Seismic triggering of landslides and turbidity currents offshore Portugal

Transcription

1 Article Volume 12, Number December 2011 Q12011, doi: /2011gc ISSN: Seismic triggering of landslides and turbidity currents offshore Portugal D. G. Masson National Oceanography Centre, Southampton, European Way, Southampton SO14 3ZH, UK R. G. Arzola National Oceanography Centre, Southampton, European Way, Southampton SO14 3ZH, UK Atlantic Petroleum, 26/28 Hammersmith Rd., London W6 7BA, UK R. B. Wynn, J. E. Hunt, and P. P. E. Weaver National Oceanography Centre, Southampton, European Way, Southampton SO14 3ZH, UK [1] Sediments in deep water basins often include turbidites that record sediment input from adjacent continental margins. In seismically active areas, where turbidity currents are triggered by earthquakes, the basinal turbidite sequence may thus contain a record of palaeoseismicity, which can be used to infer the frequency of earthquakes affecting the margins of the basin. This is particularly useful where large earthquakes have a recurrence interval than is greater than the historical record. However, turbidity currents can be triggered by several processes, and it is often difficult to trace individual turbidites to their precise source areas and to assign a definite trigger to a particular turbidite. Here, we demonstrate that turbidites emplaced at 6600 and 8300 Cal yr BP in the Tagus Abyssal Plain, off Portugal, correlate with erosional hiatuses in two submarine canyons on the continental margin. The turbidites are sourced from simultaneous landsliding in both canyons, requiring regional triggers interpreted as earthquakes. An earthquake recurrence interval for the continental margin of 4000 years is estimated by extrapolation to deeper turbidites in the basin sequence. However, the example of the 1755 earthquake, which caused widespread devastation in southwest Iberia, shows that palaeoseismic interpretations must be made with caution. The 1755 earthquake had a magnitude >8.5 and yet the associated turbidite in the abyssal plain is typically 5 cm thick, while older turbidites can be >1 m thick. Given the large 1755 earthquake magnitude, the difference in turbidite thickness is unlikely to be related to the relative size of triggering earthquakes. Instead, we suggest that the offshore location of the 1755 earthquake, coupled with low sedimentation rates during the Holocene, may have limited the size of the associated turbidite. Components: 11,100 words, 9 figures, 1 table. Keywords: palaeoseismicity; submarine canyon; turbidites. Index Terms: 3002 Marine Geology and : Continental shelf and slope processes (4219); 3022 Marine Geology and : Marine sediments: processes and transport; 3045 Marine Geology and : Seafloor morphology, geology, and geophysics. Received 17 August 2011; Revised 25 October 2011; Accepted 25 October 2011; Published 14 December Masson, D. G., R. G. Arzola, R. B. Wynn, J. E. Hunt, and P. P. E. Weaver (2011), Seismic triggering of landslides and turbidity currents offshore Portugal, Geochem. Geophys. Geosyst., 12, Q12011, doi: /2011gc Copyright 2011 by the American Geophysical Union 1 of 19

2 1. Introduction [2] Deep-water sedimentary basins contain records of sediment input from adjacent continental margins. Large-scale changes in sediment input to a margin, for example due to climate and associated sea level change, and catastrophic events such as continental slope landslides, are among the events likely to be recorded in the adjacent basin. A major aim of studies of such basins is to decipher what these records mean in terms of margin evolution. The value of basin sedimentary records has been proven in areas such as the Moroccan turbidite system where a dateable record of volcaniclastic turbidites related to Canary Island flank collapses, which are difficult to date in proximal settings near the islands, is preserved in the Agadir and Madeira Abyssal Plain basins [Weaver and Kuijpers, 1983; Weaver et al., 1992; Masson et al., 2002; Wynn et al., 2002] Turbidite Palaeoseismology [3] The concept of turbidite palaeoseismology was originally developed for the Cascadia continental margin off the western United States [Adams, 1990; Goldfinger et al., 2003]. The basic principle of this concept is that turbidity currents that are generated simultaneously in different areas along the margin, and flow through separate channels to amalgamate on the basin floor, cannot be generated by purely sedimentological processes, but must have an external trigger. In seismically active areas, this trigger is likely to be a large regional earthquake, as demonstrated elsewhere by the sequential downslope failure of submarine cables impacted by turbidity currents generated by historical earthquakes [e.g., Heezen and Ewing, 1952; Hsu et al., 2008]. [4] The field of submarine palaeoseismicity has expanded rapidly over the last 20 years, with published studies from a wide variety of locations around the world (see Goldfinger [2011] for a complete list of locations). The major challenge facing this discipline, especially for prehistoric turbidites where there is no independent record of an earthquake, is finding criteria that can be used to separate turbidites generated by earthquakes from those generated by other triggers that include slope loading by sedimentation, flank collapse of island volcanoes, storm waves, tilting due to margin subsidence or salt movement, changes in methane hydrate stability or hyperpycnal flows [Masson et al., 2010a; Goldfinger, 2011]. Thus most studies have been undertaken on recent sedimentary sequences, where turbidites can be positively correlated with their triggering earthquakes. For example, multiple coarse fraction pulses, with or without contrasts in mineralogy, have been identified in turbidites that were generated by historical earthquakes [Nakajima and Kanai, 2000; Goldfinger et al., 2008]. This has been interpreted as a signal of multiple turbidity currents from different sources triggered simultaneously by a large earthquake affecting a relatively broad area. However, turbidites with multiple coarse bases are also associated with volcanic island flank collapse [Wynn and Masson, 2003]. Although the mineralogy of the multiple volcanic bases is very similar, in some cases they can be distinguished using geochemical techniques, because the individual flows, although originating from the same source, eroded geochemically distinct areas of the volcanic edifice [Hunt et al., 2011]. This could give a false impression of multiple sources. The converse situation, where multiple earthquake-triggered failures give rise to a single fining-upward turbidite in a distal basin is also possible, especially where the initial failures occur over a relatively limited area (e.g., 1929 Grand Banks turbidite [Piper et al., 1999]). In addition, where at turbidite is generated from a single landslide, there is no obvious reason why the triggering mechanism would have any influence on the resulting deposit, typically a single finingupward Bouma sequence turbidite. A good knowledge of the geology of the source area is thus a prerequisite for concluding that the turbidite had a seismic trigger. [5] Gorsline et al. [2000] showed that earthquake generated turbidites offshore southern California could be separated from storm generated tubidites on the basis of their greater volume and areal extent. However, it is not clear to what extent these criteria are dependent on the local sedimentary environment (e.g., sediment input rates and processes, slope stability factors) from which the turbidites originated. It seems unlikely that these distinguishing criteria can be extrapolated with confidence to all ancient basins, as suggested on the basis of the California study [Gorsline et al., 2000]. Complex crossbedding and an increase in terrigenous sediment content relative to other turbidites in the same basin have been attributed to the earthquake triggering of a turbidite associated with the January 2010 Haiti earthquake [McHugh et al., 2011]. However, as with the California example, it is difficult to understand which of the characteristics of this turbidite are a direct consequence of seismic triggering, and which are related to its depositional environment, in this case in a small enclosed basin. Complex cross 2of19

