Chapter Four High Mud-Supply Shores

CHAPTER FOUR

High Mud-Supply Shores

4.1. Introduction Although wave-dominated clastic shores are much more commonly composed of sand (Chapter 5), gravel or a mixture of both (Chapter 6), specific sediment supply conditions can result in predominantly muddy, open wave-exposed shores. These shores differ from the classical, relatively sheltered, mudflat and marsh systems in that they are commonly associated with significant mud supplies, and, sometimes, relatively high levels of wave energy that undergo considerable dampening within mud bank and mudflat substrates. This high mud supply generally constitutes the overarching geological control, commonly leading to rapid progradation and growth of a muddy clinoform. Muddy shores occur along several open coasts associated with the accumulation or longshore dispersal of high fine-grained river discharge, the most important of which is the 1600 km-long coast between the mouths of the Amazon and the Orinoco Rivers, South America, which is bounded by the longest stretch of muddy shoreline in the world. Other notable examples include parts of the Texas–Louisiana coast to the west of the Mississippi, the Yangtze and Red River delta shores in Asia, and the West African coast between Guinea-Bissau and northern Sierra Leone. Such shores are generally flanked by marshes and bare mudflats several hundreds of metres to several kilometres wide. They may also form thick muddy lithosomes under conditions of progradation associated with still-stand conditions (e.g. Allison & Nittrouer, 1998; Walsh & Nittrouer, 2004). The mud may occur in various stages of concentration and consolidation, ranging from very high SSCs (1–10 g l1), through fluid mud, to settled mud which, in turn, ranges from underconsolidated (r650 g l1) to over-consolidated beds (Z750 g l1). Fluid mud has concentrations at which the settling velocity starts to be impeded by inter-particle interactions, and has been described by Mehta (2002) as an energy-absorbing slurry with typical densities ranging from 10–300 g l1. Fluid mud concentrations may be organised into distinct mud banks more or less attached to the shore, as on the South American coast between the Amazon and the Orinoco. The mud banks on the coast of Kerala, in India, are different, and appear to be self-organised forms that undergo a seasonal cycle of in situ dynamic changes involving no longshore or cross-shore dispersal. This coast is fronted by long sandy beaches with no rivers that are liable to supply mud, and Narayana, Jago, Manojkumar, and Tatavarti (2008) have suggested that the Kerala mud banks are palimpsest, marshy, lagoonal deposits rich in organic matter and derived gas, that were submerged after a marine transgression. Mud banks evincing morphodynamic and sediment interchanges with the shore are derived from specific fine-grained sediment accumulation processes on the shoreface in the vicinity of river mouths 131

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that serve as purveyors of such sediment. In the case of the Amazon, the mud banks (E1.0–1.5  106 tonnes yr1) and associated highly turbid suspensions (E106 tonnes yr1) dispersed alongshore account for 15–20% of the fine-grained sediment (E1.2  109 tonnes yr1) supplied by this river (Eisma, Augustinus, & Alexander, 1991). These mud banks, an example of which is depicted in Figure 4.1, are currently spaced at intervals of 15 to 25 km, are up to 5 m thick, 10 to 60 km long and 20 to 30 km wide (Gardel & Gratiot, 2005).

Figure 4.1 A SPOT image showing the location of the Macouria mud bank in September 2003 on the French Guiana coast. This mud bank is one of several (up to seven) banks migrating along the 200 km-long coast of French Guiana at any time, en route from the Amazon towards the Orinoco. Adapted from Anthony et al. (2008a), with permission from Elsevier.

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4.2. Coastal ‘Mud Streams’ and Processes of Mud Bank Formation Shorefaces in the vicinity of large river mouths may exhibit bed and sediment characteristics directly influenced by estuary-mouth and deltaic processes. In these conditions, the delta front normally exhibits a progressive, offshore decrease in the bulk grain size, associated with gradually thinning and fining sandy beds, and a corresponding increase in the thickness of the mud interbeds (Dalrymple & Choi, 2007). In strongly tide-influenced deltas, such as the mouths of the Amazon (Allison & Nittrouer, 1998) and the Fly (Dalrymple et al., 2003; Walsh & Nittrouer, 2004), and of the numerous interlinked river mouths on the muddy West African coast between northern Sierra Leone and Guinea-Bissau (Anthony, 2006a), fluid mud beds may form from high-discharge sedimentation, and may constitute the basis for the formation of inshore mud banks. Dalrymple and Choi (2007) suggest that such mud beds are structureless and unbioturbated, except in the upper layers, but the intervening sand layers may be deposited more slowly and exhibit a higher degree of bioturbation. Geostrophic forcing may lead to the alongshore advection of river plumes that detach from the coast (e.g. Liu et al., 2006; Warrick et al., 2007). Under conditions where such plumes are rich in suspended sediment, there may develop a longshore dispersal of sediment (Figure 2.6), sometimes over very long distances in coastal ‘mud streams’. Mud bank formation is essentially related to rapid and sustained finegrained sediment concentration and trapping associated with fresh water–saltwater interaction and front activity over the shoreface. In this shore system, the classic processes of mud trapping in the ETM (Section 2.2.1) shift seaward onto the shoreface, generally as a result of high river discharge, where they form seasonally constituted fluid muds that migrate alongshore.

4.2.1. Front Dynamics With distance from the mouth, a river outflow may evolve to stratified conditions across a frontal transition, the position and extent of which depend on the bathymetry and tidal conditions of the receiving waters (Geyer et al., 2004). The frontal zone may become broader at greater distances from the mouth under the influence of strong tidal currents and shallow receiving waters, the extreme example being the Amazon River, in which the frontal zone is 150 km seaward of the river mouth (Geyer & Kineke, 1995). The width of the front depends on the more complicated factors of the tidal mixing intensity and bottom slope in the vicinity of the front. The gentle slopes and strong tidal currents associated with high mud-discharge deltas may result in frontal zones stretching over tens of kilometres, as in the case of the Amazon, the frontal zone of which extends over 50 km in the cross-shore direction (Geyer & Kineke, 1995). As Geyer et al. (2004) have remarked, the Amazon frontal zone has a cross-shore salinity distribution, dynamics and kinematics similar to that of an estuary, except for the absence of lateral boundaries, thus permitting transport in the along-front direction. From a dynamic point of

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view, inner-shelf frontal zones can be regarded as estuaries that have been displaced onto the shelf due to the combination of strong river outflow and ebb-tidal currents (Geyer et al., 2004). These authors have illustrated the potential efficacy of the frontal zone as a sediment trap, due to the same mechanisms that make an estuary an effective sediment trap, i.e. convergence of near-bottom flow and separation of the outflow from bottom-generated turbulence. On the landward side of the front, sediment distribution occurs throughout the water column, maintained in suspension by vigorous bottom turbulence. Across the front, the increase in stratification suppresses turbulence in the upper part of the water column, even if tidal currents are strong. The sediment-charged fresh water is advected over the saline layer into the plume, but with the shutting off of bottom turbulence, the settling is no longer balanced by resuspension, and sediment begins to settle from the plume. Enhanced flocculation may occur in the frontal zone due to a decrease in turbulence, which increases the settling velocity and further promotes trapping of sediment. Salt-induced flocculation may, however, not be the only mechanism involved in rapid river-mouth sedimentation. Thill et al. (2001) found evidence for only a minor role by salt-induced flocculation in the Rhoˆne river mouth.

