The nature of the Ionian Sea crust has been the subject
of scientific debate for more than 30 years, mainly because seismic imaging
of the deep crust and upper mantle of the Ionian Abyssal Plain (IAP) has not
been conclusive to date. The IAP is sandwiched between the Calabrian and
Hellenic subduction zones in the central Mediterranean. A NNE–SSW-oriented
131 km long seismic refraction and wide-angle reflection profile, consisting
of eight ocean bottom seismometers and hydrophones, was acquired in 2014.
The profile was designed to univocally confirm the proposed oceanic nature
of the IAP crust as a remnant of the Tethys and to confute its
interpretation as a strongly thinned part of the African continental crust.
A P-wave velocity model developed from travel-time forward modelling is
refined by gravimetric data and synthetic modelling of the seismic data. A
roughly 6–7 km thick crust with velocities ranging from 5.1 to 7.2 km s-1, top to bottom, can be traced throughout the IAP. In the vicinity of
the Medina seamounts at the southern IAP boundary, the crust thickens to
about 9 km and seismic velocities decrease to 6.8 km s-1 at the crust–mantle
boundary. The seismic velocity distribution and depth of the crust–mantle
boundary in the IAP document its oceanic nature and support the
interpretation of the IAP as a remnant of the Tethys lithosphere with the
Malta Escarpment as a transform margin and a Tethys opening in the NNW–SSE
direction.
Introduction
Starting back in the Cretaceous the ongoing convergence between the African
and Eurasian plates results in a complex tectonic setting in the central
Mediterranean Sea. This complexity includes segmented and strongly curved
subduction zones, shortening of overriding plates, subduction rollback, the
formation of extensional back-arc basins, and the existence of several
proposed continental or oceanic microplates (e.g. Dewey et al., 1989;
Faccenna et al., 2001, 2014; Stampfli et al., 2002; Polonia
et al., 2016; Barreca et al., 2016, and references therein). However,
seismicity and geodesy measurements show that strain is concentrated in
narrow bands bounding the microplates that are moving independently from the
overall convergent motion (Faccenna et al., 2014). The lithosphere of the
Ionian Sea constitutes one of these microplates, for which the long-lasting
contradictions in the interpretations of the nature of the Ionian Sea crust
(“oceanic” vs. “thinned continental”) result from a lack of conclusive
imaging of the deep crust and upper mantle, preventing a clear
characterization. Of special difficulty is the Messinian evaporite unit,
which massively impedes seismic energy penetration and hence limits the data
quality of geophysical imaging methods. The challenges in imaging have
ignited a long-standing debate about the deeper structure and the nature of
the crust and lithosphere of the Ionian Abyssal Plain (IAP). The
interpretations range from continental or hyperextended continental
lithosphere (Finetti and Morelli, 1973; Cloething et al., 1980; Baldi et
al., 1982; Makris et al., 1986; Ferrucci et al., 1991; Cernobori et al.,
1996; Mantovani et al., 2002; Hieke et al., 2003; Roure et al., 2012) to
oceanic or atypical oceanic lithosphere (Finetti, 1981, 1982, 2003; Finetti et al., 1996; Makris et al.,
1986; Leister et al., 1986; de Voogd et al., 1992;
Faccenna et al., 2001, 2004; Catalano et al., 2001; Gallais
et al., 2011, 2012; Speranza et al., 2012; Dellong et al., 2018). A new
feature was added to the discussion by Polonia et al. (2017) interpreting
diapiric structures to be constructed of serpentine, similar to findings in
the back-arc basin of the Tyrrhenian Sea (Prada et al., 2016). If this is
true, serpentine should be present in the subducting Ionian plate.
Finetti (1982) used geological, geophysical, and drilling exploration data to
infer that the crust in the IAP is of oceanic type adjacent to continental
crust. The IAP was interpreted to consist of oceanic type crust by a
two-ship deep refraction and reflection seismic experiment (de Voogd et al.,
1992). Oceanic crust east of the Malta Escarpment is inferred from the deep
reflection seismic CROP lines M23A and M3 (Finetti, 2003). Furthermore, time
and pre-stack depth-migrated reflection seismic data (Gallais et al., 2011,
2012) and magnetic anomaly data (Speranza et al., 2012) indicated oceanic
crust within the IAP. Contrasting interpretations are based on echo sounding
and seismic reflection data (Hieke et al., 2003) and a paleo-geographic
analysis of faults (Roure et al., 2012), interpreting the crust of the IAP
to be of continental type. While the majority of studies today infer an
oceanic type of crust in the IAP (LeBreton et al., 2017; Dellong et al.,
2018), robust information on the crustal structure of the IAP, including Moho
depth and seismic velocities, to confirm the oceanic nature of the crust are
still sparse.
Understanding the geodynamic evolution and intricate interplay between
continental and oceanic fragments hence requires exact knowledge of the
crustal and lithospheric structures. Given its location in the central
Mediterranean Sea, the Ionian Sea is a key element in reconstructing the
kinematic evolution of the central–eastern Mediterranean (Finetti, 1982). To
fill this gap, RV Meteor cruise M111 in 2014 targeted the crustal and
lithospheric structure of the Ionian Abyssal Plain (Fig. 1). Along line
DY-05, modern seismic refraction and wide-angle reflection data were
acquired using four ocean bottom seismometers (OBSs) and four ocean bottom
hydrophones (OBHs). The aim of this work is to provide information on the
seismic velocity distribution and crustal structure to assess the debate on
the nature of the lithosphere in the IAP.
