1. INTRODUCTION - References Map
Christchurch Bay comprises a shallow embayment (average depth of -7 m CD) bounded by Hengistbury Head (Photo 1) to the west and Hurst Spit (Photo 2) to the east. In comparison to Poole Bay, the shoreface is wider and shallower, promoting wave shoaling and refraction. The landward margin is pre-dominantly a currently or previously eroding cliffed coast of 20-30m in height e.g. Photo 3, and includes Christchurch Harbour confined by low spit beaches near its western extremity (Photo 1). The study area extends several kilometres offshore to depths of between -6 and -14 m CD to include the submerged rock platform of Christchurch Ledge and the sediment accumulations of Dolphin and Shingles Banks. The bay is of relatively recent origin, formed by coastline retreat during the mid to late Holocene transgression, although its contemporary geomorphology has been influenced by earlier events.
1.1 Coastal Evolution
Pleistocene evolution was dominated by the Solent River which flowed across the floors of Christchurch and Poole Bays and eastwards through the Solent (Everard 1954, Dyer 1975, Velgrakis et al 1999). During climatically-controlled fluctuating sea-levels of this period, superimposed upon possible neotectonic uplift (Maddy, et al, 2000), the river and its tributaries deposited a sequence of progressively wider gravel-floored channels and terraces which mantled the Tertiary deposits of Christchurch Bay (Nicholls 1987, Allen and Gibbard, 1993). A critical factor in the evolution of Christchurch Bay was the breaching of the Chalk ridge which previously extended between the Needles and Handfast Point, Purbeck. Opinion is divided as to the date of this event. Until recently, the prevailing view favoured breaching in the early Holocene (eg Everard 1954). Recent studies of the rates of Chalk erosion; remnant buried and infilled palaeo- channels located in a west to east sequence offshore; the chronology of the lower terrace sequence of the Stour and Avon; early human occupation sites around Christchurch Harbour, and the depth, relief and inclination of the planation surface which truncates the Purbeck-Wight Chalk ridge indicate an early to mid-Devensian breach (West 1980, Wright 1982, Nicholls 1985, Nicholls 1987, Allen and Gibbard, 1993, Brampton et al 1998, Velegrakis et al 1999, Maddy, et al 2000). An extension of the River Frome may have cut a gap through the western part of the ridge by the mid-Devensian, but later (Holocene) denudation of this feature appears to have utilised cols created by the headward erosion of rivers originally flowing southwards from its crest. It is suggested that much of Poole Bay was eroded during the mid to late Devensian, but Christchurch Bay was largely protected by the resistant strata (including ironstone concretionary seams) of a previously much more extensive Hengistbury Head. Velegrakis (1994) and Brampton et al (1998) describe a 2 m thick gravel ridge approximately 12 kms south of the Needles that both tentatively interpret as a relict barrier structure. This might define the position of the ancestral coastline in mid-Devensian times, when sea-level was between -40 and -60 m OD, prior to the final breakthrough of the chalk ridge. It was not until the early to mid-Holocene sea-level transgression (12-5000 years BP) that the Chalk ridge was removed, and the barrier between the two bays (now forming Christchurch Ledge) was cut back, and rapid erosion of Christchurch Bay proceeded (Wright 1982, Bray and Hooke, 1998b). Excavation of Christchurch Bay was also facilitated by its connection with the Western Solent between 8400-6500 years BP, which created strong tidal current scour in the eastern part of the bay (Nicholls and Webber 1987a). Simultaneously, wave and tidal current transport both out of, and towards, the West Solent entrance initiated the Shingles Bank; this has grown subsequently to its present volume of 40-60 million m3.
Both Christchurch and Poole Bays have a log spiral/zeta curve planform and concepts of crenulate bay formation have been applied, based on studies conducted in the United States (Wright, 1980, 1982). This analysis reveals that Christchurch Bay initially had an immature form which was an unstable configuration relative to a wave climate characterised by dominant waves from the south-west. An equilibrium plan shape has still to be achieved, hence continuing retreat. It is also apparent that Hengistbury Head has performed an important role in "anchoring" the planform of both Poole and Christchurch Bays, so geological controls are also important. (Halcrow, 1999; Brampton, et al, 1998). By comparison to Poole Bay, most of Christchurch Bay is more exposed to swell waves, and has a higher energy wave climate, despite the effects of wave refraction.
Cliff and shoreface erosion of Poole and Christchurch Bays, coupled with tidal scour of the western approaches to the Solent, released very substantial quantities of sediment, including gravel from river channels and valley terrace deposits as sea-level rose. It is thought that much of this material was transported eastward by littoral drift to form a succession of prototype forms of Hurst Spit (Nicholls and Webber 1987a, Halcrow, 1999). The spit is probably not a classic multi-recurved form developed solely by littoral feed, but has evolved in a complex manner involving barrier transgression over a series of earlier Pleistocene gravel terraces and Holocene estuarine deposits, which themselves supplied sediment (King and McCullagh, 1971, Nicholls 1985, Nicholls and Webber 1987a, Bradbury, 1998). The spit is not a sediment sink because rapidly increasing depth and strong tidal currents at the distal end remove material. It therefore results from a complex interaction between sea level rise, sediment supply, storm overtopping events and the substratum (ie its capacity to support and contribute to the spit). Longitudinal extension has probably been controlled by offshore water depths and rates of sediment loss. Hurst Spit has probably receded relatively uniformly over the past 4000-5000 years since sea-level approached its present position (Nicholls and Webber 1987a). Due to the complexity of controlling factors (and the fact that it is not a sink), it is particularly sensitive to change. Its response has been to vary its rate of recession, with periods of more rapid retreat being associated with phases of diminished sediment supply and high magnitude, low frequency storm surges. (Nicholls 1985; Bradbury 1998).
Two possible sinks are envisaged for coarse sediments released from Christchurch Bay, both involving net eastwards littoral drift to Hurst Spit and subsequent removal offshore. Some sediment may be transported from Hurst Spit to the West Solent (Dyer 1971), but the major sediment flux appears to be south-west from Hurst Spit to feed the Shingles and Dolphin Banks and Dolphin Sand via the Needles Channel (Dyer 1970, Nicholls 1985, Velegrakis, 1994). The thickness of some of these accumulations has recently been determined by Velegrakis (1994) enabling approximate volumetric calculations. It is concluded that Christchurch Bay has been a virtually closed bedload transport system since the late Holocene except for probable (but diminishing) littoral drift input from Poole Bay and relatively small losses to both the outer bay and the West Solent (Nicholls 1985). Recent budget changes have resulted from human activity, including beach and offshore aggregate mining, cliff stabilisation, reduction and control of longshore transport and beach replenishment. These have had both positive and negative effects on the quantities and rates of coastal sediment circulation, (as discussed in subsequent sections), but the system remains a quasi-isolated system, with minimal fresh input and relatively small losses in the context of the overall budget (Velegrakis, 1994). SANDFLOW model simulation confirms this (Brampton et al, 1998).
1.2 Waves and tidal conditions
Wave and tidal current action are the dominant sediment transport mechanisms and the small tidal range serves to concentrate nearshore wave energy into a narrow zone.
(a) Wave Climate
An offshore wave climate was compiled for the more energetic conditions in east Christchurch Bay using a hindcasting technique based on 15 years of Portland wind data and direct measurement from the BMT tower, offshore Milford-on-Sea (Hydraulics Research 1989a and b). Prevailing direction was southwesterly and waves exceeding 1.0m were predicted for 31% of the time, and those exceeding 3.0m for 2.6% of the time. Extreme value analysis for waves of longest fetch (225-255EN) revealed offshore maximum significant wave height of 5.9m for a 1 year return period, and 7.2m for a 20 year return period. All waves exceeding 6.0 m in height moved in from the south-west or west-south-west. HR Wallingford (1999a) using both actual and synthetic wave data obtained from the above, and subsequent, studies calculated that 28% of all waves approached from the south west, 13% from the south-south-west, and 11% from the south-east. Wave height and direction are modified inshore by shoaling and refraction; studies of these processes were conducted by Henderson (1979) using a five year wave- rider record (offshore Southbourne, in Poole Bay) for calibration. These studies and HR Wallingford (1999a) indicated high concentrations of wave energy on Hengistbury Head and at Barton. By comparison, the western sector of Christchurch Bay is less affected by south-westerly swell waves, but more exposed to waves approaching from the south-east. The complex offshore bathymetry of Christchurch Bay (particularly the shoals of Christchurch Ledge, Dolphin Bank and Shingles Bank) exerts a major refraction and focusing effect on incoming waves so that the resultant nearshore wave climate is spatially complex and difficult to model. Wave energy divergence occurs between Mudeford Spit and Highcliffe. (Henderson and Webber 1979, Henderson 1979). Gao and Collins (1994b) state that immediately offshore The Run, Christchurch Harbour, the most frequently occurring significant wave height is 0.6 m, with 50% of recorded wave heights less than 0.25 m. HR Wallingford (1999a) calculate that maximum wave height offshore The Run, for a one year recurrence, varies between 2.64 m (155o approach) and 4.50 m (215o approach). Halcrow (1999) state that the maximum inshore significant wave height varies between 4.1 m (one year return frequency) and 5.3 m (one in 50 years recurrence). Both inshore and offshore wave climate affecting the shoreline at Milford-on-Sea and along Hurst Spit have been determined by numerical modelling, calibrated by data obtained from the BMT wave rider tower for specific periods (Hydraulics Research, 1982, 1989a and b; Bradbury, 1998; Halcrow, 1999). For Milford-on-Sea and the proximal end of Hurst Spit, Hydraulics Research (1989b) calculate mean inshore significant wave height to be between 2.47 m (1 year return frequency) and 3.76 m (1 in 100 year recurrence). This was based on using the INRAY model to transform data from the tower site 1.5 km offshore (in a mean water depth of 8m). In a later study, based on combined numerical and physical modelling, HR Wallingford (1993) concluded that mean inshore significant wave height, with a 1 in 10 year frequency, was 4.12 m; Wimpey (1994) used a higher resolution version of INRAY to recalculate the same data set; this study proposed that a 1 in 100 year maximum wave height of 5.98 m might be experienced. The equivalent value for offshore wave conditions, obtained using OUTRAY, was 6.15 m (240o approach direction). Hague (1992) proposed a range of 6.5-7.00 m for approaches between 210-270o . His work, which used previous data sets supplemented by subsequent records (Hydraulics Research, 1989b), proposed a maximum offshore significant wave height of 3.82 m, with an annual frequency. This study revealed that the prevailing mean significant wave height outside the nearshore zone was between 0.5 and 1.00 m (approaching from between 240 and 270o ). A recent re-application of the HINDWAVE model in combination with TELURAY was calibrated using all available data sets on wave and water level conditions (HR Wallingford, 1999a). For mean offshore significant wave heights with a one year recurrence, the derived values were: 1.38 m (155o ); 3.60 m (245o ) and 4.15 m (215o ). Bradbury (1998) determined maximum 1 in 1 year maximum wave height to be 5.91 m (135-240o ). This value was derived from previous models, but was calibrated using improved digitised hydrographic data.
The comprehensive scheme of protection and stabilisation of Hurst Spit, completed in 1996 (see Section 4.3) generated specific studies of the variation of wave climate along its proximal to distal ends. These are set out, and analysed, in Bradbury (1998), who proposes a 1 in 100 year mean offshore significant wave height of 4.14 m, approaching from 240o, for the proximal 800-1000 m sector. For inshore waves, with an annual frequency, maximum significant wave height varies between 3.57 m (240o) and 2.89 m (210o). Halcrow (1999) used models previously employed by HR Wallingford (1989a; 1993) but applied them to UK Meteorological Office synthetic wave data. For the proximal sector of Hurst Spit, this study determined that the most frequently occurring nearshore wave heights are between 0.1 and 1.0 m for all approach directions combined.
For the distal end of the spit (Hurst Point and Castle), Bradbury (1998) determines that the 1 in 100 year mean significant wave height (240o) is 3.10 m. For inshore conditions, extreme wave heights with a probability of recurring once a year are 2.10 m (210o) and 2.68 m (240o). For the 240o approach direction, Halcrow (1999) deduced mean wave heights to vary from 1.6 m (1 in 1 year) to 2.60 m (1 in 50 years). Some 75% of all waves arriving at the distal end, from all directions, have a mean height of 0.1 to 1.5 m.
It is therefore apparent that there is a progressive north-west to south-east reduction in nearshore wave energy along the main "corridor" of Hurst Spit. Taking into account all previous studies, Bradbury (1998) suggests that mean maximum nearshore wave height declines from 4.1 m to 3.1 m. The main reason for this is the attenuating/dissipating influence of Shingles Bank and North Bank, setting up complex refraction and wave train "crossover". Wave shoaling and breaking (at low water states) induced by the complex bathymetry of the banks and channels seawards of the distal sector reduces the height of offshore waves by almost precisely one third (Bradbury, 1998). It is probable that nearshore wave conditions have varied as the offshore banks, shoals and channels have evolved. Contemporary monitoring (New Forest District Council, 1992; 1997-2001) reveals constant fluctuations in bank morphology, but specific impacts on wave climate are probably fairly limited over the medium timescale (1-10 years). More important is the impact of high magnitude, low frequency storm and tidal surges on water levels, and therefore water depths, on shoaling and refraction. This is apparent from analysis of the 1 in 100 year recurrence storms of October and December 1989 (Bradbury, 1998). The latter generated a higher water level, but prevailing inshore wave heights (2.90 m at the neck of Hurst Spit) were less than two months previously (3.60 m) because of reduced wave height peaking. Monitoring of the frequency of maximum nearshore wave heights since 1996 (New Forest District Council, 1997-2001) will allow further analysis of this factor.
For the Bay as a whole, dominant waves approaching from the west to southwest are therefore subject to spatially and temporally variable refraction with the result that nearshore wave conditions vary both along the coast and with time. Halcrow (1999), using simulated wave data derived from the UK Meteorological Office model and local bathymetric data concluded that wave climate at all points along the shoreline was principally a function of wave approach direction.