4.2.2. Sedimentation An important criterion that favours intense muddy sedimentation over the shoreface is the relatively short sediment transport distance, before settling, relative to the width of the frontal zone. In the case of the Amazon, the frontal zone exhibits intense frontal trapping of fine sediment that reflects the following conditions (Geyer et al., 2004): characteristic horizontal velocities of 1.5 m s1, settling velocities of 1 mm s1 and a frontal zone depth of 10 m, and a horizontal scale for settling of sediment of 15 km, compared to a 50 km width of the frontal zone. The trapping of sediment in the frontal zone of the Amazon is so important as to generate concentrations high enough to produce the fluid mud evoked earlier. The Amazon mud stream extends, in both the highly concentrated form of mud banks and highly turbid suspensions, over 1600 km to the mouth of the Orinoco. Allison, Lee, Ogston, and Aller (2000b) showed from seismic profiles, sediment cores and water column measurements near the mouth of the Amazon that the mud banks are translating over a modern shallow (o5 m) inner shoreface mud wedge. Initial mud bank development occurs over intertidal and shallow subtidal mudflats associated with an alongshore-accreting clinoform feature. The sediment trapping is controlled by strong water column stratification produced by the Amazon freshwater discharge on the shelf. In an earlier study, Allison, Nittrouer, and Faria (1995) had shown that the 350 km shoreline adjacent to the Amazon River mouth in the area of mud bank formation is one of significant mud recycling that comprises three distinct dynamic/morphosedimentary types: erosional mud, accretionary sand and accretionary mud. Sand bodies supplied by the local rivers in the vicinity of the mouth of the Amazon are up to 5 m thick and overlie erosional mud shorefaces. Elsewhere, muddy aggradation and progradation take place on under-consolidated, low-gradient tidal flats backed by mangrove swamps. 210Pb and 14 C geochronology of vibracores from the mudflats indicate that sediment

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accumulation is rapid (0.24–2.0 cm yr1) landward of the 2 m isobath, supplied from a thick (50–150 cm) seasonal surface layer. Shoreface progradation is episodic and separated by decadal hiatuses. These mudflats receive fine-grained suspended sediment flux from the Amazon and minor amounts of sand and mud from the local rivers. Shore-normal tidal currents and solitary waves rework the surface mud layer, preferentially transporting available sand landward onto the mangrove fringe, and producing very fine-grained accumulation on the tidal flat (10–12 j mean grain size). Despite the fluidity of these muds, cores show some degree of layering indicative of grain-size sorting, and orientation of clay minerals is common (Rine & Ginsburg, 1985; Allison et al., 1995). Rine and Ginsburg (1985) identified alternations of beds of massive structureless mud up to as much as 2 m thick with parallel, wavy and lenticular laminations and, rarely, even micro cross-lamination. Laminae of silt and fine sand alternate with more clay-rich laminae. The role of diatom-supported biofilms in such laminations has been highlighted by Debenay et al. (2007). In the Gulf of Papua system, Walsh and Nittrouer (2004) identified four stratigraphic facies: (1) supratidal to high-tidal muds; (2) mid-tidal sandy muds; (3) low-tidal sand and mud; and (4) subtidal channel sands. Supratidal and high-tidal sediments of mangrove forests are typically muddy, but can have high sand contents (W30%) in areas exposed regularly to ocean waves. These muds can be homogenous, but commonly laminated and bedded, probably reflecting high rates of sediment accumulation relative to bioturbation. They may have low water contents due to infrequent inundation and, because of their cohesive nature and abundant roots, can form steep intertidal banks and erosional scarps. Mid-tidal and low-tidal sediments are thinly laminated (o2 mm), but the latter also contain thick (W2 mm) laminations, likely deposited from fluid mud transport. Channel sediments have thin and thick mud and sand laminations, but are identified by thick (W5 cm) sand beds. Walsh and Nittrouer (2004) proposed, from these facies, a stratigraphic model for the Fly River delta system. In deltaic settings, such as that of the Fly River, the large supply of sediment may lead to dilution of organic matter levels and preclude peat development. Such peat development may be more typical of ‘B-type mangroves’ that depend essentially on authigenic sediment supply.

4.2.3. Sediment Recycling The shallow depths of the frontal zone are subject to significant sediment remobilisation due to tides, wind events and waves. The sediment trapped in frontal zones may, thus, be remobilised several times before its ultimate burial. On the Amazon–Guianas shoreface, Aller, Heilbrun, Panzeca, Zhu, and Baltzer (2004) employed a broad range of tracers such as 234Th (t1/2 ¼ 24 days), 210Pb (t1/2 ¼ 22 years), seasonal Cl profiles, and non steady-state diagenetic models of pore water concentrations and oxidant–reductant relationships to demonstrate that the mobile mudbelt is a zone of extraordinarily intense sedimentary and biogeochemical recycling, greatly exceeding stable coastal systems, such as salt marshes, in material exchange with the sea. The upper 0.1–1 m of deposits are reworked and exchanged with overlying water on timescales of o10 days to seasonally. In such areas, the seafloor, thus, acts as a massive suboxic batch reactor, entraining and processing

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reactive marine plankton, regenerating Fe, Mn oxides, exchanging metabolites and nutrients with the oxygenated water column, and generating suites of non-sulfidic authigenic minerals (Aller et al., 2004).

4.3. Mud Bank–Shore Interaction Processes The foregoing sub-section on the shoreface sedimentation processes and largescale shoreline development encapsulates a number of finer-scale processes that directly determine patterns of mud bank interaction with the shore. Both field and remote sensing approaches are progressively throwing light on these interactions which involve wave energy dampening by mud, mud bank liquefaction, cross-shore and longshore mud advection that control patterns of mangrove colonisation and elimination.

