Submarine Geomorphology☆ NC Mitchell, University of Manchester, Manchester, UK ã 2015 Elsevier Inc. All rights reserved.
Introduction Technology Preservation of Morphology and Erosion Mid-Ocean Ridges Seamounts Volcanic Islands Continental Slopes The Continental Rise and Sedimentary Fans Convergent Margins and Trenches Abyssal Plains Continental Shelf Seas References
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Introduction Geomorphology (the study of shape of Earth’s surface and the processes that have shaped it) is more challenging beneath than above sea level, because visual observations are limited by water turbidity. The development of the subject of submarine geomorphology has therefore more closely followed the development and deployment of new technology than subaerial geomorphology. Thus, whereas in subaerial geomorphology, hypothesis and inquiry have been prompted by structures that have always been visible, in submarine geomorphology, advances of understanding have tended to occur when new features have been discovered or imaged with instruments at higher resolution. Also, because of the limited accessibility, our knowledge of the marine geologic processes creating these morphological features has relied mostly on a forensic type of analysis (reconstructing events based on analyses of samples or other geophysical data) rather than by monitoring of processes more directly, although monitoring has been possible in some instances, in particular in the more accessible shelf seas. Beneath the oceans, erosion can be important locally such as in submarine canyons but is not as important as generally as it is in subaerial geomorphology where rainwater runoff and glaciation can cause radical changes in landscapes. Some features created by flows (of lava, debris, evaporites, etc.) and tectonics can therefore remain unmodified for long periods, making interpretation of process in some respects and in some locations easier than on land. These factors are illustrated in the following text with examples from a variety of physiographic settings.
Technology The general shapes of the ocean basins and the existence of major features such as the Mid-Atlantic Ridge became clear only in the mid-nineteenth century with the introduction of long lead sounding lines and their application mainly in surveying routes for telegraph cables. Since then, the oceans have been mapped both increasingly comprehensively and in finer detail. Besides mapping with echo sounders, radar altimeters on satellites have been extremely valuable in revealing the shape of the ocean floors to about 10 km resolution. This is because the ocean is a surface of nearly constant gravitational potential energy and because seabed topography modifies the local gravity field, owing to the density contrast between seawater and the rocks or sediments making up the bed (water tends to flow downhill – in this context flowing to areas of stronger gravity field). The resulting map (Figure 1(b)) provides an excellent view of the regional morphology of the deep oceans. Compare it with, for example, that of Heezen and Tharp (1961) in Figure 1(a). Although Figure 1(a) is stylized, it represents the most comprehensive bathymetry data available during the early 1960s. Many more tectonic features are now observable than were previously. Sonars reveal the ocean floor at finer resolution, in particular multiple-beam echo sounders. Other types called side-scan sonars produce images of seabed echoes (acoustic backscattering) and reveal seabed morphology more indirectly, although some sidescan sonars can also record bathymetry. At finer scale, sonars have been deployed on deeply submerged vehicles, such as autonomous underwater vehicles (AUVs) and remotely operated vehicles (ROVs). At finer scale still, submersibles and cameras/ videos have been used for many years to observe or image the seabed more directly. Around coasts where the water is not too turbid, aircraft fitted with LiDARs have been used to map the seabed down to 10s of meters. Efforts by the oil and gas industry have led to extensive seismic reflection datasets being collected in shelf, slope, and other deep-water environments. The first reflection ☆
Change History: April 2015. NC Mitchell updated the text.
Reference Module in Earth Systems and Environmental Sciences
Figure 1 Two views of the same area of the central North Atlantic bathymetry based on different generations of depth data. (a) Physiography based on sparse echo soundings of research vessels (Heezen and Tharp, 1961). (b) A modern version (18.0) of the bathymetry for the area in (a) obtained by interpolating between echo soundings with the gravity field derived from sea-surface shape from satellite altimeters (Becker et al., 2009; Smith and Sandwell, 1997). Annotation in (a): A, abyssal plain; R, continental rise; S, continental slope; CS, continental shelf. Arrows in (a) highlight the deepening of the seafloor with distance from the Mid-Atlantic Ridge crest caused by upper mantle cooling.