(b) Tidal Regime
The tidal range of Christchurch Bay is the lowest along the south-central Channel coast. At the entrance to Christchurch Harbour, the mean spring range is 1.4 m, reducing to 0.8m during neap cycles, although there is some attenuation of wave form as it moves through the narrow channel connecting the harbour with The Run immediately seawards.
Tidal currents are extremely rapid at Hurst Narrows (up to a maximum 3ms-1 at the surface and 2.5 ms-1 close to the seabed) and are capable of significant transport of coarse sediment (see also unit on West Solent and Webber (1980a)). Currents diminish offshore and longshore so that dominant south-eastward currents on the southwest face of Hurst Spit attain peak velocities of 0.8-0.9ms-1 (Nicholls 1985). Metering surveys show that currents generally diminish to less than 1ms-1 to the west of Shingles Bank (Velegrakis 1994) and are below 0.5 m sec-1 in the central area of the Bay (Brampton et al, 1998). In the vicinity of Highcliffe and Barton, dominant ebb tidal currents flow eastward at up to 0.3-0.4ms-1 (Lacey 1985) to give a weak near-seabed clockwise tidal gyre, defined mathematically from tidal flow residuals, for the Bay as a whole (Riley et al, 1994; Riley, 1995, Halcrow, 1999). Rapid tidal currents are thus restricted to the eastern extremity of the bay, except for Christchurch Harbour entrance, where ebb flow coupled with fluvial discharge can create maximum velocity currents approaching 2.5ms-1 (Tosswell 1978, Gao and Collins, 1994a and b, 1995c; HR Wallingford, 1999b). Here, the duration of the flood tide is 6.4 hours, whilst the ebb phase occupies 6.01 hours. Gao and Collins (1994b) determined that the threshold velocity for the entrainment of fine sand on the bed of the harbour entrance channel is between 1.7 and 1.9 m s-1. Mean tidal range within the harbour is lower than on the adjacent open coast because of attenuation effects set up by The Run (Hydraulics Research, 1987). In both Christchurch Bay and Harbour, a "double high water" effect (ie a high water long stand, followed by a second high water peak) is discernable, though less well pronounced than in Poole Bay. This contributes to the ebb-dominant asymmetry of the tidal regime of Christchurch Harbour.
2. SEDIMENT INPUTS
2.1 Marine Inputs - EO2 F1 F2 F3 References Map
As Christchurch Bay is widely regarded as a largely self-contained sediment circulation system (Nicholls 1985, Bray and Hooke, 1998b, Halcrow 1999), it is difficult to envisage any significant externally-derived marine input. This view is supported by a variety of evidence:
(a) Chart comparisons of the Pot Bank dredging area south of the Needles covering the period 1937-72 revealed seabed lowering of 1m (Hydraulics Research 1977). The surveyed seabed change was only 18% of that removed by dredging during this time (4 million m3), thus providing indirect evidence that material was leaving the area. Despite this net loss, analysis suggests that Pot Bank is a sediment sink receiving sediment from the Needles Channel and Shingles Bank, with no return transport (Hydraulics Research 1977, Brampton et al, 1998). This conclusion is partly based on modelling of tidal residuals, and near sea-bed current velocities operating below gravel mobilisation thresholds.
(b) Offshore survey involving echo sounding, oblique asdic sonar and grab sampling have revealed an abrupt transition in seabed morphology and materials south of Dolphin Bank and Sand (Dyer 1970, Velegrakis 1994). Both banks are characterised by predominantly sandy glauconite-rich sediments with bedforms indicating a net overall westward transport. Further offshore, in deeper water, gravel bedform features indicate southward transport, and limonite-rich sediments suggest possible feed from the nearshore/offshore zone of southwest Wight. Sediment transport between the two areas appears extremely limited, and thus output from Christchurch Bay appears much more likely than input.
(c) Christchurch Ledge is widely recognised as a barrier to net onshore bedload sediment transport into the western part of Christchurch Bay from further west and south-west (Dyer 1970, Wright 1982, Brampton et al, 1998), although evidence is not conclusive.
(d) Littoral drift has been calculated for Poole and Christchurch Bays using an energy flux technique based on wave refraction analysis (Section 3). Differences in drift between adjacent transport compartments are attributed to both onshore and offshore transport. Although potentially unreliable, both the initial analysis (Henderson 1979) and improved versions (Lacey 1985, Halcrow 1999) predict net offshore transport from specific beaches in Christchurch Bay, although this is acknowledged to be temporarily and spatially inconstant. However, it is uncertain if onshore gains elsewhere counterbalance offshore losses.
(e) Chart comparisons covering the whole of Christchurch Bay for the period 1880-1968 revealed overall loss of 505,000 m3a-1 of sediment (Lacey 1985). Beach profiles and nearshore hydrographic survey covering the period 1974-90 also indicated loss of sediments (Webber 1980b, Halcrow 1980, Lacey 1985, Velegrakis, 1994). This information again suggests net sediment output rather than input. Although the accuracy of hydrographic surveying is frequently insufficient to determine changes over wide areas for short time periods, the long-term bed erosion identified by Lacey (1985) is equivalent to uniform lowering of 0.39m, which is probably within the limits of survey accuracy.
(f) Analysis of suspended sediments in Christchurch Bay using Landsat imagery revealed a marked seasonal trend corresponding to peak cliff erosion input during winter storms (Lacey 1985). The majority of suspended sediments are probably derived from this source and no evidence of any independent/supplementary marine input is available. However, research has been insufficient to draw any firm conclusions.
Potential marine input can nonetheless be identified, namely:
Feed from the West Solent (eo2) (see intro to marine inputs)
Ebb tidal currents are of shorter duration and more rapid at Hurst Narrows (Webber 1980a), thus a dominant southwest transport pathway extends along the Needles Channel (Dyer 1970, Nicholls 1985, Velegrakis 1994). Previous research suggests net eastward transport into the West Solent, but this does not fit with the energetics of the tidal regime. Thus it is postulated that tidal current-moved sediment output is probable from the West Solent, thereby feeding the vigorous southwest transport path at the extreme eastern margin of Christchurch Bay. (See Section 6 for a full account based on bathymetric, sedimentological and morphological data).
F1 Hengistbury Head (see intro to marine inputs)
Littoral drift analysis revealed a major mismatch between eastward longshore supply to Hengistbury Head and drift rates along Mudeford Spit (Photo 1). This was attributed to onshore feed of sand and gravel between 476,000m3a-1 (Lacey 1985) and 559,000m3a-1 (Henderson 1979). These estimates are of low reliability as they are based solely on theoretical calculations of longshore wave energy flux with no consideration of site-specific littoral processes or sediment availability. Serious problems were encountered with wave refraction and energy diffusion in the area of complex bathymetry set up by Christchurch Ledge, and sediment transport equations have been calibrated with extremely limited data. Seabed drifter studies, beach profiling and aluminium pebble tracer experiments all indicate net offshore transport on Solent Beach one kilometre to the west (Watson 1975, Wright 1976, Webber 1980b, Halcrow 1980, Wright 1982, Hydraulics Research 1986a, Brampton et al, 1998; Halcrow, 1999). (See section on Poole Bay for more detailed discussion).
F2 Mudeford and Highcliffe (see intro to marine inputs)
Several studies have been conducted in Christchurch Bay using Woodhead Seabed Drifters (Clark and Small 1967, Watson 1975, Tyhurst 1976, Turner 1990); a total of nine experiments were undertaken between 1966 and 1975. These clearly revealed that Mudeford-Highcliffe was a major receptive area for drifters, indicating a potential for onshore transport to this coastal segment. Recoveries were generally between 38 and 75%, but as they relied upon the general public there were problems with the quantity and quality of information. This was only available for time and location of injection and recovery, so the intervening pathways, as well as the mechanisms of transport, are conjectural. It is uncertain exactly what drifters measure, because it has been shown that they are not necessarily reliable indicators of movement at or close to the seabed (Collins and Barrie 1979). The experiments probably indicated onshore currents at the seabed, but the significance of this to sediment transport is uncertain; thus this pathway is of low reliability. Limited support is provided by study of the historical behaviour of Mudeford Spit using map and chart comparisons (Robinson, 1955; HR Wallingford, 1999) from which it is tentatively concluded that it has been sustained by onshore rather than longshore supply. (See Section 5.2 for detailed discussion).
F3 Milford on Sea and Hurst Spit (see intro to marine inputs)
One study using seabed drifters indicated a potential for onshore sediment supply at Milford and Hurst Spit (Clark and Small 1967), but subsequent drifter experiments indicated an offshore southeasterly trend in this area (Watson 1975, Turner 1990). These studies are therefore contradictory and inconclusive; an assessment by Dobbie and Partners (1984) and Nicholls (1985) concluded that significant net onshore feed to Milford-Hurst was unlikely, particularly of gravel. Offshore survey by side-scan sonar, echo sounding and sediment sampling revealed asymmetric gravel ripples which indicated a net onshore transport vector over the western part of Shingles Bank (Velegrakis 1994). This pathway supplies gravel to a dominant north-westwards sediment flow in North Channel (Dyer 1970, Velegrakis 1994), but the ultimate sink is uncertain. Coarse sediments are therefore highly mobile offshore Milford-Hurst, although an onshore transport supply to North Channel has been recognised (Velegrakis, 1994). This may not be a long-term feature.
2.2 Fluvial Input - FL1 References Map
FL1 The Rivers Stour and Avon
Combined discharge of the Stour and Avon to Christchurch Harbour averages 30 m-1 with a minimum flow of 7.5m3s-1 and a maximum of 220m3s-1 (Tosswell 1978, Gao 1993, Gao and Collins 1997a). It is reported that suspended sediment load is low due to catchment supply from Chalk aquifers and interruptions to flow, e.g. weirs (Murray 1966, Tosswell 1978, Rendel Geotechnics and University of Portsmouth, 1996), but is periodically supplemented by river channel dredging to alleviate flood risk. Sampling surveys revealed sand and gravelly sandbanks adjoining the river channels in close proximity to their points of discharge into the harbour, and these may provide some supply (Gao 1993). Annual siltation is reported in navigable channels in the harbour and accelerated rates during the winter of 1990-91 on Grimsbury Marsh were alleged to result from dredging of the lower Stour. Mean fluvial discharge is approximately equal to the tidal prism, which indicates clear potential for sediment supply (Tosswell 1978). Rendel Geotechnics and the University of Portsmouth (1996) calculate a maximum combined bedload input of some 150 tonnes a-1, but actual quantity is probably less than 10% of this total due to numerous weirs and sluices on both rivers. Gao and Collins (1995b) calculate bedload input as 320-640 m3a-1, based on a largely theoretical calculation. Suspended sediment discharge potential is close to 70,000 tonnes a-1, but actual delivery is unlikely to exceed 10,000 tonnes a-1. It is uncertain how much of this material is transported through the harbour entrance; some is understood to be retained within the estuary system, where it contributes to inner harbour mudflat and marsh sedimentation (Gao and Collins 1995a).
There is no sediment delivery via Chewton Bunny, as it discharges via a culvert. Other streams to the east, notably Beckton Bunny, contribute only very small quantities of predominantly fine sediments, despite their well-incised valleys and erodible catchments.
2.3 Coast Erosion - E1 E2 E3 E4 E5 E6 References Map
Much of Christchurch Bay is backed by rapidly eroding cliffs of up to 30m height, which provide a major sediment input where they are not stabilised and protected. Local rates of erosion are determined by geology, the longshore distribution of wave energy, the degree of cliff/coastal slope protection afforded by the natural beach or protection structures and cliff stabilisation measures. The cliff-forming strata comprise Tertiary sands and clays which dip 0.5 to 1o towards the east-north-east and strike nearly parallel to the coastline, so that progressively younger beds outcrop from Hengistbury Head (Hengistbury Formation) to Hordle and Milford (Headon Formation). There is also a slight inland stratal dip. These substrate materials are overlain by a mantle of Pleistocene Plateau Gravels and thin Holocene brickearth deposits. The nature of sediments supplied is therefore controlled by cliff lithology and it is possible that this has altered throughout the Holocene as coastal recession has truncated lithologically variable deposits. The Pleistocene gravels are particularly variable in composition, and it is suggested that previous cliff sections included deep infilled channels (Nicholls 1985), contributing a gravel supply that was at least twice the volume that prevailed after coast protection. Lithological variations of 'solid' substrate also affect the rate of erosion where sandstones (permeable) overlie clays (impermeable) with critical groundwater pore pressures facilitating slope failure (Barton 1973, Barton and Coles 1984, Rendel Geotechnics, 1991, 1998). Nicholls (1985) has proposed that over the past 2-3000 years, up to the introduction of protection and stabilisation measures, the supply of sand from cliff erosion was eight to ten times the volume that is received today. Lacey (1985) calculated that between 48% and 54% of all material derived from cliff erosion is fine grained, and is rapidly transferred permanently offshore via suspended transport.
Long-term recession has created a shore platform that widens progressively eastwards in response to exposure to wave energy. Posford Duvivier and the British Geological Survey (1998) provide some largely hypothetical figures for contemporary annual shoreface erosion (both vertical erosion and volume of yield).
E1 Hengistbury Head (see intro to coast erosion)
Map comparisons covering the period 1896-1976 have revealed cliff top erosion at 0.42-0.77ma-1 to the north-east of the Long Groyne due to recurring slips (Photo 4). The higher rate was located almost adjacent to this structure and attributed to wave energy focusing caused by it (Hydraulics Research 1986a). An armoured revetment and four rock groynes were constructed in the late 1980s and now protect the cliff toe, although input is still possible due to erosion by high energy storm waves approaching from the east/south-east. Small quantities of sediment are yielded from these cliffs, which comprise sandy clays capped by thin valley gravels. Artificial removal of talus is undertaken to maintain exposure of the internationally important stratigraphy and palaeontology of this GCR/SSSI site (Bray et al, 1996; Bray and Hooke, 1998).
