Marine Geology 177 (2001) 271±293
www.elsevier.com/locate/margeo
Landslides and the evolution of El Hierro in the Canary Islands
Martin J.R. Gee a,*, Anthony B. Watts a, Douglas G. Masson b, Neil C. Mitchell c
b
a
Department of Earth Sciences, Oxford, UK
Southampton Oceanography Centre, Southampton, UK
c
Department of Earth Sciences, Cardiff, Wales, UK
Received 30 March 2000; accepted 19 March 2001
Abstract
Seismic and sonar data have been used to evaluate the extent and characteristics of giant landslides on the ¯anks of El Hierro
in the Canary Islands. As the youngest and most southwesterly of the Canary Islands, El Hierro has experienced rapid growth
and destructive events in its 1.12 million year history. At least four giant landslides (El Golfo, El Julan, San Andres, and Las
Playas) have modi®ed ,450 km 3 of El Hierro during the last 200±300 thousand years, with each landslide event removing
around 3% of the total edi®ce volume. The extent of landsliding indicates that it is the main process of decay. We characterise
¯ank morphology around El Hierro and distinguish between rugged, unfailed ¯ank, failed ¯ank and steep gullied ridge. Flanks
affected by landsliding have downslope long pro®les with distinctive b coef®cients and exponential forms. The El Golfo
landslide is the most recent (15 ka), best described and clearly de®ned landslide in the Canary Islands. The El Julan landslide
(SW ¯ank) has an estimated volume of 130 km 3, an age of .200 ka and is characterised by gravitational slumping. On the SE
¯ank, two new landslide events are reported. The younger landslide (Las Playas) occurred 145±176 ka, has a narrow, steepsided embayment and a corresponding blocky debris avalanche deposit. The older landslide (San Andres) is recognised on the
basis of a highly chaotic seismic facies offshore and reduced upper ¯ank gradients. Its lack of an upper ¯ank embayment and
offshore blocky debris avalanche lead us to interpret that the landslide involved gravitational slumping, possibly a series of
events, which reduced upper ¯ank gradients, but did not catastrophically collapse to produce a debris avalanche. q 2001
Elsevier Science B.V. All rights reserved.
Keywords: Oceanic island; Landslide; Debris avalanche; Slump
1. Introduction
Large-scale landsliding is now known to be one of
the most important and effective processes involved in
the destruction of oceanic islands. Failure of 10±
1000s of km 3 of material from oceanic island ¯anks
results in high seacliffs, arcuate embayments and
* Corresponding author. Present address: Ocean Mapping Group,
Department of Geodesy and Geomatics Engineering, University of
New Brunswick, P.O. Box 4400, Fredericton, NBE3B5A3, Canada.
E-mail address: [email protected] (M.J.R. Gee).
related offshore landslide deposits. Menard (1956)
introduced the term `archipelagic apron' to describe
the deposits of gravity driven processes such as
slumps, debris avalanches and turbidites surrounding
volcanic oceanic islands. He argued that for many
islands, the volume of volcanic material present in
the archipelagic apron may exceed the volume of
the subaerial and submarine part of the island itself.
Giant landslides have been reported from the
Hawaiian Islands (Lipman et al., 1988; Moore et al.,
1989; Moore et al., 1994), Reunion Island (Lenat et
al., 1989; Labazuy, 1996; Ollier et al., 1998) and the
0025-3227/01/$ - see front matter q 2001 Elsevier Science B.V. All rights reserved.
PII: S 0025-322 7(01)00153-0
272
M.J.R. Gee et al. / Marine Geology 177 (2001) 271±293
Fig. 1. Location map of El Hierro in the Canary Islands. Solid lines show locations of short-streamer seismic data (Line 2 and Line 3) with
letters A±B, C±D and E±F indicating seismic pro®les presented in Figs. 8,9 and 11. Dashed line shows other tracks of Darwin cruise CD108.
Dotted lines show smoothed bathymetric contours at 1 km intervals. Sea¯oor surveyed with TOBI 30 kHz sidescan sonar is grey shaded.
Canary Islands (Watts and Masson, 1995; Urgeles et
al., 1997; Urgeles et al., 1999; Carracedo et al.,
1999a,b). The strongest evidence for these landslides
comes from the offshore regions, or lower ¯anks
where sonar systems have proved effective in imaging
large landslide deposits. Often these deposits can be
traced upslope into large coastal embayments, for
example El Golfo on the northern ¯ank of El Hierro
in the Canary Islands (Masson, 1996). While standard
seismic re¯ection pro®ling systems are not considered
capable of adequately resolving subsurface information of landslides (Moore et al., 1989) we report a
limited success with the pro®ling system used here.
Much of what we know about giant landslides on
oceanic islands comes from sidescan sonar and
morphological data, which gives information only
on the landslide surface. Seismic re¯ection data
presented here allow some of the internal structure
of the landslide deposits to be imaged and gives
valuable insights into the landsliding process.
In a study of the Hawaiian ridge, Moore et al.
(1989) identi®ed two general landslide types, slumps
and avalanches. The slumps are `slow moving, wide
(up to 110 km) and thick (up to 10 km) with tranverse
blocky ridges and steep toes. The debris avalanches
are fast moving, long (up to 230 km) compared to
width, and thinner (0.05±2 km).' Debris avalanches
commonly have well-de®ned amphitheatres at their
head and hummocky terrain in their lower part
(Moore et al., 1989). The debris avalanches described
here, from the ¯anks of El Hierro, are analogous
(although smaller) to those described from the
Hawaiian ridge. Our data only provide evidence for
slump activity on El Hierro to be imaged in the upper
few hundred metres, although we suggest it may affect
the ¯ank to a greater depth.
M.J.R. Gee et al. / Marine Geology 177 (2001) 271±293
273
Fig. 2. Map of El Hierro island showing the approximate location of the three rift arms orientated at 1208 to each other (thick grey lines, as
interpreted by Carracedo (1994)). The three embayments of El Golfo, Las Playas and El Golfo are located between the rift arms. The white
dashed lines indicate the concealed and exposed collapse scars and the solid white line shows a simpli®ed form of the San Andres fault system
(Day et al., 1997).
2. Geological setting
The Canary Islands consist of seven major islands
which are located close to the northwest African
margin (Fig. 1). The islands, extend for 450 km westward from the African margin. Tenerife, with its
3718 m high Volcano of Pico de Teide is the world's
third highest oceanic island after the two Hawaiian
Islands of Mauna Kea and Mauna Loa.
The Canary Islands are generally believed to have
originated in the early Miocene (Schmincke, 1982) as
the African plate moved slowly over a mantle hotspot
(Burke and Wilson, 1972). Unlike the Hawaiian
Islands, however, the Canary Islands do not follow a
simple age progression, although there is a general
age progression of the oldest volcanic rocks from
Fuerteventura in the east (.20 Ma) to El Hierro in
the west (,2 Ma). La Palma and El Hierro, the
westernmost islands, are both younger than 2 Ma
(Ancochea et al., 1994; Guillou et al., 1996) and
Tenerife is less than 7.5 Ma (Ancochea et al., 1990)
in age, indicating that the focus of present day hotspot
activity is in this region, and that the island chain is
extending towards the west. La Gomera is an anomaly
in that it is at least 12 Ma in age and located between
these three younger islands. However, all the Canary
Islands, apart from La Gomera, have been volcanically active at 5 ka (Schmincke, 1982). Carracedo et
al. (1998) reports that activity in the eastern islands is
much less frequent than in the west. According to
Carracedo et al. (1998) the islands of El Hierro, La
Palma and Tenerife are currently in an initial `shield
building-stage', with Gran Canaria, Fuerteventura and
Lanzarote being in an `erosional post-shield stage'. La
Gomera is interpreted as being in a `post-shield gap'.