Geological setting
The Ionian Sea and its abyssal plain are sandwiched between the Calabrian
(CSZ) and Hellenic (HSZ) subduction zones in the central–eastern part of the
Mediterranean basin (Fig. 1). The IAP comprises an asymmetric basin of
approximately 600 km length and 330 km width with water depth reaching
∼4100 m in its central part. Its lithosphere is actively
subducting underneath Eurasia along both subduction zones, causing a high
potential for devastating earthquakes and tsunamis. Both subduction zones
are characterized by large accretionary prisms (Calabrian Arc to the north
and western Mediterranean Ridge (WMR) to the east), which have advanced into
the Ionian Sea and cover vast parts of the IAP. Imaging of the Ionian
lithosphere is difficult because of the voluminous sedimentary cover of the
accretionary prism (Dellong et al., 2018), which furthermore is underlain by
a thick sequence of Messinian evaporites (Ryan et al., 1982). Only a
relatively small “window” of approximately 100 km length and 60 km width
remains undisturbed by the highly deformed accreted sequences advancing from
the north and east or by the Medina seamount cluster of volcanic origin
found to the south (Finetti, 1982). To the south and west, the Ionian Sea is
bounded by the continental platforms of Libya and the Malta–Hyblean plateau,
respectively. The transition from the deep-ocean environment of the IAP to
the shallow-water carbonate platform of the continental Malta–Hyblean
plateau is marked by the Malta Escarpment (Micallef et al., 2016). This
distinct, 290 km long and 3.2 km high scarp is described as an inherited
transform margin from the early Mesozoic (Argnani and Bonazzi, 2005;
Micallef et al., 2016; Dellong et al., 2018) that traces the transition from
the oceanic domain of the Ionian Sea to the Tertiary–Quaternary continental
foreland domain of the Pelagian platform (Barreca et al., 2016).
Map of the study area showing the location of the OBS profile
DY-05 (red line) and the OBS locations (yellow dots). White lines present
MCS data (Gallais et al., 2011, 2012; Polonia et al., 2011; Gutscher et al.,
2017). The green star marks the location of the ESP5 (de Voogd et al., 1992)
and the location of DSDP site 347 (Cita et al., 1978; Hsü et al., 1978)
next to it. The ESP shot line is indicated by the green dashed line. Black
lines indicate thrust faults (Gallais et al., 2011). The black-and-white map
shows the regional structures around the study area: AE – Apulian
Escarpment, CA – Calabrian Arc, CSZ – Calabrian
subduction zone, HSZ – Hellenic subduction zone, IAP – Ionian Abyssal Plain, ME – Malta Escarpment, MHP – Malta–Hyblean
platform, MS – Medina seamounts, LP –
Libyan platform, WMR – western Mediterranean Ridge.
Direct sampling of the sediment cover is available from ODP–DSDP sites (Ryan
et al., 1973), while sub-basement structure was mainly inferred from
potential field and refraction data. Early studies (Locardi and Nicolich,
1988; Nicolich, 1989; de Voogd et al., 1992; Scarascia et al., 1994)
considered the crust to be 15–20 km thick. Later investigations refined this
value to 8–10 km (Catalano et al., 2001; Gallais et al., 2011; Dellong et
al., 2018). Important insight has come from heat flow measurements, which
revealed that very low values prevail in the eastern Mediterranean compared
to the western Mediterranean (Jiménez-Munt and Negredo, 2003), indicating
distinct differences in the age and thickness of the lithosphere in these
realms (younger, warmer lithosphere in the west; older, colder, and thicker
lithosphere in the east). Low heat flux values of 30–40 mW m-2 (Della Vedova
and Pellis, 1989) from the Ionian basin underpin its oceanic origin of at
least Mesozoic age, coherent with the Pangea breakup and rifting of the
Neo-Tethys Ocean during the late Paleozoic or early Mesozoic (Ricou, 1994;
Stampfli et al., 2002; Frizon de Lamotte et al., 2015). In addition, the
relatively high Bouguer gravity anomaly of the IAP with values exceeding
200 mgal (Morelli et al., 1975) suggests a shallow Moho boundary (Dellong et
al., 2018). The sea-floor magnetic pattern of the IAP indicates that the
Ionian Sea lithosphere was formed around 220–230 Ma during Triassic times,
and it is considered a remnant of the Tethys Ocean (Speranza et al.,
2012).
The IAP lithosphere is subducting beneath the Calabrian Arc, which is
proposed to have formed in the late Miocene during the opening of the
Tyrrhenian Sea (Faccenna et al., 2001; Chiarabba et al., 2008). Clear
Wadati–Benioff zones are imaged below the CSZ in the west (Selvaggi and
Chiarabba, 1995) and below the HSZ in the north (Hatzfeld et al., 1993), and recent
studies interpreted the slabs to be composed of subducted oceanic crust
(Chiarabba et al., 2008; Pearce et al., 2012). The Ionian basin is bounded
by the Apulian Escarpment (AE) in the north and the Malta Escarpment (ME) in
the southwest. The escarpments are interpreted to be the conjugated passive
margins separated during the opening of the Tethys (Catalano et al., 2001;
Chamot-Rooke et al., 2005). In contrast, Gallais et al. (2011) and Frizon de
Lamotte et al. (2011) interpreted the ME differently and see it as part of a
continental transform margin separating the continental domain west of the
ME from the deep Ionian basin. The oceanic nature of the Ionian lithosphere
supports the hypothesis that the Adriatic microplate, which comprises the
Ionian Sea in its southernmost portion, was a rigid promontory of Africa
(e.g. Channell et al., 1979; Dewey et al., 1989). However, according to
Gallais et al. (2011, 2012), Polonia et al. (2011), and Roure et al. (2012)
seismic data show thrust faulting and inversion structures in the IAP (Fig. 1), which are indicative
of active deformation and would contradict a
totally “rigid” connection to Africa. Roure et al. (2012) favour the IAP to
be a thinned part of the African Plate. The occurrence of active
deformations in the IAP (Gallais et al., 2011, 2012; Polonia et al., 2011)
would support the idea of an independent Adria microplate (i.e. Ustaszewski
et al., 2008; Handy et al., 2010, and references therein) allowing it to move
relative to Africa, at least in Neogene times (LeBreton et al., 2017).
During Neogene times Ustaszewski et al. (2008) determined a ∼20∘ counterclockwise rotation of Adria relative to Eurasia, while
only ∼2∘ of rotation has been accommodated between
Europe and Africa.
Stacked and bandpass-filtered (f=20, 30, 80, 120 Hz)
multi-channel data (MCS) along profile DY-05. Yellow triangles mark the OBS
locations along the MCS profile. Green bar marks the projected location of
DSDP site 374 on the profile DY-05.
Data and methodology
During RV Meteor cruise M111 wide-angle refraction seismic data were
acquired simultaneously with multi-channel seismic reflection seismic (MCS) data.
