North West Isle of Wight
1. INTRODUCTION - References Map
The north-west coast of the Isle of Wight forms the southern margin of the West Solent, with which it has evolved contemporaneously. The West Solent occupies part of the valley of a formerly more extensive Pleistocene river system, the Solent River, which has experienced a complex history of change. (refer to separate text on Quaternary History of the Solent for full details). Three critical stages can be recognised in the evolution of the West Solent seaway, namely:
In conjunction with these spatially variable hydraulic influences, the major factors influencing coastal morphology are geology and topography. The narrow Chalk ridge exposed along the south of Alum Bay is relatively resistant to erosion and forms high cliffs, rising to 100m (Photo 5). The remainder of the coast comprises Tertiary (Eocene and Oligocene) strata, a sequence of poorly consolidated sands, silts and clays interbedded with thin and mostly soft limestones. Strata immediately succeeding the Chalk to the north dip almost vertically so that the Reading Clay and Thames Group formations have extremely limited outcrops in Alum Bay. Younger Palaeogene strata dip more gently towards the northeast and these comprise the main geological formations outcropping on this coast between Headon Hill and Old Castle Point. The coastal topography is undulating with high points at Headon Hill (120m - see Photo 3), Bouldnor Cliff (61m - see Photo 6), Burnt Wood (57m) and Gurnard Cliff (45m - see Photo 7). Small estuaries are developed in former tributaries of the Solent River that have been inundated by the Holocene transgression. These comprise the Western Yar (Photo 2), Newtown Harbour (Photo 8) and the Medina. Other minor tributaries have been truncated by post mid-Holocene recession of the coast and form short, steep gradient coastal valleys e.g. Alum, Brambles and Widdick Chines, or the marshy valleys of the Gurnard and Thorness streams (Photo 9). The latter have been partly blocked, or deflected, by the eastward growth of small gravel spits.
The combination of relatively non-resistant rock material and a spatially varied exposure to waves and currents has resulted in the formation of a predominantly eroding coastline characterised at several locations by well-developed cliffs and landslides. Headlands occur on more resistant strata that also outcrop on the foreshore to form protective ledges or platforms. In places the prominence of headlands has been accentuated by nineteenth century construction of forts and associated coast protection structures e.g. Fort Victoria, Fort Albert abd Warden Point (Photo 10). The shoreline exhibits a varied and complex sediment transport pattern due to both coastal configuration and hydraulic regime. Transport sub-cells on the open coast are separated by headlands, and each of the three estuaries has distinct, albeit small scale, circulation patterns (Halcrow, 1997).
2. SEDIMENT INPUTS
2.1 Marine Input - F1 F2 F3 F4 References Map
F1 Coarse Input at Hurst Narrows
Entry of coarse sediments into the West Solent from Christchurch Bay is normally restricted by tidal conditions at Hurst Narrows. Examination of tidal curves for Lymington, Yarmouth (Isle of Wight) and Totland reveal marked asymmetry, because the ebb flow is concentrated into a shorter time period than the flood (Webber 1980). The ebb flow is therefore considerably more rapid than the flood and transport of coarse bedload sediments (sand and gravel) is therefore likely to be in a net southeastward direction, parallel to the shoreline between Fort Albert and the Needles, determined by peak current velocities.
Coarse sediments may enter Hurst Narrows during exceptional conditions. A combination of high wave energy and a storm surge from the southwest coincident with peak flood tide velocities can be sufficient to transport pulses of coarse sediment into the West Solent against the prevailing net transport direction. This would certainly explain the growth of recurves and the extension of Hurst Point and may also supply materials to the main channel. Such a process is unlikely to operate on the Isle of Wight shores of Hurst Narrows due to shortage of mobile gravel.
The extent to which these transport pathways are significant sources of supply of sediment to beaches between Fort Albert and Alum Bay remains uncertain. Studies of the Pot Bank dredging area by Hydraulics Research (1977) identified significant coarse sediment circulation from Hurst Narrows offshore to feed Shingles Bank and Dolphin Sand in Christchurch Bay and, to a lesser extent, Pot Bank. Although much of the analysis, involving comparison of successive editions of Admiralty hydrographic charts, concentrated on Pot Bank (located south-west of the Needles) it was concluded that sediments from this offshore directed pathway from Hurst Narrows did not directly feed the beaches of the north-west Wight coast. Evidence is not conclusive because sediment throughputs may occur with no net alteration in seabed levels. A general survey of the Isle of Wight coast revealed that in this sector beaches were generally depleted, and thus concluded that there was little supply of coarse material from offshore (Barrett, 1985). A study of the potential effect of dredging of the Shingles Bank (Bradbury et al. 2003) also did not identify any onshore supply of sediments to these beaches, although it did highlight the important function of the Shingles Bank in providing shelter against waves approaching from the west.
F2 Suspended Sediment Input at Hurst Narrows
Net suspended sediment transport is likely to be into the West Solent at Hurst Narrows due to the greater duration of the flood current. Thus, it is likely that fine marine sediments and suspended clay sediments derived from cliff erosion of the west Isle of Wight and Christchurch Bay coasts become drawn into the West Solent. Remote sensing studies of suspended sediments within Christchurch Bay and the Western Solent support these conclusions (Strisaenthong, 1982; McFarlane 1984).
F3 Onshore Transport to Newtown Spits
Twin gravel spits flank Newtown Harbour entrance and comparison of a time series of both OS maps and Admiralty hydrographic charts revealed significant changes in morphology, as well as shoreline retreat, over the period 1879-1951. The sediment source for periods of spit growth was attributed to net onshore supply, involving complex sediment circulation between Solent Bank, Newtown Gravel Banks and Newtown Spits (Hydraulics Research 1977). Possible transport mechanisms and pathways are poorly understood because a phase of spit recession between 1914-1951 occurred at the same time as major growth of Solent Bank. Significantly increased bed levels over Newtown Gravel Beds between 1963 and 1973 accompanied diminution of the size of Solent Bank (Hydraulics Research 1977). This evidence suggests the following:
Other studies have revealed beach and associated nearshore changes which may indicate complex sediment transfers both on and offshore, involving possible bedload transfer of coarser sediment grades. Trott (2001) records late Iron Age and Roman artefacts that have accumulated on a gravel bank close to maximum low water at Bouldnor. As there is no evidence for the derivation of this material from cliff erosion, the tentative conclusion is that there is considerable mobility of coarse material in the inter-tidal zone. Aerial photographs (Photo 8) also reveal various gravel bars and other morphological features within the intertidal zone that could be indicative of shoreward migration of gravel from channel deposits.
F4 Suspended Sediment within the West Solent Channel
There is a significant flux of fine-grained sediment, moved in suspension, by both the flood and ebb tidal currents within the Wesstern Solent Channel. Net suspended sediment input to the West Solent is indicated by tidal conditions at Hurst Narrows (see F2), so some of this must derive from sources external to the West Solent, but there is no quantitative data available (Halcrow, 1997; 1998). Erosion of the local soft clay cliffs of the NW Isle of Wight coast is also likely to contribute suspended sediments to the channel. Much of the lower foreshore between Newtown Harbour and Egypt Point comprises fine muds and it is possible (but not proven) that these are of external marine origin (Posford Duvivier, 1999).
Tidal regimes at the mouths of estuaries and inlets in the West Solent are characterised by a rapid short duration ebb current and a more pronged lower velocity flood (MacMillan, 1955, 1956; Webber 1969, 1980; Price and Townend, 2000). This regime favours net input of suspended sediments into inlets, so that tributary estuaries and creeks flanking the West Solent are subject to progressive infilling and are flanked by mudflats and accreting saltmarshes (Photo 2 and Photo 8).
A sequence of dominantly fine-grained estuarine sediments, up to 14m thick, has been described for the western Yar estuary (Devoy, 1987; Tomalin, 2000) representing pulsed (unsteady) sediment input over the past 7000 years of sea-level transgression. This may have a marine source, but no mineralogical analysis has been undertaken to confirm this. Dredging at 4,000m3a-1 is required to maintain the approaches to Newport Harbour on the Medina estuary, which strongly suggests that suspended sediment input remains a contemporary process (The Harbour Master, Newport 1991). Dredging has also been undertaken in response to slow but progressive siltation in Yarmouth Harbour (MacMillan, 1955; Western Yar Liaison Committee, 1998), although in this case the tidal prism, which has been reduced by piecemeal land claim since medieval times, provides a possible explanation.