One possible explanation is a single hotspot, which
has been de¯ected beneath all seven major islands
(Hoernle and Schmincke, 1993).
El Hierro is the youngest and most southwesterly of
the seven Canary Islands (Fig. 1). The oldest subaerial
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M.J.R. Gee et al. / Marine Geology 177 (2001) 271±293
Fig. 3. Grey shaded image of bathymetry and topography of El Hierro with contours shown at 200 m intervals. Landslide boundaries are shown
as white lines. Black lines locate the three geological pro®les in Fig. 12.
rocks have been dated at 1.12 ^ 0.02 Ma, although
much of the island is covered by more recent lavas
which record the rapid growth of the island (Guillou et
al., 1996). El Hierro has an estimated total edi®ce
volume (subaerial and submarine) of 5500 km 3 and
rises about 5500 m from its base in 4000 m of water
(Schmincke, 1990). On the basis of magnetic anomaly
studies, El Hierro is believed to be located on Jurassic
oceanic crust aged around 156 Ma (Roeser, 1982;
Klitgord and Schouten, 1986; Roest et al., 1992).
El Hierro displays some of the most recent and
dramatic evidence of landslides in the Canaries.
Onshore, the island morphology is characterised by
three large embayments separated by three ridges
(Fig. 2). The term `landslide' is used here in a general
sense without implication of the failure process. The
onshore morphology, characterised by high seacliffs
surrounding valleys, has been the subject of controversy. The El Golfo embayment was originally
interpreted as the remains of a great central caldera
by Von Knebel (1906) and later by Blumenthal
(1961). One of the early authors to attribute El
Golfo to a landslide origin was Hausen (1972),
although at that time he had no knowledge of debris
avalanche deposits which resided offshore. Others
have proposed non-landslide origins for similar
M.J.R. Gee et al. / Marine Geology 177 (2001) 271±293
features on Tenerife, such as creation by ¯uvial
erosion (Palacios, 1994), or caldera collapse involving
mainly vertical and some lateral collapse (Ridley,
1971; Marti, 1998). The most unambiguous evidence
for landsliding comes from offshore sidescan sonar,
and to a lesser extent seismic re¯ection data (Holcomb
and Searle, 1991; Masson, 1996; Urgeles et al., 1997;
Gee, 1999). On the basis of submarine data the three
embayments all result from landslides occurring
within the last few hundred thousand years (Holcomb
and Searle, 1991; Masson, 1996; Urgeles et al., 1997;
Gee, 1999). At least four giant landslides, relating to
the three embayments, can be recognized offshore
(Figs. 2 and 3) (Holcomb and Searle, 1991; Fuster
et al., 1993; Masson, 1996; Urgeles et al., 1997; Day
et al., 1997; Carracedo et al., 1999a; Gee, 1999). Clear
subaerial evidence for an older north±west directed
landslide (Tinor) was shown by Carrecedo et al.
(1999a,b), although no related offshore deposit are
observed. The youngest landslide occured around
15 ka and created a huge debris avalanche on the
north ¯ank (Masson, 1996; Urgeles et al., 1997).
The avalanche deposit is connected to the subaerial
El Golfo embayment by a smooth `chute' bounded by
lateral scarps, however, subaerial evidence indicates
a much greater age for the El Golfo embayment
(Carracedo et al., 1999a).
On the basis of GLORIA sidescan sonar and bathymetry data, Holcomb and Searle (1991) reported a
landslide on the south western ¯ank, offshore of the
El Julan embayment, with an estimated volume in the
region of 100 km 3. This embayment and related landslide is investigated in more detail using new swath,
seismic and sonar data. In addition, two new landslides (San Andres and Las Playas) are reported on
the southeastern ¯ank (Fig. 3).
Landslide deposits from the ¯anks of oceanic
Islands tend to deposit into surrounding ¯exural
moats. Tenerife represents a signi®cant volcanic
load, resulting in ¯exure of the underlying crust by
between 2.5 and 3 km (Watts et al., 1997), however,
there is no evidence to show that a ¯exural moat has
formed around Las Palmas and El Hierro. Las
Palmas and El Hierro may be too small and too
young to have a signi®cant loading effect on the
underlying crust. In accordance with a hotspot
model, the islands of Las Palmas and El Hierro
may be located above a hot mantle plume and
275
therefore experiencing thermal buoyancy, rapid
growth and landsliding
3. Data acquisition and processing
This paper presents EM12 multibeam, seismic
re¯ection and TOBI 30 kHz sidescan sonar data
obtained during the 1997 R.R.S. Charles Darwin
cruise 108 (Fig. 1). The Simrad EM12 system is a
hull-mounted multibeam echo sounder capable of
recording detailed bathymetric and backscatter data
with a resolution of tens to hundreds of metres in
the horizontal and less than 5 m in the vertical direction
(Hammerstad et al., 1991). Operating at a frequency of
13 kHz, the EM12 system uses 81 beams spread over
1208 to image a swath of sea¯oor 2±2.5 times the water
depth in width. Bathymetric data was merged with topography digitised from Spanish Instituto Geographica
maps and a ®nal grid produced at 0.1 £ 0.1 mins (i.e.
183 £ 183 m) using the Generic Mapping Tools (gmt)
software (Wessel and Smith, 1991). A grey-shaded
image of the bathmetry (constructed using the ®nal
bathymetric grid) is shown in Fig. 3.
More than 600 km of seismic re¯ection data were
obtained over the southern ¯anks of El Hierro. The
seismic pro®ling system consisted of a 200 m long
streamer with four active sections spaced 50 m apart
and a single 300 cubic inch airgun ®tted with a wave
shape kit. Sampling rate was 0.1 ms. Processing (of
the seismic re¯ection data) was carried out using the
promax Version 6 processing package. Processing
steps included stacking, predictive and spiking deconvolution, amplitude recovery, ensemble balancing,
bandpass ®ltering, migration and trace mixing.
The TOBI 30 kHz sidescan sonar system is a high
resolution mapping device developed by the UK Institute of Oceanographic Sciences (Murton et al., 1992).
When towed 400 m above the sea¯oor, TOBI can
insonify up to 3 km of sea¯oor on each side of the
vehicle and resolve features as small as a few metres
across. Processing of TOBI data was carried out using
ERDAS imagine and the Southampton Oceanography
Centre's prism software. The same software packages
were also used to process backscatter data extracted
from the EM12 multibeam system. Navigation during
CD108 used standard global positioning system
(GPS) with an accuracy of ,35 m.