detected in 3-D seismic data can provide interesting views of the seabed (e.g., Bulat and Long, 2005). Such data also reveal some of the stratigraphic history leading up the present day, which can help in understanding how the morphology arose (e.g., Peakall et al., 2000; Posamentier and Kolla, 2003). Many other technological developments have been important for investigating the origins of morphological features, such as better seabed sampling and imaging methods and developments in analytical science. Although the investigations mostly use forensic approaches, there have been some important developments in monitoring to study processes more directly (Mitchell, 2012). For example, information on sedimentary flows is becoming available from acoustic Doppler current profilers (ADCPs) suspended over canyons (Xu, 2011), monuments tracked acoustically (Paull et al., 2010), and monitoring of bedforms with repeat surveys (Smith et al., 2005). In volcanology, there have been successful attempts to record deformation leading up to eruptions with seabed pressure sensors (Chadwick et al., 2012b) and seismometers (Dziak et al., 2012) and the lava flows produced by eruptions with repeat surveys (Caress et al., 2012; Figure 2). Eruptions of arc volcanoes have been monitored with hydrophones and repeat surveys (Chadwick et al., 2008a, 2008b, 2012a; Watts et al., 2012; Wright et al., 2008). In the field of tectonics, the modest depth accuracy of techniques available to us and the cost of the deployment of bottom sensors have meant that there has not been much information on how, for example, the earthquake cycle contributes to building morphology as is available on land (Stein et al., 1991). Nevertheless, bottom-moored acoustic transponders have been used to measure some movements, for example, Sato et al. (2011) described displacements immediately overlying the March 2011 To¯hoku–Oki (Japan) earthquake epicenter, derived using acoustic ranging from ships. Phillips et al. (2008) recorded vertical movements in the Hilina slump of Kilauea with bottom pressure sensors. In some shallow coastal waters, where tidal currents create sand banks and dunes, currents are more easily measured with moored current meters, ADCPs on ships, and Doppler radar systems for measuring surface currents, so monitoring is more extensive, as is information on sediment texture and bedform migration. Given the superior accessibility and commercial interests on the continental shelf, the literature in this area is mature (e.g., Dyer and Soulsby, 1988).
Figure 2 Lava flows on the Juan de Fuca Ridge mapped with an autonomous underwater vehicle (AUV) by Caress et al. (2012). (a) Change in bathymetry during eruption and (b) bathymetry after the eruption.
Preservation of Morphology and Erosion Some of us may have been taught as undergraduates that there is little erosion underwater because there is no rainfall. Erosion is not absent everywhere in the oceans because deep (>1 km) exhumation can occur physically such as by turbidity currents and mass movements (Hampton et al., 1996; Mitchell, 2014; Shepard, 1981) and chemically by dissolution of carbonate rocks (Paull et al., 1990). Nevertheless, in places where erosion has not occurred, features can be preserved well so that the submarine versions of various types of flows can sometimes appear less modified underwater than they are on land. For example, compare the surface morphologies of evaporite flows in the Red Sea (Figure 3; Mitchell et al., 2010) with namakiers in Iran, where rainwater dissolution has significantly altered their surfaces (Talbot and Aftabi, 2004) (the submarine evaporite flows are protected from dissolution by covering hemipelagic mud (Ross and Schlee, 1973)). The physiographic maps of Heezen and Tharp (1961) were developed using widely spaced lines of echo soundings. It is remarkable that they were able to reveal successfully the main features, such as transform and axial valleys in Figure 1(a). This success arises from the lineated morphology of the seabed resulting from plate tectonics and because erosion has not greatly transferred rock mass between areas, as it might have if these areas were exposed to continental erosional processes. The observations of systematic patterns in the Heezen and Tharp (1961) maps were made during the developing seafloor spreading hypothesis, so the lack of erosion may have been ultimately helpful for the early acceptance of plate tectonic theory. A world above sea level without biology would be significantly different from that on the present Earth (Dietrich and Perron, 2006). Biology probably also plays an important role in the development of seabed morphology. Crabs and other organisms burrow into consolidated or indurated sediments within canyons (Dillon and Zimmerman, 1970; Malahoff et al., 1982; Paull et al., 2005b; Valentine et al., 1980; Warme et al., 1978), preparing the beds of canyons for erosion by energetic sedimentary flows. Sediment resuspension by organisms, leaving the suspensions to deposit under gravity or be carried by currents, may also help to shape the seabed too. For example, the uppermost continental slope is commonly upward convex in profile, which mimics a decline observed in modern currents with water depth (Csanady et al., 1988). This characteristic convex profile shape has been variously explained as originating from how bed shear stresses due to currents decline with depth and modulate sediment deposition (Mitchell and Huthnance, 2007; Pirmez et al., 1998). However, over this same depth interval, rapid changes in biological mixing rates have been determined from radiometric tracers in sediment cores (Anderson et al., 1988; Henderson et al., 1999; Middelburg et al., 1997; Schmidt et al., 2002; Soetaert et al., 1996). The curved shape could therefore also partly originate from greater biological activity in the shallower water. More generally, pelagic sediment deposits in the deep ocean commonly have simple curved surfaces as though transported downslope according to a diffusion transport model scheme, such as from repeated resuspension and deposition during downslope movement of those suspensions (Mitchell, 1995). Resuspension by
Figure 3 Three-dimensional view of multibeam sonar data from the Red Sea showing giant flows of evaporites invading the spreading center (N. Augustin (pers. comm., 2015) based on multibeam data from Ligi et al. (2011).)