2.2 Fluvial Input - FL2 References Map
FL2 Fluvial input to the Western Yar Netown Harbour and Medina Estuaries
Rivers on the Isle of Wight are small due to limited catchments and therefore contribute negligible sediment to the coast. Rendel Geotechnics and the University of Portsmouth (1996) estimate that all of the rivers discharging sediment to this coastline potentially contribute some 2,450 tonnes a-1 of suspended load and 740 tonnes a-1 of bedload material. However, various barriers and regulation of flows reduce the delivery volume very substantially. The River Medina has a mean flow of 0.5m3s-1 and this comprises only 0.67% of the tidal volume entering at the mouth during a corresponding tidal period (Webber 1978). Thus, marine sediment input to estuarine mudflats and saltmarshes must be the dominant source of supply and fluvial sources are considered to be relatively insignificant. Several small coastal streams, e.g. Gurnard and Thorness, have been partially or wholly infilled behind spits that have grown across their mouths. It is not clear if this represents marine or river-derived sediment. If present day spits are the product of breaching of medieval or earlier barriers then there could have been a significant earlier phase of trapping fluvial sediment (Tubbs, 1999). Conversely, it could be that spits increasingly grow across inlets when marine infilling has reduced the flushing effect of their tidal exchange.
2.3 Coast Erosion - E5 E6 E7 E8 E9 E10 E11 E12 References Map
Much of the northwest Wight coast is subject to active erosion, but its morphology varies spatially from simple high-angle cliffs, as at Colwell Bay, to compound slopes with multiple scarps and intervening degradation zones, e.g. Headon Hill. This is principally related to the mechanisms of mass movement and slope failure. A coastal landslide can be regarded as a flow of sediment from an area of elevated topography to the foreshore. Slope instability and a semi-continuous sediment cascade is maintained by basal erosion (e.g. Photo 11) which can act in two ways: (i) degraded materials are removed from the base of the slope, which prevents a stable slope angle being achieved; (ii) basal erosion of in-situ strata can undercut the cliff toe so that the slope is steepened to a greater repose angle than would naturally be maintained by the ground-forming materials. From a coastal viewpoint the result is the same, in that sediment is supplied to the littoral zone, and, assuming it is removed thereafter, the coast retreats.
The type and rate of coastal slope retreat is controlled by the geology and hydrogeology of outcropping strata, and antecedent topography (height of the coastal slope), thus promoting slope failure through various slide and slip mechanisms (Hutchinson and Bromhead, 2002). All these factors vary spatially, so rates of retreat and volumes and grades of sediment input are also non-constant. Reports of past coastal erosion and landsliding reveal similar rates of activity and landform development to the present day situation (Norman, 1887; White, 1921; Colenutt, 1938; Moorman, 1939). Thus, it is likely that this coast has retreated throughout much of the late-Holocene period following the establishment of interconnection between the West Solent and Christchurch Bay. Evidence of this is provided by recognition of an ancient landslide deposit, extending up to 100m offshore, from a foreshore lobe of boulders off Brickfield Farm (Munt and Burke, 1987).
Spatial variation in sediment yield from eroding cliffs is also, in part, a function of the contrast in hydraulic regime east and west of Fort Albert. To the east, dominant waves are fetch-limited, whilst westwards the more open coast receives attenuated and refracted swell as well as locally propagated waves. There is no routinely monitored data on incident wave heights and periods (New Forest District Council, 1998-2000) on which to base any quantitative comparisons. H.R. Wallingford (1999) undertook numerical modelling of modified swell waves for Totland Bay, using HINDWAVE applied to synthetic data. For an annual return period, Hs (mean) was computed to be between 0.22 and 1.71m, depending on wave approach. For a 1 in 10 year frequency values are between 0.33 and 2.05m.
Overall, the longer-term retreat of this cliffed coastline has widened the West Solent estuarine channel and contributed a substantial input of fine sediment to its tributary estuaries. It is probable that much of the finer grained sediment stored in the West Solent itself comes from the same source, but nothing is currently known about residence times and supply pathways. Posford Duvivier (1997) estimated that the eroding cliffs and platforms between Sconce Point and Gurnard Bay currently yield 150-200,00 m3a-1 of fine sediment, very little of which is vailable to littoral transport, but which may provide (or provided) a source of supply to estuarine mudflats and saltmarshes in the Western Solent. By contrast, the annual yield of coarse sediment is considered to be less than 500m3.
E5 Chalk Cliffs at Alum Bay (see introduction to coast erosion)
These cliffs (Photo 5) comprise the northern face of the Chalk ridge, which terminates at the Needles. The Chalk is significantly more resistant than other geological units outcropping further northeast but is nevertheless subject to erosion, albeit at slow mean rates in the order of 0.1ma-1 (May, 1966; Halcrow, 1997; Posford Duvivier, 1999). It should be noted that the recession processes is episodic with major cliff falls and long intervening periods of little activity. Erosion takes place by basal undercutting followed by periodic localised falls that generate temporary accumulations of scree at the cliff toe. The cliff face then retreats very slowly by sub-aerial processes until marine erosion removes the debris at the toe and another cycle of undercutting can begin. Several large falls have occurred in recent decades causing localised recession of up to 10m within single events (Photo 5)
The significance of the Chalk is that it contains insitu flint nodule bands, which are released as irregular gravels that become abraded to form beach pebbles. However due to the short fromtage and modest retreat rate the overall supply is quite small. An estimated shoreface erosion rate of 3mma-1, combined with the above recession value, would yield approximately 100m3a-1 of coarse flint debris (Posford Duvivier, 1999). The relative absence of other durable lithologies in the cliffs between the Needles and Warden Point make them the most important gravel source for local beaches (Lewis and Duvivier 1962, 1973), especially in Alum Bay.
E6 Alum Bay and Headon Hill (see introduction to coast erosion)
In the south of Alum Bay, Reading Beds and London Clay dip steeply (75 degrees to 85 degrees), but the outcrops of the Bracklesham and Barton Groups are wider because of a rapid reduction in dip angles as the Isle of Wight monocline fold levels out northwards. All strata in Alum Bay are soft and easily eroded, comprising clays, sandstones and occasional grit and pebble horizons. The near vertically inclined strata in the south of the bay are primarily sandy and form relatively steep simple cliffs that fail by rockfall. Exceptions are the Reading and London Clay outcrops immediately north of the Chalk where mudslides and a wider less steep degrading coastal slope has formed. These materials are supplied to the foreshore by cliff falls, flows and mudslides (Hutchinson, 1965; Hydraulics Research, 1977) and gulleying (Gifford and Partners, 1994). Northwards, alternating sands clays and limestones form units of differing resistance and permeability generating deeper seated landslides and giving to a wide degradation zone incorporating benches and scarps towards and around Hatherwood Point on the western flanks of Headon Hill. Headon Hill rises to 120m and is underlain by Oligocene age Headon Beds, Osborne Beds, Bembridge Limestone, Bembridge Marls and a thin cap of Pleistocene Plateau Gravels. The varying resistance and permeability of these strata have led to development of a complex coastal slope, with mudsliding over a series of partially concealed scarps and both translational and deep seated failures, especially towards the cliff top (Hutchinson, 1965, 1983, Hutchinson and Bromhead, 2002). The cliff top and toe environments are partially "decoupled" by the interposition of the degradation zone.
A wide range of sediment grades is supplied to the shore by these processes. Little quantitative work has been undertaken, but analysis of the lithology of Headon Beds yielded a composition of 20% sand, 20% limestone and 60% clay (Lewis and Duvivier 1973). The other beds are predominantly clays and sands with a major limestone unit and small quantities of gravel from the superficial drift deposits. The limestones are of significance for they break down into joint-controlled boulders and thus provide some protection to the toe of the coastal slope (Hydraulics Research 1977). There has been no quantitative estimation of their residence time, but this is probably limited due to the relatively low durability of these limestones. Map comparisons covering the period 1868-1963 revealed long-term cliff retreat at Alum Bay and Headon Hill of between 0.2-0.5ma-1 (May, 1966). Corresponding estimates by Halcrow (1997) for 1909-95 were 0.24ma-1 for Alum Bay and 0.69ma-1 for Headon Hill. Posford Duvivier (1997; 1999) give a rate of between 0.35 and 1.1ma-1 for the sector between Widdick and Alum Bay Chines. Total erosion yield is calculated at 110,000 m3a-1 of which 22,500 m3a-1 is estimated to be sand, gravel and limestone boulders. It should be noted that the value for coarse materials is not based on field sampling and is rather uncertain, although 500 m3a-1 is quoted for flint gravel from superficial deposits that cap the hill.