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M.J.R. Gee et al. / Marine Geology 177 (2001) 271±293
Fig. 4. 3-D views of three ¯anks affected by landsliding (landslide boundaries shown as dashed line). (A) El Golfo landslide viewed from the
NW, (B) San Andres and Las Playas landslides viewed from the SE and (C) El Julan landslide viewed from the SW.
M.J.R. Gee et al. / Marine Geology 177 (2001) 271±293
4. Data analysis and observations
4.1. Swath bathymetry and topography
On the basis of the swath bathymetry maps and the
subaerial topographic data, the ¯anks of El Hierro can
be divided into three contrasting morphological types:
(i) rugged, unfailed volcanic ¯ank, (ii) steep gullied
ridge, and (iii) broad ¯at ¯oored valleys with relatively low slope gradients, leading to debris avalanche
deposits. `Rugged, unfailed ¯ank' describes the
morphology of the NE and NW ¯anks of El Hierro
which extend (from onshore) to a depth of about
3200 m offshore (Fig. 3). Slopes average 158,
although the rough ¯ank morphology, characterised
by large pinnacles, ridges and gullies typically 150±
300 m in amplitude, gives rise to large variations in
slope. There is a distinct break of slope at around
3200 m depth which marks the base of the exposed
volcanic edi®ce (Fig. 3). Below this depth, slope
angles rapidly decline from a mean gradient of ,15
to ,68. Onshore, the ridge zones are de®ned by a
three-armed pattern of fractures, dykes and volcanic
centres (Pellicer, 1977; Fuster et al., 1993). Offshore,
however, there are no discrete ridge systems. Instead,
the submarine ¯ank morphology is rugged over a
broad area, characterised by pinnacles and irregular
gullies (Figs. 3 and 4). Tentatively, we suggest that
within each NW and NE offshore rugged sector, there
might be a bifurcation of rift zone activity (Gee et al.,
2001).
The Southern Ridge of El Hierro is characterised by
deeply gullied slopes with (slope) gradients of 10 to
.308, which give way to smoothly sedimented
sea¯oor at depths of between 3400 and 3700 m
(Figs. 3 and 4). This type of morphology is distinct
from the other submarine ¯anks and is interpreted as
the eroded remnant of a volcanic edi®ce (Gee et al.,
2001) probably older than the main part of El Hierro.
Broad, relatively ¯at-¯oored valleys, with gentler
slopes (5±158), characterise island ¯ank regions
between the three constructional sectors (Figs. 3 and
4). These valley areas contrast with the other ¯ank
types by being much smoother, with a notable absence
of a distinct slope break around 3200 m water depth.
The valley on the northwest ¯ank of the island,
immediately downslope from the El Golfo embayment, has the smoothest morphology, characterised
277
by a relatively ¯at ¯oor and steep lateral scarps
about 400 m high. In this valley, the mean slope
decreases from 108 at 1200 m to 58 at 3000 m, with
a smaller variability in slopes compared with the
rugged, unfailed slopes. A large area of blocky
morphology, interpreted as debris avalanche deposit,
is located downslope of the El Golfo embayment in
water depths greater than 3200 m, indicating that
landsliding is the dominant process in valley formation (Masson, 1996; Urgeles et al., 1997). Individual
debris avalanche blocks from the El Golfo landslide
are irregular in shape, up to 1.2 km across and 200 m
high, and appear to be randomly distributed (Masson,
1996).
The SW ¯ank (El Julan), has a large embayment
with a smooth morphology, similar to the El Golfo
valley, although it has less well-de®ned lateral scarps.
A broad bathymetric bulge or swell is seen across the
lower part of the island ¯ank at depth .3000 m
(Fig. 3). It is also characterised by the presence of a
few large pinnacles, each several hundred m wide and
up to 300 m high (e.g. at 28833 0 N, 18813 0 W, Fig. 4).
Compared with the irregular debris avalanche blocks
within the El Golfo landslide, pinnacles within the El
Julan valley are larger, fewer in number and have
more regular conical shapes. They show some degree
of downslope alignment. Similar `pinnacles' occur
within the rugged, unfailed ¯anks indicating that
they are probably submarine volcanic features related
to igneous dykes radiating from the main island
edi®ce.
The SE ¯ank lacks an obvious broad valley like El
Julan and El Golfo. Instead, it has a complex head
wall region, which includes a narrow, steep-sided
embayment (Fig. 4). This occurs at the southern end
of the San Andres fault system, (see Fig. 2) where part
of the SE ¯ank is displaced t300 m seawards along
steep normal faults that run parallel to the shore (Day
et al., 1997). The fault system was interpreted by (Day
et al., 1997) as evidence for an aborted landslide. At
the southwestern end of the fault system, they interpreted the narrow steep-sided embayment as a series
of eroded strike±slip faults oriented perpendicular to
the shoreline. Our new multibeam data indicate that
this embayment is the headwall of a landslide (Figs. 3
and 4). Further downslope, EM12 multibeam, 3.5 kHz
and seismic re¯ection data show evidence for a blocky
debris avalanche deposit and an older, larger slump
Fig. 5. EM12 13 kHz backscatter sonar of El Hierro (A) with geological interpretation (B). Dark grey tones are low backscatter and light tones are high backscatter (A); `SP' shows
the speckled patterns typical of debris avalanches, `CL' shows high backscatter, curving lineations over the southwestern ¯ank. The geological interpretation (B) is based on all
seismic and sonar data presented in this paper and shows landslides interpreted as either debris avalanches or slumps. Onshore locations of rift zones and landslide boundaries are
simpli®ed from Carracedo et al. (1999a). Bathymetric contours are annotated at 1 km intervals.
278
M.J.R. Gee et al. / Marine Geology 177 (2001) 271±293
M.J.R. Gee et al. / Marine Geology 177 (2001) 271±293
event extending ,50 km offshore (Fig. 3). In contrast
to the El Julan landslide there is no broad bathymetric
bulge or swell across the lower ¯ank. The SE submarine ¯ank appears to be generally more rugged
than the other landslide headwall regions although
the slopes are more gentle than rugged, unfailed ¯ank.
4.2. EM12 backscatter data
EM12 backscatter data is a useful reconnaissance
aid in the search for debris avalanches. Where backscatter is shown as dark grey, the (hemi) pelagic sediment cover is probably greater than a few metres to
attenuate the 13 kHz acoustic signal (Mitchell, 1993).
The steep volcanic ¯anks of El Hierro are characterised by high backscatter, extending to around
3200 m water depth (Fig. 5). In deeper water, backscatter levels are generally lower with some areas
characterised by a distinctive speckled pattern. This
`speckled pattern', also recognized from other sonar
data (e.g. GLORIA and SEAMARC II) around oceanic
islands, is a signature typical of debris avalanches
(e.g. Holcomb and Searle, 1991; Labazuy, 1996)
and results from blocks of volcanic debris standing
above (hemi) pelagic sediments accumulating
between blocks. The northern ¯ank of El Golfo
shows the clearest evidence for debris avalanche
deposits, with a large areas of high backscatter
`speckled pattern' corresponding to the El Golfo
debris avalanche on its lower slopes (Masson,
1996). The SE lower ¯ank is characterised by an elongated lobe of `speckled backscatter', approximately
25 km wide and 50 km in length, extending downslope from the Las Playas embayment (Fig. 5). This
indicates the presence of debris avalanche deposits. In
contrast, EM12 backscatter data show the El Julan
lower ¯ank region to be dominated by gently curving,
high backscatter lineations which trend along strike,
correlating with a broad bathymetric bulge in the
lower ¯ank (Fig. 3). This area was previously mapped
as the location of the El Julan landslide (Holcomb and
Searle, 1991).