4.3.1. Wave Energy Dampening by Mud and Mud Bank Liquefaction by Waves Mehta (2002) proposed a sequence of events involved in muddy shore response to wave forcing. Increase in bottom fluid stress initiates erosion of a relatively rigid bed surface. This, in turn, generates turbidity and an increase in sediment concentration towards the shoreline. The process may result in the dislodging and transport of large mud clasts away from the breaker zone, with liquefaction of the bottom in due course, and penetration of wave orbital motion into the fluidised mud layer. The formation of fluid mud leads to significant and complex interactions between the bottom and waves. The concept and sequences of wave dampening have been synthesised by Winterwerp, de Graaff, Groeneweg, and Luijendijk (2007). The cyclic pressures induced by the incoming waves start by generating small elastic deformations within the seabed. As these stresses exceed bed strength, internal failures commence, resulting in the inception of bed liquefaction, a process reported by these and other authors (e.g. De Wit & Kranenberg, 1997) to be very rapid, of the order of tens of seconds and up to a few minutes at most. As further waves come in, they generate internal waves at the liquefied mud–water interface that are dissipated by internal friction within the mud layer. Winterwerp et al. (2007) state that although more sediment may liquefy below this fluid mud layer, the process may not be likely because of wave and, therefore, stress dampening within the bed, and because the stress history tends to limit the thickness of the bed that is sensitive to liquefaction. In reality, aspects of stress and liquefaction should depend on the intrinsic properties of the mud, its degree of consolidation and the wave climate, with potential feedback effects on fluid bed thickness. The mud bank regime on the Kerala coast of India has been reported to be an essentially in situ phenomenon controlled by seasonal monsoonal wave energy variations. Narayana et al. (2008) reported that the surficial sediment is annually entrained during the monsoon, but erosion is limited by the formation of a benthic fluid mud layer, which attenuates wave-generated turbulence. Although some fine

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sediment disperses alongshore and offshore, most is returned to the seabed as the monsoon declines. Tatavarti and Narayana (2006) carried out measurements of waves, cross-shore currents and longshore currents over a period of approximately one year that showed significant differences in the nearshore hydrodynamic regime (i.e. wave and current characteristics) between events of wind and wave activity during the non-monsoon and monsoon seasons (Figure 4.2). The non-monsoon season, relatively free of suspended sediment loads, was characterised by progressive edge waves in the infragravity frequency band with weak reflections, while the hydrodynamic regime of the monsoon season was marked by the presence of far infragravity waves, infragravity waves (leaky modes and trapped edge wave modes) coupled with strong shoreline reflections and undertow. The non-linear wave–wave interactions were noticed to be more pronounced in the upper water column, progressively diminishing vertically down toward the seabed and horizontally toward the shore. These intense wave–wave processes during the high-energy monsoon season are responsible for the formation and sustenance of mud banks. During the monsoon, the mud bank regime is characterised by enhanced turbidity and by the benthic fluff layer, due to shear stresses exerted by swell during the early monsoon period, probably with gas being forced into the surficial sediments either by wave pumping or by seaward-flowing sub-bottom freshwater derived from monsoonal rains (Narayana et al., 2008). With the waning of the monsoon, the benthic fluid mud layer rapidly disappears and the seabed returns to its pre-monsoon state as suspended sediments are redeposited. Wave refraction over the shoreface bathymetry may lead to shore-fast mud bank formation at known locations of wave energy concentration, but such mud is restored back to the shoreface mud bank reservoir during the non-monsoon season (Mathew & Baba, 1995). Under significant wave action, fluid mud can, thus, be advected shoreward against gravity by wind-induced onshore surface currents and by Stokes’ drift ( Jiang & Mehta, 1996; Rodriguez & Mehta, 1998, 2001; Gratiot et al., 2007; Winterwerp et al., 2007). Waves maintain the fluid mud in suspension but the wave height decreases dramatically with distance shoreward due to energy absorption by mud, and the effect of breaking becomes much less significant than during the initial phases of wave motion. This important energy dampening effect of thick mud beds on waves has been demonstrated in a number of studies (e.g. Wells & Kemp, 1986; Mathew, Baba, & Kurian, 1995; Jiang & Mehta, 1996; Sheremet & Stone, 2003), and has been apprehended (Figure 4.3) in a numerical wave model (Winterwerp et al., 2007). Wells and Kemp (1986) measured a dissipation rate that grew from 88% to 96% for wave heights at three muddy shoreface locations off the Surinam coast over a distance of about 7 km in water depths that decreased from about 7 to 3 m. They highlighted the dissipation of both short- and longer-period waves, although the latter underwent greater dampening. Jiang and Mehta (1996) calculated energy loss, energy storage and wave attenuation coefficients for three typical mud densities and found a clear relationship between energy absorption, wave damping and bottom liquefaction. Wave attenuation levels can be extremely high in association with fluid mud, with values of up to 96% attenuation over a mud bank in Surinam (Wells, 1983). Sheremet and Stone (2003) compared wave dissipation rates over sandy and muddy portions of the Mississippi delta shoreface and

138 Shore Processes and their Palaeoenvironmental Applications

Figure 4.2 Sketches of mud bank extension during the non-monsoon and monsoon seasons on the Kerala coast in India. The calm sea area corresponds to one of quasi-total dissipation of wave energy. Adapted from Tatavarti and Narayana (2006), with permission from the Coastal Education and Research Foundation.

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Figure 4.3 Computed wave spectrum at three locations of the Amazon-in£uenced Demerara coast of Surinam showing extreme wave dampening shoreward over the muddy shoreface and changes in spectral frequency. Solid and dotted lines show, respectively, conditions with and without locally generated waves. Note the di¡erent scales. Adapted from W|nterwerp et al. (2007), with permission from Elsevier.

observed wave heights 70% lower over the muddy bed, and attributed this to enhanced attenuation. They also noted that dampening affected the entire wave frequency, thus suggesting that even the short waves were affected by the muddy bed. Gratiot et al. (2007) reported typical ‘bank’ and ‘inter-bank’ profiles and corresponding mud densities, and a three-month record of changes in the thickness of the fluid mud layer in an estuarine navigation channel, as well as an 80-day record of bed-level changes in the intertidal zone using a pressure transducer. Waves did not deviate significantly from the 2nd order Stokes theory up to about 5 m water depth (11–13 km offshore), but are totally dampened at water depths less than 1 m (6–8 km offshore). This wave action on mud mobility has also been highlighted by

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Froidefond et al. (2004) from estimates of suspended particulate matter using SPOT satellite imagery. The episodic nature of high wave-energy events generally results in the formation of mud bar features (further discussed in Section 4.4) from the shoreward mobilisation of gel-like fluid mud, in good agreement with a scenario proposed by Jiang and Mehta (1996) for wave–bank interaction. Because of thixotropic properties, the cyclic pressure gradients liquefy the 1–3 m thick mud layer which is then transported en masse by shoreward wave drift due to wave asymmetry, and alongshore, due to the wave height gradient associated with oblique wave incidence. The effect of waves on mud is particularly marked following long periods of low energy, and especially at the onset of the high wave-energy season (October to May), when even moderate wave-energy events can lead to significant mobilisation of mud. The transport of sediment under these circumstances becomes due largely to fluid mud flow, as opposed to upward entrainment of sediment into the water column, because most of the suspended mass tends to reside near the bottom. Once wave action ceases, the profile may become hardened once again through gelling and self-weight consolidation but, these processes are relatively long and involve drying out of the substrate (Fiot & Gratiot, 2006).