E2 Highcliffe (see intro to coast erosion)
The cliffs here are composed of Highcliffe Sands to the west (Photo 3), with an increasing thickness of underlying impermeable, overconsolidated and fissured Barton Clay to the east. The latter contains lenses of sand, which supply groundwater to the cliff face and thus facilitate landsliding. The succession is overlain by Plateau Gravels varying from 1.5 to 7.6m in thickness, creating strong seepage and gulleying at its junction with the underlying sandstones.
The cliffs at Highcliffe have a long history of instability and retreat. Historical records indicate that retreat may once have been extremely rapid, and a possibly unreliable rate of 6ma-1 is quoted for the period 1760-1830 (Nicholls 1985, Bray et al, 1996). This is supported by an account suggesting loss of up to 1km (10ma-1) between the late 1700s and late 1800s (Mockridge 1983). Subsequently, the cliffs become significantly more stable, probably due to improved drainage and toe protection by an accreting beach fed by the extension of Mudeford Spit (Mockridge 1983). Mean recession of 0.18ma-1 was recorded over the period 1908-59 by comparison of maps and a cliff-top survey in 1958 (Wise 1959). Rates were spatially and temporally variable with maximum recession for a short central cliff section of 0.95ma-1 for 1931-39 (Wise 1959). The cliff toe periodically advanced and retreated over this period as basal landslide debris intermittently surged seaward and was then eroded (Barton 1973). This phase of relative stability ended in the 1950s following the retreat and breakdown of the formerly extended form of Mudeford Spit, and thereafter the Highcliffe cliffline was characterised by increased frequency of landsliding and cliff-top retreat rates which increased to 0.68ma-1 for the period 1965-1975 (Barton 1973, Mockridge 1983, Tyhurst 1985a, 1986).
The cliffs formed a bench and scarp profile with failure of the upper scarp by relatively deep-seated slumps and degradation, and transport of debris over the undercliff benches by mudslides (Barton 1973, Mockridge 1983). These processes supplied significant quantities of sediment to the foreshore between 1932 and 1968; Lacey (1985) calculated a supply of 4,400m3a-1 of material greater than 0.08mm diameter, regarded as being stable on the beach and lower foreshore. This supply has now largely ceased due to the effectiveness of coast protection and cliff stabilisation schemes commencing in the late 1960s. These include construction of a timber revetment and groynes, cliff drainage and stabilisation in 1973/74 and 1978/79; vegetation establishment over the undercliff to enhance slope stability, and combined beach nourishment and rock groyne construction in 1985 and 1991 (Photo 5). Although the cliffs between Friar's Cliff and Highcliffe Castle are unprotected, from 1986 onward cliff drainage was inserted further west towards Highcliffe Castle (Mockridge 1983, Tyhurst 1986, Christchurch Borough Council 1991; Halcrow, 1999). Cliff input is not possible whilst these measures are effective, but the talus store in the vicinity of Friar's Cliff is potentially available as an input source.
E3 Chewton Bunny to Barton-on-Sea (see intro to coast erosion)
This segment comprises naturally and rapidly eroding cliffs, in contrast to the now protected and stabilised adjoining cliffs at Highcliffe and Barton (Photo 6). It thus provides one of the few remaining natural input sources of beach sediment in Christchurch Bay (Mackintosh and Rainbow, 1996). The cliffs are composed of eastward dipping Barton Sand and Clay overlain by 1.5-3.0m of Plateau Gravel (Barton and Coles 1984). It is reported that former exposures of these gravels included Pleistocene buried channel deposits up to 16-18m thick, but diminished as cliff retreat cut further landward (Nicholls 1985). Lithological variations within the Barton Clay cause development of three preferred bedding plane shear surfaces marked by benches in the cliff profile. The rear scarp fails by deep-seated rotational slides extending to the uppermost shear surface and periodically down to the intermediate surface. The cliff top therefore continues to recede rapidly via recession heads and "breakaways", with single events removing up to 2m. The benches are sites of degradation and downslope sediment transfer due to small scale slips, bench sliding, slumping and mudsliding (Barton and Coles 1984). Bench sliding is considered the dominant process, possibly accounting for over 90% of slope movement (Halcrow, 1999). Groundwater seepage at lithological junctions is a prime factor facilitating superficial flows and slides.
A long-term retreat of approximately 1ma-1 has been characteristic of these cliffs according to map comparisons over the period 1867-1959 (May 1966, Nicholls 1987, Hooke and Riley 1987). This conceals spatial variations of short-term rates of up to 5 ma-1. Retreat over the period 1869-1939 was 0.3ma-1, increasing to 0.4ma-1 by the late 1950s, 1.3ma-1 (1960s), 1.9ma-1 (1970s) and 2.4ma-1 (1990s) (Mackintosh and Rainbow, 1996). Following construction of groynes updrift at Highcliffe, this sector retreated by 60m or 3ma-1 1970-1985 (Halcrow, 1999). This accelerated retreat is attributed to sediment starvation due to outflanking set up by the terminal groyne effect at Chewton Bunny (Bray, et al, 1996; Mackintosh and Rainbow, 1996). Sparshott (2001) derived a mean rate of cliff top recession for Naish cliffs of 1.12ma-1, 1976-1998. Maximum and minimum rates of 1.55ma-1 and 0.45ma-1, respectively, were determined for different sectors. Cliff toe recession was calculated to be 0.65ma-1 over the same period, with a maximum of 1.22ma-1 recording the rapid evacuation of a debris fan. This research, which was based on contour plotting from DEMs based on three sets of vertical aerial photography, revealed a reduction in the gradient of the coastal slope,. It also demonstrated that there is close adjacency of relatively stable and unstable sectors of the cliff face, but that the zone of maximum morphological change moved progressively eastwards between 1976 and 1991. Together with other studies (Rendel Geotechnics, 1991; 1998; Mackintosh and Rainbow, 1996), active slope degradation was identified as a largely winter season activity.
Integration of information on sediment composition exposed in these cliff sections with the above quoted mean retreat rates have enabled sediment yield to be calculated. Total supply of sediments >0.08mm diameter was 9000m3a-1, in 1984 (Lacey, 1985). Sampling of the Plateau Gravel deposits revealed that 46% was gravel >8mm diameter (Indoe 1984). Combining this with a mean deposit thickness of 2m, and a prevailing retreat rate of 2ma-1 (Barton and Coles 1984) over the 1300m eroding length of cliffline gives an approximate gravel supply of 2400m3a-1. Yield of fine sediment, from cliff erosion, is estimated at 58,000 m3a-1 (Posford Duvivier and British Geological Survey 1998). Using SURFER, Sparshott (2001) calculated volume loss, below 30mOD, to be a mean of 179,000 m3a-1, a figure which includes clay, sand and gravel. This applies to the entire sector between Chewton Bunny and the western end of the stabilised cliffs at Barton. Barton and Coles (1984) proposed that a total annual yield of 5740 m3 derived from a 200 m segment of this cliffline in the early 1980s.
Recession of this cliffline between two protected frontages has created an embayment with an immature log spiral plan shape. Crenulate bay concepts can be applied to model the evolution of such features and indicate that a further 140m (Lacey 1985, Webber 1980b, Halcrow 1980) to 240 m (Halcrow 1999) of erosion is required at the western extremity before a stable plan shape is established.
E4 Barton-on-Sea to Beckton Bunny (see intro to coast erosion)
Historically, this segment was also extremely unstable and subject to rapid cliff top recession, measured at 0.75ma-1 for the period 1867-1959 (Hooke and Riley 1987). Phillips (1972) calculated mean rates of 1.1 ma-1 (1868-1907); 0.9 ma-1 (1908-1936) and 0.5 ma-1 (1937-1958) with rates as high as 3.0 ma-1 for individual years. Cliff form and processes were generally similar to those now exhibited in the unprotected segment to the west except for an important lithological distinction. Eastward dip of the stratigraphical succession gradually brings the impermeable Barton Clay down towards beach level at Beckton Bunny, thereby increasing the extent of cliff outcrop of the overlying permeable Barton Sands and Plateau Gravels. Groundwater is directed towards the cliff-face at the interface between the sandy and clayey strata separating (i) the Upper and Middle Barton Beds and (ii) the Plateau Gravel and Upper Barton Sand, promoting seepage and gulleying (Daley and Balson, 2000). The sandy materials above this interface tend to fail by deep-seated rotational slides extending down to this shear surface/zone of weakness. (Barton, 1973; 1998; Barton and Coles, 1984; Fort et al, 2000). This surface formed a bench or undercliff characterised by degradation and seaward transport of sediment towards a steep sea-cliff cut in the underlying Barton Clay (Clark, Ricketts and Small 1976). Increased cliff top recession of 2.0-2.4ma-1 for 1959-66 (Barton 1973), despite groyne construction between 1939 and 1954 and a temporary increase in beach width due to the longshore dispersal of sediments composing Mudeford Spit after its breaching in 1935, endangered cliff top properties. This led to a more comprehensive coast protection scheme which was completed between 1966 and 1968 (Photo 7 and Photo 8). This involved slope re-modelling and regrading, and toe protection comprising groynes and a flexible timber revetment with retaining rock armour (Phillips, 1972; 1974, Wright, 1992). Cliff stabilisation over a frontage of 1.75 km included steel sheet piling driven through the undercliff to intercept groundwater flow along the principal clay/sand interface, and seaward drainage of groundwater thus re-directed (Phillips, 1974; Clark, Ricketts and Small 1976, Fleming and Summers, 1986; Wright, in Bray and Hooke, 1998). These measures were only partially effective because deep-seated failures in the undercliff in the winter of 1974/75 penetrated beneath the sheet piling causing its rotation, bowing and splitting (Clark, Ricketts and Small, 1976). It is probable these failures were along previously undetected slip planes within the Barton Clay. This would be similar to the landslide process west of Barton where movements are frequent on an upper slip plane, but intermittent on a lower shear surface (Barton and Coles 1984). The result was two-fold; first, mudslides surged over the cliff and revetment and supplied sediment to the beach; second, cliff-top recession averaged 5ma-1 over a 150m front in 1975 (Clark, Ricketts and Small 1976, Indoe 1984, Wright, 1992). Various emergency and remedial regrading and drainage works and beach stabilisation using five 400m spaced rock groynes/strongpoints (Photo 8) and beach nourishment have been largely successful in preventing further major slides (Wright, 1992; Bray, et al, 1996), although detailed site monitoring revealed continuing ground and sub-surface movements (Rendel Geotechnics, 1998). This (ongoing) investigation has established that rainfall exceeding 80 mm per month is sufficient to initiate movement, usually during the winter. A major multi-plane slip at the western end of this sector in 1993 displaced the toe revetment 8m seawards and required the construction of new drainage and a major rock revetment to maintain slope stability by toe weighting (Photo 9).
The eastern part of the cliffline towards Beckton Bunny remains largely unprotected (Photo 10) and retreated at 1.63ma-1 from the late 1970s to early 1980s (Lacey 1985). Supply capacity is high because Plateau Gravel deposits attain a thickness of 5-9m (Nicholls 1985) and the thick Barton Sand supplies material capable of contributing to the lower foreshore. Total supply >0.08mm diameter was calculated as 13,000m3 for 1984 (Lacey 1985), although prior to the 1966-8 protection scheme it was much greater.
Substantial movements over all parts of this section were experienced during the exceptionally wet winter of 2000/01, but quantitative estimates of additional sediment yield are not yet available.
E5 Beckton Bunny to Hordle (see intro to coast erosion)
The cliffs along this segment are composed of the predominantly sandy Upper Barton Beds and exhibit a steeper upper profile than further west. Nevertheless, lithological variations occur and the lower part of the cliff has high clay content resulting in a characteristic scarp and bench profile regulated by frequent mudflows. Historically, retreat averaged 0.85ma-1 over the period 1869-1959 (Hooke and Riley 1987), although a more rapid retreat of 1.5ma-1 was recorded for the period 1932-68 (Lacey 1985). Since approximately 1960, retreat has accelerated significantly to a mean of 2.0 ma-1, with between 3.6 ma-1 and 4.2ma-1 recorded for the period 1968-82 (Nicholls 1985, Lacey 1985, Dixon, et al 1998, Halcrow, 1999). This acceleration is partly attributed to accentuated terminal scour resulting from reinforcement in 1970 of the Beckton Bunny outfall (Photo 11), which created an impermeable groyne effect (Nicholls 1985), but also to reduction of beach width. The zone of cliff failure reactivation would appear to be migrating eastwards (Dixon et al, 1998). Cliff erosion supply was calculated at 15,000m3a-1 (>0.08mm diameter) for 1984/85 which greatly exceeds that prevailing before the early 1970s (Lacey 1985). For a 1.8 km length of shoreline, Dixon, et al (1998), calculated from high resolution aerial photography, 1967-1993, that total sediment yield for this period was 1,400,000 m3 (approx 52,000 m3a-1). Immediately downdrift of Beckton Bunny, the major morphodynamic zones of the cliff face shifted landwards, but an overall steady state balance of upper and lower profile processes was maintained. Nicholls (1985) calculated gravel input of 7000m3a-1 from cliff erosion west of Hordle, the majority derived from the Beckton-Hordle segment, because the Beckton outfall has operated as a partial, but effective barrier to input from further updrift.
Between Taddiford Gap and Hordle Cliff, retreat is slow, averaging 0.12ma-1 for the period 1809-1969 with negligible retreat recorded between 1931-69 (Hooke and Riley 1987). This is due to accretion of a substantial gravel beach (Photo 12), which protects the cliff toe and prevents marine erosion and fresh sediment input (Nicholls 1985, Halcrow, 1999), and which is linked in an as yet undetermined way to a persistent nearshore bar below maximum low water (Halcrow, 1999). Posford Duvivier (1997) calculate that the long-term yield of fine sediment has been approximately 57,000 m3a-1 from this sector, but that coarse sediment release is no greater than 1,500 m3a-1.