The remainder of the cliff input comprises fine sands, silts and clays that are susceptible to rapid suspended transport offshore. Only coarse sands, gravel and limestones can contribute to beach volume in the long-term and the potential availability of these materials in the cliffs is limited. Posford Duvivier (1999) conclude that the 250m wide and 10m deep shoreface is scoured to a depth of between 14 and 44mma-1, yielding 15,800m3a-1 of fine sediment. Most of this is removed offshore by suspended transport.
E7 Totland and Colwell Bays (see introduction to coast erosion)
Totland Bay has historically been subject to basal and cliff-top erosion at mean rates of 0.1-0.3ma-1 (May, 1966) and a maximum of 0.56ma-1 was recorded for the period 1907-1961 (Lewis and Duvivier, 1962). Historical map comparisons by Halcrow (1997) indicate a long-term mean of 0.38ma-1 for the period 1909-1961 immediately preceding sea wall construction. A series of cliff falls in 1960-61 led subsequently to sea wall protection of the cliff toe throughout the bay from Widdick Chine around Warden Point to Colwell Chine. Installation of cliff drainage at selected points to prevent or reduce future cliff top instability has been in progress since 1925, with final completion of more comprehensive cliff stabilisation in 1998, (Lewis and Duvivier, 1973; Posford Duvivier, 1989, 1991, 1993; HR Wallingford, 1999). It should be noted that significant instability continues within some cliff sections and results in occasional extension of debris lobes across the esplanade e.g. winter of 2000/01.
The southwest part of Colwell Bay has been fully protected by a seawall since 1993. The Headon and Osborne Beds, which form the cliffline in the remainder of the bay, are subject to active erosion at their toes. The geological units of the cliffs comprise gently northward dipping sands and clays with occasional soft limestones, which promote seepage erosion and landsliding. In the south, cliff profiles and regarded and vegetated, but north of Colwell Chine simple steep eroding profiles are characteristic with a tendency for increased landsliding and wider degradation zones towards Fort Albert. Cliff morphology may follow a cyclic pattern of response to marine undercutting of the toe that results in cliff failure. Marine processes must then excavate protective basal debris produced by failures before another cycle of toe undercutting and cliff failure can begin. Rising topography and increasingly clayey lithological units of the Cliff End Member of the Headon Hill formation complicate conditions towards Fort Albert.
A variety of estimates are available for the mean long-term (100-120 year) recession rate: 0.3-0.6ma-1 (Hutchinson, 1965), up to 0.45ma-1 (Hydraulics Research, 1977), 0.10-0.60ma-1 (Lewis and Duvivier, 1962; 1981), 0.5ma-1 (Lewis and Duvivier, 1986; Posford Duvivier, 1989), 0.6ma-1 (Barrett, 1985) and 1ma-1 for cliff top retreat (McInnes, 1994). Historical map comparisons by Halcrow (1997) indicate a long-term mean of 0.32ma-1 in southern Colwell Bay for the period 1866-1975 covering the period prior to full protection. A mean of 0.52ma-1 is indicated for the central bay (1909-1975) with 0.93ma-1 for the section at Fort Albert.
Differences are due to measurement accuracies and the various time periods covered by map analysis, but all indicate consistent long-term retreat with faster rates operating towards the north of the bay. Recent erosion rates suggest faster than average recession in the Brambles Chine area and especially towards Fort Albert (Posford Duvivier, 1991). Retreat between 0.5 and 1.0ma-1 was recorded for the period 1970-85 (McInnes, 1994; Posford Duvivier, 1997) and maximum short-term retreat of the cliff-top was recorded at 1-2ma-1 (Lewis and Duvivier, 1986; Posford Duvivier 1989, 1991). The fine sands and clays yielded have little stability on the beach and much of the estimated cliff erosion input (approximately 5,000 m3a-1) is rapidly lost offshore (Posford Duvivier, 1999). An additional shoreface erosion rate of 17mma-1, yielding 7,000 m3a-1 of fine sediment is also proposed.
E8 Fort Albert to Fort Victoria (see introduction to coast erosion)
This coastal sector comprises a relatively low angle coastal slope degrading primarily by mudsliding in lower parts with some upper parts thickly vegetated and relatively inactive (Hutchinson, 1965; Lewis and Duvivier, 1973; Posford Duvivier, 1990b; Halcrow, 1997). A sea wall protects the cliff toe for 200m to the northeast of Fort Albert, but there is considerable instability of the slopes behind. Along the remainder of this unit, the soft clays at the cliff toe appear to be eroded faster than the rate of supply of material from mudslides, thus some lower slopes are oversteepened and controlled by shallow failures (Halcrow, 1997). Serial map comparisons do not indicate any discernible cliff-top erosion, possibly due to the thickly vegetated and complex morphology of the upper slope (Lewis and Duvivier, 1973). Despite this, long-term toe erosion at 0.5ma-1 has been calculated (Lewis and Duvivier, 1981; Posford Duvivier, 1989, 1990b, 1997; Halcrow, 1997). It would appear that aggressive toe erosion is leading to progressive reactivation of relict landslides upslope, so that the scale of landsliding is likely to increase in future as the full slope becomes active.
The geology of the coastal slope is obscured by vegetation and disturbed by landsliding, but White (1921) and geological maps indicate Headon and Osborne beds overlain by Bembridge Limestone and Marls, so cliff erosion input must be predominantly clays with some sands and soft limestones (Halcrow, 1997). Posford Duvivier (1997) estimate an annual cliff erosion yield of 5,000m3. It is reported that small quantities of gravel are also supplied (Lewis and Duvivier, 1973, 1981). This coast is more sheltered from wave erosion than areas to the west, but is swept by rapid tidal currents of Hurst Narrows so relatively little beach material accumulates. The shoreface between these Fort Albert and Fort Victoria is some 250m wide and 20m deep; given an estimated 0.5ma-1erosion rate, the yield of fine sediment is approximately 7,000 m3a-1 (Posford Duvivier, 1999). For the shoreface between Fort Victoria and Bouldnor, the respective values may be in the order of 1mma-1 and 3,000m3a-1.
E9 Bouldnor and Hamstead Cliffs (see introduction to coast erosion)
Between Bouldnor and Hamstead Ledge, cliffs rise to 61m at Bouldnor Cliff (Photo 6) and 35m at Hamstead Cliff. The coastal slope is underlain principally by gently northward dipping clay-rich Hamstead Beds of the Bouldnor Formation (White, 1921; Daley and Insole, 1984). It exhibits complex morphology and degrades by deep-seated rotational slides at the backscar and by mudsliding within extensive mid and lower mudslide dominated terraces. Morphology comprises a steep upper cliff, with several embayments feeding small mudslides. These move over a series of terraces formed by more resistant limestone and converge to form a major mudslide lobe (Photo 8), which periodically surges up to 100m across the foreshore (Munt and Burke, 1987). Mudsliding is long established and is recorded back to at least 1913 (White, 1921; Hutchinson, 1965; Moorman 1939 and Posford Duvivier 1995). These landforms have been classified as relatively shallow, multiple translational slides (Bromhead, 1979). Mudslide movement is seasonal and controlled by precipitation, groundwater availability and enhanced porewater pressures generated by undrained loading at the head of the mudslide (Hutchinson and Bhandari, 1971; Bromhead, 1979) and seepage erosion. It has been postulated that enhanced porewater pressures has greater effect on initiating a slide than toe erosion by marine processes (Bromhead, 1979). This could explain the historical and present day instability and rapid mudsliding despite limited wave energy available for toe erosion. However, active undercutting of the cliff toe operates in many places (Photo 11) and mudslide instability is maintained by marine erosion of lobes as they extend seaward.