4.3.2. Mud Bank Mobilisation and Alongshore Mud Diffusion The migration of highly turbid suspensions and of distinct mud banks is an important mechanism of longshore diffusion of mud to shores located considerable distances away from source zones. Once mud banks are formed through the processes evoked above, their migration is generated by waves, winds and tidal currents, with waves playing a determining role in maintaining sediment suspension in the coastal zone. Following Wells and Coleman (1978, 1981), a number of theoretical efforts on mud transport by waves and a few field investigations have suggested a leading role for wind-generated waves in mud bank mobility ( Jiang & Mehta, 1996; Rodriguez & Mehta, 1998; Tatavarti & Narayana, 2006; Gratiot et al., 2007; Chevalier, Froidefond, & Devenon, 2008). Wave liquefaction of mud, the mechanisms of which have been discussed above, comprises an alongshore transport component that is fundamental to mud bank migration. Apart from variations in sediment supply from the Amazon, changes in the intensity and direction of the trade winds and their effects on waves have also been held responsible for temporal variability in mud bank migration rates (Eisma et al., 1991; Allison et al., 1995, 2000b). Eisma et al. (1991) used the angle of shore incidence of synoptic winds as a surrogate for assessing temporal variations in the intensity of longshore drift, and, hence, mud bank migration rates. This approach was further used by Augustinus (2004) to explain both changes in rates of mud bank migration and in the lengthening of mud banks. A 44-year record (1960–2004) of the ERA-40 wave dataset generated by the European Centre for Medium-Range Weather Forecasts (ECMWF) was used by Gratiot et al. (2007), together with complementary field investigations in French Guiana, to define both event-scale and longer-term patterns of mud mobilisation induced by waves. From analyses of muddy bed profiles, fluid mud layer thickness and mud loading and their relationship with the wave data, these authors

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highlighted a close relationship between wave energy and fluid mud mobilisation, and singled out the ratio H 30 =T 2, combining wave height H and period T, and the angle of wave incidence, as the most relevant parameters for describing wave forcing. Gratiot et al. (2007) showed that significant phases of increased wave energy are attended by higher long-term (annual) rates of longshore mud bank migration (Figure 4.4), but the correlation was rather poor between the wave forcing parameter H 30 =T 2 and migration rates because stronger wave forcing is generally associated with low angles of wave incidence. This suggests a complementary role of other hydrodynamic mechanisms, such as geostrophic and tidal currents, in longshore mud bank migration (Chevalier et al., 2008; Gratiot et al., 2008). Mud bank migration rates can vary significantly. Along the South American coast, they exhibit low multi-annually averaged rates (0.2–1.8 km yr1) in the early 1980s and high rates (1.8–3.0 km yr1) from the mid-1990s to 2005 (Gardel & Gratiot, 2005). The mean mud bank migration rate of 1995–2000 was twice higher than that of 1979–1984, for instance, while the wave forcing parameter was only 4/3 higher. A first source of divergence is oblique wave incidence, which provides a mechanism for longshore streaming of mud liquefied by strong wave action. A second set of factors involves local irregularities such as nearshore bedrock outcrops and rocky headlands (Anthony & Dolique, 2004), and river channel mouths and river discharge (Gardel & Gratiot, 2005), all of which are expected to affect significantly the migration or stabilisation of mud banks. Closely related to this is the large-scale plan shape of the coast itself, which, in many areas, comprises alternations of mild capes and embayments that should affect wave drift gradients alongshore, especially during inter-bank phases. The overall dynamics underlying these alternations of capes and bays are, however, not known (Lakhan & Pepper, 1997), although the high-angle wave instability mechanism of Ashton, Murray, and Arnoult (2001) may be an explanatory one. A third source of divergence is the rheology of the mud banks. The rheological behaviour of the mud shows a strongly non-linear and thixotropic response to stress (Fiot & Gratiot, 2006). Beyond a threshold forcing, the apparent mud viscosity decreases considerably, and this would, in turn, induce an increase in mud bank migration rate, due to the increase of wave forcing. Finally, mud bank migration must also be conditioned by a combination of other lower-order forcing mechanisms, notably geostrophic forcing associated with the Guianas current, tidal currents propagating northwestwards, density currents due to the Amazon freshwater discharge, the effect of impinging wind stress on the shore and the generation of compensatory northwestward flows due to north to northeasterly winds during the active trade wind season. Currents generated by wind stress would depend not only on wind velocities and incidence relative to the coast but also on shoreline morphology.

4.3.3. Mud Banks and Shoreline Dynamics: The Roles of Waves and Intertidal Drainage Where these mud banks come into contact with the shore, it has been shown that their surface may be characterised by marked topographic heterogeneity (Anthony et al., 2008a). There is a clear differentiation, however, between a lower intertidal

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Figure 4.4 Mud bank migration rates and wave dynamics in French Guiana: (a) longshore mud bank migration rates between Cayenne and Kourou (Figure 4.1), from 1979 to 1983 (based on aerial photographic interpretations by Froidefond, Pujos, and Andre´ (1988)), and from 1992 to 2002 (based on satellite image interpretation by Gardel and Gratiot (2005)); (b) bank and inter-bank mangrove shoreline evolution trends between Cayenne and Kourou from 1988 to 2002 (based on satellite image interpretation by Gardel & Gratiot, 2005); (c), (d) inter-bank and mud bank pro¢les and schematic wave attenuation patterns. MWL is the MeanWater Level and MTR the Mean T|dal Range deduced from tidal signal series; (e), (f ) associated sediment surface concentration pro¢les; the circle diameter is representative of the vertical error bar. Adapted from Gratiot et al. (2007), with permission from Elsevier.

zone (below MWL) characterised by relatively regular linear bar features, and an upper intertidal zone (above MWL) exhibiting a much more intricate topography of highs and lows associated with significant drainage channel activity (Figure 4.5). Lefebvre, Dolique, and Gratiot (2004) had highlighted earlier, from a combination