4.3. Numerical characterisation of ¯ank morphology
(of El Hierro) and criteria for distinguishing ¯ank
types using bathymetric data
A series of 2-D bathymetric pro®les, arranged in a
radial fashion along the average direction of maxi-
279
mum dip around the island, were constructed in an
attempt to characterise the landslide and non-landslide morphologies on the ¯anks of El Hierro (Fig.
6, see Fig. 7 for location). In total, 47 pro®les were
constructed from the EM12 swath bathymetry with
depth sampled at 100 m intervals along each pro®le.
Each pro®le is 40 km long, with a common depth
origin at 1081 m (limited by the extent of data in
shallower water). Each pro®le was spaced with care
to avoid as far as possible aligning with gullies or
groups of pinnacles. No pro®les were sampled over
the Southern Ridge as the direction of dip varied by
over 908 between the upper and lower ¯anks. No
pro®les of suf®cient length could be sampled in the
NE quadrant (between 40 and 1108) due to insuf®cient
multibeam data coverage (Fig. 7A).
The gross form of each curve is concave upwards
with a generally rugged, steeper upper section and a
smoother, more gently sloping lower section (Fig. 6).
It was noted that some pro®les, especially those
within the landslide sectors of El Julan and El Golfo
(Fig. 7b), showed very good ®ts to exponential curves,
while pro®les from other ¯ank regions did not. In
order to characterise the degree to which each bathymetric pro®le can be expressed by an exponential
function, the root mean squares difference (rms
residual), correlation coef®cient (r) and exponential
coef®cient (b term in Eq. (1)) were calculated. The
form of the exponential curve to which distance and
depth values were ®tted is given by the equation:
y ˆ …ae2bx † 1 c
…1†
where y is the water depth (m), x the horizontal
distance (m), a scales the exponential term, b scales
the distance over which the exponential declines
(m 21) and c de®nes horizontal asymtotical value (m).
Pro®les representing a range of ¯ank types are
illustrated in Fig. 6 (see Fig. 7A for location). Pro®les
20, 31 and 42 sample sections interpreted as `rugged,
unfailed ¯ank'. They have steep, rugged upper
sections above a distinct break-of-slope at around
3200 m, below which the topography is subdued.
P20 shows a second, less well-de®ned break-ofslope at around 2250 m, with steeper slopes above
2400 m. Pro®les from unfailed ¯ank are ®tted with
exponential curves with large `b' coef®cient values
(Eq. (1)) (Fig. 6). They show good ®ts with
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M.J.R. Gee et al. / Marine Geology 177 (2001) 271±293
Fig. 6. Selected topographic pro®les perpendicular to the ¯anks of El Hierro (see Fig. 7A for pro®le locations). Pro®les re¯ect the range of rms
(root mean square residual) and b coef®cient values measured. Dotted lines show exponential curves ®tted to the pro®les and values above each
pro®le give the rms and correlation coef®cient of the ®t and the exponential coef®cient (b).
M.J.R. Gee et al. / Marine Geology 177 (2001) 271±293
281
Fig. 7. (A) location of 47 pro®les around El Hierro used in slope analysis. (B) Prediction of sectors affected by landsliding using rms residual
and the exponential coef®cient `b' as criteria. Landslide outlines interpreted from sonar and seismic re¯ection data shown as solid lines. (C) plot
of rms versus bcoef®cient for all 47 pro®les. Note that data from the landslide regions of El Golfo, San Andres, Las Playas and El Julan tend to
have rms residual values ,50 (m) and `b' coef®cient values between 20.06 and 20.12. Data from constructional sectors tend to have higher
rms residuals and higher `b' coef®cients. Notable exceptions include pro®les 27 and 26 which are not located within known landslide regions
and pro®le 13, which is located within the El Julan landslide region.
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M.J.R. Gee et al. / Marine Geology 177 (2001) 271±293
Fig. 8. Interpreted seismic section (A±B) crossing the southwestern ¯ank of El Hierro (see Fig. 1 for location). Internal re¯ection geometries
within the landslide boundaries are very chaotic, in contrast to the area outside, where more strati®ed re¯ection geometries are seen. An
unconformity is shown east and west of the landslide (u/c) which indicates the onset of mass wasting associated with the early seamount growth
phase of El Hierro.
exponential curves below 3600 m but large deviations
of up to 250 m above the slope break at around
3200 m. This is re¯ected in the higher rms residual
values relative to pro®les from the smoother morphology of the landslide regions (Fig. 7C).
Pro®les 9, 15 and 36 were sampled from landslide
regions and show good ®ts with exponential curves (and
thus low rms values, see Fig. 7C). These pro®les have a
more gentle gradient above 3000 m depth, a lower rate
of change of slope (curvature), and are smoother than
pro®les from unfailed ¯ank areas. They have no
obvious breaks of slope along the pro®le.
The calculated values of rms residual and `b' coef®cient around El Hierro suggests that these two parameters, taken in conjunction, can be used to de®ne
areas affected by landsliding (Fig. 7). Within the El
Golfo, El Julan and San Andres/Las Playas landslide
sectors where giant landslides have previously been
documented (or are at least suspected to occur),
pro®les tend to have low rms residuals (,50 m) and
`b' coef®cient values in the range 20.06 to 20.12
(Fig. 6). A plot of rms residual against `b' coef®cient
shows the areas we interpret as landslide and unfailed
¯anks separated into two distinct groups (Fig. 7C).
Some apparent anomalies are seen, e.g. the two
pro®les 26 and 27 (Figs. 6 and 7). These two pro®les
have both the `b' coef®cient and low rms values of
typical `landslide ¯ank' (e.g. pro®le P27, Fig. 6), but
appear to be situated on an area of unfailed slope (see
discussion). P43 is also located on apparently unfailed
¯ank but has an rms residual value similar to pro®les
from landslide regions (ie ,30 m), although a `b'
coef®cient which is much higher. Within landslide
sectors, some pro®les have higher than expected rms
M.J.R. Gee et al. / Marine Geology 177 (2001) 271±293
283
Fig. 9. Interpreted seismic re¯ection section showing scarps and chaotic facies within the El Julan landslide (see Fig. 1 for location). `CF' ˆ
chaotic facies, `a' to `f' ˆ scarps discussed in text. Lower pro®le shows 3.5 kHz data, illustrating scarps (a±d) and contrast in facies where
seismic chaotic zones exist. Pro®le E±F crosses the scarps obliquely and therefore their apparent dips are which is less than their true dips.
residuals, for example pro®le 13 from the El Julan landslide (Fig. 6). This pro®le crosses a large pinnacle at
,2000 m water depth, giving rise to a higher than
expected rms residual, although the `b' coef®cient has
similar values to other pro®les within landslide
regions.
4.4. Seismic re¯ection pro®les
Seismic re¯ection pro®les were collected on the SE
and SW ¯anks of El Hierro (see Fig. 1 for location).