E6 Hordle to Milford-on-Sea (see intro to coast erosion)
Cliffs at gradients of between 20 to 45o composed of sands, clays and marls of the Headon Formation decrease in height from 20-25 m at Hordle Cliff to beach level at Milford. Map comparisons indicate a mean rate of retreat at 0.7 ma-1 over the period 1869-1969, but increased recession was recorded during the latter part of this period, at a mean of 1.08 ma-1 (Hooke and Riley 1987, Nicholls 1985). Since the early 1970s, erosion rates of up to 1.8 ma-1 have been calculated (Halcrow, 1999) following up-drift cliff stabilisation and reduced beach volumes. The cliffs are capped by Pleistocene gravels, attaining a maximum thickness of 7m at Rook Cliff. These are regarded as remnants of a formerly more substantial infilled tributary channel of the River Solent (Nicholls 1987). Evidence is based on historical records of uncertain reliability and an observed trend elsewhere for inland thinning of gravel deposits. Parts of this cliff segment are colonised by vegetation, and have been protected by groynes and sea walls first installed in the early 1930s, upgraded in the late 1960s and protected by a fronting rock revetment in the late 1990s - see Photo 13 (Dobbie and Partners 1984, Halcrow, 1999). Input supply of 3422m3a-1 (>0.08mm diameter) is calculated for 1932-68 compared with 3000m3a-1 for 1984/85 (Lacey 1985). Posford Duvivier (1997; 1999) calculate a total yield of 12,000 m3a-1 (58% sand, 40% clay and 2% gravel). Continued supply, despite protection, can be attributed to rapid recession of unprotected parts, some continuing superficial instability of stable slopes and progressive loss of beach volume. The latter appears to have accelerated to a rate now exceeding 0.5 ma-1 since the late 1980s (Bray, et al, 1996; Bray and Hooke, 1998, Halcrow, 1999).
Shoreface Erosion - References Map
Posford Duvivier and the British Geological Survey (1998) provide some very approximate calculations of wave and tidal current-induced erosion of the shoreface of Christchurch Bay. This feature has a width that is generally less than 1000 m, narrowing to less than 350 m at Milford and is mostly developed in basement Eocene sands and clays. The following data is given:
The reliability of these figures is subject to considerable uncertainty, as they derive from the application of simplistic formulae to base data of undetermined accuracy.
Interpretation and Summary
2.4 Beach Nourishment - References Map
Artificial beach nourishment constitutes a significant sediment input and has been practised since the late 1970s at several locations including Mudeford Spit, Highcliffe and Barton (May 1990), and frequent emergency topping up/beach restorations at Hurst Spit, which culminated in comprehensive reprofiling/recharge in 1996.
Hengistbury Head and Mudeford Spit
Small-scale beach nourishment has been undertaken since the late 1970s both to the immediate north- east of the Hengistbury Long Groyne (Hydraulics Research 1986a) and on Mudeford beach (Tyhurst 1985a), the latter involving three separate operations each of 1000 tonnes of coarse sand.
Falling beach levels in the 1970s prompted an initial sand recharge, followed later by a sand and gravel nourishment scheme to compensate for ongoing volume losses. The artificial beach profile was designed by two-dimensional modelling in a wave flume which simulated storm wave and surge conditions (Hydraulics Research 1984, 1986b; Tyhurst 1985b). Optimum beach configuration was subsequently achieved using 55,000m3 chert and flint gravel from an inland source. Beach retention was facilitated by existing timber groynes and two newly constructed downdrift rock groynes. This was completed in March 1985 and was subject to the regular monitoring of 30 transverse profiles measured between October 1984 and June 1986. Analysis revealed a 14% increase in volume (6700m3) for the period March 1985-February 1986, although the initial volume of the artificial beach was uncertain because its packing density was not measured (Hydraulics Research 1986b). The measured volume increase may thus be partially an artefact of the monitoring technique. Further profile monitoring in February and March 1987 indicated an overall volume loss of 5% (Hydraulics Research 1988). This information is difficult to compare with previous monitoring because volumes were calculated above Ordnance Datum, whilst the previous survey calculated volumes above -0.5m OD. If both surveys are accepted as accurate, net sediment gain of 6700m3 is estimated for March 1985 to February 1986, followed by loss of 8600m3 between February 1986 and March 1987. Profile monitoring and analysis up to June 1990 revealed diminution of beach volume from 57,130m3 (March 1985) to 45,100m3 (above -0.5m ODN), a total loss of 21% (2400ma-1) over the five year period (Hydraulics Research, 1991). Accepting the measured gain in the first year followed by rapid loss in the second year as an expected result of renourishment, loss between March 1987 and June 1990 is estimated at 3,100m3a-1. Because the effects of compaction and the quantity of natural beach feed from onshore transport is uncertain, it is difficult to determine losses accurately, although the pattern of rapid initial drawdown followed by a decreasing rate of loss in subsequent years is characteristic of gravel nourishment schemes As a result of these losses, an improved scheme entailing replenishment with 18,000 m3 of gravel and replacement of existing timber groynes with rock groynes was undertaken in 1991, with subsequent maintenance. Details of its performance are not available, but progressive loss of volume is reported (Coates and Bona, 1997). However, the rate of loss has been lower than that which is characteristic of most other south coast gravel recharge schemes.
Small-scale gravel renourishment has been carried out in connection with the construction of rock groynes in the mid 1980s, to provide bay-shaped beaches (Fleming and Summers, 1986; Wright, 1992).
Milford on Sea
Periodic renourishment has occurred since the mid 1970s, though details of quantities are not available. The largest volume was apparently placed in 1985.
Hurst Spit has been declining in volume, and its foreshore receding, since at least the late nineteenth century (Hooke and Riley, 1987, Bradbury, 1998). Rates of erosion have undoubtedly been accelerated by the interruption to, and reduction of, the rate of littoral transport due to updrift protection works in Christchurch Bay, whilst losses to the offshore zone at the distal end have been more or less constant over this period. The consequence has been increasingly severe erosion of this barrier structure, with overtopping, overwashing, crest cut-back and breaching occurring under conditions of storm surge (Bradbury, 1998; Bradbury and Kidd, 1998 - further morphological and morphodynamic description and explanation is given in Section 5).
To prevent breaching at the root, or proximal end, a 600 m length of rock armour was emplaced between 1967 and 1968. Renourishment at an annual average of 1000 m3 was undertaken between 1980 and 1985. Storm erosion in 1984 resulted in widening, reprofiling and recharge over a 450 m length of the spit beyond the earlier rock armour (Dobbie and Partners, 1984). Severe storms in October and December 1989, with a 1:100 year recurrence, caused overwashing, crest flattening over some 800 m of the spit, up to 80 m of "rollback" (landward migration), and the displacement of some 50,000 tonnes onto the leeward saltmarsh of Mount Lake. Unquantified offshore losses also occurred. The immediate response was to recover some 25,000 tonnes of gravel from Mount Lake, and import an equal quantity from inland sources, to rebuild the structure, now some 12 m set back from its pre-storm position. Some recharge was also carried out following major storm-induced erosion in February 1991 and April 1994. (See Section 4.3).
In response to the impact of the 1989 storms, and because imported recharge sediments from inland sources proved to be too small to be retained over the longer-term, New Forest District Council undertook in-house and commissioned a series of field-based and model investigations, together with ongoing routine monitoring of beach behaviour, to determine a long-term, optimum scheme of protection (HR Wallingford, 1993; Mackintosh and Rainbow, 1995; New Forest District Council, 1990, 1996; Wimpey Environmental Ltd, 1994; Bradbury and Kidd, 1998). The outcome, constructed in 1996, involved a series of measures (see Section 5.3), which included the recharge of the most vulnerable 800 m length of the proximal and mid-sectors of the spit (Photo 14). This was obtained from the nearby Shingles Bank offshore, as it is the natural source of supply of gravel to the spit and functions as a store for material removed from it. It thus provided a source whose size grading was very close to the prototype beach material. The effect of this recharge was to almost double the previous volume of the spit. Crest level and width were both increased, though declining eastwards in conformity to the reduction in wave climate severity from proximal to distal ends. Allowance was also made for subsidence, as the basement support of the spit is composed of low yield strength salt marsh deposits, over which it has migrated.
3. LITTORAL TRANSPORT (BEACH DRIFT) - LT1 LT2 LT3 LT4 LT5 LT6 LT7 References Map
The overall net pathway of bedload transport is eastwards, with alongshore wave energy approximately proportional to drift rates. Wave approach direction is considered more important that inshore wave heights in determining spatial variations in longshore transport efficiency. Since the mid-twentieth century, the rate of longshore movement of sediments has been slowed by the presence of groyne compartments at Highcliffe and Barton. Cliff stabilisation and protection has significantly reduced the volume of sediment available for longshore transfer. Drift rates are spatially variable, currently in the range of 5-20,000m3a-1. The proportion of sand declines from west to east, indicating that finer materials are winnowed offshore. Drift supply to Hurst Spit is almost entirely of gravel.
LT1 Hengistbury Head (see intro to littoral transport)
(See also the relevant section in the account of Poole Bay)
Littoral drift at Hengistbury Head has varied over the past 150 years, mostly due to human activity. Between 1848 and 1870, ironstone boulders were mined from the foreshore at Hengistbury and from Christchurch Ledge. This is believed to have caused marked acceleration of west to east littoral drift, which was previously partially blocked by the larger salient form of the headland (Tyhurst 1985a, Halcrow, 1999). Drift - especially of gravel - was again very substantially intercepted by construction of the Long Groyne in 1938 (Lelliott 1989). Two critical questions must be asked: (i) what is the natural potential for littoral drift?; (ii) to what extent is the Long Groyne a barrier to drift?
A substantial gravel and sand beach has accumulated to the west of the Long Groyne, indicating net eastward drift. Analysis of volumetric change at this barrier indicates drift of 600m3a-1 (1968-78) to 900m3a-1 (1938-68) but this only relates to material coarse enough to be retained (Webber 1980b). Fluorescent sand tracer experiments on Solent Beach indicate net eastward drift of 30,000-50,000m3a-1 but these figures are basically conjectural as they are unlikely to be temporally representative (Webber 1980b). A wave refraction analysis based on Portland wind data for 1977 was employed to determine longshore wave energy flux. Use of appropriate sediment transport equations calibrated for sediment size enabled a littoral drift calculation (Henderson 1979). A rate of 86,000m3a-1 was calculated at the Long Groyne, with a potential rate accelerating to 645,000m3a-1 northward to Mudeford Spit (Henderson 1979). Similar analysis using improved sediment transport equations indicated drift of 96,000m3a-1 to the immediate west of the Long Groyne and 575,000m3a-1 northward to Mudeford Spit (Lacey 1985). A further improved longshore wave energy flux technique suggested drift of 45,000m3a-1 towards the Long Groyne (Hydraulics Research 1986a). Several conclusions can be drawn from these estimates. The wide range of figures not only suggested that some of the analytical techniques were subject to substantial error but that different approaches measured different attributes. Beach volume analysis primarily measured coarse sand and gravel accumulation, whilst wave energy flux calculations were biased towards sand moving as both bedload and suspended load. However, it is apparent that the littoral drift rate accelerates markedly north-east of the Long Groyne, although it is probable that the values of Henderson (1979) and Lacey (1985) are large overestimates. This is because the refraction technique performed poorly in areas of complex bathymetry, eg. Christchurch Ledge, and did not adequately model the local diffracting/reflecting effects of the Long Groyne (Hydraulics Research 1986a).
The barrier effect of the Long Groyne is therefore difficult to assess. Some authors suggest that it prevents littoral drift (eg Wright 1982), whilst others state that it is only a partial boundary (eg Tyhurst 1985a). Limited evidence available favours the second explanation, with almost total interruption of gravel transport, but only partial interception of sand. A physical model study of the Long Groyne and co-adjacent areas at 1:80 scale employed 1.5mm diameter coal particles and 0.37mm sand grains to simulate natural (prototype) gravel and sand respectively. Operation of the model over a variety of representative wave conditions accurately replicated observed beach behaviour; furthermore, perspex in suspension was carried eastward by longshore currents and was pushed offshore by the Long Groyne. Transport was eastward or east-south-east towards Christchurch Ledge, with no evidence of a return (onshore) feed to Mudeford Spit. By contrast, no coal particles were carried around the Long Groyne (Hydraulics Research 1986a). Quantitative information could not be obtained as it was not possible to test a sufficient number of wave conditions, but mathematical model studies indicated net eastward or north-eastwards longshore bedload movement of approximately 45,000m3a-1 (Hydraulics Research 1986a); actual sand bypassing the Long Groyne may be at a similar order of magnitude. Reliability of this conclusion is uncertain because problems of model scale cause difficulty in selecting experimental material that is representative of the size and density of natural sediment. However, ability of the model to simulate observed beach changes probably indicates overall results to be of medium reliability. The model predictions broadly agree with previous conjecture and some field observations, eg it was reported that eastwards sand flow around the Long Groyne fed Highcliffe rather than Mudeford (Wise 1959). This was generally supported by the results of sea-bed drifter studies, which indicated a net offshore tendency west of Hengistbury Head, north-eastward transport over Christchurch Ledge towards Mudeford Spit and marked onshore transport between Mudeford Quay and Highcliffe (Tyhurst, 1976, Turner 1990).