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a

b

c

Figure 4.5 Digital elevation models (DEMs) and representative topographic pro¢les of the Macouria mud bank in French Guiana: (a) SPOT image source, (b) Lidar image, (c) ¢eld survey. MHWL: Mean high water level, MLW: Mean water level, MLWL: Mean low water level. Adapted from Anthony et al. (2008a), with permission from Elsevier.

of aerial photographs and field monitoring, the presence of a narrow linear topographic high over a mud bank near the border with Brazil. The linear features are clearly identifiable from SPOT images and low-flying aircraft and are commonly dissected by channel networks (Figure 4.6). As incoming waves are progressively dissipated by the mud, the fluid mud deposits are pushed shoreward over the neap– spring cycle, forming the distinct bar-like accumulation features. Gratiot et al. (2007) showed from concurrent time series of wave heights, fluid mud thickness and in situ bed-level changes from a pressure transducer from two

a

b

Figure 4.6 A SPOT image taken on 17 October 2006 (a) and an oblique aerial photograph taken on 18 December 2006 (b) showing linear bar features characterising wave-dominated mud bank topography on the trailing edge of the Macouria bank. These bar features are drained by channels sourced by tidal discharge and dewatering of mud.Triangle in (a) delimits photo in (b) which also corresponds to the Lidar survey area of Figure 4.5 and highlights erosion of the trailing edge of the bank, resulting in a £at consolidated bed and an o¡set between the distal edge of the bar and the eroding proximal edge and the terrestrial shoreline. Note the more numerous drainage channels cutting across this outer bar feature compared to the earlier (2003) Lidar image in Figure 4.5. Adapted from Anthony et al. (2008a), with permission from Elsevier.

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mud bank sites that the bar-like features on these South American mud banks are formed from gel-like fluid mud in the intertidal zone at locations where wave dampening is complete. With the shoreward cessation of wave forcing through progressive energy dampening, mud reworked from the subtidal and lower intertidal zones forms the afore-mentioned linear gel-like bodies. At the same time, the wave processes may leave behind a flat consolidated, furrowed bed. The successive bands of linear bar-like features may reflect successive phases of wave-induced shoreward transport of mud under seasonal variations in wave energy, in combination with neap–spring tidal range variations. These linear shore-parallel bar-like accumulations are typical of wave-formed shore bodies, such as those commonly found in sandy beach environments. The fundamental difference here, however, is that these features are formed from wave drift of suspended gel-like mud that rapidly settles out in areas of complete wave dissipation. Although these linear features generally occur as shore-parallel bodies in the inner mud bank areas near the terrestrial shoreline, as depicted by the large-scale DEMs, bar-like features with an angular offset relative to the terrestrial shoreline are observed at the eroding trailing edges of mud banks, where they are reworked by the obliquely incident northeasterly trade wind waves affecting the Guianas coast. The more complex mud bank profiles reflect a primary control by waves, with subsequent closely-related influences resulting from topographic feedback on patterns of mud settling during the tidal excursion, consolidation processes by evaporation and dewatering, and dissection by intertidal drainage channels. The linear bar-like features are formed from gradual accumulation of fluid mud inshore under the influence of incident waves, with a predominantly tidal modulation of the vertical excursion of wave activity. They show marked cross-shore variations in the degree of consolidation that reflect three factors: (1) position within the intertidal frame, (2) trapping of fluid mud in depressions between bar-like features, and (3) mud remobilisation and fluidisation by wave activity. Once in the upper intertidal zone, the bar-like features become immobilised over fairly long phases of low wave energy, and, thus, progressively dry out, through changes in physical parameters involving yield stress, water loss and pore water salinity under conditions of both evaporation and dewatering (Fiot & Gratiot, 2006). This drying and compaction phase is associated with the formation of dessication cracks (Figure 4.7). Wetting and drying cycles have been shown to vary considerably with elevation within the intertidal frame (Fiot & Gratiot, 2006). These authors carried out laboratory and field investigations on a wide range of fluid to desiccated muds. Changes in physical parameters, such as sediment erodability, water loss and pore water salinity indicated progressive mudflat compaction as well as fluctuations related to the successive wetting and drying cycles. Mud cracks constituted a spectacular feature representative of the contractional stress. These ephemeral features were found to (re)open after a few days of dewatering and to (re)heal during the subsequent wetting. These observations also suggest that the bar-like features have a feedback influence on subsequent patterns of fluid mud accumulation and channel development. During the tidal excursion, the troughs isolated by the bar-like features trap very high suspended mud concentrations that progressively consolidate,

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Figure 4.7 Mud cracks on a mud£at in West Africa following successive wetting and drying cycles, the former involving pore water expulsion through evaporation. Photograph E.J. Anthony.

protected from direct wave remobilisation. Finally, mud remobilisation can lead to marked spatial variability in fluid mud concentration levels that are especially well expressed by the bar-like features in the lower intertidal zone still subject to onshore mobility. The bar-like features are also dissected by channels that serve as drains for ebb-tidal waters, discharge from dewatering of the fluid mud as it becomes consolidated, and rainfall. The combination of such drainage networks and variations in fluid mud consolidation can generate decimetre-scale variations in the elevation of the mud bank, while channel dissection results in the substitution of the linear bar forms in the upper intertidal zone by more complex topography.

4.3.4. Wave-Dominated Open-Shore Mud Banks and Mangrove Dynamics Mangrove-colonised shorelines in wave-exposed, highly dynamic shores, such as those under the influence of the Amazon (Gardel & Gratiot, 2005; Gratiot et al., 2007, 2008), or in open estuary mouths along the muddy West African coast (Anthony, 2006a) may fluctuate at significant short-term (order of months to a few years) rates of several metres to several kilometres in the cross-shore direction as shown in Figure 4.4b. These massive fluctuations depend on the longshore distribution and dynamics of ‘bank’ (dissipative) and ‘inter-bank’ (erosive) phases, notably along the Amazon-affected coast of South America. These phases are discussed in Section 4.4. In French Guiana, mangrove seedling establishment has

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Figure 4.8 Cross-shore pro¢le assemblage across the Macouria mud bank surface, from the three methods depicted in Figure 4.5, synthesising patterns of mud consolidation and mangrove colonisation. Levels of consolidation are derived from ¢eld observations in the light of both published data (Fiot & Gratiot, 2006) and unpublished data provided by Sandric Lesourd.Variations in relative mud consolidation in the lower intertidal zone (below MWL) re£ect the preponderant role of wave remobilisation and £uidisation, while consolidation and mud concentration levels in the upper intertidal zone (above MWL) re£ect both trapping of mud spilling over into troughs and depressions and in situ drying out and consolidation processes. Adapted from Anthony et al. (2008a), with permission from Elsevier.