Seismic pro®les crossing the strongly lineated region
of high backscatter on the SW (El Julan) ¯ank of El
284
M.J.R. Gee et al. / Marine Geology 177 (2001) 271±293
Fig. 10. Interpreted seismic re¯ection section A±B crossing the SE ¯ank of El Hierro (see Fig. 1 for location). Two landslides can be identi®ed;
the Las Playas debris avalanche and the San Andres slump. Chaotic and irregular geometries in the centre of the section are interpreted as debris
avalanche deposits. Contrasts between the chaotic facies and an underlying (partly) strati®ed facies are interpreted (small black arrows) as a
possible surface over which the debris avalanche moved. The older landslide (San Andres) also has a chaotic/irregular seismic geometry which
contrasts with more strati®ed facies to the SW and NE. Landslide deposits average 80±100 ms (TWTT) in thickness.
Hierro show a highly chaotic facies characterised by
numerous short and often steeply dipping, high amplitude re¯ections (Fig. 8). This chaotic facies is interpreted as the deposit of the El Julan landslide
(Holcomb and Searle, 1991). In contrast, at the NW
and SE ends of the pro®le, the internal facies is more
strati®ed and characterised by laterally continuous
re¯ections. In these areas, an unconformity at about
100 ms (TWTT) below sea¯oor, marks the base of a
set of high amplitude, strong and laterally continuous
re¯ection events (Fig. 8). These `events' cannot be
traced across the region characterised by chaotic facies.
The eastern boundary of the landslide, directly
southwest of the Southern Ridge, is easily recognized
by a transition from strati®ed to chaotic internal facies
and a distinct morphological break of slope in the
sea¯oor pro®le. In contrast, the western boundary
has no clear morphological expression. Instead, it is
characterised by a more gentle gradient (oblique to the
pro®le shown in Fig. 8) and a series of small downslope-facing scarps. These scarps extend from the
western margin into the centre of the main landslide
deposit. They are tens of metres high and spaced
2±5 km apart (Fig. 9). These scarps correlate directly
with the broad, high backscatter, along-slope lineations observed on EM12 backscatter data (Fig. 5).
The internal geometry of the chaotic seismic facies
consists of short, steeply dipping re¯ections extending
to a depth of around 100±150 ms (TWTT). These
re¯ections consistently dip towards the island, in the
opposite direction to the general dip of the sea¯oor
(Fig. 9). 3.5 kHz data over the area of chaotic facies
show numerous hyperbolae, indicating small-scale
roughness (Fig. 9). We interpret these scarp features
and related subsurface dipping re¯ections as evidence
for listric style faulting within the landslide mass.
M.J.R. Gee et al. / Marine Geology 177 (2001) 271±293
285
Fig. 11. TOBI 30 kHz sidescan sonar data showing saw-tooth scarps running along strike of the El Julan landslide (see Fig. 1 for location). `1',
`2' and `3' indicate small scarps. `A' and `B' are landslide blocks.
Pro®les crossing the lower part of the SE ¯ank, also
exhibit a chaotic seismic facies (Fig. 10). This ranges
from a chaotic, crudely strati®ed facies which
obscures deeper re¯ections to a chaotic facies about
100 ms (TWTT) thick overlying a strati®ed, apparently folded subsurface. The central part of the seismic section (Fig. 10) is characterised by a distinctive
chaotic facies, in®lling irregular but apparently folded
strata. This correlates with the speckled, lobate pattern
observed on EM12 backscatter data (Fig. 5) and is
interpreted as the debris avalanche deposit of Las
Playas. The boundaries of the older San Andres landslide are de®ned by a change from chaotic to more
strati®ed facies, with over 400 ms (TWTT) of strati®ed facies recognized outside the avalanche deposit
(Fig. 10).
4.5. TOBI 30 kHz deep-towed sonar data
A single TOBI 30 kHz sidescan sonar swath, 6 km
wide, was obtained over the El Julan ¯ank (Fig. 11,
see Fig. 1 for location). This sonar swath shows a
series of high backscatter, saw-toothed scarps orientated along slope in water depths of 3200±3800 m
(Fig. 11). They are spaced approximately 1±2 km
apart and range from around 10 m to over 20 m
high. The saw-toothed geometry of these features is
interpreted as erosional degradation of the scarps.
TOBI 30 kHz data also show several positive topographic features which cast acoustic shadows (Fig.
11). Two of the larger examples (labelled A and B
in Fig. 11) have estimated heights of between 25 and
30 m. From the steep sides and angularity of the two
features, it is likely that they are landslide blocks.
5. Discussion
5.1. Landslide structure and emplacement processes
The El Golfo landslide, on the northwest ¯ank of El
Hierro, is the most recent example of a large scale
¯ank collapse in the Canary Islands (Masson, 1996;
Urgeles et al., 1997; Masson et al., 1998). As it has a
clearly de®ned embayment, landslide valley and offshore debris avalanche deposit, it is a useful reference
for the recognition of similar morphological landslide
features on the other ¯anks. There is evidence to
suggest that the onshore embayment might be older
than the offshore deposits (see Carracedo et al.,
1999a), however, there is only evidence for one
offshore debris avalanche, dated at ,15 ka (Masson,
286
M.J.R. Gee et al. / Marine Geology 177 (2001) 271±293
1996). We only use offshore data in our numerial
analysis of morphology.
Compared to typical Hawaiian landslides, the El
Julan landslide is similar in that it lacks a well-de®ned
amphitheatre. However, given the age of El Julan
(probably .200 ka, see below) constructive and
destructive processes have had time to remove or
bury all subaerial evidence for a landslide erosion
scar. In particular, lava ¯ows of the younger Rift
series (Guillou et al., 1996) are presumed to bury
the landslide scar. The El Julan landslide deposit
appears as a highly chaotic unit on seismic pro®les
and, based on the size of the broad lower ¯ank swell
observed on EM12 swath bathymetry data and on
seismic pro®le data, is up to 300 m thick. The dominant features of this landslide are the numerous scarps
aligned along strike, seen on EM12 backscatter (Fig.
5) and TOBI images (Fig. 11) and seismic pro®les
(Fig. 9). These scarps and their related steeply dipping
subsurface re¯ectors are evidence for widespread
gravitational sediment failure involving at least the
upper 100 m of the landslide deposit. In contrast to
the El Golfo and Las Playas landslides, EM12 multibeam sonar shows less of the blocky and speckled
facies that characterizes most debris avalanches.
This indicates that the emplacement of the El Julan
landslide might have involved slower, super®cial
slump-like behaviour, rather than a catastrophic
collapse to form a debris avalanche such as the El
Golfo landslide. We have no data to constrain the
depth of ¯ank affected by slumping, and thus cannot
accurately estimate the volume. The volume quoted
(130 km 3) is based only on the size of the bathymetric
swell over the lower ¯ank.