benthic organisms and their effect on the threshold of motion of sediment have been known for some time to be important sediment transport agents (Eckman and Nowell, 1984; Jumars and Nowell, 1984; Nowell et al., 1981) but how they affect the larger-scale geomorphology of the seabed is still poorly known in a quantitative sense.
Mid-Ocean Ridges Where Earth’s tectonic plates diverge, the rising mantle is locally hot and less dense. Subsequent cooling of the upper mantle from the surface causes it to contract and subside with seafloor age. These boundaries therefore form low-relief ridges, such as illustrated along the northern boundary of Figure 1(a). The subsidence follows well-known trends expected of cooling of the lithospheric mantle by conduction (Parsons and Sclater, 1977). The Mid-Atlantic Ridge created by this process is arguably the largest morphological feature on Earth, extending from Iceland to near Bouvet Island in the South Atlantic. At finer scale, the spreading ridges formed where plates are diverging rapidly, such as in the Pacific, contain a smaller-scale ridge where lava is erupted (the neovolcanic zone). At slower-spreading ridges, the neovolcanic zone is instead usually contained within a deep valley, such as that of the Mid-Atlantic Ridge (Figure 1). This valley is probably a consequence of the colder and hence more rigid lithospheric mantle under slow-spreading ridges (Tapponier and Francheteau, 1978). Once formed, the new oceanic lithosphere strengthens, so that loads placed on it cause the seabed to flex like an elastic plate, as seen around fracture zone valleys
(Sandwell and Schubert, 1982), rift valleys (Watts, 1982), and volcanic islands (Watts et al., 1980). In Figure 1(b), the principal east–west-trending structures are fracture zones produced by strike-slip motions at the ridge crest (Wilson, 1965). In contrast, the lineaments running oblique to the major fracture zones are the ends of volcanic spreading cells, which leave paired oblique trails on the adjacent plates as they migrate along the ridge (Briais and Rabinowicz, 2002). The more detailed work on mid-ocean ridge geomorphology has addressed the faults left by rifting, volcanism, mass wasting, and hydrothermal venting. Tectonic extension was originally considered to occur on moderate-angle faults, as this was the impression from low-resolution bathymetry and dips of nodal planes in earthquake fault plane solutions derived from teleseismic recordings (Thatcher and Hill, 1995). However, higher-resolution sonar data have revealed corrugated surfaces suggesting that low-angle faults also occur on the seabed (Blackman et al., 1998; Cann et al., 1997), where the corrugations were potentially imposed on footwalls by irregularities in the hanging walls (Spencer, 1999). Owing to their shallow dips, those surfaces are only weakly eroded, whereas the steep valley wall faults are easily degraded by mass wasting. On the Mid-Atlantic Ridge, a deep landslide embayment was identified in earlier multibeam data by Tucholke (1992), and deep-tow high-resolution bathymetry data have revealed the progressive degradation of fault scarps over millions of years (Tucholke et al., 1997). Escarpments of different lithologies can have systematically different seabed gradients (Mitchell et al., 2000), which deep-tow data from the valley walls now suggest are explainable by slides in ultramafic rocks and more granular-type flows of lava breccias (Cannat et al., 2013). High-resolution images within the axial valleys from deep-tow side-scan sonars have helped to reveal neovolcanic ridges and cones, which are apparently genetically connected; some cones may represent the final eruption phase after fissure eruptions have formed the ridges (Head et al., 1996) or the cones and terraces are fed by lava tubes from eruption sites (Smith and Cann, 1999). Inflation and deflation structures associated with emplacement of lavas were observed from submersible in the FAMOUS and other early expeditions (Bryan and Moore, 1977). Modern deeply deployed sonars have revealed these much more extensively and allow them to be studied in similar ways to subaerial flow fields (Deschamps et al., 2014). Data from spreading centers in marginal seas have shown that they are similar to the mid-ocean ridges but can be covered with terrigenous and other sediments. For example, much of the axial valley of the Juan de Fuca Ridge off the western United States and Canada is flooded with turbidites so that the axial valley faults offset the turbidites (Davis and Lister, 1977). Figure 3 shows a view of bathymetry from the central Red Sea, where Miocene evaporites deposited around the ridge are flowing into the spreading center (Augustin et al., 2014; Ligi et al., 2011; Mitchell et al., 2010).