3 bedding is common in turbidites generated by a variety of processes in many different sedimentological environments; not all can be related to seismic triggering (e.g., cross-bedding in volcaniclastic turbidites in the Madeira Abyssal Plain [Rothwell et al., 1992]. [6] Goldfinger [2011] argues that virtually all palaeoseismic studies invoke the synchronous triggering test, even when additional sedimentological arguments for seismic triggering are proposed [e.g., Nakajima and Kanai, 2000; Goldfinger et al., 2008]. This appears to be a reasonable appraisal of submarine palaeoseismic studies at the present time. Studies based on sedimentological criteria [e.g., Gorsline et al., 2000; McHugh et al., 2011] may be applicable in a local, wellunderstood context, but are difficult to extrapolate for use in global recognition of seismically generated turbidites Turbidites of the Tagus Abyssal Plain and Iberian Continental Margin [7] The most recent turbidite in the distal Tagus Abyssal Plain (TAP), off Portugal, has been correlated with the 1755 earthquake that devastated much of SW Iberia [Thomson and Weaver, 1994]. This suggests that the sequence of thick turbidites in the TAP [Lebreiro, 1995] might contain a prehistoric record of seismicity on the Portuguese margin [Vizcaino et al., 2006; Gràcia et al., 2010]. Furthermore, Gràcia et al. [2010] have shown that all the major earthquakes affecting SW Iberia in the last 2000 years left their mark in the marine record, albeit with variable expression in terms of magnitude and extent. A better understanding of the way in which turbidites in the marine record reflect the event that triggered them would thus be a major step forward in understanding the geohazard threat posed by earthquakes along that margin. However, interpreting the turbidite record has been difficult because to date it has only rarely been possible to correlate individual turbidites in the basin with their source areas along the margin [e.g., Goldfinger et al., 2008; McHugh et al., 2011], and without knowledge of the source it is hard to distinguish between different possible turbidity current triggers. For example, turbidity currents could be derived from landslides resulting from high sediment input to the margin, with or without a seismic trigger. In this paper, we demonstrate that turbidites deposited in the TAP at 6600 Cal yr BP and 8300 Cal yr BP have the same age as erosional events in the Setubal and Cascais Canyons offshore Portugal. The widespread extent of this erosion strongly suggests that the erosional events had external triggers of regional extent, most likely significant earthquakes in the vicinity of the Portuguese margin near 38 N Study Area [8] The area discussed in this paper lies immediately to the north of the Africa-Eurasia plate boundary southwest of Iberia (Figure 1). The exact location and structure of the plate boundary in this area is the subject of ongoing discussion [e.g., Zitellini et al., 2009, and references therein]. It appears to be represented by a broad and complex deformation zone dominated by NE trending thrust faults and NW trending strike-slip faults; an accretionary wedge has also been identified under the Gulf of Cadiz (Figure 1) [Gutscher, 2004; Zitellini et al., 2009]. The epicenter of the 1755 earthquake is located within this complex deformation zone and modeling of the arrival times of the associated tsunami at various places around the N. Atlantic gives a best estimate of 200 km off Cape St Vicente [Baptista et al., 1998b]. The fault that generated the tsunami has not been positively identified, although the Horseshoe Fault has been suggested as the most likely candidate [Stich et al., 2007]. [9] Setubal and Cascais Canyons cut the Portuguese continental margin between 38 and 39 N (Figure 1). Cascais Canyon cuts the steepest part of the margin. Consequently it is relatively short and has steep axial gradients that typically range from 5 to 20 [Lastras et al., 2009]. It has a v-shaped cross section for most of its length but broadens and becomes flat floored beyond 4500 m before merging with the TAP between 4700 and 4800 m. [10] Setubal Canyon, and its tributary Lisbon Canyon, deeply incise the continental shelf in the vicinity of the Sado and Tagus Rivers respectively. Setubal Canyon is twice the length of Cascais Canyon but crosses the same depth range, hence axial gradients are much lower, rarely exceeding 5 [Lastras et al., 2009]. Although some sections of the upper and middle canyon (as defined by Lastras et al. [2009]) are v-shaped, much of the canyon has a distinct flat-floored axial channel ranging from a few hundred to a thousand meters in width. Below 4200 m depth, the canyon widens abruptly into a broad flat-floored channel that gradually increases in width from 2 to 5 km over a distance of about 25 km, where the canyon merges with the TAP between 4800 and 4900 m. 3of19

