continental rise likely cause excess pore pressures in relatively shallow
sediments resulting in under‐consolidated weak layers. We assume
that this is the main triggering mechanism for the large buried debris
flows in the abyssal plain.
5.4.4. Earthquakes
Many researchers consider the earthquakes as the major triggering
factor for the submarine mass failures (see Owen et al., 2007 and
references therein). The slides in the area are located very close to the
limit of extensional tectonics of the Black Sea Basin. The extension is,
however, inactive today and the active tectonic setting of the area is
compressional tectonism along the Pontides thrust belt to the
south (e.g., Robinson et al., 1996). This can produce earthquakes
with moderate magnitudes (Barka and Reilinger, 1997) and the
largest instrumentally recorded earthquake was occurred in 1968
offshore of Bartın. In addition, the most important seismic activity is
related to the NAF located at >130 km south of the AMFZ (see Fig. 1a
for location), which can produce strong earthquakes. It is proposed
that all different types of sliding in the area including the larger slides
in AMFZ can be triggered by both the moderate earthquakes of the
region as well as the effective seismic activity of NAF. In addition, the gravitational loading could be a contributing factor to the seismic triggering
for the small‐scale sliding on the SSZ in the southernmost area.
5.4.5. Shallow gas accumulation, gas seepage and gas hydrate dissociation
Submarine fluid flow such as shallow gas accumulations of either
thermogenic or biogenic origin can promote slides. In some cases,
the fluid flow can be supplied by decomposition of gas hydrates
above the BSR reflections. Gas hydrates may cement sediments and
modify the sediment shear strength preventing its normal compaction
process under continuous sediment load (Sultan et al., 2004b; Owen et
al., 2007). The volumetric methane to water ratios in aqueous solution
and hydrate phases are approx. 1 to 150, respectively. Once dissociated,
dissolved gas in the pore spaces can lead to excess pore pressures since
the amount of released methane is far beyond the solubility of the dissolved gas in the aqueous solution. Based on the small scale theoretical
and laboratory investigations, Grozic (2010) suggested that gas hydrate
dissociation can produce significant excess pore pressures in shallow
sediments providing that the dissociation is rapid or fluid flow is
restricted. Failure is then possible if a coincidence between slide scars
and the base of gas hydrate layer, e.g., the BSR reflection, exists
(Grozic, 2010). Therefore, the base of gas hydrate stability zone can be
considered as a potential cause of geohazard. Recent studies indicate
that giant submarine landslides may be triggered by the existence of
submarine fluid escape of the gas hydrates, e.g., Storegga Slide on the
Norwegian margin (Sultan et al., 2004b; Mienert et al., 2005), and the
Brunei slide on the Borneo Margin (Gee et al., 2007). In addition, if
failure is triggered by another factor such as seismic loading, gas
hydrate dissociation may play an important role in the propagation of
the failure (Grozic, 2010).
Küçük et al. (2011) mapped acoustic blanking zones and BSR distribution offshore of Amasra between 1300 and 1950 m water depths,
and inferred their observed BSR reflections to gas hydrate accumulations. The BSR reflections in the area are located between 180 and
200 mbsf, and most of them have acoustic turbidity zones and
enhanced reflections below attributable to the free gas accumulations, which indicates that gas hydrate layer is acting as a cap rock
(Figs. 8 and 9). The origin of the gas forming the gas hydrates is not
known. Based on the MCS data and the production wells nearby,
however, Küçük et al. (2011) proposed a possible thermogenic origin.
Distinct coincidence between BSR reflections with possible acoustic
turbidity zones and the slide scars exists just below the AMFZ slide
scars (Figs. 8 and 9). Furthermore, MCS data indicate gas chimneys
which coincide with the headwall zones of ASL2 and ASL3 slides
(Fig. 8b). We do not know if the gas associated with these chimneys
is directly related with gas hydrate dissociation above the BSR. If
257
this is the case, a possible mechanism for the gas hydrate dissociation
in the Black Sea would be a consequence from either (1) a pressure
decrease due to sea‐level variations, or (2) a temperature increase
within the gas hydrate stability zone.