Seamounts Volcanic seamounts (submarine volcanoes) are the most common forms in the deep oceans, although seamounts can also be created by faults. Guyots are flat-topped seamounts that were originally islands, now subsided beneath sea level, their flat tops representing surfaces produced by surf erosion or coral that died abruptly during submergence (Menard, 1984). Volcanic seamounts form sublinear chains left by mantle hot spots as the overlying plates migrated (Morgan, 1972; Wessel and Kroenke, 2008), or they may be related to major fissures or lines of weakness in the plates (Winterer and Sandwell, 1987). The Louisville guyots, recently sampled by the Integrated Ocean Drilling Program, illustrate many of these features (Figure 4); they form a chain created by a mantle hot spot and have flat summits and steep slopes containing chutes formed by landslides. Seamounts typically are truncated cones with profiles that vary in flatness (Smith, 1988). In plan view, small seamounts are circular to elliptical, whereas seamounts taller than around 2–4 km are more irregular (Mitchell, 2001). Seamount summits can contain collapse pits or craters (Fornari et al., 1984). The use of a hull-mounted multibeam sonar with improved resolution led Clague et al. (2000b) to suggest that seamounts grow by central eruptions rather than from ring dikes suspected previously. Small island arc volcanoes are commonly simple cones, suggesting they grow by erupting from central vents with cone sides at the angle of repose of talus (young lavas on steep slopes have been observed around Hawaii to rapidly disaggregate (Sansone and Smith, 2006) so steep young lavas are unlikely to survive for long). The effects of repeated eruptions of Monowai volcano (Kermadec arc) have been captured with repeat surveys since 1998 (Chadwick et al., 2008b; Watts et al., 2012; Wright et al., 2008). The results show not only gain of material around the central vent of 10s of meters but also loss of >100 m caused by large slope failures, representing the mechanism by which magma erupted in the center is redistributed. Similar observations have been made of a Mariana volcanic arc volcano, where hydrophones captured acoustic emissions produced by slope movements (Chadwick et al., 2012a) and hydrophone and video images have captured details of eruption activity at the vent (Chadwick et al., 2008a).
Volcanic Islands The submerged parts of volcanic islands are far larger volumetrically than their parts above sea level. In volcanically active islands, their submarine slopes typically continue their subaerial slopes beyond a narrow shelf produced by surf erosion though with steeper gradients, particularly just below the shelf edge where slopes are commonly close to the angle of repose of talus (Lee et al., 1994; Mark and Moore, 1987; Mitchell et al., 2008). Between these two areas lies a narrow shelf. Quartau et al. (2012) demonstrated with modeling how the shelf of Faial Island (the Azores) is explained by surf erosion combined with varied Pleistocene sea
levels. Limited high-resolution multibeam data from the shelves of active islands have revealed lava flows emplaced on them (Figure 5). The ridges above sea level of volcanic rift zones (Carracedo, 1994; Dieterich, 1988) typically continue as ridges beneath sea level. A detailed survey of the Puna Ridge of Hawaii found volcanic cones, ridges, and terraces that are like those of mid-ocean ridges (Smith et al., 2002). A morphometric study of cones in the Azores (Mitchell et al., 2012b) found some cones in water shallower than 400 m were unusually flat, possibly a result of spreading of eruption columns at the ocean surface (Cashman and Fiske, 1991). Another study of unusually flat but deeper cones in the Hawaiian Islands concluded that they were likely formed from small lava shields (Clague et al., 2000a). Studies of submarine eruptions close to islands have provided clues to how cones are created. During an eruption near Terceira island (the Azores) in 1998–2001, lava balloons were observed inflating and rising to the ocean surface where they subsequently cooled and became waterlogged before subsiding (Gaspar et al., 2003). Multibeam sonar data collected afterward reveal the
Figure 4 Guyots (flat-topped seamounts) of the Louisville chain in the South Pacific in multibeam data collected by Peter Lonsdale (from front cover of March 2010 issue of the journal Oceanography).