4 Geosystems 3 G MASSON ET AL.: SEISMIC TRIGGERING OF TURBIDITY CURRENTS /2011GC Figure 1. Location of Setubal and Cascais Canyons and the Tagus Abyssal Plain, offshore southern Portugal. Regional bathymetry (pale colors and 100 m contour interval) from GEBCO. More detailed bathymetry of canyons (dark colors) based on EM12 multibeam data. Labeled cores are those described in detail in this paper, except for MD that is described by Gràcia et al. [2010] and cores and , described by de Stigter et al. [2011]. Other coring sites used to delimit the sandy canyon mouth lobe are also shown; note the failed coring sites (red symbols) that prevented the acquisition of a complete core transect between the canyons and the abyssal plain. Inset top left shows location of study area, inset bottom right summarizes the tectonic setting; in both cases the rectangle shows the area of the main figure. Locations of major thrusts (blue), strike-slip faults (red) and the limits of the accretionary prism in the Gulf of Cadiz (green) are from Zitellini et al. [2009]. Locations of historical earthquakes with magnitudes >6.0 (black dots with dates) are from Gutscher [2004]. The likely location of the 1755 earthquake is after Baptista et al. [1998b]. [11] The Tagus Abyssal Plain (sensu lato) is defined as the broad basin that extends west from the base of the Portuguese slope between 37 and 38.5 N (Figure 1). The northeastern part of this area slopes gently toward the southwest with an average gradient of 0.1 ; this area is underlain by a sandy depositional fan related to the two canyons [Lebreiro, 1995] (Figure 1). The deepest and flattest part of the TAP lies to the southwest, with an area of 18,000 km2 enclosed by the 5100 m contour. 2. Materials and Methods 2.1. Sediment Cores [12] Piston cores used in this study were collected during RRS Discovery Cruise 187, RRS Charles Darwin Cruises 157 and 179 and RRS James Cook Cruise 27 (Figure 1 and Table 1). In the canyons, a large number of cores were collected in order to gain an understanding of the heterogeneous sedimentary environment. Cores were collected mainly from terraces adjacent to but elevated above the canyon axis. These terraces typically preserve a record of predominantly muddy sediments that is relatively easy to core, more or less continuous, and contains dateable material (planktonic foraminifera). In contrast, the canyon axis is typified by sand, gravel and landslide deposits that are difficult to sample and contain little in situ dateable material [Arzola et al., 2008]. Core sites in the canyons were chosen on the basis of TOBI (Towed Ocean Bottom Instrument) 30 khz deep-towed side-scan sonar data, which cover all of the canyon floor and most of its walls (Figure 2) [Arzola et al., 2008]. Only relatively wide and flat terraces were chosen for coring, because of possible errors of up to 200 m in absolute positioning of the side-scan data coupled with possible offset of the corer on the seabed 4 of 19

5 Table 1. Core Locations a Core Location Latitude Longitude Water Depth (m) Canyon Axis Depth (m) Height Above Thalweg (m) CD56820 Cascais Canyon N W CD56837 Cascais Canyon N W CD56822 Setubal Canyon N W CD56825 Setubal Canyon N W <10 CD56843 Setubal Canyon N W CD56414 Setubal Canyon N W D11931 Tagus AP N W 5065 D11951 Tagus AP N W 5080 a For cores taken on terraces in the canyons, water depth at the core site and in the adjacent canyon axis is estimated from multibeam bathymetry data. Note that the multibeam data may underestimate the canyon axis depth in areas where the axis is narrow and the topography steep. relative to the ship. This ensured successful targeting of the intended terraces. Terraces between 10 and 170 m above the apparent canyon floor were sampled, although studies using a Remotely Operated Vehicle (ROV) suggest that the canyon axis depth is often underestimated in v-shaped canyon sections when measured by shipborne multibeam systems [Masson et al., 2011]. Figure 2. TOBI side-scan sonar images and superimposed bathymetry data (100 m contours) showing the locations of the canyon cores discussed in this paper. High backscatter (light) on the side-scan images generally equates to steep slopes, or hard substrates or coarse sediments (gravel) at outcrop. Low backscatter (darker gray) represents smooth sediment covered areas. All cores, with the exception of CD56414, are located on terraces elevated above the canyon axis, in sediment covered areas. Core CD56414 sampled a landslide deposit in the canyon axis. 5of19

6 Figure 3. Examples of core photographs. In all cases the top of the core is top left and the sections follow in sequence down core. All are 10 cm diameter cores except D11951, which has a 7 cm diameter. In general, hemipelagic sediments are paler (typically pinkish to pale olive green) and turbidites darker (dark olive green to gray). Note that colors are not exactly comparable between cores because photographs were taken at different times in variable lighting conditions immediately after opening the cores on board ship, to record the colors before they faded on exposure to the air. Stars show location of dated samples. Core CD56837 (Cascais Canyon) and CD56822 (Setubal Canyon) show pale hemipelagic sediments with rare interbedded turbidites (upper sequence) overlying thin bedded dark gray turbidites and paler gray hemipelagic sediments (lower sequence). Note the thin clast rich debris deposit (section 1, cm) that occurs at the sequence boundary in core CD Core CD56825, consisting of pale hemipelagic sediments interbedded with darker turbidites, sampled the upper sequence in Setubal Canyon; gravel size clasts were recovered from the core catcher in this core, indicating that it reached the upper/lower sequence boundary. Core CD56414 contains pale hemipelagic sediment overlying landslide debris consisting of contorted thin-bedded turbidites derived from the levee. Core D11951 is from the Tagus Abyssal Plain. Each dark bed is a single thick muddy turbidite. Paler beds are hemipelagic sediments. The trigger core shows how the top of the main piston core has been compressed by the coring process. [13] In the TAP, cores were obtained only from the southern and western parts of the basin. Low acoustic penetration on 3.5 khz profiles and a series of failed coring attempts suggest that the northeastern part of the plain is underlain by sandy sediments that cannot be sampled by conventional piston coring techniques (Figure 1) [Lebreiro, 1995]. [14] Cored sediments were initially separated into turbidite and hemipelagic facies on the basis of visual logging using color, grain size, sedimentary structures and the presence or absence of planktonic foraminifera in muddy units (Figure 3). Turbidites are recognized by sharp changes in color and grainsize at their base, the presence of well sorted sand or silt bases often showing normal grading, and an overlying fine-grained (mud) cap lacking in foraminifera. Hemipelagic sediment is typically fine-grained but poorly sorted, with planktonic foraminifera, usually clearly visible on the cleaned core surface, scattered randomly through each interval. Turbidite mud typically contains a higher percentage of terrigenous sediment than the more carbonate rich hemipelagic mud and is thus usually darker in color (Figure 3). Landslide deposits are characterized by contorted bedding and chaotic facies assemblages Mineralogy [15] Mineralogical analysis was performed on selected turbidites from the two TAP cores as an aid 6of19