The sea‐level variations in the Black Sea and their associated timings
are still controversial. Lericolais et al. (2009) proposed a sea‐level curve
of the Black Sea for last 15000 years before present (BP). The curve
includes a forced regression period between 14,000 and 11,000 years
BP after the Last Glacial Maximum (LGM), followed by a rapid transgression period between 8500 and 7150 years BP (Fig. 10a). The period
between forced regression and rapid transgression periods is known as
Younger Dryas (YD). During LGM, the Black Sea was an enclosed lake
with a sea‐level of −120 m, and the sea‐level increased to –40 m at
the end of LGM. During Younger Dryas, sea‐level of the Black Sea was
100 m lower than the present day sea‐level and a cool and drier climate
prevailed (Lericolais et al., 2009).
At the end of Younger Dryas period (approx. 8500 years BP), the
sea‐level of the Black Sea dropped 60 m (from −40 to − 100 m).
However, a sea‐level drop of 60 m would not move the gas hydrates
into the two-phase region, and the considered BSRs still remained
in the gas hydrate stability zone since they are located deeper waters
between 1300 and 1950 m in the area. Therefore, the pressure
decrease due to the sea‐level variations during the Holocene cannot
be an agent alone for the gas hydrate dissociation.
An alternative explanation might be an increase in the temperature
along the BSR level. The increase in the temperature can occur in two
ways; either due to the rapid sedimentation (e.g., a possible change in
the geothermal gradient), or water bottom temperature increase during
the sea‐level changes (e.g., due to the warmer Mediterranean seawater
input). Poort et al. (2005) investigated the effect of Mediterranean seawater input to the Black Sea during the Holocene and suggested a 2 to
5.5 °C increase in the water bottom temperature, which then resulted
in 1.1–4.6 Gt of methane carbon release from the gas hydrates.
The high sedimentation rate in the continental rise can also
change the temperature conditions around the existing base of gas
hydrate stability zone (BGHZ), which leads to a dissociation of the
gas hydrates at the BGHZ. We suggest a possible evolution model
and corresponding schematic pressure-temperature (P-T) curves of
the gas hydrate stability for the slides in AMFZ. The model is based
on the temperature changes in the BGHZ and its possible effects on
gas hydrate dissociations immediately below the slide headwalls.
According to this model, a relatively stable gas hydrate zone existed
at approx. 180 to 200 mbsf (Fig. 10b) until the end of Younger Dryas
where the sea‐level of the Black Sea was −100 m. After connection
of the Black Sea with Marmara via Bosphorus sill, the sea‐level started
to increase drastically during the rapid regression period due to the
warmer Mediterranean seawater invasion. Together with the rapid
sedimentation, this resulted in a temperature increase in the sub‐bottom
sediments as well as along the BSR. As a consequence, the BSR level
moved slightly upwards, which resulted in the dissociation of gas
hydrates at the BGHZ (Fig. 10c). Gas hydrate stability zone (GHSZ)
moved upwards and thermobaric conditions of the gas hydrates
relocated at shallower subsurface depths by ΔPg due to the increase in
temperature distribution. The effect of temperature increase due to the
warmer Mediterranean water is omitted on the temperature curve in
Fig. 10c. Dissolved gas from gas hydrates then migrated towards the seafloor producing gas chimneys and gave rise to excess pore pressures in
the uppermost sediments. As a result, top of the gas hydrate zone became
unstable and a massive slide occurred on the seafloor (Fig. 10d). It is also
likely that the sliding itself can further modify the gas hydrate stability
since considerable amount of material is removed during sliding event,
which results in additional methane release into the water column and
then into the atmosphere.
Anyway, the connection between the change in the depth of GHSZ
and its effects on the slope stability in the continental rise of the western
Black Sea margin needs further investigation and gas hydrate stability
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D. Dondurur et al. / Global and Planetary Change 103 (2013) 248–260
Fig. 10. Schematic illustration of the proposed slide formation in AMFZ (left) and corresponding schematic pressure–temperature (P–T) curves for gas hydrate stability (right).
(a) Sea‐level curve of the Black Sea for last 15,000 years before present (BP) showing a forced regression (FR) period after the Last Glacial Maximum (LGM) to the end of Younger
Dryas (YD) and a rapid transgression (RT) period (dashed line between 8500 and 7150 years BP). The sea‐level curve is adapted from Lericolais et al. (2009). (b) At the end of LGM,
a stable gas hydrate zone exists in the Quaternary sediments. (c) During the rapid transgression phase, warmer Mediterranean seawater input and rapid sedimentation result in an
increase in the temperature at the base of gas hydrate stability (BGHZ) and cause dissociation of the gas hydrates. Gas hydrate stability zone (GHSZ) moves upwards and
thermobaric conditions relocate at shallower subsurface depths by an amount of ΔPg due to the change in temperature distribution. Dissolved gas migrates upwards producing
gas chimneys and results in an unstable over‐pressured zone in the uppermost sediments. (d) A massive slide on the seafloor forms in the over‐pressured shallow sediments.