Figure 5 Surface morphology of the seabed immediately adjacent to where an a’a’-type lava flow entered the sea in AD 1718, revealing elaborate dendritic lava flow fronts. Image was generated from multibeam sonar data collected on the north coast of Pico Island, the Azores (Mitchell et al., 2008). Top right inset locates the dataset with respect to the island.
location of the site on a volcanic ridge emanating west of Terceira island (Chiocci et al., 2013). In contrast, an eruption south of El Hierro (the Canaries) in 2011–12 was located off the volcanic ridge running south of the island and was accompanied by a band of seismicity also offset from the ridge (Lopez et al., 2012; Martı´ et al., 2013). Although volcanic ridges are considered to be the pathways for magma (Vogt and Smoot, 1984), the El Hierro eruption suggests that magma may not always be supplied along the ridge axes. The widespread presence of cones around such islands also suggests that dikes may transfer magma to the surface over a wider range of locations. Repeated sonar surveys of the El Hierro site after the onset of the eruption revealed that a cone developed over a preexisting valley and its summit reached 90 m depth (Rivera et al., 2013). An eruption in 1996 of Lo¯’ihi volcano adjacent to Hawaii coincided with the development of a 300 m deep pit crater on its summit (Garcia et al., 2006), mimicking the behavior observed in the caldera of Kilauea, where rift zone eruptions or intrusions often accompany caldera floor subsidence. The construction of volcanic islands produces steep and unstable slopes, which can fail in landslides that are among the largest on Earth (Carracedo, 1999; Holcomb and Searle, 1991; Masson et al., 2002; Moore et al., 1989). Multibeam and side-scan sonar data played a crucial role in the acceptance of these features as landslides; without such data, the large subaerial embayments could instead potentially be explained by other processes, such as caldera collapse. The causes of landslide are often unclear because the conditions at the time of failure of these prehistorical events are difficult to reconstruct (Keating and McGuire, 2000). Nevertheless, a slump movement captured in continuous Global Positioning System (GPS) measurements of south Kilauea coincided with intense rainfall, suggesting that pore pressures are potentially important in reducing effective stress on failure surfaces (Cervelli et al., 2002). Landslide embayments are largely absent in submarine volcanoes and islands less than 2.5 km in relief (Mitchell, 2003), suggesting also a volcano height threshold involving changing permeability structure and pore pressures. Besides these spectacular events, the flanks of islands also develop channels from a variety of smaller-scale failures (Mitchell et al., 2003). Repeat multibeam surveying of the 2002 landslide of Stromboli island showed how the slope was repeatedly affected by deposition and failure of young deposits (Chiocci et al., 2008).
Continental Slopes Seaward of the continental shelves lie continental slopes. In profile, the slopes have a variety of shapes, though commonly sigmoidal (Schlager and Adams, 2001; Figure 6). The uppermost slope is convex upward because of declining energy of the environment with depth, which allows increasingly steep deposition (Mitchell and Huthnance, 2007; Pirmez et al., 1998). Slope sediments are commonly muddy and Cacchione et al. (2002) had suggested an intriguing mechanism by which deposition of fine particles steepens the slope until the gradient reaches a critical value at which oceanographic internal waves break, much like surface waves on a beach, preventing further steepening. Where continental slopes are underlain by deformable shale or rock salt, the overlying strata can be folded and faulted (Morley et al., 2011) like in accretionary prisms of convergent margins described in the succeeding text. In sonar data, canyons and gullies can form an erosive-like morphology similar to fluvial networks (Belderson and Kenyon, 1976; Belderson and Stride, 1969a; McGregor et al., 1982; Mitchell, 2004; Figure 6). Pratson and coworkers (Pratson and Coakley, 1996; Pratson et al., 1994) had investigated how this morphology arose from downward-eroding sedimentary flows, such as initiated by distributed slope failures. Harris and Whiteway (2011) developed a global inventory of canyons, which suggests a relationship between canyon incidence and the supply of sediment from the adjacent continents (canyons are more common where there are glacial or fluvial sediment sources and on active margins where the shelf is commonly narrow). Slope channel networks can have properties similar to mountain bedrock river networks, such as comparable branching patterns (Pratson and Ryan, 1996) and confluences obeying Playfair’s rule (tributaries commonly join main channels at the same elevations without intervening steps; Mitchell, 2004), longitudinal profiles are commonly upward-concave (Mitchell, 2004), and channel segments that are steepened can show knickpoint retreat and a variety of knickpoint morphologies as do rivers (Adeogba et al., 2005; Heinio¨
Figure 6 Three-dimensional view of multibeam sonar data from the continental slope off the US Atlantic coast. From data originally provided by the NOAA (Mitchell, 2004).