Compared to the El Julan landslide, the San Andres
and Las Playas landslides (see Fig. 3 for location)
appear to have much smaller volumes. The younger
of the two landslides (Las Playas) is the smaller of the
two in terms of both surface area and volume. The
blocky nature of this landslide deposit and a wellde®ned erosional scar are similar to the El Golfo landslide (Fig. 4) and therefore we interpret this event as a
debris avalanche. The older San Andres landslide is
clearly different and may involve several separate
landslide events. It lacks features typical of many
other Canary landslides, such as an arcuate upper
¯ank embayment, or a region of smooth morphology
within a landslide valley leading to an area of blocky,
slightly raised sea¯oor on the lower ¯anks. Our analysis shows that rms residual and `b' coef®cient values
from pro®les within the San Andres and Las Playas
landslides have similar values to pro®les from the El
Golfo and El Julan landslides. The upper ¯ank region
of the San Andres landslide is more rugged than the
corresponding El Golfo and El Julan upper ¯ank
regions, however, pro®les 1±5 mostly avoid the
downslope trending ridges which create this `ruggedness'. Seismic re¯ection data over the lower ¯ank of
the San Andres landslide, indicate that ,100 m of
landslide deposits exist (Fig. 10) although EM12
backscatter data indicate debris avalanche deposits
are restricted to a narrow (,25 km) central part. We
interpret the lower curve gradients (`b' coef®cients
between 20.06 and 20.12) of the San Andres landslide 1 and the ,100 m of chaotic seismic facies in the
subsurface of the corresponding lower ¯ank as
evidence for slumping, in contrast to a debris
avalanching process such as created the El Golfo landslide. We interpret the subaerial mapping of the San
Andres fault system on El Hierro (Day et al., 1997) as
indicating the SE ¯ank may be associated with slumping of which the onshore fault system is part of the
slump headwall.
As with the El Julan landslide, we have no data
allowing us to image the deeper ¯ank structure and
to understand the nature of slump activity. The apparently folded, semi-strati®ed seismic character underlying much of the chaotic facies over the SE ¯ank
might have been generated by compressional or
shear forces imposed on the underlying sediments
during slump movement. However, we have little
constraint on the orientation of apparent fold features.
An alternative explanation is that such `apparent folding' relates to compaction effects over an irregular
basement. In favour of the latter hypothesis is the
interpretation of several high amplitude `sills' in the
subsurface (see Fig. 10), over which sediments experienced differential compaction.
Three geological cross-sections (Fig. 12) were
constructed, based on our interpretations of sonar
and seismic re¯ection data and subaerial mapping
by Carracedo et al. (1997) The El Golfo landslide
has a clearly de®ned headwall region and corresponds
to landslide deposits in water depths .3 km. The
maximum thickness of the landslide measured from
seismic re¯ection records is 220 m (Urgeles et al.,
M.J.R. Gee et al. / Marine Geology 177 (2001) 271±293
287
Fig. 12. Geological cross sections through the El Golfo (A), El Julan (B) and Las Playas (C) landslides. Landslide basal surfaces are estimated
from morphological and seismic re¯ection data. The geology of El Hierro is based on work by (Carracedo et al., 1997).
1997). This agrees with a thickness measured by
extrapolating a gently curving landslide base (Fig.
12A). The El Julan landslide has an inferred maximum thickness of approximately 300 m (based on
the height of the broad lower ¯ank swell observed
on EM12 swath bathymetry data) in addition to an
inferred landslide scar buried by in®lling lavas. The
Las Playas landslide is shown as a thinner deposit
(,100 m) as interpreted from seismic pro®les (Fig.
10). In the case of the older San Andres landslide,
288
M.J.R. Gee et al. / Marine Geology 177 (2001) 271±293
Table 1
Landslide parameters. The parameter A/(V 2/3) is from (Dade and Huppert, 1998)
Landslide
Age
(ka)
Area
(km 2)
Volume
(km 3)
A/V 2/3
Height
(km)
Length
(km)
H/L
Process
El Golfo
El Julan
San Andres
Las Playas 2
10±17
.200
176±545
145±176
1500
1800
1700
950
150±180
130
?
,30
47±53
70
4.7
4.6
4.5
4.5
65
60
51
52
0.072
0.076
0.088
0.086
Debris avalanche
Slump/debris avalanche
Slump?
Debris avalanche
.98
the offshore data do not clearly indicate a debris
avalanche such as the El Golfo debris avalanche.
The chaotic/semi-strati®ed seismic facies (Fig. 10)
could result from some form of partial landslide
deformation, rather than debris avalanche originating
from the upper ¯ank. The rugged ¯ank morphology in
water depths ,3 km shows that disintegration and
subsequent formation of a smooth landslide valley
with steep bounding scarps did not develop.
5.2. Flow ef®ciency
A common method of expressing the ef®ciency of
gravity driven landslides is by measuring the distance
travelled and the subsequent decrease in elevation of
the landslide mass. Scheidegger (1973) developed a
friction based model in which the elevation difference
(h) divided by the total runout distance (l) was equivalent to a friction coef®cient acting at the base of the
landslide. This crude but simple model is useful when
applied to giant landslides, giving some indication of
the relative ef®ciencies of different landslides. Landslides on El Hierro have h/l ratios of between 0.072
and 0.088 (see Table 1), which are similar to ratios for
other landslides both in the Canaries and around the
Hawaiian chain. Dade and Huppert (1998) developed
scaling arguments for the runout behaviour of large
rockfalls and landslides which involved the radial
spread of the deposit. Using the parameter A/(V 2/3)
(from Dade and Huppert, 1998) where A is area and V
volume, values of 47±53, 70 and .98 were calculated
for the El Golfo, El Julan and Las Playas landslides
respectively (A reasonable estimate for the volume of
the older San Andres landslide could not be made, see
Table 1). Higher numbers indicate greater ef®ciency,
re¯ecting how the landslide mass spreads over the
lower ¯ank. It is interesting to note that the Las Playas
landslide has the higher h/l ratio (i.e. higher apparent
coef®cient of friction) but also the higher A/(V 2/3)
ratio. The ratio A/(V 2/3) takes into account the spreading of a landslide mass and is clearly the more reliable
indicator of landslide ef®ciency in the case of the Las
Playas landslide. We therefore conclude that the smaller Las Playas landslide was more ef®cient than the El
Golfo and El Julan landslides. The El Golfo and El
Julan landslides have similar A/(V 2/3) and h/l ratios,
indicating that both landslides might be comparable
events in terms of ef®ciency. However, if the El Julan
landslide was a slump rather than a debris avalanche,
then the volume might be considerably higher than
quoted here. This would mean the El Julan landslide
was less ef®cient relative to the El Golfo landslide.
5.3. Constraining the ages of the landslides
In the absence of samples from the offshore landslide deposits, the ages of the El Julan, San Andres
and Las Playas landslides cannot be determined
directly. However, by examining onshore geological
evidence and offshore 3.5 kHz data, some constraints
can be placed on the likely age of landsliding.
The age of the El Golfo landslide is currently in
debate (see Carracedo et al., 1999a). Offshore it has
been dated ,15 ka by its relationship with the Canary
debris ¯ow and a turbidite (see turbidite `b' in
Masson, 1996). Over the region of the landslide
deposit, 3.5 kHz data and TOBI 30 kHz sidescan
sonar data indicate little or no sediment drape
(Masson, 1996), which would be expected for a relatively young landslide. Onshore, there is a marine-cut
platform with overlying lavas, within the El Golfo
embayment. Dating of these post-landslide lavas indicates a minimum age of 21 ^ 3 ka (Guillou et al.,
1996), although lavas onlapping higher up within
the embayment indicate an age of 134 ^ 3 ka (Carracedo et al., 1999a). This lack of onshore-offshore
M.J.R. Gee et al. / Marine Geology 177 (2001) 271±293
correlation may be reconciled by interpretation of an
older collapse event to create the subaerial embayment, and a younger offshore event occurring around
15 ka (Carracedo et al., 1999a).