7 Figure 4. Visual logs, 14 C dates and multisensor logs (MSCL) for cores in Setubal Canyon. Cores were collected from terraces between 10 and 145 m above the canyon axis. Error bars on 14 C dates are typically in the order of 100 yr (see Table S2). Correlation between cores is based on all available data. The upper/lower sequence boundary is typically marked by a distinct change in both lithology (see Figure 3) and MSCL values. An increase in density is seen at the boundary in most cores and the magnetic susceptibility is relatively smooth above the boundary and much more spiky below it. In two cores where dates bracket this boundary, it is clearly an erosional hiatus. in correlation between cores. Samples were dried and the mm fraction separated by sieving. All analyses were conducted on this size fraction in an attempt to minimize differences between turbidite sands introduced by hydrodynamic sorting of minerals during transport. A total of grains (dependent on the available material in the size fraction) were analyzed for each sample and the relative abundances of quartz, lithic fragments, biotite, muscovite and opaque minerals calculated (Table S1 in the auxiliary material) MultiSensor Core Logging [16] For the canyon cores, visual facies identification was supported with physical properties measurements made with a Geotek MultiSensor Core Logger (MSCL). Magnetic susceptibility, gamma ray density and p-wave velocity data were obtained at 1 cm intervals (Figures 4 and 5). Magnetic susceptibility and p-wave velocity data for the older TAP cores were taken from Lebreiro [1995] (Figure 6). Magnetic susceptibility was measured using a Bartington Instruments MS 2 magnetic susceptibility meter at 2 cm intervals down core; p-wave velocity with a Geotek MSCL at 2 cm 1 Auxiliary materials are available in the HTML. doi: / 2011GC intervals. In general terms, turbidite sands and silts correspond to peaks in density and p-wave velocity. The magnetic susceptibility signal is more variable, possibly dependent on the maturity of the sediment. The coarsest beds (typically well-sorted medium sand) are predominantly composed of quartz and typically have low susceptibility (Figure 7). Higher susceptibility occurs in some, but not all, silts and fine sands, which tend to be richer in micas and opaque mineral, although there is no obvious correlation between susceptibility and the proportion of such minerals (Figure 7) Radiocarbon Dating [17] The strategy for radiocarbon dating was first to target sedimentary deposits that could potentially be correlated across more than one core (e.g., individual turbidites or major changes in sediment type; Figures 4 6). On the basis of visual and MSCL evidence, this was possible within both the canyon area and within the TAP; however there are no marker horizons that obviously correlate between the two areas. The radiocarbon dating was then used to establish a temporal correlation between them. Only hemipelagic sediment with sufficient foraminifera was sampled, avoiding landslide deposits and turbidite mud units. Samples were collected, where possible, directly below redeposited 7of19

8 Figure 5. Visual logs, 14 C dates and multisensor logs (MSCL) for cores in Cascais Canyon. Cores were collected from terraces 150 and 170 m above the canyon axis. Error bars on 14 C dates are typically in the order of 100 yr (see Table S2). The only correlatable horizon is the upper/lower sequence boundary, identified as described for Figure 4. Turbidites are generally absent from the upper sequence except for several thin muddy laminae near the top of the core. units, taking care to avoid sampling the redeposited units themselves. Most of the turbidites are fine grained, with silt or fine sand at the base, so it is assumed that there was minimal erosion associated with their emplacement [Thomson and Weaver, 1994]. This assumption is supported by a study of coccolith assemblages found in turbidites that were emplaced on the Madeira Abyssal Plain by turbidity currents from the NW African margin [Weaver, 1994]. The main conclusion of that paper was that these turbidity currents did not contain an excess of sediment from the contemporary seafloor, and thus that once formed were virtually non-erosional over travel distances of hundreds of kilometers. The possible exception to this is core D11931, which contains several thick turbidite sands (Figure 6). However, the consistency of dates between cores D11931 and would seem to indicate that even in this case, any erosion due to turbidite emplacement was minimal. [18] Approximately 1 4 cm 3 of sediment was sampled per horizon and washed through a 63 mm sieve mg of mixed planktonic foraminifera (1000 individuals) were picked under a binocular microscope. The main planktonic species found were Globigerina bulloides, Globorotalia scitula, Neogloboquadrina pachyderma, Globorotalia truncatulinoides, Globigerinoides ruber, Globorotalia inflata, Globoritalia hirsuta and Orbulina universa. Picking mixed assemblages assured that enough specimens were available in all samples selected for dating. The picked samples were dated by the AMS (Accelerator Mass Spectrometry) method at the NERC Radiocarbon Laboratory in Scotland and at Beta Analytic in Florida (Table S2). Conventional radiocarbon ages were converted to calibrated radiocarbon ages using the MARINE04 database and IntCal04 calibration curve [Hughen et al., 2004]. This calibration uses a 400-year reservoir age for a deep-sea setting, which is valid up to 21,786 Cal yr BP. For older ages, the Fairbanks0107 calibration curve with a 255-year reservoir age was used instead [Fairbanks et al., 2005]. 3. Results and Interpretation 3.1. Canyon Sediment Cores [19] Cores collected from canyon terraces all contain an upper pale-colored, 1 3 m thick, finegrained sequence composed mainly of hemipelagic 8of19