and Davies, 2007; Mitchell, 2005; Pirmez et al., 2000). However, these systems differ in a few key respects, for example, the sedimentary flows gain or lose power as they incorporate or deposit sediment, respectively, because of the effect on flow density relative to ambient water and thus driving stresses (Parker, 1982); the effect on flow density relative to air is much more muted in river flows. Furthermore, continental slopes typically aggrade over time from the accumulation of particles shed from the continents, in contrast with subaerial mountain landscapes that are denuding. Gerber et al. (2009) attempted to model some of these effects to reproduce the forms of slope channels. If time-averaged erosion rate were locally related to the bed shear stress or flow power imposed by sedimentary flows passing through them as suggested for bedrock rivers (Howard, 1994), there might be expected to be some relationship between the relief of canyons and the volume of sediments that have passed through them (Mitchell, 2014), but Normark and Carlson (2003) showed that there appears to be none, possibly suggesting that large landslides may be important in developing canyons. The larger sedimentary flows that do contribute to excavating canyons are very rare (Paull et al., 2005a) and are not captured in modern monitoring data. Nevertheless, information on the structures of smaller flows has been obtained using ADCPs, such as over Monterey Canyon (Xu, 2011). Multibeam sonar data from the United States and other continental slopes have revealed abundant landslides (Harders et al., 2011; Hu¨hnerbach and Masson, 2004; McAdoo et al., 2000; ten Brink et al., 2009; Twichell et al., 2009). A modest tsunami was created during the Grand Banks earthquake and landslide (Piper et al., 1999), and concerns of potential tsunami risk to the US East Coast have been expressed based on possible geophysical evidence for incipient cracks in the outermost continental shelf (Driscoll et al., 2000). In detail, landslides have varied morphologies, for example, Micallef et al. (2007) described the giant Storegga Slide off Norway as a spreading type. Because the continental slope comprises a broad dipping surface on which sediment layers have accumulated, leading to widespread dipping planes of weakness, landslides can be much larger than those on land where continuous slopes are generally smaller (ten Brink et al., 2009). As with the volcanic island landslides, the causes of the continental slope landslides are uncertain. Geotechnical evaluations are complicated by creep behavior of some clay-rich deposits (Silva and Booth, 1984), so short-term behavior in laboratory geotechnical tests may not represent behavior in situ over longer periods. Gas hydrate (a solid form of methane and water) is stable in many shallow sediments, and gases released during their dissociation (e.g., during ocean warming) have been thought to be a potential cause of slope failure. The Storegga Slide was originally considered to be caused by hydrate dissociation but dissociation is now thought to have been only a minor influence; more likely, failure was caused by an earthquake (Bryn et al., 2005). Nevertheless, Hu¨hnerbach and Masson (2004) showed that North Atlantic landslide headwalls intriguingly cluster around 1000–1300 m water depth.
The Continental Rise and Sedimentary Fans The continental rises are low-relief aprons to the continental slopes, extending to, and merging indistinctly with, the abyssal plains (Figure 1(a)). They comprise debris shed from the continents by sedimentary flows and fallout of particles from suspension (hemipelagites). Off the US East Coast, in the uppermost rise, pits up to 75 m deep occur below slope gullies and have been interpreted as produced by impacts of debris flows or hydraulic jumps of sedimentary flows emerging from the gullies (Lee et al., 2002). The rise has a more subdued relief than the continental slope but with irregularities due to deposits of debris flows (Pratson and Laine, 1989; debrites). In other continental margins, the rise may contain different sediment types but in general usually has subdued relief away from the sedimentary fans. Sedimentary bedforms created by deep-ocean currents commonly occur, as described under ‘Abyssal Plains.’ Sedimentary fans occur below the continental slopes, where sedimentary flows leaving slope canyons become less constrained laterally. Fans form some of the largest sedimentary features on Earth, for example, the Bengal Fan extends for 3000 km. Channels on fans can appear similar to river channels, as they are also sinuous and flanked by levees (Damuth et al., 1988). However, more detailed analysis has revealed important differences. For example, seismic reflection data suggest that such channels do not migrate laterally (swing) with time as do rivers, perhaps a result of internal pressure gradients setting up different helical movements in passing flows (Peakall et al., 2000). The channel floors can lie above the elevation of the surrounding fan. Where the levees have failed, allowing a new channel to form in the levee break, the new channel can have a steep gradient, producing a knickpoint that can migrate upstream with erosion like a river channel knickpoint (Pirmez et al., 2000). Given the importance of fans as oil and gas reservoirs, the industry has collected enormous amounts of 3-D seismic data in these areas, which are extremely useful for research on channels (Posamentier and Kolla, 2003).