The nature of landsliding varies spatially, with a zone of highly developed and active mudslides at Bouldnor Cliff, repetitively triggered deep-seated rotational failures to the west, and less well developed, superficial mudslides to the east. The level of relatively resistant units at the base of the Hamstead Member and the top of the Bembridge series are identified as the critical control of this variability (Halcrow, 1997). Rock dips are locally reversed by faulting so that underlying Bembridge Marls and Bembridge Limestone rise to beach level in the northeast of this sector (White, 1921). Thus, at Bouldnor Cliff this horizon lies at 1 to 3m above mean sea-level and provides optimum conditions for mudsliding. This persistent tendency for shallow mass movement has apparently increased in both magnitude and frequency of events here since the mid 1980s (Posford Duvivier, 1995). To the west the resistant horizon is not exposed and the soft clays exposed at beach level are rapidly eroded at rates in excess of mudslide supply. The coastal slope becomes oversteepened, facilitating deep-seated failures. To the east, resistant strata rise well above beach level and increase the resistance of the base of the slope to marine erosion so that recession of is less rapid and mudslides less well developed (White, 1921; Hutchinson, 1983).
Mean long-term cliff-top retreat over the period 1868-1963 was 0.61ma-1 (May, 1966; Posford Duvivier, 1997), but a high rate of 3ma-1 was recorded for a part of the Bouldnor Cliff complex over the period 1922-1962 (Hutchinson, 1965). Historical map comparisons by Halcrow (1997) indicate long-term (1909-1995) mean cliff top recession of 1.13ma-1 for western and central Bouldnor and 0.84ma-1 for Hampstead Cliff. Although, map comparisons covering the period 1908-1971 indicated locally rapid recession of mudslide lobes toe at rates of up to 1.6ma-1 (Webber, 1977), it appears that cliff top recession has been more rapid than recession of mean high water at the toe leading to an overall flattening of the slope profile (Halcrow, 1997).
Cliff recession yields significant sediment volumes, but much is clay and silt so only a small proportion, estimated at 15% (Bray and Hooke, 1997), of total cliff input is stable on the beach. Posford Duvivier (1997) estimate a total annual sediment yield of 65,000m3, of which less than 500m3 is gravel. Some gravels are supplied from Pleistocene cliff-capping coarse deposits (Hydraulics Research, 1977; Posford Duvivier, 1995; Halcrow, 1997) and Moorman (1939) reported gravel scree beneath the steep upper cliff. Mapping and sediment sampling of the gravel outcrops has not been undertaken so exact contributions remain unquantified although they could be significant on this low drift coast. The erodible shoreface materials may be scoured to a depth of 0.12ma-1, yielding some 23-25,000m3a-1 of fine sediment (Posford Duvivier, 1999), which is transported offshore as suspended load.
E10 Newtown Harbour to Pilgrims Park (see introduction to coast erosion)
For some 2km eastward of Newtown Harbour there are steep, but low eroding cliffs with basal landslide debris and fallen trees on the beach (Hydraulics Research, 1981). Cliffs increase slightly in height eastward and landsliding rather than rockfall becomes increasingly evident as the major cliff recession process. Historical map comparisons by Halcrow (1997) indicate a long-term mean of 0.73ma-1 for the period 1909-1995.
Further east, the coastal slope rises in height to 57m near Burnt Wood. At this location there is active shallow translational landsliding and transport of debris in mudslides (May, 1966; Hutchinson, 1965; Halcrow, 1997). The lower part of the coastal slope at Burnt Wood is composed of the relatively more resistant Bembridge units, while the upper slopes are composed of the clayey Hampstead Member of the Bouldnor Formation and capped by Plateau Gravels. Retreat is generally slightly less rapid than at Bouldnor to the west, perhaps due to the outcrop of the Bembridge strata at beach level (similar geological sequence as at Hampstead). Some areas of localised severe erosion were nonetheless reported by Hutchinson (1965). Mean cliff top retreat of 0.36ma-1 was measured from map comparisons covering the period 1868-1963 (May, 1966). Posford Duvivier (1999) propose a higher rate of 0.6ma-1. Historical map comparisons by Halcrow (1997) indicate a long-term mean of 0.99ma-1 for the period 1909-1995. These different estimates reflect considerable spatial and temporal variation in the recession process and also some uncertainty in the exact cliff top position due to presence of woodland and scrub.
Material supplied is predominantly clay, but a limited gravel input is also reported (Lewis and Duvivier, 1981, Halcrow, 1997). The latter is probably limited to a deposit of Pleistocene Plateau Gravel at Burnt Hill, although it may also derive from erosion of in situ Pleistocene gravel-bearing deposits on the foreshore (Lewis and Duvivier, 1981). These materials may be similar in composition and age to those recognised offshore of Brickfield Farm (Munt and Burke, 1987). Posford Duvivier estimate a total cliff erosion sediment yield of 75,000m3a-1 for the sector between Newtown Harbour and central Gurnard Bay. Estimates suggest that less than 500m3 is coarse material, although mapping and sampling of the gravel outcrops has not been undertaken so exact contributions remain unquantified, although they could be significant on this low drift coast. A shoreface erosion rate of 12mma-1, yielding 11,000m3a-1 of fine material, has been calculated by Posford Duvivier (1999).
E11 Whippance Farm to Gurnard; Thorness Bay (see introduction to coast erosion)
The entire coast between Whippance Farm (Thorness Bay) and Gurnard displays evidence of coast erosion, with cliffs up to 45m in height, much active mudsliding and shallow translational slides that supply and debris accumulations on the foreshore - see Photo 12 (Hutchinson, 1965; Hydraulics Research 1977, 1981; May 1966; Halcrow, 1997; Posford Duvivier, 2000, Moore and McInnes, 2002). The landform assemblage is comparable to that at Bouldnor and Burnt Wood, but smaller in scale. Recession has been measured at 0.36ma-1 for the period 1868-1963 (May, 1966) and 0.6ma-1, 1862-1938 (Hydraulics Research, 1977). Some basal protection afforded by Bembridge Limestone ledges at Gurnard Ledge, and to the east, results in some increased cliff stability and slower retreat rates slower to the northeast of the Ledge compared to the cliffs to the south. These ledges eroded by 0.6ma-1 to 1.2ma-1 over the period 1862 to 1938 which suggests that their protective capacity is limited (Hydraulics Research, 1977; Posford Duvivier, 1997; 1999). Historical map comparisons by Halcrow (1997) indicate long-term (1909-1995) mean cliff top recession of 0.48ma-1 for the cliffs to the south of the Ledge (Photo 12) and 0.18ma-1 for those to the northeast.
Map and field evidence indicates that cliff erosion supplies material from (i) the Bembridge and Osborne Beds; (ii) Plateau Gravels, which cap the high cliffs immediately south of Gurnard Ledge (White, 1921). The solid strata contribute predominantly clay sediments that are transported offshore but also some limestone boulders, which temporarily remain on the foreshore as boulder arcs that mark the seaward, limit of former mudslide surges. Posford Duvivier (1997; 1999) estimate a total sediment yield of 75,000m3a-1 for the sector between Newtown Harbour and central Gurnard Bay. Estimates suggest that less than 500m3 is coarse material, although mapping and sampling of the gravel outcrops has not been undertaken so exact contributions remain unquantified, although they could be significant on this low drift coast. The rate of inter-tidal shoreface abrasion is calculated at between 4 and 24mma-1 (Posford Duvivier, 1999), providing a yield of rapidly removed suspended sediment of 2,500 to 14,000m3a-1.
E12 Gurnard Marsh to West Cowes (Cowes Castle) (see introduction to coast erosion)
North of the small valley occupied by Gurnard Marsh, a partly active wooded coastal slope up to 35m in height is protected by revetments and sea walls, currently in generally poor condition (Photo 13). The slope continues east to West Coves, but to the east of Gurnard slipway, it becomes less steep, and is protected at its toe by seawalls and an esplanade (Photo 14). Slope morphology comprises numerous irregularities, which indicate past and active seepage erosion and the presence of relic deep-seated and shallow landslides (Posford Duvivier, 2000; Isle of Wight Centre for the Coastal Environment, 2000, Moore and McInnes, 2002). Although an average rate of cliffline recession of between 1.5 to 3.0ma-1 between approximately 1850-1950, is suggested by Hutchinson (1965), present conditions do not support such rapid recession of the entire cliff. It could be that the rates quoted relate to local areas where inactive landslides have rapidly reactivated upslope.