Rock groynes constructed on Mudeford Spit between 1988 and 1990 and upgraded in 2000/01, rapidly filled with sediment (Photo 15) and indicated strong northward net drift fed by possible onshore supply (Tosswell 1978, Tyhurst 1987; Nicholls and Wright, 1991; Halcrow, 1999; H R Wallingford, 1999b;). Historically, this phenomenon has been marked by a highly dynamic phase of spit extension and punctuated breaching between the mid 1800s and 1938, which has been variously attributed to onshore transport (Robinson 1955) and littoral drift (Tosswell 1978) operating together (see Section 5.2). By assuming that subsequent spit regression resulted mostly from abruptly reduced littoral drift, a rate of 107,000m3a-1 has been calculated for the preceding period of most rapid documented change, ie between 1932-39 (Lacey 1985). Based on a limited experiment using aluminium pebbles, the bedload littoral transport rate immediately south of Christchurch Harbour entrance was computed to be 53,000m3 a-1 (Nicholls and Wright, 1991). H R Wallingford (1999b) calculated the prevailing rate to be in the order of 40,000m3a-1, based in part on numerical modelling of wave climate modified for local wave refraction. Most of this takes place in the nearshore zone with sand interception by the present system of rock groynes providing an estimated beach drift rate not greater than 16,000m3a-1. (Halcrow, 1999).
It must be concluded that a significant difference in littoral drift potential exists either side of the Long Groyne. West of the groyne, drift has been estimated by a variety of techniques and an approximate order of magnitude established. To the east, drift estimates are affected by complex refraction and diffraction effects. Despite this, a sharply accelerated rate of drift is indicated. Aeolian transport of sand along and across Mudeford Spit has been largely neglected, although HR Wallingford (1999b) estimate a gross rate of approximately 11,000m3a-1, with losses of some 9,000m3a-1 due to onshore and offshore movements; the net longshore throughput is therefore no more than 2,000m3a-1.
The Long Groyne is an effective impediment to littoral drift of gravel, although limited quantities may overtop the backshore part of this barrier during storms (Hydraulics Research 1986a). Sediment shortage is the main contributory factor to low rates of gravel movements with the sources of supply for this pathway uncertain. Onshore transport (F1, F2) from Christchurch Ledge has been proposed (Tosswell 1978) and may be supported by diver observations of mobile sand and gravel (Collins and Mallinson 1986; Gao and Collins, 1994c, 1995a). Halcrow (1999), however, infer that gravel is likely to be retained by the 'troughs' developed in the seabed bedrock outcrop of Christchurch Ledge.
LT2 Mudeford - Chewton Bunny (see intro to littoral transport)
This shoreline sector is suggested as a littoral drift sub-cell, because prior to major protection in the late 1960s, marked accretion was recorded in the vicinity of Highcliffe Castle. By contrast, the segment immediately west of Chewton Bunny exhibited erosion, so the original sub-cell boundary was located between these areas (Nicholls 1985). A similar distinction is made by Lacey (1985) based on analysis of wave energy flux, and attributed to variable offshore bathymetry, which causes complex wave refraction and longshore variation in littoral drift rate potential (Henderson 1979, Lacey 1985). It must also be recognised that coast protection structures intercept drift and generate artificial littoral drift boundaries. Construction of groynes at Highcliffe in and after 1971 intercepted littoral drift and therefore created new transport compartments at either end of the protected section; this was further enhanced by beach nourishment and construction of rock groynes further east in 1985. Subsequent monitoring of the beach involving measurements at 30 profiles over a 5 year period revealed a significant eastward shift of beach sediment identifiable in the first year (Hydraulics Research 1986b) and continuing in subsequent years (Hydraulics Research 1988; 1991). It was concluded that eastward drift was mainly responsible for the measured diminution of the nourished beach. The groynes were therefore not a total barrier to littoral drift but only intercepted sediment on the upper beach. Extended beach profiles showed that the nearshore and offshore zones were dynamic and it can therefore be postulated that offshore-onshore transport of sand during storm-swell conditions also involves a net eastward transport component which carries sediment around these structures (Lacey 1985). A mean drift rate could not be determined from beach profile measurements because inputs to the protected segment were not measured and nearshore-offshore information was limited. Calculations based on longshore wave energy flux and transport equations calibrated by measurement of sand loss on a replenished beach (Bournemouth) indicated an eastward potential drift of 191,000m3a-1 (Henderson 1979) and 195,000m3a-1 (Lacey 1985). These estimates may be unreliable due to difficulties with the refraction technique over the complex bathymetry of Christchurch Bay (Henderson 1979). Despite these problems, drift direction is correctly predicted according to observations by Robinson (1955), Tosswell (1978) and Tyhurst (1987), and all analysis indicates that drift is less rapid than between Hengistbury Head and the entrance to Christchurch Harbour.
Littoral drift input from Mudeford Spit to the Mudeford - Highcliffe sub-cell is possibly achieved by two mechanisms. Between the mid 1800s and early 1900s, growth of Mudeford Spit caused periodic lengthening of "The Run", building up a hydraulic gradient across the spit and facilitating breaching (Burton, 1931; HR Wallingford, 1999b). The detached spit was then driven onshore to supply Mudeford and Highcliffe (Robinson 1955). This process has not been operative since the 1930s, although it is believed that bypassing of the inlet is facilitated by sand drift via storage on an offshore bar (Tosswell 1978 - see Section 5.2 for further discussion). The relatively healthy sandy beaches at Mudeford (Photo 16) and Friars Cliff (Photo 3) are indicative of a process of continued onshore supply otherwise they would become depleted by the west to east net drift.
LT3 Chewton Bunny to Beckton Bunny (see intro to littoral transport)
Construction of groynes and a timber revetment at Barton-on-Sea as part of the 1964-69 protection schemes divided the shoreline into two littoral drift sub-cells: a natural western sub-cell with unimpeded transport and an eastern sub-cell subject to interrupted transport by groynes (Nicholls 1985; Bray and Hooke, 1998). Compartmentalisation of the beach has subsequently been increased by construction of four rock groynes/strongpoints.
West Barton beach was monitored by intensive cross-section profiling involving 47 surveys along 12 survey lines over the period 1976-78 (Webber 1980b, Halcrow 1980). These revealed net accretion of 16,000m3a-1 over the period, but probably were not representative of long-term trends because net intertidal erosion of 3000m3a-1 and nearshore erosion of 20,000m3a-1 was recorded from a larger data set covering the period 1974-82 (Lacey 1985). These data are not easily converted to littoral drift volumes because these processes may operate with no change in beach volume. Estimation of drift is possible by integration of all known inputs and outputs, but has not been undertaken for this sector. Profile analysis of the protected East Barton beach revealed intertidal accretion of 2,000-4,800m3a-1 and nearshore accretion of 3,000-63,000m3a-1 over the period 1976-84 (Lacey, 1985). Although it is difficult to determine longshore transport, a distinct tendency exists for net beach erosion in the western (unprotected) part and accretion in the east where groynes intercept material (Mackintosh and Rainbow, 1996). Profiles indicated that onshore-offshore exchanges of sand were related to development of summer (net onshore transport) and winter (net offshore transport) were much greater than longshore changes (Halcrow 1980, Webber 1980b). Additionally, it was apparent that the sand component was more dynamic.
Sand transport was examined by fluorescent tracer experiments in 1976 and 1977 (Babbedge 1987a and b, Webber 1980b, Halcrow 1980). Reliable drift estimates were only obtained for the first two tides of the main experiment, thereafter the technique proved inaccurate. These rates were combined with mathematical calculation of longshore wave energy flux to calibrate sand transport equations, and were subsequently used to calculate net drift from an annual wave climate. Drift volumes of 109,000m3a-1 and 220,000m3a-1 were obtained with the first and second tides respectively. These are not representative annual drift volumes because experimental data covered very restricted wave conditions untypical of overall wave climate. Better estimates of littoral drift were obtained by including refraction effects to generate longshore wave energy flux and combining these with previously established sediment transport calibrations. Analysis of this type for west Barton indicated drift of all sediment types combined between 167,000m3a-1 (Henderson 1979) and 308,000m3a-1 (Lacey 1985), the latter employing an improved sand transport calibration. Rates were determined at regular intervals along the coast (ie not based on natural transport boundaries), thus the drift rates quoted are applicable from just west of Chewton Bunny to the approximate boundary between Barton West and East beaches. Much lower drift rates of between 2,000m3a-1 (Henderson 1979) and 21,000m3a-1 (Lacey 1985) were determined for the sector covering East Barton beach. This was attributed to diminished longshore wave energy flux due to a local refraction effect, and lower drift efficiency due to increased sediment retained on the beach. Although incomplete understanding of the relative importance of factors and processes precludes reliable evaluation, the predicted eastward reduction of drift rates is corroborated by visual evidence of beach and nearshore accretion.
The eastern transport boundary is marked by Beckton Bunny outfall and approximately coincides with the artificial boundary utilised in the refraction analysis of Henderson (1979). Eastward drift of both sand and gravel has been substantially intercepted at Beckton Bunny between 1971 (when the outfall was reinforced and a rock and sheet-pile strongpoint was constructed) and 2001. Cliff input analysis suggested that gravel drift past Beckton Bunny may have been 7,600m3a-1 for the period 1867-1971, thereafter diminishing to almost zero with outfall modification (Nicholls 1985). This figure is within the range of the energy flux derivations and provides a fair degree of corroboration.
LT4 Beckton Bunny to Hordle (see intro to littoral transport)
Estimates of littoral drift using the energy flux technique range from 2,000m3a-1 (Henderson 1979) to 40,000m3a-1 (Lacey 1985), the difference resulting from sediment size calibration; the lower rate assumes an all-gravel beach, as opposed to a 54:46 sand: gravel beach used for the upper rate. Theoretically, the latter should be more accurate because sediment size calibration was based on beach sampling (Lacey 1985). An alternative method based on historical changes of the beach, cliff and nearshore zone determined the minimum littoral drift necessary to produce the measured change. This analysis showed that gravel drift diminished from between 8,400-13,800m3a-1 for 1939-68 to 7,000m3a-1 for 1969 to 1982 (Nicholls 1985). The reduction was attributed to the Barton Cliffs protection programme and interception of drift by the Beckton Bunny outfall. Long-term beach gravel accretion in the form of a small cuspate foreland with a volume between 0.5 and 1.5 million m3 is recorded at Hordle Cliff from map comparisons, at rates of 2,600-2,900m3a-1 for 1867-1939, 8,400m3a-1 (1939-68) and 5,900m3a-1 for 1958-82 (Nicholls 1985, Halcrow, 1999). This location is widely recognised as a transport discontinuity, with accretion occurring as the littoral drift rate reduces to near zero (Nicholls 1985, Nicholls and Webber 1988a, Halcrow, 1999). The effect is believed to be due to local wave refraction which causes a relatively low longshore wave energy flux (Nicholls 1985). Map analysis indicated that this accretion zone migrated 500m to the west between the periods 1867-98 and 1938-68, reflecting supply from the west at a more rapid rate than output to the east. After 1968, this differential was reduced; westward migration ceased and a reverse trend for eastward migration was initiated (Nicholls 1985). Beach profiling between 1981 and 1982 revealed a loss of 24,000m3 of predominantly coarse sandy sediments at Hordle, but it was concluded that survey duration was too limited to distinguish any long-term trend (Nicholls 1985). Halcrow (1999) indicate that beach erosion has been prevalent, albeit at a slow rate, since the early 1980s.
LT5 Hordle to Milford-on-Sea (see intro to littoral transport)
The Hordle drift boundary has not acted as an absolute barrier to littoral transport, but its effect has been progressive, allowing 65% of longshore movement to pass in 1867-98, 61-62% in 1908-1939 and 0-36% in 1939-68 (Nicholls 1985). Due to the intercepting effect of this foreland, drift between Hordle and Milford diminished from a maximum of 4,900m3a-1 during 1939-68 to 1,400m3a-1 between 1969-82 (Nicholls 1985). It is reported that timber groynes have been present on the Milford frontage since 1937; three rock groynes or strongpoints were constructed at the eastern extremity in 1983 (Dobbie and Partners 1984). It is likely that these structures have been effective in intercepting material so that gravel transport downdrift of Milford (ie towards Hurst Spit) is now extremely limited (Nicholls 1985). Longshore wave energy flux analysis indicated potential drift of 7,000m3a-1 of gravel in the west, reducing to 2,000m3a-1 at Milford (Henderson, 1979) and 71,000m3a-1 of both sand and gravel reducing eastward to 21,000m3a-1 (Lacey 1985). The higher values are theoretically more accurate, but their discrepancy with historic rates suggests three possible sources of error: (i) Henderson's analysis was only for 1977, which may be unrepresentative of long term conditions; (ii) wave refraction analysis was inaccurate over complex bathymetry; (iii) the sediment transport calibrations employed are not suitable for this site, with Henderson (1979) failing to represent the movement of sand. This is clearly a larger volume than gravel.
LT6 Hurst Spit Beach (see intro to littoral transport)
Littoral drift has been studied on Hurst Beach using a variety of techniques which have included historical planform changes, beach profile analysis, aluminium pebble tracer experiments and numerical modelling. (See Section 5.3). A 500m segment at the landward root (proximal sector) of the spit opposite Sturt Pond has been protected by rock armour since the late 1960s. The effects of this protection reduced natural drift because of interference with longshore energy flux (Bradbury 1998), but was modified in 1996 when this armouring was reinforced and its gradient reduced.
Analysis of historic changes by map comparison yielded eastward drift of 7,500-18,300m3a-1 over the period 1867-1968 (Nicholls 1985). Beach profiles surveyed on 27 occasions over the period 1980-82 indicated drift of 11,000-13,000m3a-1 along Hurst Beach, increasing to 13,000-15,000m3a-1 at the distal end (Hurst Castle) (Nicholls 1985). Direct measurement of littoral drift was undertaken by means of aluminium tracer experiments. Quantitative analysis was difficult due to low tracer recoveries (1-25%) caused to some extent by burial and rapid longshore transport, but especially by inferred offshore transport (Nicholls 1985, Nicholls and Webber 1987b). Reliable results were only obtained for a 7 day period before recoveries diminished to unacceptably low levels. Wave conditions over this period were insufficiently representative for accurate calibration of transport equations. Experiments were therefore of limited use for volumetric calculation of drift, but yielded valuable information on particle sorting. These were most clearly sorted according to their 'C' axes, with larger material preferentially transported either towards Hurst Castle (distal end) or moved down the beach face to the lower foreshore and nearshore areas (Nicholls 1985, Nicholls and Webber 1987b).