9 Figure 6. Visual logs, 14 C dates and multisensor logs (MSCL) for cores in the Tagus Abyssal Plain. MSCL data are redrawn from Lebreiro [1995]. Error bars on 14 C dates are typically in the order of 100 yr (see Table S2). The top of core D11951 was compressed by the coring process and has been reconstructed using information from the trigger core (see Figure 3); the MSCL logs for this core are based only on the main core. Correlation between the predominantly sandy turbidites of core D11931 and the finer grained turbidites of core D11951 is based largely on the 14 C dates. However, this is supported by the magnetic susceptibility record and mineralogical data (Figure 7). Turbidites younger than about 18,000 Cal yr BP have no susceptibility signal, while the coarse bases of older turbidites each correspond to a distinct susceptibility peak. sediment but interspersed with a few muddy turbidites, overlying a lower sequence composed of thin bedded turbidites, mass wasting deposits or sand and gravel (Figure 3) [Arzola et al., 2008]. The upper sequence is between 0.8 and 3.3 m thick. A reddish-brown layer a few centimeters thick at the top of most cores shows that the oxidized seabed layer had been recovered. In cores missing this layer (e.g., CD56414; Figure 3), the sediment surface was probably lost during coring. [20] In the upper sequence, MSCL data have limited usefulness in identifying turbidites and in correlation between cores. Some turbidites show gradual density increases toward the base (e.g., core CD56825; Figure 4), but in most cases the basal silt/sand laminae are too thin to be detected by this method. Similarly, a few turbidites correspond to magnetic susceptibility peaks, but an equal number show no change compared to the background hemipelagic sediment. One potential problem in identifying muddy turbidites in the upper sequence is that some of these beds are so thin that they may have been destroyed by bioturbation. Magnetic susceptibility data from cores CD56822 and CD56825 provide some evidence for this. Peaks in magnetic susceptibility at 2 m in core CD56822 and 1.5 m in core CD56825 appear to correlate on the basis of position in the cores and radiocarbon dates (Figure 4). In the latter, a double peak clearly correlates with two turbidites. In the former, a more diffuse peak corresponds only to a section of greyer sediment in the core, with no individual turbidites identified. We interpret this as indicating bioturbation of thin muddy turbidites in Core CD [21] The lower sequence consists of several distinct facies: interbedded turbidites and hemipelagic sediments on terraces (e.g., core CD56822, Figures 3 and 4), and mass wasting deposits (e.g., core CD56414) or sand/gravel in the canyon axis [Arzola et al., 2008]. The terrace turbidite sequences consist of closely spaced, thin-bedded sands/silts interbedded with fine-grained sediment that appears to be a mixture of hemipelagic material and turbidite mud. Similar fine-grained sediments in nearby Nazaré 9of19

10 Figure 7. Comparison of mineralogical data and magnetic susceptibility profiles from correlated turbidites in cores D11951 and D11931 from the Tagus Abyssal Plain. Mineralogical data were taken from sieved mm sediment fractions in all cases. Crosses on core logs mark sample locations. In core D11931, quartz dominates in basal sand layers. However, the relative proportions of biotite, muscovite and opaque minerals in the turbidite silts vary between turbidites and correlate well between the two cores. Canyon have been interpreted as the product of quasi-continuous sedimentation from nepheloid layers mixed with background hemipelagic deposition [Masson et al., 2011]. Planktonic foraminifera up to 250 mm in size, too large to have been transported by gravity currents with the fine-grained sediment, are distributed through the fine-grained sediment and are the primary evidence for the hemipelagic input and thus the quasi-continuous nature of the nepheloid layer deposition. The mass wasting deposits consist of contorted clasts of thin bedded turbidites reworked from the canyon slopes mixed with sand and gravel from the canyon floor (Figures 3 and 4) [Arzola et al., 2008]. In one core, remobilized sediments occur in a distinct layer between the upper and lower sequences (Core CD56837; Figures 3 and 5). Sand and gravel occur widely, beneath a thin veneer of upper sequence sediment, on the canyon floor. They have been described in detail by Arzola et al. [2008] and are not further discussed in this paper. 10 of 19