Geomorphological and ground behaviour mapping reveals several active mudflows and landslides that have intermittently extended downslope and surged across the foreshore of the southwestern sector of this frontage (Hydraulics Research, 1981, Posford Duvivier, 2000, Moore and McInnes, 2002). Cliff input to the sediment transport system is clearly indicated, and comprises clays and limestones from the Bembridge and Osborne Beds together with a limited quantity of Plateau Gravel and remobilised relict landslide debris.
Between Egypt Point and West Cowes the upper coastal slopes exhibit evidence of instability ,but the toe has been protected by an esplanade and sea wall since 1894, so no contemporary sediment supply occurs (Hydraulics Research, 1977; Hutchinson, 1965; Halcrow, 1997; Posford Duvivier, 2000) so long as it maintains its function. A programme of reconstruction of defence structures is in progress at Gurnard Bay, which will reduce the historical recession rate of slightly less than 0.1ma-1 (Posford Duvivier, 2000). A low shoreface erosion rate of 1,300m3a-1 (Posford Duvivier, 1999) is a function of protection from high-energy waves. It should be noted that increases in winter rainfall (effective precipitation) that are likely to result from future climate change could have serious implications as it would raise groundwater levels, potentially causing more widespread reactivation of the coastal slope along this frontage (Halcrow Maritime et al, 2001).
2.4 Beach Nourishment - References Map
Limited beach nourishment has been undertaken at several locations in response to falling beach levels so as to temporarily prevent undermining of coast protection structures and reduce the historical trend of inter-tidal narrowing (Halcrow, 1997). In all cases, volumes are small and designs governed by the perception of critical losses rather than through and systematic long term monitoring of beach profiles and volumes. The main sites are:3. LITTORAL DRIFT - LT4 LT5 LT6 LT7 LT8 LT9 LT10 LT11 LT12 LT13 LT14 References Map
Both the potential for, and actual rates of, littoral drift vary along the North-West Wight coast due to spatial changes in wave climate and the role of tidal currents. Between the Needles and Fort Albert, the coast is subject to obliquely approaching refracted Atlantic swell waves, modified by the shallow water of the western English Channel and Christchurch Bay, especially the Shingles Bank. Drift potential is thus high (New Forest District Council, 1998). North-east of Fort Albert, the coast is sheltered by Hurst Spit and the mainland so that most incident waves are fetch limited (rarely in excess of 1.3m) and of relatively low energy (Webber, 1978; Posford Duvivier, 1990; Halcrow, 1997). Thus, actual volumes and rates of drift are well below potential rates (Brampton et.al 1998). Despite this, transport potential is not uniformly low along this coast, for ebb tidal currents are rapid within the West Solent (Webber, 1980). Meandering of ebb and flood flow brings these tidal streams close to the shore at certain points so sediment transport is possible by tidal currents augmenting wave action (Dyer, 1971; Halcrow, 1997; Brampton et.al, 1998). Ebb and flood currents at the mouths of the estuaries have created local transport sub-cells (Bray, Carter and Hooke, 1995).
Beach monitoring at Alum, Colwell and Totland Bays from 1997-2002 (programme is ongoing), has revealed considerable seasonal fluctuation of profile form and volume, but relatively modest net morphological changes (New Forest District Council, 1997; 1998-2000, Bradbury et al. 2003). Trends have followed expected adjustments to changes in wave energy, although a general increase in beach cross-section areas has been identified.
LT4 Alum to Totland Bay (see introduction to littoral drift)
All authors agree that net littoral drift is from southwest to northeast, so flints released from the Chalk (E5), sand and limited quantities of gravel from Eocene rocks and cliff top Quaternary sediments are transported from Alum Bay and Headon Hill (E6) towards Totland Bay (Lewis and Duvivier, 1962, 1973, 1981; Hydraulics Research, 1977; Barrett, 1985; Posford Duvivier, 1989; Halcrow, 1997). Boulder aprons on the foreshore at Hatherwood Point and beneath Headon Hill, appear to intercept drift significantly so that only relatively small quantities of coarser materials appear to reach Totland Bay where they are intercepted by substantial groynes installed in 1993.
The transport discontinuity at Hatherwood Point, appears to confine the well-defined gravel upper beach in Alum Bay (Photo 3) where as predominantly sandy beaches occur in Totland Bay. Net offshore loss of fine sand in Alum Bay is suggested by Brampton et.al, (1998) and beach profile monitoring has revealed an overall loss of beach material over the period 1996 to 2002 (Bradbury et al, 2003).
Net drift is believed to operate from south to north within Totland Bay although very little gravel is available and only a low gradient intertidal sandy foreshore is present. Observations suggested that beach depletion was the dominant trend in Totland Bay between 1960 and 1990, but the first consistent programme of beach monitoring has revealed a gradual increase in beach volume over the period 1996-2002 (Bradbury et al, 2003). The profile analysis revealed that the beach within Totland Bay varied significantly seasonally with a greater volume being evident in summer. Its profile was also considerably more volatile than at corresponding locations in Colwell Bay and it was reported that an equilibrium profile did not form. These latter features are believed to result from the depleted state of the beach and are indicative of interaction with the seawall (Bradbury et al, 2003).
Warden Point at the northern extremity of Totland Bay is a natural headland resulting from an outcrop of resistant strata on the foreshore to form Warden Ledge, which partly intercepted littoral drift prior to sea wall and esplanade construction in the early 1990s (Photo 10). The prominence of this headland has been accentuated by the protection structures and the nearshore seabed has been lowered by beach drawdown, so that deep water now extends directly to the sea wall (Lewis and Duvivier, 1981; Barrett, 1985). Due to limited material availability it is probable that north-eastward drift of gravel into Colwell Bay is now totally intercepted (Lewis and Duvivier, 1981; Barrett, 1985; Halcrow, 1997). It is uncertain whether the same holds true for sand or whether northward transfer into Colwell Bay is possible.
LT5 Colwell Bay (see introduction to littoral drift)
Colwell Bay no longer receives coarse sediment input from Totland Bay by longshore drift, due to depletion of the latter and Warden Point acting as a drift barrier. Within Colwell Bay, net movement is from southwest to northeast as indicated by beach accumulations against groynes (Posford Duvivier, 1989, 1993; Halcrow, 1997), reflected by beach sediment grading. The beach in the southwest corner became severely depleted, an effect starting in the 1940s, whilst central parts maintained a relatively stable shingle and sand beach (Lewis and Duvivier, 1973; Posford Duvivier, 1989). This trend led to reinstatement of the beach by nourishment and the rebuilding of retaining groynes between 1966 and 1977 (Barrett 1985), but these latter structures now severely restrict drift. The first consistent programme of beach monitoring has revealed a gradual increase in beach volume over the period 1996-2002 (New Forest District Council, 1997; 1998-2000; Bradbury et al, 2003). The profile analysis revealed that the beach within Colwell Bay varied seasonally with a lower volume being evident in winter, but otherwise maintained a slowly increasing profile volume and an equilibrium form.
The northeast extremity of Colwell Bay is marked by Fort Albert, which was constructed in the mid-19th century. Subsequent coast recession and foreshore lowering has created a prominent salient here with deep water adjoining the fort. It is probable that this artificially strengthened headland almost completely prevents northeastwards drift of coarse sediment and thus promotes downdrift foreshore lowering (Lewis and Duvivier, 1981; Halcrow, 1997). Evidence therefore suggests that Colwell Bay is now an isolated pocket beach, which may only receive sediment from local cliff, or shoreface erosion and possible onshore transport. Although potential littoral transport is likely to be towards Fort Albert, negligible accretion has occurred against this barrier and sediments are concentrated in the central part of the bay (Lewis and Duvivier, 1973). Two possible explanations exist: (i) there is no net drift in Colwell Bay, except for a tendency for sediment to move away from the headlands; (ii) net drift is indeed north-eastward, but sediment is lost offshore in the vicinity of Fort Albert entrained by strong tidal currents generated at Hurst Narrows. Insufficient information is currently available to test these possibilities, although recent beach profile monitoring enables a careful check to be maintained on beach volumes.