Convergent Margins and Trenches The downgoing plates at trenches bend into subduction zones. This causes flexing of the plates, so that at typically 100 km from the trench, the plate rises before starting its descent into the trench (Watts, 2001). In the intervening zone where the curvature of the plate is largest, bending stresses may exceed the strength of the lithosphere, creating normal faults displacing the seabed (Watts et al., 1980). The depression formed between the two plates is typically filled with turbidites and other gravity-transported sediments derived from the adjacent slope, islands, or continent (von Huene and Scholl, 1991). Within the trench, the sedimentary rocks on the plate being overridden (subducted) and turbidites filling the intervening trench are scraped off, creating stacks of folded and faulted strata called accretionary prisms (von Huene and Scholl, 1991). The pattern of
Figure 7 (a) Bathymetry of the lower Gulf of Alaska continental slope collected with a multibeam sonar of National Ocean Service ship Surveyor (Mitchell, 2006). Depth contours are plotted every 50 m (numbers on bold contours are depths in kilometers). White bold dashed line is survey line of R/V Lee, along which US Geological Survey scientists collected seismic reflection data in 1981 (lines ‘13’ and ‘14’ of Fruehn et al. (1999)). (b) Interpretation of seismic reflectors adapted from Fruehn et al. (1999) for line Lee-13.
folds and thrust faults is illustrated by the bathymetry in Figure 7(a), where the northeast–southwest-trending ridges are anticlines of the Gulf of Alaska accretionary prism. Figure 7(b) shows the geometry of the strata deformed by the convergence (Fruehn et al., 1999). Such margins also commonly contain channels excavated by sedimentary flows initiated higher up in the slope. In Figure 7(a), a major channel can be observed running between the anticlines. Near the range front (to the southeast), the channel has a knickpoint, highlighted by the white circle in Figure 7(a). This knickpoint and another in a piggyback basin also highlighted lie upstream of the range front, as expected if knickpoints produced by fault movements have migrated upstream as a result of erosion by sedimentary flows similar to knickpoint migration in rivers (Gardner, 1983). A series of arcuate bedforms are observable where the channel opens out at the base of the slope into the trench. As a result of steepening of the continental slope by change in convergence rate or friction on the underlying faults, a slope channel with a sinuosity developed on an originally low-gradient slope can be preserved (Soh and Tokuyama, 2002) and landslides initiated (Kawamura et al., 2012). Seamounts and other prominent basement features on the subducting plate can locally indent and steepen the accretionary prism and cause landslides (Harders et al., 2011).
Abyssal Plains Between the ridges, continental slopes, trenches, and islands lie the very low-relief abyssal plains. They are commonly underlain by oceanic crust or, less commonly, by highly extended continental crust. Within oceanic crust lie abyssal hills, which were formed from the development of normal faults and volcanism at the original ridge crest where the crust was created. Abyssal hills provide a topographic fabric or grain, such as can be seen in Figure 1(b), running approximately north–south and parallel to the Mid-Atlantic Ridge. As also clear in that figure, the abyssal hills become progressively covered by sediment so that their relief becomes increasingly subdued with distance and age from the ridge, finally becoming fully obscured by turbidites originating from the continental margins.
Figure 8 Shaded relief bathymetry showing elongated depressions and ridges in pelagic sediments in the abyssal equatorial Pacific (Mitchell and Huthnance, 2013). In (a), the features wrap around two seamounts that have small moats produced by accelerated bottom flow at their bases. In (b), the features run across the abyssal hills and in two cases emanate from depressions created by carbonate dissolution from waters emanating from basement. Sun illumination directions are from (a) N300 E and (b) N330 E. Dashed lines represent the local orientation of abyssal hills.
Although these areas have low relief, sonar, seismic, and sediment profiler data have revealed many bedforms produced by ocean currents. Flood (1988) described mudwaves, with spacings between crests of up to 6 km and heights up to 100 m, created by abyssal currents. Furrows are longitudinal bedforms created by helical secondary currents and are aligned with the main current (Flood, 1983). These features and sediment drifts can help in working out the pattern of abyssal flow (Lonsdale, 1981). Figure 8 shows shaded relief views of multibeam data from the northern central Pacific, with elongated sediment drifts and furrows suggesting the local direction of bottom water flow (wrapping around two seamounts in Figure 8(a) and running obliquely across abyssal hills in Figure 8(b)). Long-term current meter data in the equatorial region show the deep-water movement varying over many periods, though with an important oscillation of 6-month to annual period, so such features are not necessarily produced by steady flow in one direction. Over time, the sediments carried by such currents can be deposited in giant rounded drifts called contourites (Fauge`res et al., 1999). Also shown in Figure 8(b) are small depressions of the sediment surface. They tend to overlie buried abyssal hills and have been suggested to arise from dissolution of carbonates by warm pore waters emanating from the basement (Bekins et al., 2007; Michaud et al., 2005; Moore et al., 2007).