Longshore wave energy flux determined for Hurst Beach by wave refraction analysis predicted westward net drift (Henderson 1979). This contradicts reliable field observation and historical information and refutes the long accepted theory of spit elongation by eastward feed (Lewis 1938, King and McCullough 1971). This analysis is therefore clearly in error, probably resulting from failure of the refraction analysis to accurately predict waves travelling over the complex bathymetry of the Shingles Bank, North Shoal and North Channel. A more refined refraction analysis was employed to determine the inshore wave climate (Halcrow 1982). Littoral drift was calculated from longshore wave energy flux, using adjustments to simulate variable tidal range and differing transport threshold conditions for the various sediment sizes. This technique yielded a net eastward drift of 15,000m3a-1. Brampton (1993), however, calculated that the mean drift rate for 1974-1990 was slightly over 36,114 m3a-1 (with a maximum of 51,600 and a minimum of 17,000 m3a-1). This work was based on numerical modelling of the inshore wave climate, using a bulk sediment transport formula applied to monthly averaged wind speed data input into a wave hindcasting approach. Highest monthly rates, at 14,000 m3a-1, occurred consistently in February of each year. No evidence for drift reversal was detected. These figures are conisdered an exaggeration of actual rates, as they are essentially based on theoretical reasoning.
The wide range of techniques used to determine drift on Hurst Beach produced results that indicated that this pathway has been established with high reliability, but there is some large measure of uncertainty regarding actual rates.
LT7 Hurst Castle (Hurst Point) to Point of the Deep (North Point) (see intro to littoral transport)
Much of the predominantly coarse gravel moving northwards between Hurst Castle and North Point is lost offshore to Hurst Narrows, but a small quantity is transported northeast around the distal point and westwards along the recurve. The drift rate is estimated to be 2,000 to 2,700m3a-1 at Hurst Point, declining to 900m3a-1 at the northern recurve tip (Halcrow, 1982, Nicholls 1985).
Interpretation and Summary
4. SEDIMENT OUTPUTS - References MapTwo types of output can be recognised:
4.1 Transport In The Offshore Zone - O1 O2 O3 O4 References Map
O1 Shingles Bank
Research into present-day sediment mobility has utilised: (i) side-scan sonar and echo sounding to map bedforms; (ii) sediment sampling, to infer pathways of movement; self-generated noise of gravel particle collisions, and measurements of near-bed tidal current and wave-induced stresses (Dyer, 1970; Velegrakis and Collins, 1992, 1993 and 1994; Velegrakis, 1994; Voulgaris, Workman and Collins, 1999). In addition, there has been some limited use of numerical modelling to derive information on probable sediment movements by waves and residual tidal currents (HR Wallingford, 1994; Brampton, et al, 1998; Halcrow, 1999).
All researchers identify strong offshore-directed (ie north-east to southwest) movement of sand and gravel in the Needles Channel due to the dominant ebb tidal current. The effectiveness of tidally-induced scour is demonstrated by the sharp boundary between the Chalk platform north of the Needles headland and the adjacent channel. In places, this sweeps clean the bedrock surface (Velegrakis, 1994). Gravel waves, with sharp crests, have an asymmetry that clearly indicates sustained south-westerly transport of coarse material along the easterly flank of the Shingles Bank (Velegrakis, 1994). The movement of sand in the same direction was deduced by Dyer (1970) from patterns of sand ribbons. HR Wallingford (1994) identified a separation of the south-west moving ebb current, with strongest flow (an average of 1.8 m sec-1 just above the seabed) along the western boundary of the Needles Channel. This was derived from application of TELEMAC, a depth-integrated finite element predictive model.
Sediment movements over the Shingles Bank are considered to result from stresses due to both waves and tidal currents, with waves effective to depths of at least 8 m. Planar areas of the crest of the bank may be due to wave abrasion, probably under storm conditions; Velegrakis and Collins (1993), however, suggest that they are gravel-armoured surfaces from which sand is winnowed by tidal currents. This would be a year-round effect, thus the presence of sand patches overlying gravel - which appears to be confined to the winter months - indicates wave transport of sand. It is possible that tidal currents effect some abrasional scour, particularly on the eastern flank of the Shingles Bank (Velegrakis and Collins, 1993; 1994). Nicholls (1985) reported that parts of the crestal area build up to above maximum low water during periods of moderate wave energy, and are lowered during storms - ie sediment is dispersed under high wave energy conditions. Analysis of bathymetric and volume changes in the area of the 1996 dredging led to the conclusion that most transport was induced by waves (New Forest District Council, 1997-2001).
Voulgaris, Workman and Collins (1999) undertook an experiment that measured self-generated noise resulting from mutual collisions of gravel particles (median grain size of 1.70 cm) at the seabed, at a mean water depth of 8 m,. This was conducted over six consecutive days in an area near the north-western edge of the Shingles Bank, with well-developed gravel waves. Bedload transport rates remained nearly constant despite periodic fluctuations in the velocity of tidal currents, and variations in prevailing wave height. Waves were therefore considered to be the primary transportation mechanism, with transport rates correlating closely with particle weights. The relative dominance of wave action is likely to increase with water depth, ie southwards and southwestwards.
Historically, erosion has been the dominant trend along the western flank of the Shingles Bank, with accretion (progradation) of its eastern margin (Velegrakis, 1994). Thus, a broadly west to east transport pathway has prevailed, involving the movement of both gravel and sand. Velegrakis and Collins (1993; 1994) report that the eastern flank exhibits an eastwards-thinning wedge containing fine to medium sand some 10-15 m in height. This is better-sorted than the sand fraction present along the western flank, and may indicate selective transport by tidal currents. Waves are considered the principal mechanism for moving coarse sand and gravel along this pathway.
Velegrakis (1994), revealed ambiguous evidence for net northwards or northwestwards movement of both sand and gravel along the western flank and in the area of North Head Shoal in the early 1990s; this was confirmed by New Forest District Council (1997-2001) on the basis of repeated bathymetric surveys which show the emergence of a distinct extension to North Head. Bedforms in this area lack any "steady state" asymmetry, tending to reverse with tidal current direction (Velegrakis, 1994). Thus, a combination of both tidal and wave-assisted transport is implied. North Channel was considered by H R Wallingford (1994) to be a product of tidal scour, though maximum (ebb stage) current velocities do not exceed 0.9 m s-1.
02 North Channel to Hordle: west-central Christchurch Bay
Asymmetrical gravel waves recorded by sediment sampling and side-scan sonar on the floor of North Channel indicate net north-westwards transport (Dyer, 1970; Velegrakis, 1994). A drifter experiment (Clarke and Small, 1967) also suggested movement in this direction, as well as onshore; however, other drifter studies (Watson, 1975; Turner, 1990) have given contrary results, leading to some confusion. Velegrakis and Collins (1994) note the sharp boundary between the western flank of the Shingles Bank and the adjacent area of the sea bed of Christchurch Bay. This may be the result of efficient "sweeping" by tidal currents, which helped to create a re-entrant feature. This boundary may be further evidence of sediment moving north west, towards the nearshore zone between Milford and Hordle, which might provide sand for nearshore accretion. However, both Lacey (1985) and Nicholls (1985) identify a trend of net erosion immediately off Milford for the period 1880-1970, thus this sand may be moved westwards, via bar and trough topography, to the offshore pathway. A small-scale, and apparently closed, anticlockwise sand transport system is thus inferred.
Velegrakis and Collins (1994) report that over most of this area - as well as seawards - sand transport rates are low; sand may only be mobilised when peak tidal currents and waves with heights greater than 1 m combine. Where water depths are less than about 9 m, slope gradients rarely exceed 0.5o, and are less than 0.1o close to the position of maximum lower water (Halcrow, 1999). The inshore zone is demarcated seawards by a distinct break of slope, with water depths increasing to up to 20 m immediately north of Dolphin Bank. Sediment transport in this deeper area has not been systematically researched, but there is no direct evidence that sand moved offshore by the W03 pathway reaches this area.
O3 Feed to Dolphin Bank
Offshore survey by echo-sounding, side-scan sonar and sediment sampling revealed sand megaripples to the south of Dolphin Bank which indicated northward transport onto and across the bank (Velegrakis 1994), in addition to westwards movement, also determined by Dyer (1970). Southward sand transport is indicated by bedforms several kilometres to the north of Dolphin Bank which suggests that a sediment sink should exist in the central - east part of Christchurch Bay. However, sub-bottom profiling has revealed only a thin sediment cover and much bedrock exposure in this area. It is therefore concluded that sand is highly mobile in Christchurch Bay and transport pathways may vary seasonally in their relative strengths (Velegrakis 1994, Brampton et al, 1998). Survey was completed during calm summer conditions and some of the bedforms identified may not be representative of long-term transport pathways, thereby explaining the partial disparity between inferred sediment transport directions and patterns of sediment accumulation. These indicated transport pathways are therefore of low to medium reliability and require verification. By contrast, survey in the vicinity of Shingles Bank (Velegrakis and Collins 1993; Velegrakis 1994) confirmed the transport pathways previously identified by Dyer (1970). This work indicated a capacity for long-term stability of the sediment transport pattern in this area (described in Section 6).
04 Transport from Dolphin Bank to Shingles Bank
Dyer (1970) discerned, from the evidence of the asymmetry of sand waves resolved by echo-sounding, side-scan sonar and sediment sampling, the westward transport of progressively finer sands from the southern Needles Channel to Dolphin Bank. Velegrakis (1994), using similar survey techniques and applying sediment sorting co-efficients, concluded that there was a net westwards, and possibly south-westwards, pathway of sand movement over only the western and central areas of the Bank. Sediment movement along its northern flank, and at its eastwards extremity, was found to be moving north and north-eastwards. This implies a bedload transport parting close to the central crest. Velegrakis and Collins (1993) argue that superficial sand deposits that accumulate during the winter months over parts of the Shingles Bank derive from the eastern part of Dolphin Bank, where water depths are less than 8 m. This material is mobilised and transported by waves, with the flood tide current a possible auxiliary mechanism. Once sand has moved across the crest line of The Shingles Bank, it is temporarily deposited on its eastern flank. The finer texture and better sorting of sand on this flank, compared to its western margin, is taken as a strong implication of a west to east transport pathway. To maintain this transport system, Velegrakis (1994) argued that there must be a supply of sand, moved by tidal currents, in the southern Needles Channel. Brampton et al (1998) consider that as this source would supply only very fine sand, the composition of Dolphin Bank supports the probability of additional supply from the south (03). Some transfer from Dolphin Bank to Dolphin Sand is also considered likely, on the basis of analysis of bedforms in the intervening area (Brampton et al, 1998).
Near sea-bed current metering, undertaken during both the flood and ebb stages of neap and spring tidal cycles, indicated that tidal current velocities were sufficient to account for all observed movements of sand on, and marginal to, Dolphin Bank. This is confirmed by the modelling of tidal residuals from TELURAY and the application of critical flow velocity thresholds to that data (Brampton et al, 1998).
The contribution of wave action to sand movement has not been determined, but is probable under storm wave conditions. This is especially likely where northward-moving sand is transported along the westward flank of The Shingles Bank, and its north-westerly salient.
Both Nicholls (1985) and Lacey (1985) determined, from comparisons of hydrographic charts published between 1880 and 1968, that there was a net loss of sediment in the area immediately north of Dolphin Bank. This was estimated to have been approximately 60,000 m3a-1 over this period. Velegrakis (1994) carried out a similar analysis, covering 1979-1990 and identified a continuing erosional trend. Sediment, mostly fine sand, moving offshore from Hordle (W03) cannot therefore be directed to this area; its ultimate destination is not clearly understood, but may be contained in a closed anticlockwise sub-cell circulation system.
Survey and research analysis has been largely confined to the offshore area in the central and western parts of Christchurch, and in particular the Shingles Bank. There is, as yet, no conclusive evidence in favour of a confined, anticlockwise circulation pattern for bedload transport of either sand or gravel. Velegrakis (1994) and Brampton, et al (1998) state that several discrete areas of smalls and bedforms located on the seabed immediately east of Christchurch Ledge indicate a net north-east to south-west transport pathway. If this is a permanent feature, there can be no supply of sand back to this area from Dolphin Bank and Dolphin Sand. Hydrographic chart analysis indicates net erosion of the seabed of Christchurch Bay north of the Outer Banks, thus implying that sediment is leaving the system. In the nearshore zone, both net onshore (F1-F3) and net offshore (W01-W03) pathways have been recognised. These may be components of a sequence of small, anticlockwise-circulating systems of transport; Halcrow (1999) regard the onshore pathways as "sand bridges" that help to sustain the temporally variable pattern of nearshore sand ridges and troughs.
Tidal currents are considered the main mechanism of sand transport, but wave-induced stresses over shallower areas of the seabed are likely to promote onshore movement, especially during the winter period.
Gravel transport is apparently confined to the Needles Channel and The Shingles Bank. Survey data and analysis of bathymetric change indicate both tidal and wave-induced mobility of gravel, with complex on and offshore patterns of movement. The Shingles Bank - and arguably the Dolphin Bank - constitute a large ebb delta established at the exit of the West Solent. It is less certain if the Needles Channel provides a significant pathway for gravel supply to Pot Bank.