LT6 Fort Albert to Fort Victoria (see introduction to littoral drift)
Net eastward drift of gravel is indicated along this segment by accumulation against sea walls at Fort Victoria (Lewis and Duvivier 1973, 1981), although the morphological evidence is only partial. There is a wide sandy foreshore, but corresponding sand accumulation is absent at Fort Victoria (Lewis and Duvivier, 1973). It is therefore possible that sand is progressively lost offshore to tidal currents and is transported eastward (Halcrow, 1997). Alternatively, there may be no net drift of sand, so that it becomes evenly distributed along the foreshore. Coast protection structures severely restrict drift transport at Fort Victoria, but it has been suggested that limited eastwards movement of coarse sediment was possible around the fort before it was halted by construction of two groynes over the period 1870-73 (Lewis and Duvivier, 1973). This coastal segment has therefore functioned as a self-contained unit since the pathway around Fort Victoria was denied.
LT7 Fort Victoria to Yarmouth Harbour Entrance (see introduction to littoral drift)
Drift direction is presumed to be eastward, but beach levels are low and transported volumes are extremely limited (Lewis and Duvivier, 1973). Nourishment programmes have supplied beach material immediately east of Fort Victoria and at Norton Spit, but groynes have been constructed here to retain sediment and thus drift quantities are small or non-existent (Lewis and Duvivier, 1981; Barrett, 1985; Posford Duvivier, 1989; Halcrow, 1997). The alignment of Norton Spit indicates that historically net drift has been eastward (Hydraulics Research, 1977; Dyer, 1980; McInnes, 1994), although inspection of sediment distribution against groynes fails to reveal a preferred drift direction.
LT8 Westward Drift at Yarmouth (see introduction to littoral drift)
Morphology of the mouth of the Western Yar estuary indicates littoral drift towards the inlet on both sides (Dyer 1980; Halcrow, 1997). This suggests a weak net westward drift over the sector to the immediate east of the inlet mouth, at variance with the eastward-directed littoral transport pathway that operates for most of the rest of this shoreline. A littoral transport divergence is thus implied, but it is difficult to locate precisely because of the small volume and rate of sediment movement. As it may not operate for fine-grained sediments, it is therefore a partial, and probably transient, boundary. This interpretation is based on limited evidence and is therefore of low reliability and requires verification.
LT9 Eastward Drift east of Yarmouth (see introduction to littoral drift)
It is generally accepted that net transport within the boundaries of this unit is eastward, although, quantitative evidence is lacking (Hydraulics Research, 1977; Lewis and Duvivier, 1981; Posford Duvivier, 1989). Beach levels are extremely low along this frontage and groynes are frequent (Hydraulics Research, 1977; Lewis and Duvivier, 1981) so it is likely that actual drift is currently nearly zero (Halcrow, 1997).
LT10 Bouldnor to Newtown Harbour (see introduction to littoral drift)
Supply to this segment from updrift (westwards) is negligible and most beach sediment is derived from local cliff and foreshore erosion. Consequently, beaches are often little more than a patchy veneer of gravel and coarse sand overlying an erosional surface cut into substrate materials (Photo 11). Net drift is eastward and sand and gravel supplied by this pathway is believed to have created the western spit (Hampstead Duver) at Newtown Harbour entrance (Photo 8). Eastward alignment of this spit is regarded as evidence of net eastward drift (Dean, 1995; Dyer, 1980; Hydraulics Research, 1977; Lewis and Duvivier, 1981; McDowell, 1990a and b; Posford Duvivier, 1989; McInnes, 1994; Halcrow, 1997). Observations of a major mudslide lobe, which temporarily extended across the beach beneath Bouldnor Cliff (Moorman, 1939), revealed beach accretion on its western side, and erosion of gravel and boulders from the mudslide toe, a process that continues to operate (Photo 6). It was stated by Moorman (1939) that these materials were transported eastwards from the lobe. Mud or landslide debris lobes or barriers therefore periodically impede transport across the inter-tidal zone. Small scale "surges" take place when they break down. Although recording was over a limited time period, this evidence corroborates other contextual information suggesting net eastward drift.
Hamstead Duver has shown significant morphological and planform variation according to analysis of maps and charts covering the period 1879-1973 (Hydraulics Research, 1977). Shorewards recession and recurvature into the harbour has been the dominant feature, although there are two features indicative of long-term gravel accretion. First, a relic spit is located in the harbour entrance behind the active one (Photo 8) and secondly, a gravel foreland has formed at Hamstead Point in front of low inactive cliffs. Such features would be consistent with accretion/erosion cycles at the shore caused by variation in littoral drift supply. Drift rates could have reduced recently due to a variety of reasons: (i) increased coast protection and correspondingly reduced supply along the updrift coast; (ii) temporary blockage of the foreshore by mudslides and debris accumulations between Bouldnor and Hamstead; and (iii) variation in cliff erosion input at Bouldnor and Hamstead Cliffs (Halcrow, 1997). Since the cliffs have been increasingly active in recent decades is likely that supply to the shore has increased, although there may be a lag for materials to be released from mudslide lobes and contribute to drift towards Hamstead Duver. It should be recognised that other factors may also affect dynamics of this spit, such as the tidal regime of the estuary and possible onshore-offshore sediment transfers involving gravel banks in the West Solent (Hydraulics Research 1977, 1981; Dean, 1995; Tubbs, 1999).
LT11 Westward Drift at Brickfield Spit (see introduction to littoral drift)
The spit to the east of Newtown River entrance is aligned westward, which tentatively indicates a low volume net westward drift (Dyer, 1980; Lewis and Duvivier, 1981; McDowell, 1990a and b; McInnes, 1994). It is suggested that westward drift is a local phenomenon associated with the hydraulics of the inlet entrance. Thus, there is a conjectured transient drift divide offshore Brickfield Farm. However, human modification of this coastline, involving previous attempts at land claim, may account in part for the present structure (Tubbs, 1999). The spit has a history of sediment depletion and has receded landwards over saltmarshes that subsequently became exposed and eroded in the seaward face (Halcrow, 1997). Timber groynes and revetments have been installed in past attempts to stabilise the spit, but recently it has breached to form a small new inlet subject to tidal flows at high water (Photo 8).
LT12 Brickfield Farm to Gurnard (see introduction to littoral drift)
Net north-eastward drift is indicated along this segment by eastward deflection of stream mouths by small, mixed sediment bars at Thorness and Gurnard (Hydraulics Research, 1977; Dyer, 1980; Posford Duvivier, 2000; Tubbs, 1999). Posford Duvivier, 2000. Drift is fed by local cliff erosion, with only a small proportion of sediment yield retained by beaches in front of cliffs on this frontage. A considerable quantity of gravel is stored on the upper and mid foreshore within Thorness Bay (Photo 9 and Photo 12), where it has formed a barrier across the stream and its low marshy valley. It is uncertain whether all of this material could have been supplied by drift from local eroding cliffs, or whether material could have arrived as small barrier beaches, or swash bars that have moved onshore, fed from relic gravel sources in the West Solent. Gurnard Ledge certainly functions as a partial impediment to drift tending to assist coarse sediment retention within Thorness bay, causing depletion of the beaches to its northeast.
LT13 Gurnard to West Cowes (see introduction to littoral drift)
Weak net eastwards littoral drift is reported along the depleted beach from Gurnard (Photo 13) around Egypt Point (Posford Duvivier, 1990a). Concrete rubble groynes at Egypt Point selectively intercept sediments (Photo 14), but quantities are small because of the presence of protection structures and a lack of available material (Halcrow, 1997; Posford Duvivier, 2000). Renourishment of part of this sector is currently under consideration.
LT14 Old Castle Point to East Cowes (see introduction to littoral drift)
Westwards directed, but very weak, littoral drift occurs between a drift divergence at Old Castle Point and the Shrape breakwater. The latter prevents input into Cowes Harbour. Falling beach levels and lack of significant accretion against the breakwater indicate low drift rates, which have necessitated some recent beach nourishment. The paucity of supply is due to the small source area and the impact of protection structures in reducing cliff erosion (Posford Duvivier, 1994). Cowes Harbour entrance therefore represents a drift convergence boundary, although the very small quantities of sediment moved by littoral transport towards the Medina entrance, together with the Shrape breakwater, makes this little more than a notional feature.