been observed to be particularly dependent on wave-induced topographic changes, with very subtle elevation changes in the upper intertidal zone (order of a few centimetres) having a determining influence on successful colonisation (Fiot & Gratiot, 2006). Observations show that the innermost bar features commonly characterising mud bank topography, and more or less dissected by drainage channels, form a ‘suture’ zone with the muddy intertidal terrestrial shoreline (Figure 4.8). In this zone, the development of mud cracks, especially during neap tides, favours the trapping of floating mangrove propagules of Avicennia germinans as the tide ebbs, and these higher-lying areas are rapidly colonised by mangroves (Figure 4.9) with plant densities exceeding 30 per square metre. Once colonisation commences, extremely rapid mangrove growth (growth rates are up to 2 m yr1) leads to the establishment of a young mangrove fringe. Although large waves account for high rates of mud liquefaction and mobilisation, such waves may not necessarily have a destructive impact on mangroves (Gardel & Gratiot, 2005). These authors showed that in French Guiana, the period 1995–2000 was characterised by considerable wave energy increase. Mangroves in the inter-bank area underwent very active retreat (150 to 200 m yr1), but at the same time the mud bank area experienced mangrove colonisation. This embodies, at first sight, an apparent contradiction, because intense wave forcing should lead to strong mud bank mobilisation. A possible explanation resides in the

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Figure 4.9 Preferential colonisation of mud cracks byAvicennia germinans mangroves on a fresh mud bank substrate in French Guiana. These mud cracks are exploited by £oating mangrove propagules. Photograph E.J. Anthony.

active formation and shoreward migration of mud bars towards the intertidal zone, as a result of the wave energy increase. With cessation of wave forcing, this mud forms fluid gel-like bodies that accumulate in the form of mud bars, as shown earlier. Such bars constitute the primary substrate for pioneer mangrove formation, as Lefebvre et al. (2004) and Anthony et al. (2008a) have shown. Reworking of the ‘suture’ zone topographic ‘highs’ inherited from the linear bar forms by high-energy waves may lead to mud dispersal over the adjacent terrestrial mangrove substrates. Mud pushed shoreward and impinging on established mangrove swamps may lead to burial and asphyxia of mangrove pneumatophores (Ellison, 1998), and in French Guiana, this results in the death of ‘old’ mangrove trees (Fromard et al., 1998; Fromard, Vega, & Proisy, 2004), alongside of which are generally found opportunistic rapid-growth juveniles adapted to the new substrate topography (Anthony et al., 2008a). The relaxation of wave activity during the low wave-energy season enables the subsequent survival of the young pioneer mangroves. At the other end of the nearshore profile, mud reworking in the subtidal and lower intertidal zones involved in the formation of the linear bar-like features leaves behind a flat furrowed mud bank surface that will eventually be completely eroded by waves as the narrow trailing edge of the mud bank attains this contact zone.

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4.4. Meso-Scale Mud Bank and Shoreline Dynamics: ‘Bank’ and ‘Inter-Bank’ Phases Along any given stretch of shoreline, individual banks migrating alongshore are separated by inter-bank zones. The waxing and waning of mud bank activity is, thus, characterised by ‘bank’, ‘inter-bank’ and transitional phases (Anthony & Dolique, 2004). This rhythmic nature of bank and inter-bank phases has an overwhelming impact on the coast, inducing rapid shoreline accretion and/or erosion, and ecological changes involving frequent and large-scale mangrove colonisation and destruction. Mud bank and inter-bank profiles show a wave– seabed interaction pattern that is in agreement (Gratiot et al., 2007) with the conceptual muddy shore profiles proposed by Kirby (2000, 2002) and evoked in Section 3.2.8. Inter-bank areas are characterised by receding, low and concave erosion-dominated profiles of consolidated mud, while mud banks are characterised by prograding, high and convex accretion-dominated profiles of soft mud, albeit with marked micro-scale topographic heterogeneity as shown earlier. As the tidal excursion proceeds in inter-bank areas, waves propagate further inshore, and at high tide, breaking occurs on the shore, which is composed of either mangrove-colonised consolidated mud that may be progressively eroded (Figure 4.10) by waves, or of sand beaches commonly subject to overwash, and associated with chenier formation (see Section 3.3.2.2). Such inter-bank areas may be characterised by cheniers, where sand is locally available. The typical sandy features formed in wave-exposed muddy settings are such cheniers, several examples of which have been identified in studies on high fine-grained discharge rivers, such as the Huanghe (Saito et al., 2000), the Yangtze (Hori et al., 2001), in addition to contributions on the classic chenier coasts of the Mississippi (e.g. Huh, Walker, & Moeller, 2001; Draut, Kineke, Velasco, Allison, & Prime, 2005b; McBride, Taylor, & Byrnes, 2007). Chenier development over muddy substrates may lead to the development of unique but ephemeral beach deformation and collapse features (Figure 4.11) that have been well described from sandy bay beaches subject to active Amazon mud supply (Anthony & Dolique, 2006). These features appear to be part of the process of sand accumulation and adjustment to the underlying muddy substrate. Although the development of these features is hinged on the marked sedimentological and geotechnical differences between the sand and mud, their formation is not due to hydraulic processes at the sand–mud interface, such as sand piping or undermining, or to collapse of void space such as from encapsulated air within the sand body. Piping processes are well developed in the water exfiltration zone on the lower beach, and commonly generate deformation of the vertical collapse walls. The linear nature of the cracks in the sand and their strong longshore development for tens of metres in the mid-beach zone on cheniers and beaches in this muddy environment are not expected to result from void space collapse, since chenier and beach sands often form a surface layer of well packed sand that progressively increases in thickness. The mechanism that most likely explains these features is hydraulicallydriven adjustment of the underlying mud to sand loading (Figure 4.12). Adjustment of the mud-rich beach profile to sand loading in the intertidal zone appears to occur through a combination of downslope and longshore migration of the fluid mud and

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a

b

Freshly deposited sediment Consolidated layers Soft marine mud

Figure 4.10 (a) Large-scale shore erosion and mangrove destruction during an ‘inter-bank’ phase; Photograph Christophe Proisy. (b) Substrate layering pattern following the erosion and retreat of a consolidated muddy mangrove substrate in French Guiana. Fresh mud may be deposited over the marsh surface but net retreat leads to scarping and the formation of mud pebbles that are visible above the freshly deposited mud; adapted from Lefebvre et al. (2004), with permission from Elsevier.