11 [22] MSCL data show the lower sequence to have a markedly different character from the upper sequence with the boundary between the two sequences marked by sharp changes in parameter values and profile character (Figures 4 and 5). Magnetic susceptibility profiles from the lower sequence of thin-bedded turbidites have a distinctive spiky character, with thin coarse beds corresponding to peaks in susceptibility. These sediments also have a much higher density than the upper sequence, with an abrupt density offset of about 0.3 g cm 3 at the sequence boundary. Gravel and remobilized turbidite clasts have the highest densities observed in this study (Core CD56414; Figure 4) Core Correlation, Radiocarbon Dating and Sedimentation Rates in Canyon Cores [23] Only broad scale correlation between the canyon cores is possible on the basis of visual logging and MSCL data. A distinct boundary between the upper, mainly hemipelagic sequence and the lower, mainly turbidite and debris flow sequence is seen in all cores and in the MSCL data (Figures 3 and 4). This abrupt and consistent change would therefore appear to be correlatable both within each canyon and between canyons. Above this boundary, the cores contain variable numbers and distributions of turbidites, making correlation of individual turbidites difficult. All the cores sampled several thin turbidites just below the seabed, and while this turbidite sequence is clearly correlatable, our data is not sufficient to correlate the individual turbidites. The sequence is not identical in all cores, although differences between cores may, at least in part, be due to compression or partial loss of the core top during recovery, or to natural variability in the preservation of the turbidite record in cores in differing sedimentary settings, e.g., different heights of terraces above the canyon axis. [24] Radiocarbon dates were obtained from six canyon cores to confirm the validity of the correlation of the upper/lower sequence boundary and to investigate the possibility of correlating individual turbidites within the upper sequence (Figures 4 and 5 and Table S2). Sediments from just above the upper/lower sequence boundary give remarkably consistent dates of 6500 Cal yr BP in three of the dated cores, two from Setubal Canyon (CD58622 and CD56843) and one from Cascais Canyon (CD56837); an extrapolated date for a fourth core (CD56414) is broadly compatible with this (Figures 4 and 5). The other two cores, one from each canyon, also give broadly consistent, but older ages of 8100 Cal yr BP for the upper/lower sequence boundary (Cores CD56825 and CD56820). This sequence boundary is therefore not a simple synchronous horizon. [25] Two samples of fine-grained sediments (hemipelagic mud) from immediately below the upper/ lower sequence boundary were also dated using the radiocarbon method. For these samples, only foraminifera retained on a 180 mm sieve were included in the dated material, to try to ensure that all came from the pelagic component of the sediment, on the assumption that any sediment reworked from further up canyon and transported by nepheloid layers or weak turbidity currents would be exclusively fine grained. This approach has been tested in similar sediments from Nazaré Canyon and found to give reliable results [Masson et al., 2010b, 2011]. The two samples yielded dates of 16,550 and 20,000 Cal yr BP, significantly older than the dates above the sequence boundary, indicating that the sequence boundary is a hiatus or erosion surface (Table S2 and Figures 4 and 5). Its interpretation as an erosion surface is supported by the abrupt increase in density across the boundary seen in several cores (e.g., Figure 4). [26] Radiocarbon dates from within the upper sequence give some insight into sediment deposition and turbidite input during the last 6500 years. The two cores with the thickest hemipelagic sediments (CD56822 and 56825) both show essentially constant hemipelagic sedimentation rates through this sequence (Figure 8). All cores suggest enhanced input of (relatively) small-scale turbidites during the last 1000 years. Three cores in Setubal Canyon also contain a cluster of turbidites emplaced during a period of 2000 years between about 3500 and 5500 Cal yr BP (CD56822, CD56825 and CD56414); these are not seen in the remaining Setubal core (CD56843) or in either of the Cascais Canyon cores. Finally, a single turbidite within a thick hemipelagic interval and with an age of 2500 Cal yr BP is seen in adjacent Setubal Canyon cores CD56822 and CD56825, but is apparently absent in all other cores Tagus Abyssal Plain Cores [27] Sediment cores from the TAP sampled a sequence of relatively thick turbidites interbedded within thinner hemipelagic sediments (Figures 3 and 6) [Lebreiro, 1995]. Core D11931, located further north in the basin and some 15 m shallower than core D11951, appears to be more influenced by the input of sandy sediments from the canyons to the 11 of 19

12 Figure 8. Hemipelagic sedimentation rates for selected Setubal Canyon and TAP cores. Individual date markers include error bars. The vertical axis is the cumulative thickness of hemipelagic sediment down core, omitting turbidites. The canyon cores show that pelagic sedimentation rates have been uniform through the last 8000 years. Core D11931 from the Tagus Abyssal Plain shows a distinct change in sedimentation rate between 18,000 and 14,000 Cal yr BP. This change could also be interpreted for core D11951, but cannot be confirmed due to insufficient data. east and contains a mixture of sandy and muddy turbidites (Figure 6). In contrast, core D11951 contains a typical distal abyssal plain sequence of thick muddy turbidites, as seen in other abyssal plain basins [e.g., Weaver et al., 1992]. Individual turbidites are up to a meter or more in thickness. [28] MSCL data from the TAP cores show strong velocity peaks associated with turbidite sands and weaker peaks with some turbidite silts, although some thinner turbidites are not resolved (Figure 6). The quartz-rich sands (Figure 7 and Table S1) typically have a low magnetic susceptibility signal, indistinguishable from hemipelagic sediments. Peaks in magnetic susceptibility values are associated with some turbidite silts, but only in the lower parts of the cores, below 2 m in core D11951 and 2.5 m in core D Hemipelagic sediments exhibit uniform low p-wave and magnetic susceptibility values Core Correlation, Radiocarbon Dating and Sedimentation Rates in TAP Cores [29] Correlation between the abyssal plain cores on the basis of visual and MSCL logs is ambiguous, primarily because core D11931 sampled a more proximal turbidite sequence of interbedded sand and mud turbidites, while D11951 sampled a more distal sequence dominated by thick mud turbidites (Figure 6). Correlation between the two TAP cores therefore relies heavily on radiocarbon dating. When these are included in the analysis, it is clear that most sandy turbidites in core D11931 correlate with thick mud turbidites in core D11951, and that most of the thin mud turbidites in the more proximal D11931 appear to be absent from the distal core. Dates are very consistent between cores and show that large turbidites deposited over the whole basin were input immediately before 6700, 8400, 10,500, 19,900 and around 23,000 Cal yr BP. [30] Once established on the basis of radiocarbon dates, it can be seen that the correlation is supported by the magnetic susceptibility and mineralogical data (Figures 6 and 7). Turbidite silts older that 18,000 Cal yr BP are associated with magnetic susceptibility peaks, while younger turbidites are not. Variation in mineralogy between turbidites is limited, but in detail also supports the proposed correlation; e.g., the relative proportions of biotite, muscovite and opaque minerals in turbidite silts varies between different turbidites in each core but is relatively constant in turbidites correlated between cores (Figure 7). This probably reflects a greater input of relatively immature sediment during the sealevel lowstand associated with the last glacial period. This is also seen closer to Setubal Canyon where sedimentation rates on the northern canyon levee decreased by a factor of four after 15,000 Cal yr BP [Lebreiro et al., 2009]. It is also reflected in the hemipelagic sedimentation rate in the TAP, which decreased by a factor of two at the same time (Figure 8). The rate of turbidite input to the TAP also decreased at 15,000 Cal yr BP, at least in the more proximal record of core D11931, 12 of 19