4. SEDIMENT OUTPUTS - References Map
4.1 Offshore Transport - O1
Rapid erosion of high cliffs along much of this shoreline yields large quantities of predominantly fine sediments. These materials are not usually stable on the foreshore, thus widespread offshore transport of fine sediments can be inferred. Little direct evidence of this process is available although the relatively rapid removal of landslide debris on the foreshore is well documented (Hydraulics Research, 1977; Moorman 1939; Posford Duvivier, 1989; Halcrow, 1997). The majority of sediments are probably transported offshore in suspension, but no precise information on pathways, quantities and ultimate 'sink' areas is available. Estimates of quantities removed annually, based on approximate measurements of shoreface width and depth and cliff recession rates, are given in Posford Duvivier (1999), and quoted in Section 2.3.
Offshore sediment transport by bedload is indicated by lobes of predominantly sandy sediments which extend offshore from the zones of active cliff input at Bouldnor and Hamstead Cliffs. These features were recognised during side-scan sonar survey and were sampled (Dyer, 1971; Fishborne, 1977). Westward deflection of these lobes or tongues suggests transport by a recirculating tidal eddy; their persistence indicates continuous sediment supply because fine sediments are not stable for longer than short periods on the seabed (Dyer 1971). It is uncertain whether offshore transport is confined to fine sediments, or whether coarser materials are also involved.
4.2 Estuarine Outputs - EO2 EO3 EO4 References Map
Throughout the Western Solent the ebb tidal flow is of shorter duration than the corresponding flood (Webber, 1980). As a result, ebb currents are of greater velocity (up to 1.2ms-1) than the flood, causing net offshore transport of coarse bedload sediments at the mouths of both larger estuaries and small tidal inlets, as well as at the western exit (see Unit Report for Western Solent). Well-defined ebb tidal deltas are not reported (esxepting Newtown Gravel Banks) and it may be that deposits have been inhibited by lack of available sediment or that they have been removed by dredging.
Dominant flow is during the ebb tide and it has been estimated that its sediment carrying potential is five times that of the flood (MacMillan, 1956; Price and Townend, 2000). No measurement of sediment transport has been undertaken to verify this statement. It is reported that sand can be transported into Yarmouth Harbour by strong northerly gales, but training of the ebb flow by breakwater structures (Photo 2) is generally successful in flushing such material back offshore (MacMillan, 1956). Maintenance dredging of the harbour and approaches is infrequent and comparison of hydrographic surveys for 1980, 1983 and 1987 revealed that bed levels were stable (Brogan, 1987). It is therefore concluded that the dominant flushing effect of the ebb current rapidly removes fine-grained sediments previously transported into the mouth (Western Yar Liaison Committee, 1998). In the past, significant quantities of sediment may have been transported across the mouth to create Norton Spit, but this is now impeded by groyne and breakwater systems either side of the harbour entrance.
It is reported that sediment mobility is greatest at the entrance of Newtown Harbour, with fine silt and clay accumulating as mudflats and marsh sediments within the inner estuary (Hodgson, 1962; Hydraulics Research, 1981; Tubbs, 1999). The bed of the main channel is composed of coarse pebbles and ebb tidal currents exceeding 0.5ms-1 have been recorded (Howard, Moore and Dixon, 1988). As a result, offshore flushing of coarse sediments may occur, fed by gravel driven by wave action along the spits flanking the harbour entrance. Although this has not been experimentally proven, the opposed alignment of these spits suggests drift convergence at the harbour mouth that would feed the losses seaward (Lewis and Duvivier, 1981). Previous research has not reported the existence of an ebb tidal delta, although the Newtown Gravel Banks surveyed by Hydraulics Research (1977 and 1981) may perform this function. It is uncertain whether coarse sediments are recycled back shorewards from these banks, although several distinctive bar-like features can be observed within the intertidal zone (see Photo 8 and also F3).
Saltmarsh erosion is currently occurring (Howard, Moore and Dixon, 1988; Raybould, et al., 2000; Bray and Cottle, 2003) and the strong ebb current may remove silt released by this process. spartina anglica 'dieback' can be traced to 1935 in the Solent, but its role in trapping and subsequently releasing sediment has not been researched at this site (Tubbs, 1999). In comparison to most other Solent estuaries, spartina loss has been limited and some areas remain accreting. In Newtown Harbour s. anglica only appeared in 1932 and has spread slowly. This site is unique in the Solent in retaining a major concentration of the native s. maritima, especially around the area of Walter's Copse. Total area of all types of saltmarsh is estimated as being 120 ha. Die-back is not reported as occurring within Newtown Harbour, indeed slow colonisation by s. anglica appears still to be continuing.
A proportion of the sediment stored in inter-tidal flats and saltmarsh is presumed to derive from input by the small rivers discharging into Newtown Harbour. Most input however, is likely to have been transported by the flood tide, and originate from cliff, platform and shoreface erosion of suspended sediment from the adjacent open coastline. The tidal prism of the harbour has not been constant, as a result of piecemeal land claim in the nineteenth and twentieth centuries, and the submergence of a previously reclaimed area resulting from a storm surge in 1954 (Halcrow, 1997).
Ebb tidal flow is of shorter duration (4 hours) than corresponding flood flow (5 hours) so ebb currents are more rapid (Webber, 1969). This produces a net offshore flushing effect of sand and gravel at the harbour entrance, which was enhanced by construction of the Shrape breakwater in 1936/37. Ebb and flood tidal flow is confined to separate channels, but the ebb flow has shifted westward as a result of the construction of the breakwater. Dominant transport sand is thus out of the harbour except along the extreme west bank, where the flood current dominates and net transport is inward (Bunce, Gibbs, Goldsmith, Jones and Spence, 1987; Posford Duvivier, 1994; Carter, 1997; Webber, 1969, 1978; Pieda, 1994). Measurement of tidal currents in the adjacent area of the central Solent indicate westward flow at high water, thus ebb currents at the harbour entrance are deflected westward and sediment transport pathways shift accordingly (Bunce et al., 1987; Price and Townend, 2000).
Examination of hydrographic charts dating back to 1856 indicate that some cyclic variations of outer estuary bed morphology may have occurred prior to construction of the breakwater, but subsequently it has been very stable (Bunce et al., 1987; Webber, 1969; Carter, 1997). This can be attributed to net offshore transport of sediment, which maintains stable channel configuration and prevents siltation even in recently dredged berths (ABP Research and Consultancy, 1994; Webber, 1969). Small sand and gravel banks exist where dominant ebb and flood flows crossover; these are probably not sediment sinks but temporary accumulation zones for sediment subject to net offshore transport (Webber 1969). In the Medina estuary upstream of Cowes, bankside erosion of marginal mudflats began to replace a longer-term tendency for channel accretion in the 1980s. Banks further offshore in the central Solent, such as Prince Consort Shoal and Bramble Bank, are probable sediment sinks (Dyer 1980) in a confluence zone receiving both wave and tidally transported sediment (Velegrakis, 2000; Bray, Carter and Hooke, 1995). Prince Consort Shoal was probably previously supplied by fine sediment flushed out of the Medina, but quantities have been significantly reduced by breakwater construction and periodic maintenance dredging (Isle of Wight Development Board and Cowes Harbour Commissioners, 1990).
4.3 Dredging Outputs - References Map
Dredging for aggregate was practised at Pot Bank, Solent Bank and Prince Consort Shoal from the late 1940s until 1994. (Hydraulics Research, 1977; 1981; Webber, 1997). Pot Bank is located several kilometres southwest of the Needles and studies indicate no direct sediment supply connection with the coast of northwest Isle of Wight. Newtown Harbour entrance may have been affected by dredging of Solent Bank, but its precise contribution is uncertain (Hydraulics Research, 1977, 1981) - see unit covering West Solent mainland for full details. Dredging of coarse sands from Prince Consort Shoal has generally been at a relatively low rate. In 1977, a licence existed for removal of 60,000 tonnes a-1, but less than 30,000 tonnes a-1 were extracted (Hydraulics Research, 1977). Webber (1977) quoted mean extraction of 50,000 tonnes per annum for the late 1960s and early 1970s. Comparisons of hydrographic charts dating back to 1912 indicate relatively stable conditions on the shoal (Hydraulics Research, 1977; Webber, 1977), which is regarded as a sediment sink (Dyer, 1980; Bray, Carter and Hooke, 1995). Dredging at this site has thus had limited impact on adjacent shores, because of its small potential to supply sediment. The low energy wave climate also makes it unlikely that increased depths caused by dredging would have much effect on local sediment budgets (Hydraulics Research, 1977). Dredging close to Cowes Harbour, however, may be of more significance, because the configuration of the harbour tends to amplify and concentrate wave energy; extreme wave events coupled with high spring tides can cause flooding at both West and East Cowes (Bunce et al., 1987; Lewin, Fryer and Partners, 1995; Webber 1978, 1981).