mud dewatering. The dewatering is related to evaporation from surficial fluid mud exposed to drying at low tide, when large areas of the foreshore are exposed, and to compaction of the underlying under-consolidated mud. These two processes generate accommodation space to which the overlying sand above the water exfiltration zone responds by forming subsiding packages of non-saturated sand delimited by cracks alongshore (Anthony & Dolique, 2006). On the muddy Amazon-influenced coast, the connection of a mud bank to the shore can create an intertidal mudflat of several square kilometres in a few months, followed by very dense mangrove development in a few years, and equally rapid

Figure 4.11 Ground photographs showing beach collapse features due to subsidence of underlying mud under a cover of downdrift migrating sand on a mud-rich shore near Cayenne in French Guiana: (a, b) Re´mire beach (May, 2002); (c) Montjoly beach (February, 2003). The collapse zone is preceded downdrift (looking to the northwest in photos b and c) by a typical mud erosion zone, the eroded mud accumulating further downdrift. Numbers 1 and 2 on photos b, c, refer, respectively, to settled mud (density up to 1500 kg m3) undergoing erosion, and to freshly accumulating £uid mud (density: 350^600 kg m3) derived from updrift erosion. Adapted from Anthony & Dolique (2006), with permission from JohnW|ley and Sons.

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a

b

Figure 4.12 Schematic representation of stages of pro¢le subsidence and of formation of collapse features associated with mud burial within an intertidal beach subject to high mud in£uence. (a) Plan view: SZ: stabilised updrift beach zone, CZ: pro¢le collapse zone, ME: mud erosion, MD: mud deposition. The pro¢le collapse zone migrates downdrift (arrow) with longshore sand transport, from a high (to the left of panel) to a low wave-energy zone (to the right of panel). (b) Schematic pro¢les of the various zones shown in (a). Small black vertical and white horizontal arrows in P2 (collapse zone) indicate, respectively, subsidence and mud dewatering. MHWS: Mean high water spring tide level, MLWS: Mean low water spring tide level. Note the variations in level of the mud surface (higher mud surface in low-energy P3 area of fresh mud accumulation ^ light and dark shadings of mud in pro¢le P3 represent, respectively, freshly mobilised mud accumulating further downdrift in the MD zone, over older settled mud). Scales are approximate. Adapted from Anthony and Dolique (2006), with permission from JohnW|ley and Sons.

erosion of mangroves and their substrate (Gardel & Gratiot, 2005). These processes have created, on this South American coast, the most extensive, but also the most dynamic, mangrove coasts in the world. The large mud supply also implies that mangrove ecological dynamics are closely controlled by topographic changes brought about by mud redistribution. Stifling of older mangroves commonly occurs, for instance, as fresh mud inputs are driven ashore from the bank (Fromard et al., 1998; Anthony et al., 2008a). The shoreline progradational mechanisms discussed at length in the preceding sections are confirmed by meso-scale analyses based on radionuclide patterns. Using 7 Be, 137Cs and 210Pb signatures in sediment cores from inner (o5 m water depth) mud bank areas on the Amazon–Orinoco coast, Allison and Lee (2004) suggested that wave-generated fluid mud suspensions constitute the primary mechanism for

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delivering sediment across the intertidal zone of a mud bank, thus enabling accretion at the shoreline (Figure 4.13). Meso-scale (bi-decadal) evolution of muddy shores has been shown to be significantly affected by the nodal (18.6 yr) tidal cycle by Gratiot et al. (2008). The various effects of the nodal cycle on cyclic climatic behaviour and its various environmental spin-offs have been reviewed by Oost, de Haas, Ijnsen, van den Boogert, and de Boer (1993) who highlighted the importance of this meso-scale tidal constituent on tidal sedimentation along the Dutch barrier-inlet system.

4.5. Large-Scale Muddy Sedimentation and Clinoform Development: The Primacy of High Sediment Supply The shoreface and coastal plains of deltas with significant mud supplies provide good examples of the complexity of shore facies, a fine example being that of the Amazon-influenced coast of South America, where thick muddy shoreface accumulations have developed since the Holocene highstand, leading to muddy progradation over several kilometres. In this system, interspersed cheniers reflect periodic reworking, during inter-bank phases, of shoreface sands delivered by the smaller coastal rivers. Under the high mud-supply conditions, although sediment accumulation rates may tend to be low on the seabed beneath well-mixed plumes due to resuspension, overall seasonal to long-term deposition commonly leads to significant clinoform development (Figure 4.14). Most of the Amazon’s sediment is trapped in the frontal zone within a few hundred kilometres of the mouth, and deposited within a clinoform structure (Allison & Nittrouer, 1998), a conclusion similar to that reached by Allison and Neil (2003) for the Mississippi clinoform, and by Walsh and Nittrouer (2004) and Walsh et al. (2004) for the clinoform of the Gulf of Papua. Apart from these well-documented cases, other examples include the Ganges–Brahmaputra (Allison et al., 2003), and the Yangtze (Liu et al., 2006; Wei et al., 2007). The uppermost portion in these clinoforms is the shoreline, the aggradation of which brings the modern sedimentary deposit to sea level. The shoreline contains a succession of facies accumulating in shallow subtidal areas, intertidal mudflats and mangrove swamps. Walsh and Nittrouer (2004) have determined, from 210Pb within cores in the Gulf of Papua, variations in accumulation rates within the frame of an intertidal zonation pattern that characterises the top of such clinoform structures (Figure 4.15). Sediment accumulation in cores taken from the high-tidal zone (mangrove areas) is slower than the mid-tidal mean rate, and probably as a result of decreased sediment supply due to less frequent inundation. The largest rates were determined from cores penetrating the mid-tidal zone, while cores from the lowtidal zone suggest a reduction in sediment accumulation, probably in response to high shear stresses in this zone. Walsh and Nittrouer (2004) concluded that this accumulation pattern, in which the largest rates occur at mid-tidal depths, is consistent with the progradation of a coastal clinoform as observed on other muddy coastlines (e.g. Allison & Nittrouer, 1998).