In the area of the El Julan landslide, 3.5 kHz data
indicate 10±12 m of post-landslide `drape', suggesting an age for landslide emplacement of at least
200 ka (Gee, 1999). In addition, the distal part of
the landslide has been overrun by the Saharan debris
¯ow, an event that occurred at approximately 60 ka
(Gee et al., 1999). Onshore, all evidence for the
inferred landslide scar is buried beneath later lavas
that in®ll the coastal embayment. These lavas have
a maximum age of 160 ma, giving a minimum age
for the landslide (Guillou et al., 1996; Carracedo et
al., 1997). Additional evidence for El Julan being the
oldest landslide reported here is the presence of
`pinnacles' interpreted as post-landslide volcanic
structures which grew as dyke material was intruded
into the landslide mass. Younger landslides such as El
Golfo show no evidence for post-landslide volcanic
activity. Given that the age of the El Julan landslide
was estimated at .200 ka, then a rough calculation of
¯ank regeneration can be made. Assuming each
`pinnacle' to be approximated by a cone shape with
an average basal diameter of 3 km and height of
150 m indicates a regeneration rate of 20 km 3 per
million years. This re¯ects the fact that island growth
occurs mainly along active rift zones, and that the
island was also close to its present size before the El
Julan landslide occurred.
In the region of the San Andres and Las Playas
landslides, 3.5 kHz data show less than 10 m of
post-landslide drape, which suggests a period of sedimentation of at least 150 ka. Onshore geological data
constrain the age of movement on the San Andres
Fault System to between 176 and 250 ka (Day et al.,
1997). Displacement on this fault system, of up to
300 m in a seaward direction, was interpreted as
evidence for a aborted landslide (Day et al., 1997),
although we believe that the subaerial faults are part
of the head wall of the older San Andres landslide.
The Las Playas subaerial embayment, interpreted as
non-landslide in origin prior to the CD 108 cruise, was
re-evaluated in the light of the discovery of debris
avalanche deposits offshore during the cruise. The
submarine data show how this embayment extends
offshore and leads to the younger, blocky deposit of
289
Las Playas landslide. On the basis of volcanic
sequences which are cut by, and post-date the embayment, S.J. Day (personal communication, 1998) estimates the Las Playas landslide to have occurred
between 145 and 176 ka. A younger age for the Las
Playas landslide compared to the El Julan landslide is
consistent with both the onshore dating by Guillou et
al. (1996) and thinner observed post-landslide drape
over the Las Playas debris avalanche.
5.4. Morphology of the edi®ce
The aim of analysing the swath bathymetry (Figs. 6
and 7) was to characterize the different ¯ank morphologies, in particular the curvature and roughness of
sectors affected by giant landslides. Landslide sectors
show good ®ts with exponential curves when
compared with other ¯ank types, suggesting that
the general exponential form could be used as a
criteria for identifying giant landslides on other island
¯anks.
Adams et al. (1998) attributed the long pro®le
exponential form of submarine slopes to the `exponential decay of transport capacity, e.g. the cumulative effect of debris ¯ows, creep and small-scale
slumping with increasing distance from the shelfbreak'. However, they considered very large slump
masses to be responsible for deviations from the exponential form. On the ¯anks of El Hierro, in direct
contrast to their observations, the long pro®le exponential slope form results from the catastrophic failure
of around 100±200 km 3 of material from island
¯anks. This different interpretation may arise because
most landslides on El Hierro developed into debris
avalanches, which are highly mobile and spread
fragmented debris over a wide area, while slumps
are typically more coherent, slow moving bodies
which often preserve distinct erosional and depositional morphologies (Moore et al., 1994). The overall effect of a landslide is to decrease the gradient of
the erosional ¯ank region and smooth topography
downslope.
Pro®les which best ®t the exponential form come
from the El Julan, El Golfo and San Andres/Las
Playas landslide regions (Figs. 6 and 7). Here, pro®les
have consistent low rms residual values and fall
within a narrow band of b coef®cient values, leading
to a tight clustering of landslide pro®les on the plot of
290
M.J.R. Gee et al. / Marine Geology 177 (2001) 271±293
rms against b coef®cient values (Fig. 7C). Some small
scale roughness (wavelength ,4 km) is present over
the distal part of the El Golfo debris avalanche region
due to the occurrence of numerous landslide blocks.
The smoother morphology of the El Julan landslide is
probably due to post-landslide sediment accumulation
and reworking of avalanche debris.
Pro®les from the San Andres and Las Playas landslides, have similar b coef®cients and rms residuals to
pro®les and curves from within the El Golfo and El
Julan landslide regions (Fig. 7C). These pro®les also
suggest that the ¯ank may have been affected by
major landsliding. The pro®le statistics successfully
detect landsliding even though the SE ¯ank has a
complex morphology. For example, the Las Playas
embayment associated with Las Playas landslide is
narrower and much less obvious than either the El
Julan or El Golfo embayments, while linear ridges
and pinnacles give the ¯ank a rugged appearance
which might be interpreted as unfailed ¯ank. The
narrow embayment and related fault system onshore,
(Day et al., 1997), coupled with the complex morphology offshore indicates more than one phase of landsliding. Alternatively, the more rugged, northeastern
part of the ¯ank may have experienced a partial
collapse, a view initially put forward by Day et al.
(1997) on the basis of subaerial mapping, while the
southern part experienced a full collapse, developed
downslope into a debris avalanche.
A few pro®les from within rugged ¯ank regions
have characteristics which overlap with typical landslide values. West of the island, two such pro®les
highlight a smooth section of ¯ank between two
more rugged ¯ank regions (Figs. 6 and 7). The ¯ank
in this region is predominantly rugged and steep and
there is no indication, on bathymetric or backscatter
maps, of either a failure scarp or signi®cant debris
avalanche deposits. Thus although the more subdued
morphology of this sector could result from a small
slope failure, we prefer to interpret it as an atypical
region between a bifurcating rift system extending
from onshore.
5.5. Relationship between rift zones and landslides
The subaerial morphology of El Hierro has been
interpreted in terms of three rift zones spaced 1208
apart. These are related to, and believed to control
the locations of large landslide events between the
rift arms (Carracedo, 1994; Carracedo, 1996; Carracedo et al., 1999a,b; Day et al., 1997). Due to the
mainly unbuttressed nature of El Hierro, ¯anks are
potentially unstable, allowing landslides to develop
freely. Onshore, the rifts are manifested as three
well-developed zones of dyke emplacement and
®ssuring. Offshore however, evidence for the continuation of narrow rift zones to the northwest and
northeast of El Hierro is lacking. To the south, the
Southern Ridge has the super®cial appearance of a
narrow rift zone, but other data suggests that it is
the eroded remnant of an older volcanic feature, rather
than part of the structure of the El Hierro volcanic
edi®ce (Gee et al., 2001).