Profile DY-05 is 131 km long and crosses the IAP from SW to NE (red line in
Fig. 1). The location of the line was chosen to cover the area of the IAP
that is neither affected by the emplacement of volcanic structures (Medina
seamounts in the south) nor by the advancing thick accretionary
prisms of deforming sediment in the north and east in order to gain an
“unobstructed view” into the deep IAP crust and lithosphere. The central
part of the profile was covered by eight sea-floor stations (OBH501–OBS508)
with a spacing of 7.4 km. Shooting (946 shots) was extended for
∼40 km beyond the first and last instrument, respectively, in
order to record long offsets from shots travelling through the subsurface.
An airgun array consisting of six G-gun clusters with a total volume of 84 L
(5440 cu in) at 210 bar was fired at a 60 s shot interval. Data quality was
very good and arrivals were recorded over the entire profile length. Mantle
phases, PmP reflections, and Pn phases were recorded on all stations. A
mini-streamer with four channels spread over an active length of 65 m was towed
at a depth of 8 m between the airguns. For each channel, six hydrophones with
a distance of 0.5 m were grouped together. The streamer served two purposes:
it was used to control the correct functioning of the airguns and to
provide information on the uppermost sedimentary structures in regions where
sea-floor roughness does not cause aliasing. A medium gun delay of 78 ms
could be identified, with one gun firing out of sequence, as recognized on
the streamer data near trace 800 in Fig. 2.
Description of the multi-channel seismic reflection data
The stacked MCS data section displays the upper subsurface structure in
two-way travel time (twt; Fig. 2). The water depth along the profile
increases from ∼3570 m (∼4.7 s twt) at the
southern termination in the vicinity of the Medina seamounts to
∼4060 m (∼5.3 s twt) in the central portion of
the profile. The water depth declines again towards the northern end of the
line, which starts to cover the Ionian accretionary prism (compare Fig. 1).
North of the rough sea floor of the Medina seamounts, the data section is
dominated by the thick sedimentary and evaporitic sequences which fill the
basin. Stratified sediment layers are softly inclined along the southern
part of the profile between shot numbers 150 and 450. Towards the centre of the
basin, horizontally layered, sub-parallel sequences are onlapping these
strata between shots 450 and 550. These units are underlain by a thick sequence
of incoherent amplitudes truncating the sediment layers in the south. A
prominent strong-amplitude reflector may be traced from shot point 150 at
5.2 s to around shot number 550 at 5.7 s. Whereas a coherent layering is
visible in the units above this reflector, the units below are seismically
much more opaque. The base of these units is marked by a strong negative
reflector at 5.5 (near trace 200) to 6.3 s (near trace 800). Below this
reflector, the seismic signal is strongly attenuated and no coherent
structures may be identified. Three deep reflectors to the north at 6.6 s,
6.8 s, and at 7.15 s lose their seismic amplitude coherence beyond shot
numbers 600 and 750, respectively.
Seismic record sections of OBH501 and OBS508. Panels (a) and (b) show
processed data. Panels (c) and (d) show the travel-time picks used for the
forward modelling (please refer to phase definitions in the text).
Calculated travel times from the final model are displayed in (e) and (f). Data are plotted with a reduced time of 8 km s-1.
Description of the seismic wide-angle reflection and refraction
data
Figure 3 presents seismic shot sections, including picked and calculated
travel-time picks, of the two outermost stations, OBH501 and OBS508. In the
following we want to describe the recorded data using pick phases, pick
colours, and layer names, which can be found in Table 1. The earliest
arrival at the stations is the direct wave through the water picked in
yellow (PDirect, phase 1). Magenta picks (PsPTopME, phase 3) represent the
base of the youngest sediment units. The geometry of these sediment layers
is well imaged in the multi-channel seismic data (Fig. 2) by a strong
positive reflector (PsPTopME). Red picks (PsME, phase 2) follow the first
arrivals to an offset of ∼15 km (3.5 to 4.5 s), representing
refracted waves travelling through the evaporite layers with an apparent
velocity of 4.5 km s-1. At about 15 km offset from the station, at the limit
of the red picks (phase 2), a prominent shadow zone indicates a strong
negative velocity inversion in the subsurface including a thick low-velocity
zone. The near-offset early arrivals are offset by 1.5 s to the deeper
crustal first-arrival phases (light green picks, Pg, phase 10) that have
been recorded starting at roughly 6 s. A strong negative reflector starting
at zero offset is present on all OBS record sections and is picked in dark
blue (PsPBotME, phase 5). It marks the base of the Messinian unit at
∼3.6 s at zero offset. After a ∼1.4 s long
sequence of high-amplitude reflectors follows a low-frequency high-amplitude
reflector (dark green, PsPTopAU, phase 7) visible between 0 and 20 km offset
at ∼5 s. Underneath, a second similar reflector is picked in
dark magenta (PbP, phase 9) at ∼5.6 s. These two reflectors
originate from the base of the slow sediment unit and the basement,
respectively. At larger offsets, two crustal phases (violet (Pb, phase 8)
and light green (Pg, phase 10) picks) could be identified. At about 35–40 km
in offset the mantle reflection is observed and picked in light blue (PmP,
phase 11), with mantle-refracted phases in bright red (Pn, phase 12). Except
for OBS508, where the PmP phase was observed on the northern branch only,
all instruments show wide-angle mantle reflections (phase 11) at both
branches.
(a) Final velocity model for profile DY-05 obtained from forward
modelling. (b) Velocity–depth functions (blue, green, and orange) starting
at the basement and extracted from the final velocity model every 10 km
according to the coloured sections. Observed velocities are compared to
oceanic crust (red grey ensemble) (Grevemeyer et al., 2018), continental
crust (blue ensemble) (Christensen and Mooney, 1995), and serpentinized
mantle (dark grey ensemble) found in the Tyrrhenian Sea (Prada et al.,
2016). MS – Medina seamounts.
Methodology and modelling strategy
By developing a velocity model for profile DY-05 we targeted the structure
of the sedimentary strata, the crust, and the uppermost mantle. This was
achieved by forward modelling of the observed travel times using the
ray-tracing software RAYINVR (Zelt and Smith, 1992; Zelt and Forsyth, 1994).