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In the case of the Amazon-influenced clinoform, aggradation of 5–10 m has occurred, while progradation takes place by overlapping of northward-extending mud capes (Allison & Nittrouer, 1998). The shoreline deposits are prograding across topset strata of the modern subaqueous delta, which is the lowermost and most important part of the compound clinoform structure. The subaqueous delta extends to a water depth of 70 m, with a depositional break between topset and foreset strata at 30–40 m. Advective sediment input to the foreset region causes high accumulation rates, which control the geometry and progradation of the clinoform structure. In areas of low progradation, the offlapping mud sequences thin seaward, merging with a relict shelf surface that has been buried farther seaward by modern accumulation of shelf mud (Allison & Nittrouer, 1998). The clinoform is expressed as over-consolidated mud on the inner shelf (0 to 20 m) that forms a relict Pleistocene to Holocene bed surface (Pujos, Bouysse, & Pons, 1990) over which the mud banks comprising fluid and under-consolidated mud migrate in shallow inshore water depths of 5 to 20 m (Allison et al., 2000b). Where mud supply has been more important, the clinoform is much thicker but may exhibit a complex pattern of dissection by estuarine channels. It may be expected that in chenier-forming areas further west along the Amazon coast, this progradational mud wedge includes thin sheets of chenier sands. Sand bodies similar to cheniers are common in deltaic settings such as the Yangtze (Hori et al., 2001), Red River delta in Vietnam (van Maren, 2005), and Mississippi (McBride et al.,

Figure 4.13 Mud bank migration along the Guianas coast. (a.I) Schematic diagram of an early model of mud bank migration synthesised by Allison and Lee (2004) from previous work (see references in Allison & Lee, 2004). Mud banks migrate alongshore driven by oblique approach of solitary waves. The model suggests that mud banks are shore-connected and accretion results when the coastal plain is protected from wave attack by the wave dampening presence of the mud bank o¡shore. Erosion of the trailing edge feeds growth of the leading edge in the manner of a migrating bedform. (a.II) Schematic diagram of the inner mud bank-shoreline model proposed by Allison and Lee (2004). In this model, the mud bank is disconnected and sediment reaches the upper intertidal zone to generate shoreline accretion by £uid muds driven onshore during periods of coastal setup and £ood tide. Some of this sediment may return o¡shore during ebb tide £uid mud transport and/or mass £ows. Arrows re£ect the relative magnitude of sediment supply to the leading edge deposition on the inner mud bank. The largest quantity is derived from erosion of the trailing edge mangrove fringe, with additional material coming from erosion of the trailing edge and inter-bank intertidal^subtidal surface, and from updrift mud banks and the Amazon River. (b) Schematic diagram (with high vertical exaggeration of the o¡shore slopes) of a model for shoreline evolution in French Guiana based on remote sensing and ¢eld observations. Adapted from Allison and Lee (2004). The diagram shows a succession of nearshore cross-sections of stratigraphy (top to bottom) with the passage of an o¡shore mud bank.The eroded, relatively low porosity inter-bank surface (top panel) is succeeded (second panel) by leading edge mud bank deposition in the subtidal zone and the upper intertidal zone driven by £uid mud delivered onshore during phases of coastal setup. Accretion continues (third panel) in the upper intertidal zone as it translates seaward with mangrove stabilisation, but ceases o¡shore with passage of the leading edge of the mud bank. W|th continued consolidation o¡shore (bottom panel), wave attack resumes and the coastal stratigraphic package is partially removed. This partial removal indicates that there is a net coastal plain growth with each mud bank^inter-bank cycle. Note that sediments deposited in the intertidal zone can later be exposed in the inter-bank subtidal zone. W|th permission from Elsevier.

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Figure 4.14 Example of large-scale mud-dominated clinoform development. The diagram illustrates the mechanics of the Gulf of Papua (GOP) clinoform and the resulting sedimentary deposits. Adapted fromWalsh and Nittrouer (2004), with permission from Elsevier.

Figure 4.15 Conceptual model of sediment accumulation on an accreting mangrove mud bank in the Gulf of Papua. Adapted from Walsh and Nittrouer (2004), with permission from Elsevier.

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2007), as well as mud-rich shores associated with small multiple river mouths in West Africa (Anthony, 1989a, 2006a). These bodies differ, however, from true cheniers which are formed from sand and/or gravel winnowed out from a muddy substrate. In certain settings such as the Red River (van Maren, 2005), such sand bodies form discrete barriers separated from each other by distributary mouths. The barriers develop from deltaic channel distributary-mouth bar deposits on the inner shoreface. These bars progressively aggrade under the influence of waves, and notably swash processes, to finally isolate back-barrier spaces that are eventually filled by mud (Figure 4.16) raining out from the ETM. In their advanced infilled stages, such muddy deltaic plains associated with discrete barriers (or barrier islands) give a superficial impression of chenier development. Cheniers may develop gradually from wave reworking of sandy or shelly deposits, without the necessary input of storms or cyclones (e.g. Anthony, 1989a; Woodroffe & Grime, 1999). Rodrı´guez-Ramı´rez and Ya´n˜ez-Camacho (2008) have described a peculiar pattern of chenier formation from storm and tsunami reworking of shells above estuarine levees in inland positions within the tidal marshlands of the Gaudalquivir River in Spain.

Figure 4.16 (a) Geologic cross-section, and (b) shoreface delta-front sand bars in the Ba Lat (Red River) delta, V|etnam, that eventually form barriers behind which occurs muddy in¢ll, giving a super¢cial, but spurious impression of cheniers. Adapted from van Maren (2005), with permission from Elsevier.

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Allison and Nittrouer (1998) have suggested that the Mississippi–Atchafalaya system is fairly similar to that of the Amapa and Guyana accretionary muddy shoreface, although a storm-dominated sedimentary signature leads to more complex arrangements in the former, as McBride et al. (2007) have shown. Similar complexity is evinced by Asian deltaic systems (e.g. Saito et al., 2000; Hori et al., 2001, 2004; van Maren, 2005; Tanabe et al., 2006) where storm events, variations in river discharge and delta channel switching lead to significant imbrication of muddy and sandy deposits. Clinoform structures from these high mud-supply shores provide analogues of the geological record over muddy shorefaces. On other continental margins with different regional characteristics (e.g. more rapid subsidence) larger fractions of the clinoform structures could be preserved (Nittrouer, Kuehl, Sternberg, Figueiredo, & Faria, 1995). Ginsburg (2005) considers as possible candidates of fossil examples of such Amazon-type muddy shores the inner shelf and shoreline parts of the thick Tertiary deltaic deposits of the Gulf Coast and the mud rock sequence of the Devonian Catskill Delta of New York.

Further Work Mud bank shores are associated with high rates of mud supply to waveexposed open coasts and are associated with processes dominated by wave activity. Mud bank formation, migration and interactivity with the shore are associated with specific wave dampening processes that determine cross-shore and longshore mud bank dynamics. Research on the mechanisms of shoreline progradation and clinoform development necessitates a comprehension of frontal mechanisms of mud concentration, and has been significantly aided by remote sensing techniques complemented by field measurements, notably on mud deposition rates. There is a need, however, for more work on the micro-scale interactions at the wave–mud interface, as well as on meso-scale processes involved in mud bank cross-shore dynamics, longshore migration, and shoreline erosion. Better insight into these processes should lead to a clearer understanding of the linkages between cross-shore and longshore aggradational and progradational mechanisms involved in long-term clinoform development.