13 where it decreased from one turbidite every 1000 years to one every 4000 years (Figure 6). This change is not seen at the more distal core D Correlation Between Canyon and TAP Cores [31] Correlation between the canyons and TAP cores is based entirely on radiocarbon dates, and is related to the timing of events rather than direct correlation of deposits. Two possible correlative events are recognized. We propose that the youngest turbidite in the TAP, dated as younger than 6700 Cal yr BP, correlates with the erosional event associated with the upper/lower sequence in the canyon cores where this is dated as older than 6500 Cal yr BP. The difference between the 6500 and 6700 Cal yr BP dates above the erosion surface and below the turbidite respectively is to be expected as a function of the sampling procedure. First, the 1 3 cm thick dated sample intervals represent the average value for up to a few hundred years of sedimentation, particularly in the slowly deposited TAP sediments. Second, in order to eliminate possible contamination, the dated samples were taken at least a few millimeters away from sedimentary unit boundaries, which will lead to slightly old dates below the event and slightly young dates above. The second youngest turbidite in the TAP sequence, dated as younger than 8500 Cal yr BP, may also correlate with the upper/lower sequence boundary in the canyon cores where this boundary is dated as older than 8100 Cal yr BP. This correlation is more speculative than the correlation of the younger event, because it was only observed in two canyon cores (CD56820 and CD56825, Figures 4 and 5). We have not obtained dates from below the sequence boundary in these two cores, although one (CD56825, Setubal Canyon) terminated in a lag or debris flow deposit composed of sedimentary clasts suggestive of an erosional event. However, in the single core that penetrated the boundary, the physical properties data do not provide convincing evidence for a hiatus (Core CD56820, Cascais Canyon; Figure 5). 4. Discussion 4.1. Sedimentary Processes in Setubal and Cascais Canyons [32] The upper and lower sequences in cores from Setubal and Cascais Canyons clearly record different sedimentary environments. The upper sequence of mainly hemipelagic sediments records a period when both canyons were largely inactive as sediment conduits across the continental slope. Hemipelagic sedimentation has dominated since at least 8100 Cal yr BP, although interrupted by an erosional event seen in both canyons at 6600 Cal yr BP. With the exception of several small turbidites in the last 1000 years, cores from Cascais Canyon do not record any post 6600 Cal yr BP turbidites (Figure 5). In contrast, several turbidites are recorded in cores from the upper sequence in Setubal Canyon (Figure 4). Most of these appear to be local events that can only be correlated between adjacent cores (e.g., CD56822 and 56825), although it cannot be ruled out that they are relatively small events that traveled through the entire canyon, but were only deposited on certain terraces. The latter interpretation is supported by core CD56825, which is located almost in the canyon axis (Table 1) and records the largest number of turbidites (Figure 4). [33] A resurgence of turbidity current activity appears to have occurred in both canyons at around 1000 Cal yr BP (Figures 4 and 5). However, most of these turbidites are represented only by thin muddy lamina and different numbers of turbidites are recorded in different cores. Detailed correlation of individual turbidites within this period is not possible, because of the variable turbidite stratigraphy, low sedimentation rate, limited thickness of hemipelagic sediment between turbidites, and variable amounts of distortion (usually compression) in the upper cm of the cores. We note that the low hemipelagic sedimentation rates in all cores in this study prevents the application of 210 Pb dating techniques, which have been used successfully in areas of the canyons where higher sedimentation rates are found [Arzola et al., 2008; de Stigter et al., 2011]. One of these turbidites probably correlates with the 1755 Lisbon earthquake, but this cannot be unequivocally established from our data. However, a 1755 turbidite is recognized in the upper Lisbon Canyon where sedimentation rates are much higher (Figure 1) [de Stigter et al., 2011] and sediment erosion due to the passage of the tsunami associated with the 1755 earthquake is observed on the shelf [Abrantes et al., 2008]. However, in the context of this study, the lack of evidence for a large 1755 turbidite is perhaps the most important observation (see later discussion). [34] The dating of the lower sequence of interbedded turbidites and hemipelagic sediments indicates that these were deposited during a time of lowered sealevel during the last glacial period (Figures 4 and 5). Deposition continued until at least Cal yr BP, although the top of this sequence appears always to be truncated by erosion, and thus 13 of 19