Not to scale.
D. Dondurur et al. / Global and Planetary Change 103 (2013) 248–260
modeling. For instance, the total amount of methane released during
the rapid transgression phase in the gas hydrate zone is not known. In
addition, we did not take into account gas composition forming the
gas hydrates (which also affects the stability conditions of the gas
hydrates) in this model. The exact timing of the sliding in the AMFZ is
also not known. In order to establish an exact connection in timing
between the slides and the sea level variations for the area, dating of
the slides using ground‐truthing data is necessary. According to the
model suggested here, however, the sliding should have occurred
around the end of rapid transgression period, (e.g., 7150 years BP).
Based on the sedimentation rates given by Duman (1994), the total
sediment thickness after rapid transgression period approx. ranges from
7 to 14 m, which conforms to the thickness of the post-slide sediments
(7.5 to 12 m) in AMFZ.
As a result, we tentatively suggest that a number of possible triggering mechanisms for the slides in the AMFZ might be involved.
The seismic activity itself may be an important triggering factor for
all kind of slides in the area. In addition, the existence of submarine
fluids, as well as gas hydrate dissociations interconnected with the
Mediterranean seawater input and rapid sedimentation, would be
another important factor triggering the slides in the AMFZ. Existence
of small‐scale normal faults around the scarps can supply suitable
pathways for the fluid movement and provide suitable weak surfaces
for the sliding.
6. Conclusions
According to the MCS data, western Black Sea margin is considered
as an unstable area by means of the mass failures. Several different
types of mass wasting and sediment erosion are observed along the
continental slope and rise. Particularly, an unstable area with several
slides and buried debris lobes located offshore of Amasra is named
“Amasra mass failure zone”. Buried large debris lobes exist along the
northern toe of the slope possibly triggered by excess pore pressures
due to continuous sediment loading and submarine fluid flow.
Three different types of sliding with varying sizes and formation
mechanisms are differentiated: (1) the block‐type sliding along the
rotational faults in the steep slope zones, (2) smaller‐scale slides on
the canyon walls in the continental rise to the north, and (3) relatively
larger slides in the Amasra mass failure zone. We propose that the
sliding in the first two groups above occurs due to the gravitational
loading on the steep slope areas possibly triggered by seismic activity.
In addition to the earthquake activities, another possible triggering
mechanism for the slides in Amasra mass failure zone is the excess
pore pressures due to the existence of submarine fluids possibly associated with gas hydrate dissociation immediately below the slide
headwalls due to the change in the temperature field in the BGHZ.
Small‐scale normal faults around the scarps can contribute to the failure
ensuring the suitable pathways for the fluid flow and providing the suitable weak surfaces for the sliding.
In recent years, the study area has become a potential region for the
petroleum plays in relatively deeper waters. The geohazards and mass
wasting structures described here should be carefully investigated
before drilling operations along the margin. In addition, detailed stability
analysis could be performed by extensive coring and geotechnical analysis on the continental slope and rise, which can also provide information on the potential sliding zones.
Acknowledgments
We would like to thank the officers, crew and scientific members
of the geophysics laboratory aboard the K. Piri Reis research vessel
for their valuable effort during the cruise. We also would like to
thank Hydroscience Technologies Inc. for their valuable technical
support during the data acquisition. The MCS data were processed
by Vista software from Gedco and analyzed using The Kingdom
259
Suite Software from Seismic Micro Tech. We also express our gratitude
to Dr. Christopher Sorlien from University of California at Santa
Barbara for his valuable comments and suggestions. This research
was supported by a grant from The Scientific and Technical Research
Council of Turkey (TUBITAK, project code 108Y110) as a bi‐lateral
joint project with Source‐Sink project funded by European Science
Foundation. The acoustic systems and data processing facilities of
the seismic laboratory funded with the financial support of the Turkish
State Planning Organization (DPT, project code 2003K120360) were
used for this work.
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