4.2 Estuarine Output - EO1 EO2 EO3 References Map
EO1 Christchurch Harbour
Marine sediment input is possible at Christchurch Harbour entrance and a significant flood delta of well sorted sands and sandy gravels is located immediately inside the entrance (Murray 1966, Tosswell 1978, Gao 1993; Gao and Collins 1994a, b c). This feature is attributed to transport and deposition on the flood tide (Murray 1966, Gao 1993), but detailed current metering at the entrance has revealed that ebb flow (peak surface velocity 1.9ms-1) is significantly stronger than the corresponding flood (1.15ms-1). Bedload transport is therefore expected to be predominantly out of the harbour, unless ebb and flood flow follow different channels, or flood flow coincides with storm waves to create intermittent sediment pulses into the harbour (Tosswell 1978). Detailed sediment analysis involving sampling at 316 sites within the harbour revealed that sand is the dominant sediment type and that clay and silt content is generally low. This indicated that suspended sediments tend to be moved out of the outer harbour (Gao and Collins, 1994a and b). However, analysis of foraminferal tests (Gao and Collins, 1995a) indicated that those with an external (marine) provenance are predominant in sediments in the outer harbour, thus demonstrating that there is net accumulation here of biogenic debris introduced by tidal currents. This material compares closely with fine to medium sand in terms of transport potential.
Current metering at the entrance has revealed that ebb tidal flow is significantly stronger than the corresponding flood (Tosswell 1978). Gao and Collins (1994b) give the mean ebb current speed to be 0.55 m s-1, that of the flood 0.165 m s-1. Maximum ebb velocities can attain 2.5 m s-1. This is because mean discharge from the rivers Stour and Avon is equivalent to the spring tidal prism, calculated as 1.06 x 106m3 (Gao and Collins, 1994, a; 1995, b). These hydraulic conditions indicate strong potential for bedload transport out of the harbour, but it is likely that actual output is small due to limited sediment supply to areas near the entrance where entrainment and seaward flushing are possible (Halcrow, 1999). It has been suggested that marine inputs under surge conditions may be significant, as indicated by a predominantly sandy flood delta inside the harbour entrance (Gao and Collins, 1994a, c; 1995 b). Samples of exotic benthic foramiferal tests taken from harbour sediments (ie derived from marine sources) indicate that their percentage in comparison to estuarine species is highest close to the entrance. Thus, marine-derived sediment (tests have similar mobility to sand grains) may not penetrate in quantity beyond the flood delta. The major effect of strong ebb currents is apparently to flush offshore any sediment transported longshore into the entrance channel from the opposing spits that confine the width of the mouth of the harbour. Thus, the channel is a littoral drift barrier, but with bypassing possible via the offshore bar (ebb delta) (Tosswell 1978, Gao and Collins, 1994b). A quantity of fine suspended sediment is supplied by erosion of the harbour margins and seabed and by fluvial input (see Section 2.2). Despite this, the harbour sediments are predominantly sandy and reports suggest a trend for increasing sandiness (Gao 1993, Gao and Collins, 1994c). Hydraulic factors indicate that because the harbour has a high tidal exchange ratio there is a very thorough flushing effect on each tide (Tosswell 1978, Gao, 1993; Gao and Collins, 1994a; 1995b). Qualitative observations indicate that sediments are readily resuspended by wave action inside the estuary and are efficiently discharged by ebb tidal flow (Hydraulics Research, 1987). Incoming flood flow from Christchurch Bay is significantly less turbid (Gao 1993). Gao and Collins (1994a) applied transport formulae to the relationship between mean current speeds through the harbour entrance; water level of the entire estuary basin and freshwater discharge. Bedload transport at the entrance operated at a much higher rate when tidal and wave-induced currents acted in combination. No net transport due to the flood tidal current alone was determined, but sediment discharge of a potential 8 x 10-2m3ms-1 is considered possible when maximum ebb currents combine with high freshwater discharge. Thus, ebb transport is between one and two magnitudes greater than the flood, moving between 11,000 and 30,000 m3a-1 towards the ebb delta. This is net output - gross movement might be in the order of 100,000 m3a-1.
The shallow entrance channel (mean width of 47 m) has had a stable configuration - partly due to protection of Mudeford Quay - and cross-sectional area since at least the late 1930s (Gao and Collins, 1995b), thus indicating that it has achieved an equilibrium condition. Sediment arriving at the entrance via littoral transport is, according to both observational evidence and mathematical modelling moved predominantly offshore (Gao and Collins, 1994a and c; 1995 a and b).
Christchurch Harbour, with a mean tidal basin area of 1.9 km2 and a mean water depth of 2.0 m, is occupied by sandy gravels, sands, muddy sands and - towards its inner margins - silty muds. Average sediment thickness is 1.5 m, increasing to over 2.0 m in areas of bank accretion. There is a general consensus that fine sediments delivered by fluvial discharge are trapped within the estuary, and contribute to the prevailing sedimentation rate of 0.1 to 0.4 mm a-1 (Gao and Collins, 1994a; c). Fine sediment introduced by the flood tide may not, however, settle out, and is moved out of the basin on the ebb current (Halcrow, 1999). There is no quantitative evidence to support this assertion, but Gao and Collins (1995a, b, c) have indicated that spring tide ebb currents have substantial capacity to entrain fine to medium sand in the vicinity of the flood delta. Some internal re-distribution of sediment occurs in the vicinity of the northern coastline; this is apparent from analysis of sorting co-efficients and the evidence of small bedforms (Gao and Collins, 1994c). There is no reported evidence of erosion of mudflats and saltmarsh margins, but abrasion of sub and inter-tidal flats may nonetheless be sources of suspended sediment loads. Maximum wave heights within the harbour vary between 0.4 and 0.8 m (1 in 10 and 1 in 100 year frequency, respectively, Hydraulics Research, 1987; Halcrow, 1999) and may effect some marginal erosion. Harbour beaches are not well-developed, but the presence of gravel in places is probably derived from erosion of fluvial terrace deposits at a stage when the harbour entrance was more open to wave action, ie pre-dating the development of the present-day spits, see Section 5.2.
Both high and low saltmarsh are present, the latter colonised by Puccinellia and several other co-dominant grasses. Only small quantities of Spartina anglica are present, with no evidence of formerly more extensive occupance. The sedimentation system of Christchurch Harbour has not, therefore, been significantly affected by the spread and subsequent "dieback" of this species, as experienced in most other south coast estuaries. Any loss of saltmarsh seems to have been counterbalanced by its re-colonisation of abandoned artificial salt pans. Extensive Phragmites reed beds occupy tidal creek margins and areas of elevated marsh, where they front wet grazing meadows.
The overall sediment budget of Christchurch Harbour is positive, as evidenced by net vertical accretion. Small, but unquantified, inputs have derived in the past from the dredging of the lower channels of the Avon and Stour as part of flood alleviation schemes. Grimby Marsh, for example, is substantially composed of spoil material. It is uncertain if this artificial source of sediment accretion continues at present. In overall terms, this estuarine system is a sediment sink.
EO2 Hurst Narrows
Ebb tidal currents are shorter lived, but more rapid, than corresponding flood currents at Hurst Narrows (Webber 1980a), thus a dominant south-western transport pathway extends along the Needles Channel. Peak surface velocities of up to 3ms-1 have been recorded (Heathershaw and Langhorne 1988) and up to 1.5ms-1 on the beach face at Hurst Point (Nicholls 1985). Further offshore, peak surface and near bottom velocities of 2ms-1 and 1.2-1.4ms-1 respectively have been measured in the channel immediately east of Shingles Bank (Velegrakis and Collins, 1992; Velegrakis 1994). Hurst Spit beach between the fossil recurves and Hurst Point slopes steeply seaward and is subject to rapid tidal current scour. These factors cause significant gravel removal, where material is readily entrained and transported seaward (Dyer 1970, Nicholls 1985, Velegrakis, 1994). This was calculated at 9,200-15,400m3a-1 for 1939-68, based on map comparisons, and 8,700-11,000m3a-1 for 1980-82, based on profile information (Nicholls 1985). Surveys involving echo sounding, side-scan sonar and sediment sampling reveal the sea-floor at Hurst Narrows and westward along the Needles Channel to be composed of spreads of mobile gravel with intermittently present overlying sand and intervening patches of current swept bedrock (Dyer 1970, Velegrakis and Collins, 1992; 1993; Velegrakis 1994). Asymmetric gravel waves or megaripples have been identified throughout this area and all surveys have identified a consistent trend for south-westwards gravel transport along Hurst Narrows channel. Diminishing currents down this tidal transport pathway result in deposition of sediments, according to grain size. In this manner, some coarse material may be deposited on Pot Bank, but the majority is recirculated northward onto Shingles Bank (Dyer 1970, Hydraulics Research 1977; Velegrakis, 1994; Brampton, et al, 1998). Some sands are transported south of Pot Bank and deposited on a crescent-like feature, but the majority is transported westward towards Dolphin Bank and Dolphin Sand (Dyer 1970). There is no indication of return northward transport from Pot Bank or the southerly sand features, so the limited evidence available suggests that these are sediment sinks (Hydraulics Research 1977, Brampton, et al, 1998). A more detailed examination of the evidence for transport pathways in this area is given in Section 4.1.
EO3 Input to West Solent
Suspended sediment analysis using Landsat imagery (MacFarlane, 1984) revealed high concentrations in Christchurch Bay, with a marked peak in the winter months (Lacey 1985). This implied cliff erosion to be the major source because it too shows similar seasonal variation. Examination of suspended sediments in the Western Solent using remote sensing also yielded a pattern of peak winter concentrations and indicated significant transfer of suspended sediments from Christchurch Bay to the West Solent (Srisoengthong 1982, Lacey 1985). Net suspended transport is likely to be into the West Solent at Hurst Narrows due to the greater duration of the flood current (Webber 1980a). Sediments transported into the West Solent in this manner are deposited on the north-west shore of the West Solent and therefore represent an output from the Christchurch Bay system. Small quantities may be returned to Christchurch Bay because Hurst Spit is receding over the Keyhaven marshes and approximately 9,000m3a-1 of fine sediment is released by erosion of the seaward face (Nicholls 1985). However, the precise fate of this material is not known (Posford Duvivier, 1999). HR Wallingford (1994) undertook a brief programme of sampling of suspended sediments off the distal point of Hurst Spit during spring tides. Concentrations varied from 0.2 to 30 mg/l, with a mean value of 5.0 mg/l. The suspended flux of silt and clay over a single tidal cycle may be 2 to 15,000 kg s-1, but how much is lost to either the Western Solent or offshore is not known.
4.3 Wave Driven offshore Loss - WO1 WO2 WO3 References Map
WO1 Hengistbury Head
Sea-bed drifter studies, tracer experiments and physical model studies all indicate net offshore transport at Hengistbury Head (Watson 1975, Tyhurst, 1976; Wright 1976, Webber 1980b, Halcrow 1980, Wright 1982, Hydraulics Research 1986a). Physical model studies showed that offshore transport was predominantly of sand in the vicinity of the Long Groyne. The model covered a limited area and no return shoreward feed was recognised (Hydraulics Research 1986a). This may indicate output to a sediment store or sink to the south of Christchurch Ledge, although the consensus view has been that sediment is swept north-west off the Ledge, where it may contribute to the beaches of Mudeford Spit and Highcliffe via the sandbank system of the ebb tidal delta seaward of the entrance to Christchurch Harbour.
Examination of beach profile data covering the period 1974-82 yielded erosion of the intertidal zone of West Barton by 1,000m3a-1 and the nearshore zone, to -5m O.D., by 19,000m3a-1 (Lacey 1985). This coincided with erosion of the offshore zone by 21,000m3a-1 and lowering of 2m determined from bathymetric chart comparisons for the period 1880-1968 (Lacey 1985). Substantial offshore loss was also predicted by wave energy flux analysis, which indicated 165,000m3a-1 (Henderson 1979) and 296,000m3a-1 (Lacey 1985). This analysis may be subject to error, thus the reliability of this offshore pathway is uncertain; intensive profiling covering the period 1976-77 revealed net beach accretion and significant short-term variability due to slow onshore bar movement in calm conditions and its rapid destruction and offshore transport during storms (Babbedge 1976a, Webber 1980b, Halcrow 1980).
Beach profiling over the period 1981-82 revealed a loss of 23,500m3. Seasonal morphodynamic variation was significant, with offshore loss during storms and onshore input during calm conditions (Nicholls 1985). The long-term trend determined from hydrographic chart comparisons was for net accretion of 2,600-2,900m3a-1 for 1867-1939, rising to 8,400m3a-1 for 1939-68 but then declining to 5,900m3a-1 for 1968-78 (Nicholls 1985). Velegrakis (1994) also identified net accretion, 1979-1990, from chart analysis, although quantities are not given. Several possible conclusions may be drawn;: (i) profile data was collected over a limited time period and was unrepresentative of long-term trends; (ii) chart analysis may reflect an underlying trend for accretion at Hordle; (iii) profile data extended to -1m O.D., further seaward than the MHW line examined by chart comparisons, thus the erosion revealed may reflect sediment loss restricted to the lower foreshore and nearshore zones. Comparison of charts covering the period 1880-1968 revealed sea-bed lowering up to 3m off Milford which supports this latter possibility (Nicholls 1985, Lacey 1985). Offshore survey by echo sounder, side-scan sonar and sediment sampling revealed asymmetric sand megaripples which clearly indicated offshore transport of sand (Velegrakis 1994). The fate of this material is uncertain because although bedforms indicated net transport towards central-east Christchurch Bay sub-bottom profiling showed that sediment cover here was thin (Velegrakis 1994). A corresponding pathway was not identified by Dyer (1970). Compilation of a beach sediment budget for the eastern part of Christchurch Bay and detailed beach sediment sampling revealed progressive eastward offshore sand loss (Nicholls 1985). Thus, an offshore pathway has been indicated, if not proven, by different survey and analytical approaches.