Dredging is also periodically undertaken for navigation purposes at Yarmouth Harbour (MacMillan, 1956; Turton, 1982; Western Yar Liaison Committee, 1998), Cowes Harbour and Newport Harbour. In all cases dredged volumes are small and predominantly comprise muds and silts. At Cowes Harbour, regular dredging was necessary to offset siltation prior to construction of the breakwater in 1936/37, but subsequent sediment removal comprising maintenance dredging of the main channel, deepening of access channels and creation of new berths, has been modest (Carter, 1997; Webber, 1969). An approximate equilibrium between loss from this source and gain from flood tide sediment input may prevail
Maintenance dredging of approximately 4,000 tonnes a-1 is undertaken in the Medina estuary upstream of Cowes Harbour to maintain the channel to Newport Harbour (Newport Harbour Master, 1991). It is reported that routine dredging began in the early 1900s but reliable historical data is lacking. For the most part, the main channel upstream to Newport is self-scouring. For details of aggregate dredging of Solent Bank, see unit on the West Solent.
Dredging at estuary entrances represents a net output from the sediment budget and may result in loss of sediments that might otherwise be transported to shorelines. Furthermore, operations close inshore potentially cause drawdown that could contribute to the steepening of local inter-tidal zones.
5. SUMMARY OF SEDIMENT PATHWAYS - References Map
6. KEY COASTAL DEFENCE AND HABITAT INTERFACE ISSUES - References Map
The following habitats are of significance within this unit: soft cliffs, intertidal foreshore (West Solent frontage) saltmarsh and mudflats (within the Western Yar, Newtown and Medina estuaries) and vegetated shingle (Newtown Spits and Thorness Bay).
The West Solent shores and its tributary estuaries are included within the recent Solent Coastal Habitat Management Plan (CHaMP) produced by Bray and Cottle, (2003). The plan identifies the present distribution and status of coastal habitats and then goes on to predict future habitat changes likely to occur up to 2001, based on an assessment of geomorphological changes. It provides guidance on habitat management and identifies any habitat creation opportunities that could compensate for future losses.
As it is not anticipated that the existing existing defences are likely to be extended to protect presently undefended cliffs, so the geomorphological, geological and ecological conservation status of currently eroding cliffs is unlikely to generate immediate issues (McInnes et.al, 1998). The upgrading of defences at Gurnard, has been undertaken using a detailed environmental assessment of the ecological character of the coastal slope. Although there would be ecological, as well as sediment budget, benefits from a relaxation of structural defences at a number of formerly freely eroding cliff sites (e.g. Colwell and Totland Bays), this is probably not an option where significant cliff top residential areas remain.
Mudflat and saltmarsh erosion is currently in progress in the West Yar and Medina estuaries, due to spartina dieback. There is no quantitative evidence as to historical and contemporary rates of loss, but estimated sea-level rise, of 2-5mma-1 over the next 50-100 years will probably accelerate the magnitude and nature of inter-tidal habitat change. In Newtown Harbour s. anglica only appeared in 1932 and has spread slowly. This site is unique in the Solent in retaining a major concentration of the native s. maritima, especially around the area of Walter's Copse. Total area of all types of saltmarsh is estimated as being 120 ha. Die-back is not reported as occurring within Newtown Harbour, indeed slow colonisation by s. anglica appears still to be continuing. In the Medina estuary, particularly between the Power Station and Dodnor, bank erosion due to mudflat recession will probably intensify, and any protection measures will induce localised coastal "squeeze".
Future changes in saltmarshes are difficult to predict with confidence for behaviour differs compared to their counterparts on the Solent mainland. At present, spartina anglica is still colonising in Newtown Harbour, although die-back has been initiated in the Western Yar and could occur in future in Newton Harbour. In the short term, the overall area of spartina is likely to increase due to continuing colonisation in Newtown estuary. Under a die-back scenario losses of some 80-90% could be anticipated within the W. Yar and 50-80% within Newtown Harbour over the next 100 years. It is likely that remaining marshes would persist in sheltered inner estuary sites. (Bray and Cottle, 2003). Saltmarsh erosion within the estuaries would release considerable quantities of fine sediment. It is uncertain whether this material would be redeposited on mudflats, or whether it would be lost from estuaries.
Relatively few opportunities are available to apply managed retreat for saltmarsh and mudflat creation due to the presence of rising topography around most of the estuarine margins. Limited opportunities exist within the W. Yar involving inundation of the southern portion of the valley to Freshwater and also the Thorley Brook valley. Together, these areas could increase the estuarine area of the W. Yar by up to 40 ha. (30-40%). It should be noted that inundation of the southern valley would significantly increase the tidal prism of all areas downstream and possibly induce tidal scour of the flanks of the existing mudflats and marshes. The Thorley valley is in the outer estuary and would not induce these impacts, except at the inlet, which is already stabilized by breakwaters. To implement a realignment of defences would threaten brackish and freshwater habitats and generate needs for their creation elsewhere, which itself would be constrained by the valley side topography.
Several very small tributary valley floors limited by topography and presently protected from inundation by embankments also exist along the flanks of the inner Medina. Although they would be unlikely to increase the intertidal area of the estuary by more than 10% they could easily double the area of saltmarsh in the estuary if they became colonised by vegetation. Although these areas are extremely small in terms of the overall Solent resource they are in close proximity to existing areas of stable, mature mid and upper marsh giving maximum opportunity for generation of high quality marsh that would have a good potential for long-term survival.
Managed retreat within Thorness Bay and Gurnard bays could allow creation of small intertidal areas controlled by topography similar in scale to the present King's Quay estuary (8 to 10 ha each). Their tidal prisms would be so small as to be marginal in stability and potentially subject to periodic closure and breaching episodes unless managed. The modest intertidal gains would be achieved at the expense of existing freshwater habitats that would become inundated.
7. OPPORTUNITIES FOR CALCULATION AND TESTING OF LITTORAL DRIFT VOLUMES - References Map
The lack of significant wave energy, modest development of natural linear beaches and prevalence headlands, reefs/ledges and debris lobes mean that shorelines of this frontage are not well suited for definitive studies of drift. There are, however, opportunities to improve knowledge of drift and beach behaviour. In particular the provision of improved monitoring of beach profiles should allow calculations of changes in beach volumes from which estimates of drift can be made. Locations especially amenable to study include:
8. RESEARCH AND MONITORING REQUIREMENTS - References Map
The SMP (Halcrow, 1997) has reviewed much of the available information and made recommendations for monitoring and research. Some recommendations are in the process of implementation by the Strategic Regional Coastal Monitoring Programme, a consortium of coastal groups working together to improve the breadth, quality and consistency of coastal monitoring in South and South East England (Bradbury, 2001). A Channel Coastal Observatory has been established at the Southampton Oceanography Centre to serve as the regional co-ordination and data management centre. Its website at www.channelcoast.org provides details of project progress (via monthly newsletters), descriptions of the monitoring being undertaken and the arrangements made for archiving and dissemination of data. Monitoring includes wave and tidal recording, provision of quality survey ground control and baseline beach profiles, high resolution aerial photography and production of orthophotos, LIDAR imagery and nearshore hydrographic survey. Not all of these actions are presently planned for this unit. Data is archived within the Halcrow SANDS database system and the aim is to make data freely available via the website.
The recommendations for future research and monitoring here therefore attempt to emphasise issues specific to the reviews undertaken for this Sediment Transport Study and do not attempt to cover the full range of coastal monitoring and further research that might be required to inform management as follows:
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