It runs on the graphical user interface MODELING (Fujie et al., 2008).
Travel times were picked using the software PASTEUP (Fujie et al., 2008).
Attempts to apply tomographic inversion approaches failed due to the
presence of a thick low-velocity zone. The final forward model (Fig. 4) was
developed starting with the water layer and progressing downward layer by
layer. The geometry of the upper layers was constrained by the MCS data
(Fig. 2). The geometry of the deeper layer boundaries was optimized by using
the inversion algorithm of RAYINVR. The approach of combined forward
modelling and layer-wise inversion was chosen to find the simplest model to
fit all data without over- or under-fitting the data. The velocity and
interface nodes have been spaced irregularly. The velocities are tied to layer
interfaces at the top and bottom of each layer. Between the nodes,
velocity and interface information is interpolated (Zelt and Smith,
1992; Zelt, 1999). The root mean square (RMS) misfit and the
Chi2 value for each phase are provided in Table 1. Pick
uncertainties were assigned to avoid over- or under-fitting the data,
especially in the inversion modelling. An overall normalized
Chi2=1 in travel-time misfit should be achieved for the
final model. We receive velocity uncertainties of up to 0.05 km s-1 for the
shallow sedimentary units, while the velocity uncertainties for the low-velocity sedimentary layer are higher at
∼0.2 km s-1. At the crustal
level and deeper, uncertainties reach 0.1 km s-1. Due to the high number of
observed PmP arrivals and the dense shot spacing, the uncertainty for the
depth of the Moho is ∼1 km.
(a) The density distribution calculated for
profile DY-05 based on seismic velocities and minor adjustments to fit the
observed FAA in panels (b, c). (b) Case study for the local influence of
the Medina seamounts on the model response. (c) Case study for oceanic crust
over the entire profile length. Black dashed line shows the residuum of the
observed and calculated data.
To corroborate the seismic velocity model obtained by seismic travel-time
forward modelling, complementary two-dimensional gravity forward modelling
was performed. Especially at the profile ends with only moderate station
coverage, the seismic velocity model is highly uncertain and gravity
modelling (Fig. 5) helps to constrain and adjust the model. The free-air
anomaly (FAA) along the 2-D profile was extracted from satellite altimetry
data (Sandwell and Smith, 1997). Based on the seismic velocities an initial
2-D density model was constructed. The model response (Talwani et al., 1959)
was compared to the observed free-air anomaly data along the profile. Minor
adjustments to the model densities, all within common density–velocity
relationships (Carlson and Herrick, 1990; Christensen and Mooney, 1995),
were applied to achieve a reasonable fit between calculated and observed
FAA.
To validate the forward modelling, synthetic data based on the final seismic
velocity model were calculated and compared to the recorded data (Fig. 6). A
finite-difference scheme for the solution of the elastic isotropic wave
equation from Hestholm et al. (1994) was used to calculate the synthetic
data for a 2-D structure. The input velocity model had a grid cell size of
50 m. The seismogram was computed with a P-wave source frequency up to 30 Hz, while a 2 ms time step was chosen. No random noise was added.
Observed seismic phases with their pick number, number of
picks, pick uncertainty, RMS fit, and Chi2 to allow for an
estimation of the robustness of the seismic velocity model obtained from
forward modelling.
(a) Synthetic seismogram of OBH501 based on the final velocity
model. (b) Original bandpass-filtered seismogram of OBH501.
Interpretation and resultsShallow structures from multi-channel seismic data
The uppermost unit in the MCS data was drilled at DSDP site 374 to a depth
of 457 m (Hsü et al., 1978) and corresponds to Plio–Quaternary
sediments. The lower limit of this unit is marked by the so-called
A reflector (Finetti and Morelli, 1972), which forms the transition to the
Messinian evaporite layers (Hsü et al., 1978; Gallais et al., 2011).
This reflector corresponds to phase 3 (PsPTopME, magenta picks) in the
seismic sections of the OBH (Fig. 3).
The base of the Messinian evaporites has been termed the B reflector by
Finetti and Morelli (1972). We interpret phase 5 (PsPBotME, dark blue
picks) in the OBS seismic sections (Fig. 3) to correspond to this reflector.
The B reflector disappears in our seismic section near trace 800 towards the
north due to seismic scattering. Three reflectors identified near trace 650
at 6.6, 6.8, and at 7.15 s indicate pre-Messinian sedimentary layering.
We link the upper two reflectors to Tortonian age, while the deeper
reflector at 7.15 s is interpreted to be the so-called K reflector, as
identified by Gallais et al. (2011) in the crossing seismic line Arch21
(Fig. 1).
At trace 100 of the MCS profile the layering of the upper sediments is
disturbed by the volcanic signature of the Medina seamounts (Fig. 2). In the
northern part of the profile, starting near trace 570 (OBH503), the
A reflector of the Messinian unit gradually disappears towards the northern
profile end. The upper Plio–Quaternary sediments are gently folded near
trace 700, whereas north of trace 880, all Plio–Quaternary sediments
are strongly folded up to the sea floor. The fold amplitude progressively decreases
towards the south as observed in Polonia at al. (2011) along
the MCS line CALA-15. We interpret these folds to represent the deformation
front of the Calabrian accretionary wedge, expressed at different depth
levels. The deformation of the sediments and/or differential sediment
loading possibly create fluid pathways and may thus support the dissolution
of the evaporites below.
Sedimentary structures from seismic travel-time modelling
Layer 1 in our final velocity model (Fig. 4a) is 250–500 m thick with
rapidly increasing seismic velocities from 1.8 to 2.2 km s-1, top to
bottom. From drilling it is known that L1 is composed of unconsolidated
Plio–Quaternary sediments (Hsü et al., 1978). The underlying layer L2 is characterized
by uniform seismic velocities of 4.4 to 4.65 km s-1 and a thickness of
0.8 to 1.3 km. L2 is interpreted as the Messinian evaporite unit (layer 2
in Fig. 4a), similar to the findings along seismic profiles in the
Calabrian Arc, profiles DY-01 and DY-03 (Dellong et al., 2018). The
A reflector (phase 3, magenta picks, PsPTopME) and the reversed-polarity B reflector (phase 5, dark blue picks, PsPBotME) occur as
strong phases in the OBS data at all stations, representing the top and
bottom of the Messinian evaporite unit, respectively. The B reflector at the
base of the Messinian unit is also known as the S1 horizon from the ESP5
(Expanding Spread Profile) studied by de Voogd et al. (1992).