Continental Shelf Seas The continental shelves were produced by surf erosion (Trenhaile, 2001) and by sediments deposited as continental margins subsided. Consequently, margins that are young or tectonically mobile tend to have narrower shelves (e.g., the Californian coast or coasts along subduction zones such as the western coasts of Peru and Chile). Coasts can be broadly classified into those dominated by tidal currents and those (commonly with narrower shelves) dominated by waves (Davis and Hayes, 1984). Depending on local water depth and other factors, morphological features produced by these two influences extend onto the shelf. During glacial lowerings of sea level, coastlines were closer to the shelf edges. In order to understand their morphologies, therefore, we need to consider the range of subaerial, coastal, and submarine processes that may have affected the seabed in these areas over the Pleistocene sea level variations. Because bedforms created by sediment movements and erosion by tidal currents and surface waves themselves often influence those currents and waves, coastal seas illustrate many interesting feedbacks (Masselink and Hughes, 2003). The ocean tide interacts with the morphology of the continental shelf to produce strong currents in some areas, such as the northwest European shelf. Sediments introduced to the shelf by rivers, coastal erosion, and (in the case of northern Europe) glaciers are reworked by those currents into bedforms of varied sizes (Belderson et al., 1982). At the largest scale are long sandbanks typically oriented in the direction of the tidal currents (Dyer and Huntley, 1999; Huthnance, 1982). Superimposed on them can be trains of sand waves (popularly called dunes; Figures 9 and 10) and smaller still waves (megaripples) with their crests oriented perpendicular to the current. The asymmetry of the dunes can suggest the transport direction of the sand (onto their lee slopes; Belderson and Stride, 1969b; Bernard and Haines, 2007; van Landeghem et al., 2009a), although repeat surveys have revealed that reverse migrations can also occur (van Landeghem et al., 2012). In areas where sediment is transported away from an area, the bed is commonly bare of easily transportable sediment (Harris et al., 1995). In other areas where the glaciers reached the shelf during
Figure 9 Sand dunes imaged with multibeam sonar at the entrance of San Francisco Bay. (Bernard and Hanes, 2007)
Figure 10 Trochoidal sediment waves imaged with multibeam sonar in the western Irish Sea (van Landeghem et al., 2009a). Depths and distances in the lower graph are in meters.
the last glacial period and where there has not been much subsequent modification, many glacial features can be preserved, such as moraines, drumlins, and eskers (Loncarevic et al., 1994; van Landeghem et al., 2009b). On the Nova Scotian Shelf, glaciers stripped the drift geology, leaving the modern bathymetry showing basement structures (Loncarevic et al., 1994). The ice likely held back large water bodies during deglaciation; multibeam data from the English Channel reveal giant scours attributed to outburst of a North Sea glacial lake (Gupta et al., 2007). When sea level was depressed during the Pleistocene glacial stages, shelf areas were exposed and rivers traversed them to a receded shoreline nearer to the modern shelf edge. In some cases, the rivers may have emptied into the heads of canyons (Twichell et al., 1977), while in others, this is less clear – the shelf river mouth need not necessarily extend in the directions of canyon heads during falling sea level (Fagherazzi et al., 2004). High-resolution seismic data from the US Atlantic shelf have revealed evidence for distributed channels and estuaries that would have reached the receded shoreline during glacial times (Nordfjord et al., 2005). The US Pacific coasts are quite different. There, the shelves are narrower and adjacent rivers typically deliver finer-grained sediment, which can deposit on the shelf as a mud patch (Nittrouer and Wright, 1994). Because deep-ocean waves are not so greatly attenuated by the time, they reach shallow water, and they influence the bed more strongly than over the wide shelves. Consequently, muddy sediment kept suspended by the wave action can produce denser water, which cascades off the continental
shelf as gravity flows (turbidity currents), potentially depositing as the wave influence and bed gradient decline (Wright et al., 2002). In shallower water, wave action on sands in narrow-shelf coastlines can similarly create sand bodies (clinoforms) elongated subparallel to the coast (Herna´ndez-Molina et al., 2000; Mitchell et al., 2012a). Excavations of sediment caused by expulsions of fluids can leave depressions called pockmarks (Hovland and Judd, 1988). The features mentioned earlier are modified variously by tectonic, volcanic, and other structures depending on the local context.
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