Based on observations in Hawaii, Macdonald
(1949) suggested that oceanic islands develop along
rift zones rather than from central vents. In a detailed
study of Hawaiian rifts, Fiske and Jackson (1972)
showed that one, two and three armed rift zones
were common, with each island typically consisting
of several such systems. Fiske and Jackson concluded
that rift zones were made up of dyke swarms, which
exploited pre-existing structures in the edi®ce. Dykes
were fed from a magma reservoir at depth and some
produced spectacular `Hawaiian-style' ®ssure eruptions, while others solidi®ed at depth. A link between
the emplacement of dykes in rift zones and ¯ank
instability was ®rst proposed by Moore and Krivoy
(1964). Moore et al. (1989) observed that a many of
the numerous landslides on the Hawaiian Ridge had
moved perpendicular to rift zones of their host volcanoes. They also noted that it is unclear whether
injection of magma into the rift zone destabilizes the
¯ank, or whether movement of the ¯ank causes
tension which results in the intrusion of dykes into
the developing headwall region. Dyke intrusion is
suggested by Elsworth and Voight (1995) as a triggering mechanism for large earthquakes which could
initiate and sustain giant landslides. A similar argument
exists for the link between earthquakes and ¯ank failures. Moore et al. (1994) comments that it is not clear
whether earthquakes are the triggering mechanism for
failure, or the result of movement along landslide
basal surfaces.
Carracedo (1994, 1996) developed a conceptual
model linking rift zones and large landslides for the
Canary Islands, citing El Hierro as an example of an
M.J.R. Gee et al. / Marine Geology 177 (2001) 271±293
island that developed along three rift zones arranged
at 1208. This rift geometry was interpreted as the
`least effort' fracture pattern of an edi®ce by
upward-directed vertical loading due to a subsurface
magma body (Luongo et al., 1991; Carracedo, 1994).
Two additional factors might also contribute to ¯ank
collapse resulting from dyke emplacement; the severing of the updip attachment of the slide block and also
the general increase of pore pressures within the edi®ce
associated with emplacement of dyke magma heated
groundwater. Once landslides are initiated, rifts zone
ridges may play an important role in landslide transport processes by preventing the moving debris from
spreading until it reaches the lower ¯anks. This would
result in a more ef®cient delivery of the landslide
material from the ¯ank and a greater total runout.
5.6. Comparison of landslides
On the Hawaiian Ridge, mass wasting features have
been categorised as either `slumps' or `debris
avalanches' on the basis of seabed morphology and
structure and degree of dislocation of the landslide
mass (Moore et al., 1989). However, Moore et al.
(1994) noted that these were general groups, with
intermediate landslide forms occurring. For example,
many slumps were associated with debris avalanches,
which may have developed due to oversteepening of
parts of the slump. Moore et al. (1994) show a slump
southwest of Oahu (their Fig. 4) characterised by a
distinctive series of northwest trending scarps which
may represent normal faults bounding large tilted
fault blocks. The series of transverse scarps reported
here from the El Julan ¯ank may represent some type
of slumping process similar to that associated with the
southwestern ¯ank of Oahu, but on a smaller scale.
Seismic re¯ection data from El Julan indicate that
these faults affect at least the upper 100 m of the landslide deposit, although we have no data which would
allow the detection of the kind of deep-seated slump
activity reported from the Hawaiian Ridge (Lipman et
al., 1985; Moore et al., 1989, 1994).
In a TOBI 30 kHz sidescan sonar survey of the El
Golfo landslide, Masson et al. (1998) report a `series
of rotational or listric faults' which are remarkably
similar to those observed on TOBI data from the El
Julan landslide. Masson et al. (1998) interpreted these
faults as partial sediment failure resulting from the
291
loading of the slope sediments by the El Golfo debris
avalanche. It is not clear, however, whether the faulting observed within the El Julan landslide results from
a similar loading and destabilizing process, or from
post-landsliding slump activity. One possible scenario
is that emplacement of the El Julan landslide destabilised the in situ sediments in the substrate, resulting in
a series of faults which propagated through the
avalanche deposit. This might involve listric style
faults extending over 300 m through the avalanche
deposit. Alternatively, faults may have developed
only in the upper 100 m of the avalanche deposit
after emplacement. Events such as the failure of the
El Golfo ¯ank, involving some 150 km 3 of material
could have been associated with seismic activity
which may have triggered post-landslide slump
activity in other areas. A separate hypothesis is that
landslide emplacement was associated with frontal
compressional forces producing a series of folds and
thrust faults (i.e. the curving ridges and high backscatter levels). A similar morphology was described
over the lower ¯ank of a large subaerial landslide
reported from the Gobi±Altay, Mongolia (Philip and
Ritz, 1999). This hypothesis, however, would require
the `short, steeply-dipping re¯ections' associated with
the surface scarps to be thrust faults dipping towards
the island, as opposed to the rotated strata of the landslide material.
6. Summary
The main process of edi®ce destruction for El
Hierro is giant landsliding. The combined volume of
the El Golfo, El Julan, San Andres and Las Playas
landslides is approximately 420 km 3, covering an
area of 5000 km 2. An estimated 10% of the total
edi®ce volume of El Hierro has been removed by
just four giant landslides in the last 200 ka. Around
50% of the sea¯oor around El Hierro within 60 km of
the shoreline is covered by landslide deposits.
Landslides around El Hierro can be recognised on
the basis of a generally smooth downslope pro®le and
reduced slope gradients relative to unfailed ¯ank
sectors. Two new landslides have been discovered
on the southeastern ¯ank (Las Playas). The older landslide (San Andres) is interpreted as a slump event
which affected up to 1700 km 2 of ¯ank and occurred
292
M.J.R. Gee et al. / Marine Geology 177 (2001) 271±293
between 175 and 545 ka. The younger landslide (Las
Playas) is interpreted as a debris avalanche with an
area of 950 km 2 and volume ,30 km 3. The Las
Playas debris avalanche appears to result from the
partial failure of the upper ¯ank and is the smallest
event reported here. The El Julan landslide involves
an area of 1800 km 3 and a volume of 130 km 3 of
material which failed from the southwestern ¯ank El
Hierro .200 ka. There is evidence for slumping of at
least the upper 100 m of the debris avalanche.
The El Golfo landslide is the youngest landslide on
El Hierro with an area of 1500 km 2 and volume 150±
180 km 3. The primary morphology of its erosional
and depositional parts are well preserved and are
thus useful to compare to the older landslides of El
Julan and San Andres. There is no evidence for slump
activity associated with the El Golfo landslide.
Acknowledgements
We would like to thank the Captain, of®cers and
crew of the RRS Charles Darwin CD108 for their
invaluable help in acquiring the data. Figs in this
paper were prepared with the GMT software system
(Wessel and Smith, 1991). Thanks to Marco Ligi for
advice concerning seismic processing. This paper
bene®ted from constructive and careful comments
by S.J. Day and an anonymous reviewer. M.J.R.
Gee gratefully acknowledges NERC studentship
GT4/95/252 and the research facilities at the Department of Earth Sciences, Oxford University and Southampton Oceanography Centre. NCM was supported
by a Royal Society University Research Fellowship.
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