Below the B reflector, the OBS data show a succession of high-amplitude
reflectors similar to observations made by Gallais et al. (2011) in the MCS data
(Archimede and PrisMed01 profiles). However, it is difficult to link these
reflectors to certain horizons in our velocity model; thus, we modelled the
pre-Messinian sediments as a low-velocity layer (layer 3 in Fig. 4a) with
velocities increasing from 3.1 km s-1 at the B reflector to 3.7 km s-1 at the
base of the unit. Chamot-Rooke et al. (2005), Polonia et al. (2011), and
Gallais et al. (2011) propose that the upper part of the unit, containing the succession
of reflectors, is of Tortonian age, overlying an undifferentiated Tertiary
sequence and Mesozoic sediments. From 20 km towards south, we observe higher
velocities up to 3.7–4.7 km s-1 (top to bottom) that we relate to the
Medina seamounts. This part of the profile, however, was not covered by OBS
stations; thus, the model here is less reliable.
A high-amplitude, low-frequency reflector (phase 7, dark green picks,
PsPTopAU in Fig. 3) marks the top of layer 4, with shows a low-velocity gradient, with seismic velocities of 4.8–4.9 km s-1 (Fig. 4a). We
interpret L4 as a sedimentary unit based on the seismic velocities and the
velocity gradient within the layer. This layer is called the ambiguous unit
in Gallais et al. (2011). Thus, we observe the crystalline basement at a
depth of ∼9.5 km at OBH501 to ∼8.5 km at
OBS508.
Crustal structures from seismic travel-time modelling
Indeed, layer 5 (Fig. 4a) shows a steeper velocity gradient to a depth of
∼11 km. The seismic velocities increase from 5.1 to 6.4 km s-1, typical for upper oceanic crust
(seismic layer 2) (White et al., 1992;
Grevemeyer et al., 2018). The upper crust is approximately 2–3 km thick.
Layer 6 (Fig. 4a) again shows a lower-velocity gradient with velocities
increasing from 6.4 to 7.2 km s-1 from top to bottom, typical for lower
oceanic crust (seismic layer 3) (White et al., 1992; Grevemeyer et al., 2018). Layer 7
(Fig. 4a), with velocities higher than 7.8 km s-1, is interpreted as mantle with
a seismic Moho at ∼15 km of depth at OBH501 and at
∼17 km of depth at OBS508.
In Fig. 4b we compare seismic velocity–depth profiles along profile DY-05
with a velocity–depth ensemble for magmatic crust (Grevemeyer et al., 2018),
as well as with a velocity–depth ensemble for serpentinized mantle found in
the Tyrrhenian Sea (Prada et al., 2016) and a velocity–depth profile
typical for extended continental metamorphic crust. In the northern portion
of the profile (orange section) and at shallow depth (∼2 km
into the basement) the profile fits both ensembles for magmatic crust and
serpentinized mantle, while at greater depth in the northern portion of the
profile (orange section), lower-crust velocities typical for gabbro are
observed. Between 40 and 130 km (green and orange sections), the observed
seismic velocity field and the velocity gradients are typical for magmatic
crust.
In the northern part of the profile in the IAP, 6–7 km thick crust is
encountered. Towards the southern end, the crust thickens to at least
∼9 km. Due to the moderate resolution at the model
termination in the south, especially at greater depth, it is difficult to
identify the nature of the crust in the southern part. In conjunction with
crustal thickening we observe crustal velocities at the crust–mantle
boundary of 6.8 km s-1, which plot closer to the velocity–depth ensemble of
continental metamorphic crust (compare to Fig. 4b, blue section). The
recorded seismic data show a change in the characteristics of the mantle
phases towards the south. This could result from a change in the nature of
the crust or simply originate from the influence of the Medina seamounts
with their proposed volcanic composition (Finetti, 1982).
Again, at the edges and with increasing depth, modelling becomes less
accurate. Pick uncertainties and data fits are presented in Table 1. Table 2
lists layers L1 to L7 (used in Fig. 4a) with their seismic velocities.
Gravity modelling
The results of the 2-D gravity forward modelling are shown in Fig. 5. In
the south, possible 3-D effects of the Medina seamounts have not been taken
into account during the modelling. However, the seismic velocities of the
crust and the density of the crust decrease towards the south, roughly south
of kilometre 30 of the profile. At the corresponding location in the model,
a thickening of the crust is observed. A good fit is observed with an RMS
deviation of 1.46 mgal. Figure 5b presents a case without the shallow
density anomaly between 5 and 20 km along the profile. The overall trend
of an increasing FAA from south to north can be still observed; however, the
short-wavelength fit is inferior compared to the final model (Fig. 5a): the
RMS misfit is 2.62 mgal. This part of the model is not covered by OBS,
resulting in less well-constrained seismic velocities; however, shallow
denser material supports the observed higher seismic velocities. This is
supported as well by the MCS data in Fig. 2, for which the sedimentary
succession is deformed and reflectors are disturbed. To further test the
velocity model, a constant crustal velocity was assumed. In contrast to the
final model, densities in the crust were defined with constant values typical
of oceanic crust over the entire profile (Fig. 5c). The model response in
Fig. 5c shows a strong RMS misfit of 13.92 mgal between observed and
calculated FAA south of kilometre 60. In this scenario, the crustal densities are
too high in the southern part of the model. Additionally, assuming a
constant crustal thickness would even enlarge the misfit of the data. By
means of the gravity study, we can confirm the final seismic velocity model,
even for the portions not covered by OBS.