14 Figure 9. Evidence for widespread landslides on the flanks of the levee between the lower Setubal and Cascais Canyons. (a) Shaded relief swath bathymetry map showing topographic depressions interpreted as landslide scars. Note that several cores from the lower flank of the levee in Setubal Canyon recovered landslide debris. The irregular green line marks the northern edge of the flat floor of Setubal Canyon. (b) Side-scan sonar image of part of the levee flank confirming the landslide scar interpretation. Core CD65414 recovered landslide debris (see Figure 3). (c) Profile across the levee showing oversteepened areas of the levee flank interpreted as landslide headwalls. the transition from sedimentation dominated by turbidites to hemipelagic sedimentation in the canyon has not been observed. Regional considerations, however, suggest that the transition occurred 15,000 Cal yr BP. This is based on the observation that turbidite deposition on the Setubal Canyon levee stopped abruptly at this time [Lebreiro et al., 2009]. It may also be reflected in the hemipelagic sedimentation rate in core D11951, which decreased by a factor of two between 18,000 and 14,000 Cal yr BP (Figure 8) Potential Cause of the 6600 and 8300 Cal yr BP Erosional Events [35] The 6600 Cal yr BP erosional event is seen in most, but not all, cores. When combined with the evidence for extensive landslide deposits in the lower Setubal Canyon and landslide scars on the adjacent levee, the most likely explanation for the erosion appears to be widespread synchronous landsliding (Figures 4, 5, and 9) [Arzola et al., 2008]. The synchronous nature of landslides is not proven directly, but inferred from the apparently similar age of landslide deposits and erosion seen in several cores. In this interpretation, the absence of the 6600 Cal yr BP event at some core sites (e.g., CD56820; Figure 5) simply means that these sites were not affected by landslides at that time. This is not unexpected, since there is no reason to expect that landslides would occur everywhere on the margin. The apparent absence of the 6600 Cal yr BP event in core CD56825, in the axis of Setubal Canyon, is more difficult to explain, because the event is present in core CD56822 further up canyon, and any resultant turbidity current must have passed over site CD One possible explanation is that the turbidite between 2.5 and 2.7 m depth in core CD56825 represents the 6600 Cal yr BP event. If this is the case, the date of 7300 Cal yr BP below this turbidite would imply that the turbidity current resulted in about 20 cm of erosion of hemipelagic sediment plus any turbidites within it. [36] The evidence for a widespread erosional event at 8300 Cal yr BP is less compelling than that for the 6600 Cal yr BP event, since the former is seen in only two cores. However, we note that the post-6600 Cal yr BP sequence frequently directly overlies sediments deposited prior to 15,000 years, such that the later erosional event would have removed any evidence for the earlier event, had it been present at the core locations. 14 of 19

15 [37] Landslides mapped on the flanks of the lower canyon (Figure 9) removed up to 40 m of sediment from discrete areas of canyon slope and thus could have provided the volume of sediment required to form the extensive TAP turbidites. Because landslides are observed in two separate canyons, it is unlikely that they were triggered by purely sedimentary processes [Goldfinger et al., 2003]. It is much more likely that that all were set in motion by the same trigger. In this area, the only likely regional trigger is an earthquake Seismic Triggering of Landslides and Turbidity Currents on the Iberian Margin [38] The interpretation of seismic triggering of the 6600 and 8300 Cal yr BP turbidity currents can potentially be extrapolated to the deeper turbidite sequence in the TAP cores. On this basis, the large turbidites emplaced at about 10,500, 19,900 and 23,000CalyrBP,couldalsobeinferredtohave been triggered by large earthquakes. However, the origin of smaller turbidites recorded only at the more proximal core D11931 remains equivocal. Theses smaller events could have been generated by relatively smaller or more distant earthquakes or they could be the result only of sediment instability on the margin, particularly prior to 15,000 Cal yr BP, when sediment input through the canyons was much higher than in the post 15,000 year period because of lower sealevel and closer proximity of the canyon head to terrestrial sediment sources [Lebreiro et al., 2009]. [39] Gràcia et al. [2010] examined the occurrence of turbidites in cores from southeastern TAP, from a continental slope basin west of Cape St Vicente and from the Horseshoe Abyssal Plain (HAP) south of Gorringe Bank (Figure 1, inset). Their study is centered on the epicenter of the 1755 earthquake. They proposed a regional scheme of 14 turbidite events labeled E1 14, that could be correlated across all or parts of their study area. Their youngest events, consisting of relatively small turbidites are correlated with historical earthquakes. The results presented in this paper allow their scheme to be extended to the north along the Portuguese margin, and to the northwest into the more distal western TAP. The wider geographical spread of data that results from combining the two data sets allows new insights into possible turbidite sources, and thus earthquake possible locations. [40] Comparison with the work of Gràcia et al. [2010] shows that the three youngest turbidites with a potential earthquake source identified in this study can, on the basis of radiocarbon ages, be correlated with turbidites E8 (6690 Cal yr BP), E9 ( Cal yr BP) and E11 (10,175 10,425 Cal yr BP). The E12 event of Gràcia et al. (13,500 Cal yr BP) can also be identified in core D11931, but is absent in D11951 (Figure 6). E8 and E9 occur in the TAP and in the Horseshoe Abyssal Plain (HAP) to the south, but appear to be better developed, relative to other regional turbidites, in the TAP. E11 only occurs in the TAP. This supports our interpretation that the Portuguese margin in the vicinity of Setubal and Cascais Canyons was an important source area for these turbidites. However, multiple sources anywhere along the margin upslope of the TAP could have contributed. Indeed, the recognition of E8 and E9 in the HAP implies that slope failures must have taken place on the Portuguese margin at least as far south as 37 N, over 150 km south of Setubal Canyon. This strongly reinforces the interpretation of an earthquake trigger, based initially on the observation that these events caused erosion in both Setubal and Cascais Canyons. In contrast to E8, E9 and E11, E12 appears as a relatively minor turbidite in the TAP but is the largest of the post-glacial turbidites in the HAP. This also implies an earthquake trigger, but one presumably located further south on the Portuguese margin or in the Gulf of Cadiz. [41] Gràcia et al. [2010] suggest that E9 (our 8300 Cal yr BP turbidite) was related to a climatic cooling event at around 8200 Cal yr BP [Bond et al., 1997], specifically stating that despite being widespread, we do not regard event E9 as being seismically triggered. However, the presence of E9 as a large turbidite in core D11951, where the turbidite record appears to be relatively insensitive to changes related to sea level rise at the end of the last glacial, leads us to suggest that an earthquake trigger for this turbidite is a stronger interpretation, more consistent with the interpretation of the other large turbidites in the sequence Wider Considerations Regarding Earthquake Generated Turbidites on the Portuguese Margin [42] The size and distribution of an earthquakegenerated turbidity current is likely to be influenced by several independent factors that include the magnitude and source mechanism of the earthquake, the availability and stability of sediment deposits, and the location of the earthquake relative to that of any sediments that have the potential to be destabilized. Study of the effects of historical 15 of 19