Synthetic data
The computation of synthetic data serves to test the forward modelling
results. Figure 6 compares the synthetic seismogram based on the final
velocity model (Fig. 4) with the recorded seismogram of OBH501. The gross
features of the observed amplitudes (Fig. 6b) could be reproduced in the
synthetic record section (Fig. 6a). The near-offset waveforms (between 5 and
15 km offset and between 3.5 and 4 s) of the evaporite unit show high
amplitudes that rapidly decrease with offset. The large shadow zone,
starting at 15 km offset and at ∼4 s to both sides, is caused
by a thick unit of sediment layers with slower seismic velocities compared
to the evaporite unit. This portion of the model was kept simple during
forward modelling; thus, we observe a lack in phases between 4 and 5 s in
the synthetic data compared to the observed data in Fig. 6b. The
additional phases in the observed data indicate internal layering of the
evaporite unit as well as layering within the slow sedimentary unit below,
which were not resolved during travel-time modelling. At roughly 5 s the
reflected phase of the fast sedimentary unit above the basement and
the reflection of the basement at 5.5 s are present in the synthetic data
with weak amplitudes (Fig. 6a). The crustal phases between 15 and 30 km
offset become stronger again, while the PmP, between 40 and 50 km offset
at about 7 s, shows high amplitudes. This feature is observed in the
recorded data as well and is associated with a discontinuity at the
crust–mantle boundary. The mantle phases with offsets larger than 55 km
slowly fade. They can be better recognized in the synthetic data in which they
are not obscured by ambient noise. The apparent velocities of the main
features fit the observed data.
Overview of the layers in Figs. 4 and 5 including interpretation
of the unit and seismic velocities for each layer at OBH501.
LayerLayer type interpretedSeismic velocities(top–bottom)at OBS501L1Quaternary and Pliocene1.8–2.2 km s-1sedimentsL2Messinian evaporites4.4–4.65 km s-1L3Pre-Messinian sediments3.1–3.7 km s-1L4Carbonate sediments4.8–4.9 km s-1L5Upper crust5.1–6.4 km s-1L6Lower crust6.4–7.2 km s-1L7Mantle> 7.8 km s-1
Comparison of the data and the final velocity model from this
study (OBH501, OBH503, OBS504) with previous studies: (a) CROP M2 (Polonia
et al., 2011), (b) ESP5 (de Voogd et al. 1992), and (c) PrisMed01 (Gallais
et al., 2011). L1–L7 follow the same layer convention as in the final velocity model.
Abbreviations in panel (b): W – water, P+Q – Pliocene and
Quaternary sediments, MS – Messinian sediments,
Pre-M – pre-Messinian sediments, UC – upper crust,
LC – lower crust.
DiscussionDeformation and tectonic thickening at the front of the Calabrian
accretionary prism
Folds in the sedimentary units have been observed at different depth levels
at the sea floor within the Plio–Quaternary unit and in the Messinian
evaporites (Fig. 2); however, we do not interpret this as a result of
deformation stepping back in time towards the north. We rather interpret the
scenario as deformation occurring at all three depth levels, most likely
simultaneously, caused by the distributed thrusting at the toe of the
Calabrian accretionary wedge. Along the MCS line CALA-15, which crosses our
profile in the northern IAP (Fig. 1), frontal thrusting affecting the
sea floor was observed, as was a zone of “proto-thrusts” that deform
deeper layers, but without reaching the sea floor yet (Polonia
et al., 2011) (see their Fig. 8). The same is true for the line PrisMed-1
(Gallais et al., 2012) (see their Fig. 5a). The distribution of the
deformation fronts is most likely a function of the rheological properties
of the different layers. In fact, we observe distributed shortening and
tectonic thickening of the Messinian salt unit, resulting in an obscured and
chaotic A reflector within a short distance towards the north. While the
Plio–Quaternary sediments at greater depth show gently undulating folds, we
observe an increase in the amplitude of these folds towards the north where
they affect the entire Plio–Quaternary sediment unit up to the sea floor. We
interpret the sudden onset of strongly undulating folds expressed at the
sea floor as the signature of the outer deformation front of the Calabrian
wedge at the sea floor, while the Plio–Quaternary sediments at the bottom of
the unit are already affected by the growth of the accretionary wedge. Along
the approximately perpendicular MCS profile CALA-15 the same pattern is
observed from SE to NW (Polonia et al., 2011). The thickness of the
Plio–Quaternary sediments is predominantly constant (Figs. 2 and 4),
supporting the idea that rheological properties influence the location of
the deformation front at each layer. The MCS profile CROP M2 (Polonia et
al., 2011) crosses the DY-05 profile at OBH501 (Fig. 7a). Polonia et al. (2011) interpreted
the base of the Messinian evaporites as the detachment
acting as the plate boundary between the post-Messinian wedge (Eurasia) and
pre-Messinian sediments on top of the very old African subducting oceanic
crust. This would indeed be consistent with our observations of a
B reflector that can be traced through almost the entire profile and the
undisturbed flat-lying Plio–Quaternary sediments on top of the Messinian
evaporites, which are affected in the north by the advancing Calabrian
wedge.
Dark blue marks the presence of oceanic lithosphere below the sea floor
in the Ionian Abyssal Plain (undeformed portion) as determined by seismic
studies. Light blue areas mark the extent of buried oceanic lithosphere
below adjacent accretionary wedges as determined by seismic studies. Profiles
1–5: M111; profile 6: Makris et al. (1986) and Finetti (2003) (part of
CROP M23A); profiles 7–8: IMERSE profiles (Fruehn et al., 2001; Westbrook
and Reston, 2002; Reston et al., 2002); profiles 9 and 10: ION6 (Cernobori
et al., 1996) and ION7 (Kokinou et al., 2003). AE – Apulian
Escarpment. Further abbreviations are explained in Fig. 1.
Pre-Messinian sedimentary layers
Beneath the B reflector, the seismic signal is strongly attenuated both in
the OBS and MCS data. In the MCS data, short reflective bands can be
observed, indicating pre-Messinian sedimentary layering. The OBS data show a
sequence of steep angle reflections with high frequencies. However, a
comparison of OBS and MCS data proved unsuitable to correlate the observed
phases at the different stations. The model in this portion is highly
unreliable, especially the velocity information with absolute velocities and
the velocity gradient.
Layer 4 was the next layer that could be verified by a high-amplitude, low-frequency reflector. Based on the absolute velocity and the low-velocity
gradient it is interpreted as a sedimentary layer. While de Voogd et al. (1992) interpreted
layer 4 (Fig. 4a) as oceanic upper-crust layer 2A, we interpret this to be a
1 km thick layer of fast sediments, possibly carbonate sediments of Mesozoic
age (Polonia et al., 2011) (Fig. 7b). For seismic layer 2 of oceanic crust,
we would expect a high-velocity gradient with absolute velocities lower than
observed (< 4.4 km s-1) (Carlson, 1998; White et al., 1992; Grevemeyer
et al., 2018) or a low-velocity gradient with higher absolute velocities
(> 5.7 km s-1) for continental crust (Christensen and Mooney,
1995). Finetti (1982), LeMeur (1997), and Gallais et al. (2011)
also proposed that L4 represents a further Mesozoic sedimentary unit,
which is supported by the imaged layered facies in their studies (compare
Fig. 7c).
Nature of the lithosphere of the Ionian Abyssal Plain
Based on the results of our seismic modelling and the gravimetric modelling,
we interpret the IAP to be of oceanic nature with a thickness of 6–7 km.
Further south in the vicinity of the Medina seamounts, we observe a change
in the characteristics of the crust towards a continental type of crust with
a thickness of 9 km. We briefly resume the discussion from the Introduction
in the context of our results, and we will concentrate on the lithosphere of
the central IAP.
Thinned continental crust vs. oceanic crust. In Fig. 4b we
compare seismic velocity–depth profiles along profile DY-05 with a
velocity–depth ensemble for magmatic crust (Grevemeyer et al., 2018), as
well as with a velocity–depth ensemble for continental metamorphic crust
(Christensen and Mooney, 1995) that includes extended continental crust. The
seismic velocities within the basement obtained from the final velocity model
(Fig. 4a) plot into the velocity–depth field typical for oceanic crust
but are too fast for continental crust (Fig. 4b).
Hyperextended continental crust vs. oceanic crust. Thinned
or extended continental crust is slower than normal continental crust
(Christensen and Mooney, 1995). In a seismic refraction study in the Gulf of
California (Lizarralde et al., 2007), seismic
velocities have been shown to decrease as the continental crust becomes thinner towards the
continent–ocean transition zone. Lower-crust velocities reach 6.4 km s-1,
which is much slower than observed in the final velocity model along DY-05
(Fig. 4a).
In conclusion, based on seismic P-wave velocities, the lithosphere of the
central IAP is interpreted to be of oceanic type. This is supported by the
gravity modelling results presented in Fig. 5. To continue the discussion
we want to investigate if the oceanic domain is constructed of magmatic
crust or serpentinized mantle.
Serpentinized mantle vs. magmatic crust. In Fig. 4b we
show the velocity–depth ensemble for serpentinized mantle found in the
Tyrrhenian Sea (Prada et al., 2016). At shallow depth (∼2 km
into the basement) the profile fits both ensembles for magmatic crust and
serpentinized mantle. Based on seismic P-wave velocities only it is not
possible to discriminate between upper crust and highly serpentinized
mantle. However, at greater depth lower-crust velocities typical for gabbro
(Carlson and Miller, 2004) are observed. The data fit the ensemble for
magmatic crust as observed by Grevemeyer et al. (2018), and we exclude
serpentinized mantle on top of gabbroic crustal rocks.
We conclude that based on seismic P-wave velocities, the lithosphere of the
central IAP comprises magmatic crust with a crustal thickness of 6–7 km
(Fig. 5). Our gravity modelling along profile DY-05 (Fig. 5) shows a good
fit assuming oceanic crust in the central and northern part of the profile
with a possible change in crustal type towards the south and towards the
extension of the Malta Escarpment.
Previous and recent studies of the area are combined in Fig. 8 to show the
distribution of oceanic lithosphere in the Ionian Sea. In the following
discussion we want to summarize previous works and compare them to our
results.
Hieke et al. (2003) presented an extensive discussion of previous work and
argued that based on magnetic anomalies, missing heat flow anomalies, and
the gravity data, the IAP is a thinned part of the African continental
crust. We agree that the system seems to be in equilibrium regarding the
heat flow; however, similar low values of 38–46 mW M-2 are
expected for very old oceanic crust as well (Sclater et al., 1980). Our
interpretation of a possible change in crustal type from the central IAP
towards the Medina seamounts and the extension of the ME is supported by new
marine satellite gravity maps of the area (Sandwell et al., 2014). We can
identify the deep-seated ME as a transition zone from continental to oceanic
type of crust, further south towards the Medina seamounts.
Similar to our findings, Makris et al. (1986) observed velocities up to 7.2 km s-1 resulting from oceanic crust. However, the authors could not exclude
stretched continental crust with intruded upper mantle. The stations within
the IAP failed and their profile could only be extended to the IAP based on
shots within the IAP recorded at stations outside the IAP. This leaves their
results with a high uncertainty in this portion of the profile. This lack of
resolution was overcome by the design of profile DY-05 that has a dense shot
(∼140 m) and station (∼7.6 km) spacing.
Thinned continental crust in the IAP was interpreted by Finetti and Morelli (1973) and Finetti (1981, 1982)
based on gravity data; however, they could
not explain the gravity data without having mantle material that intruded
into the crust. In a later work based on seismic MCS line CROP M3, Finetti (2003) interpreted the crust east of the ME to be of oceanic nature, with a
crystalline basement at 7 to 8 s twt, which is, if converted into depth,
similar to our findings. As part of the work of Polonia et al. (2011) the
data were reprocessed and the authors could confirm the occurrence of
oceanic crust east of the ME. During RV Meteor M111 cruise, wide-angle
refraction seismic data (DY-01) were acquired along CROP M3 (Dellong et al.,
2018). In this study the deeper crustal structures along the profile could
be better resolved and the authors proved that oceanic crust has been
subducted into the CA.
De Voogd et al. (1992) analysed the ESP5 data from a two-ship experiment in
the vicinity of OBS502. We find similar results for the shallow part of the
models (Fig. 7b), while the deeper portions vary regarding the depth of the
layers. The crustal thickness at ESP5 is about 8 km versus 6 km at OBS502
along profile DY-05. The lower crust has a similar thickness for both
experiments. We observe a seismic mantle velocity of 7.8 km s-1 at the
crust–mantle boundary that is increasing with depth in at least the upper 2 km below the Moho, while the studies of the ESP5 experiment (de Voogd et
al., 1992) (Fig. 7b) and Makris et al. (1986) show constant mantle
velocities of 8.1 and 8.5 km s-1.
Sioni (1996) and Gallais et al. (2011) interpreted the high-amplitude
reflector at 8 s twt (in the seismic line PrisMed01) as the top of oceanic
crust. The reflectors from OBS502 and PrisMed01 lines show an ideal fit (Fig. 7c).
Polonia et al. (2017) interpreted diapiric structures in the CA to be
constructed of serpentine. In our study as well as in the seismic study from
Dellong et al. (2018) we find no evidence for major serpentine structures
as layers in the IAP lithosphere or in the vicinity of fault
systems in the CA. This does not rule out the findings of Polonia et al. (2017) for the CA region, where a fossil fracture zone might has been
subducted, allowing serpentinized material to rise.
The oceanic IAP in the frame of plate tectonics
The geophysical characteristics, for example seismic velocities, crustal
thickness, densities (Figs. 4 and 5), and heat flow, confirm the IAP to be
oceanic lithosphere. The IAP is consequently a leftover of the Tethyan
lithosphere, and the Apulia–Adriatic domain and the eastern Mediterranean
basin do not belong to the former southern continental margin of the Tethys
as suggested by Roure et al. (2012). However, the isostatic equilibrium and
magnetic anomalies exclude a sea-floor age younger than Late Triassic or
Early Jurassic (de Voogd et al., 1992; Stampfli, 2000; Speranza et al.,
2012). The Apulian Escarpment in the N and the Malta Escarpment in the SW
are bounding the Tethys basin and are most probably transform margins
(Gallais et al., 2011; Frizon de Lamotte et al., 2011) and not conjugate
passive margins (Catalano et al., 2001). They are present as steep and high
steps in the sea-floor bathymetry, which is not expected at passive margins
where sedimentation and erosion over a period of ∼200 Myr
would overprint and smoothen the topography (Mosher et al., 2017).
Additionally, the ME and the AE as passive margins would require a plate
boundary in the NW–SE direction within the Ionian basin, which has not been observed
so far. On the other hand, interpreting the ME and the AE as transform
margins, a plate boundary would be expected in the SW–NE direction, parallel to
the observed deep reverse faults beneath the IAP (Gallais et al., 2011;
Polonia et al., 2011) (Fig. 1), and could explain our results of the observed
transition from oceanic crust towards continental crust in the vicinity of
the Medina seamounts (Fig. 4).
Oceanic lithosphere is considered to be rigid and shows little
deformation compared to continental lithosphere that is mechanically weak,
especially when thinned (Polonia et al, 2016). An IAP made of Tethyan
lithosphere, due to its rigidity, would imply that the connection between
Africa and Adria was rigid since the end of oceanic spreading in the early
Mesozoic. However, deformation within the IAP lithosphere was observed in
seismic data by Gallais et al. (2011), Polonia et al. (2011) (Fig. 1), and
in this study (Fig. 2). The deformation caused by contraction is possibly
connected to subduction processes during Neogene times and subduction might
be still active. The connection between Adria and Africa during Neogene
times is consequently not particularly rigid (Gallais et al., 2011; Le
Breton et al., 2017) as proposed earlier (e.g. Channell et al., 1979; Dewey
et al., 1989), and Adria might move relative to Africa. This supports the
hypothesis of an independent Adriatic microplate that was separated during
the opening of the Tethys as proposed in reconstructions by Handy et al. (2010) and Ustaszewski et al. (2008).
Conclusions
Our new seismic velocity model images the deep subsurface of the Ionian
Abyssal Plain in more detail than previous approaches. In addition,
gravimetric modelling validates these findings. We interpret the layer
directly above the crystalline basement, interpreted earlier as layer 2A, as
a unit of seismically fast sediments, possibly carbonates of Mesozoic age.
This is in agreement with the warm environment during the formation of the
Tethys Ocean. Our data and the model indicate that the Ionian Abyssal Plain
is underlain by oceanic crust with seismic velocities increasing with depth
from 5.1 to 7.2 km s-1. The thickness of the crust in the Ionian Abyssal
Plain is 6–7 km but thickens to ∼9 km in the south. At the
crust–mantle boundary crustal seismic velocities decrease from 7.2 km s-1 in
the north to 6.8 km s-1 in the south. This change in the crustal structure
supports the interpretation of the Malta Escarpment as a transform margin,
reactivated normal faults in the SW–NE direction in the Ionian microplate, and a
passive margin in the vicinity of the Medina seamounts.
Based on the seismic and gravimetric results the crust is interpreted to be
oceanic. We consider the Ionian Abyssal Plain to be a remnant of the Tethys
lithosphere. This supports the hypothesis that the Adria was a rigid
promontory of Africa until the beginning of the opening of the Tethys in the
early Mesozoic. During the opening of the Tethys, the Adriatic microplate was
separated from Africa. Between the end of spreading in the early Mesozoic
and the formation of the Calabrian Arc in the Miocene the connection between
Adria and Africa might have been rigid. However, at least in Neogene times,
the Tethyan lithosphere was less rigid, allowing the Adria microplate to move
relative to Africa again. This places the African Plate margin much further
south than previously thought, to the Hellenic subduction zone and the
Calabrian subduction zone, and supports an Adria microplate that can move
relative to Africa and Europe.
Data availability
Seismic data are available on request from the first or second author or via
10.1594/PANGAEA.899796 (Dannowski and Kopp, 2019).
Competing interests
The authors declare that they have no conflict of
interest.
Acknowledgements
The cruise M111 of RV Meteor was funded by the Deutsche
Forschungsgemeinschaft (DFG) with additional support from GEOMAR. Our special
thanks go to the captain and crew of RV Meteor for their excellent support
at sea. We thank Alina Polonia and Eline Le Breton for their detailed and
helpful reviews to improve the paper.
The article processing charges for this open-access publication were covered by a Research Centre of the Helmholtz Association.
Review statement
This paper was edited by Patrice Rey and reviewed by Eline Le Breton and Alina Polonia.
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