Sediment eroded from beaches in Christchurch Bay does not balance with inputs from littoral drift, onshore transport and cliff erosion, so there is a presumed net loss of beach sediment offshore (Lacey 1985). A major component of this loss is progressive offshore transport of sand eastward towards Hordle and Milford (Nicholls 1985; Nicholls and Webber, 1988 and b). This process is confirmed by bedform morphology (Velegrakis 1994) and the high glauconite content of sea-bed sediments, which is characteristic of the eroding cliffs and shoreface (Dyer 1970, Velegrakis, 1994). The majority of sea bed sediments in Christchurch Bay are thus derived from long-term cliff erosion. Clays are transported out of Christchurch Bay in suspension (Lacey 1985), whereas sand is transported as bedload offshore and deposited in the outer Bay (Dyer 1970, Nicholls 1985, Velegrakis 1994). Gravel is retained on the beach, but is transported eastward by littoral drift and is subject to some abrasional wear. Coarse material reaching the distal zone of Hurst Spit is transported offshore by rapid tidal currents at Hurst Narrows, and into the Needles Channel. Hydrographic chart comparisons of the whole of Christchurch Bay revealed an equivalent net loss of 505,000m3a-1 over the period 1868-1968 (Lacey 1985). If this is accurate, it indicates that it is a site of net output, with a continuing tendency for selective offshore loss of beach sediment. It is therefore not a fully self-contained or confined transport cell.
5. SEDIMENT STORES AND SINKS - References Map
Specific details of beach morphodynamics, most of them relating to short-term and relatively recent surveys and analysis are given in the preceding section on littoral transport and are also detailed in Webber (1980b) and Nicholls and Webber, (1988a and b). The first section below provides a longer-term overview of beach behaviour, adding further information where available. The second section addresses the evolution of the spits enclosing Christchurch Harbour, and Hurst Spit, together with a summary of their contemporary morphological, morphodynamic and sedimentological characteristics.
5.1 Restrained Beaches: Highcliffe to Milford-on-Sea
Groyne construction commenced in the mid nineteenth century and for the next 80 years these structures were effective in retaining a stable beach. However, the more influential factor was the rapid growth of Mudeford Spit during this period (see section 4.2). This provided a sediment source and significantly reduced nearshore wave energy. At its maximum extent, the distal end of the spit was nearly co-incident with the position of Highcliffe Castle estate (Burton, 1931; Robinson, 1955). Repetitive breaching and a tendency to landward migration caused substantial beach accretion, particularly in the 1920s and early 1930s, at Avon Beach. Beach volume loss commenced in the early 1940s, and continued through to the late 1970s, after which a coordinated programme of beach nourishment and groyne construction was initiated. Nonetheless, there was a 21% loss of volume between 1985 and 1990 (Hydraulics Research 1991; Tyhurst, in Bray and Hooke, 1998) but modest losses thereafter. The submerged sand bar identified by Gao and Collins (1994-1995), which is further seaward of the breaker zone by comparison to similar examples in Poole Bay, is of uncertain morphodynamic status. It may indicate net onshore transport, but could be a partial product of earlier renourishments. Alternatively, it might represent in part the foundations of the former Mudeford Spit, which is now effectively a component of the complex ebb delta seaward of the mouth of Christchurch Harbour.
Throughout the 1990s, net beach accretion was promoted by further limited renourishment and the sequence of robust rock groynes that extend to Chewton Bunny. The underlying trend, however, is erosional. The terminal groyne promotes local wave diffraction, and thus immediate updrift beach retention. Gao and Collins (1994-1995) infer, from limited bathymetric analysis, that there is net offshore (seawards) sediment movement, and thus potential beach drawdown, in the main rock groyne sector. This must remain an uncertain statement, as it contradicts the observation of a probable onshore-directed feed (F3).
(b) Naish Beach
The mixed sand and gravel beach fronting the eroding cliffs of this sector has experienced nearly continuous volume loss, and recession of the position of mean low water, since the late 1960s. One cause, at the western end, is the angled terminal rock groyne that flanks Chewton Bunny, and which sets up wave diffraction (Mackintosh and Rainbow, 1996). Beach monitoring since 1988 indicates an average annual depletion of approximately 4000 m3 (Halcrow, 1999). In an attempt to address this problem, without recourse to cliff stabilisation, a proposal to combine renourishment with a set of "dynamic groynes" built of shingle was put forward. These would promote wave energy dissipation whilst also providing a downdrift sediment supply if and when eroded. Physical modelling, using a groyne spacing of 200 m demonstrated that these structures would dynamically evolve, although their efficiency might depend on periodic sand recycling and/or gravel recharge (HR Wallingford, 1995). A clear advantage of this innovative approach would be the maintenance of sediment input from cliff erosion, and the bridging of the "pinch point" between the adjacent defended frontages (Mackintosh and Rainbow, 1996; Bray et al, 1996). To date, this scheme has not been implemented.
(c) Central and East Barton (east to Beckton Bunny outfall)
Hooke and Riley (1987) and Halcrow (1999) have analysed serial Ordnance Survey maps and air photos, and identified the retreat of both MHW and MLW between 1867 and 1969. MLW retreat over this sector was almost twice as fast as that of MHW (approximately 0.80 ma-1), thus leading to beach narrowing and steepening. With the fixing of the position of MHW following the installation of a basal rock revetment along the central Barton shoreline, this trend has continued since the late 1960s.
(d) Beckton Bunny to Hordle
Immediately eastwards of Beckton Bunny outfall there has been a marked tendency towards beach steepening. Between Taddiford Gap and Hordle (Rock Cliff), both MHW and MLW moved seawards, at 0.34 ma-1 between 1870 and 1970 (Hooke and Riley, 1987). This has been ascribed to a persistent tendency for onshore bar migration and the accretion of a small foreland feature (Gao and Collins, 1994-5). Halcrow (1999) identify net beach width reduction (but no apparent steepening) commencing in the early 1970s, at rates of up to 2 ma-1 of LWM recession. The reasons for this switch from accretion to erosion is not clear, but is likely to be linked to some loss of onshore feed.
MLW retreat at 0.65 ma-1, and MHW recession at 0.45 ma-1 occurred between 1867 and 1969 (Hooke and Riley, 1987). This trend towards steepening has continued since the installation of the sea wall in 1970 (Webber, 1980b, Halcrow, 1999). Overall volume loss has also occurred during recent decades, partly as an outcome of the reflective and sediment yield reduction effects of the seawall; and partly because of winter storm erosion of a previously well developed nearshore sand bar in the 1980s (Halcrow 1999). Generally, beach levels have fallen approximately 2 m since renourishment in 1985 (Mackintosh and Rainbow, 1995).
All researchers are agreed that the beaches of Christchurch Bay are currently declining in volume, and have been doing so for several decades. The annual loss is between 10-20,000 m3 (Webber, 1980b; Nicholls and Webber, 1988a and b; Nicholls, 1985; Halcrow, 1999), a range that reflects inter-annual spatial and temporal variability of forcing conditions.
The relatively small tidal range concentrates wave energy into a narrow beach zone. The nature of the wave climate determines the fact that the highest rates of volume loss, and morphodynamic change, have been along the central-eastern sector. Wave energy is less along the far eastern and western sectors, with relatively less change (see sections 5.2 and 5.3).
Gao and Collins (1994-5) conclude that, although the overall trend for the past 80-100 years (at least) has been one of beach narrowing and steepening, it is not a consistent recent feature of all beaches fronting Christchurch Bay. However, this conclusion was based on the analysis of a relatively short temporal sequence of beach profiles.
The latter (a total of 44) also revealed the presence - especially in winter - of shore-parallel bars or ridges, extending into water depths of between -5 and -10 m COD (Gao and Collins, 1994-5). Most were presumed to be dominantly composed of sand (or sandy gravel in a few cases), and to indicate seasonal profile change rather than any steady net onshore transport in the nearshore zone. There is some, ambiguous, evidence that the persistence of these features has been declining over the past 20-25 years, especially along the eastern sector.
Although median grain size coarsens eastwards, as a function of wave climate and selective removal of the fines fraction, beach gradients are not precisely adjusted to this control variable.
5.2 Mudeford Spits - References Map
The entrance channel to Christchurch Harbour is confined by spits of unequal length. To the south, attached at its proximal end to Hengistbury Head is Mudeford Sandbank, with an approximate south-west to north-east orientation. To the north is the smaller, but wider, Mudeford Quay spit which projects into the harbour on an east-north-east to west-south-west orientation.
The above geographical terminology is confusing, and for the purposes of this account Mudeford Sandbank will be referred to as Mudeford Spit, whilst the opposing structure will be termed Haven House spit. Through the analysis of estate plans, hydrographic charts, topographical maps and other archival sources, the evolution of both spits over the past three hundred and fifty years has been reconstructed (Burton, 1931; Robinson, 1955; Lacey, 1985, Bray et al, 1996, Christchurch Borough Council, 1999; HR Wallingford, 1999b; Halcrow, 1999). Mudeford spit has been especially dynamic, although interpretation of the evidence of change is made more complex by the nearshore presence of the ebb tide delta of Christchurch Harbour (Gao and Collins, 1994a and b; 1995a, b and c). Full morphodynamic understanding of these accretion forms remains elusive (Halcrow, 1999 and forthcoming Coastal Defence Strategy Study, Halcrow, 2003).
5.2.1 Mudeford Spit
The basement sediments are mixed gravel and sand, overtopped by fine sand that has generated small dunes and sandhills up to 7 m in height. Earliest evidence for the presence of this spit dates to 1660, when an abortive attempt was made to construct an artificial channel through its narrowest point. This had infilled with sediment within 50 to 60 years, although Clarendon Rocks (extending some 250 m seawards) remain as a legacy of this intervention. Between 1760 and 1785, the spit was approximately 8900 m in length, with the harbour entrance channel more or less on its modern alignment. Rapid extension occurred between about 1850 and 1880, involving some 2,500 m of shore-parallel growth. The harbour exit in 1880 was opposite Highcliffe Castle estate, with the spit enclosing a 3000 m long narrow channel, The Run, between the old and new entrances to Christchurch Harbour. The reason for this dramatic change was anthropogenic: ironstone mining of Hengistbury Head between 1847 and 1856 caused a very rapid increase in cliff recession, providing a large quantity of sand for north-eastwards longshore transport. In addition, the by-passing of Hengistbury Head of sediment moving eastwards from Poole Bay was greatly accelerated. Much of this "surge" of sediment supply was used to build the extension of Mudeford spit in the 30 years following the cessation of opencast mining. It also increased in width over this period, from an average of 50 m in 1840 to 300 m in 1885, and acquired substantial crestal dunes. Although no further lateral growth occurred after 1880, Mudeford spit remained a substantial feature up to the 1930s. Nonetheless, it was temporarily breached under storm conditions several times, at a mean frequency of once every twelve years (eg 1895/6; 1911; 1924 and 1935). At the same time, it "rolled back" along segments between breach sites, to "weld" with Avon beach (Highcliffe) behind.
Formal protection started in 1931, concentrating on low points close to the distal point. Following the construction of the Long Groyne, at Hengistbury Head, in 1938 there was near instantaneous reduction in updrift longshore sediment supply. In response, the spit had reduced to a length of 900 m by the mid-1940s, and had acquired its present configuration by 1948. The breach of 1935 was converted to a new harbour entrance channel, thus cutting off the downdrift distal sector; this subsequently transgressed landwards, to form a wide beach that temporarily trapped a lagoon. The site of the spit, as it was prior to 1935, continued to be marked by a wide, submerged sandbank; this feature persists to the present time, confining the channel of The Run, and constitutes the major part of the ebb delta constructed by tidal current outflow from Christchurch Harbour.
Since the late 1940s, protection works have stabilised the basic morphology of the spit. Nonetheless, the submerged sandbank beyond its distal point expanded, in stages, up to 1973. Since then it has shown the reverse trend. This apparent cycle of accretion followed by erosion may have occurred during the phase of spit extension, with a periodicity of 15-20 years. Rapid erosion of the distal tip occurs under high energy south or south-easterly incident waves.
A progressive programme of installing concrete and rock groynes, and, more recently, sand renourishment has been carried out in stages between 1945 and 1996, together with restoration of longer-term recreation-induced erosion of the dunes (now only a small remnant of the dune field present in the late nineteenth century). Renourishment in the early 1990s, maintained by annual recycling, has maintained beach stability, which also benefited from a modest increase in sand supply following the recharges of Bournemouth beaches in 1974/75 and 1988/89. (The dynamics of sand by passing of the Long Groyne are discussed in detail in the chapter covering Poole Bay, and also in part 3 of this section). A net advance of MHW has been the trend in recent years. Recent modelling of beach behaviour (HR Wallingford, 1999b), in the context of proposals to both widen and heighten Mudeford Spit (Christchurch Borough Council, 1999) indicate a 70% probability of breaching under wave energy conditions with a 1 in 100 year return frequency.
5.2.2 Haven House Spit
The natural form of this spit has been substantially modified by protection structures, especially the seawall/promenade built following the breach threat in 1950. It has an expanded distal head, connected to Mudeford beach by a narrow (30 m) proximal "tail". Haven House was built around 1700, so the feature dates at least to his time. Evidence suggests it has a gravelly sand foundation, with overlying sand that was formerly a small set of dunes or sand hillocks.
In relation to Mudeford Spit, it is not only considerably smaller, but is offset to it - it is a type of apposition spit, with vestigial evidence of a distal recurve. Brampton et al (1998) suggest that it might have been constructed as a conventional spit by littoral drift moving towards the west, supplied by cliff erosion at Highcliffe and further east. This introduces the necessity for localised drift reversal, for which there is no contemporary or recent evidence. Brampton et al (1998) also suggest the possibility of "swash-bar welding", presumably similar to foreland construction by shore-normal waves. This might have occurred prior to any phase of extension of Mudford spit, and would be characterised by several sub-parallel gravel and/or sand ridges - now concealed, and possibly destroyed, by development. Robinson (1955) preferred to interpret Haven House spit as a truncated sector of an ancestral (pre-1600) Mudeford spit that extended across the embayed mouth of the Avon and Stour rivers, ie pre-dating the modern form of Christchurch Harbour. It was subsequently permanently breached either by a major storm surge, or by hydraulic pressure set up by exceptionally high river discharge. Indeed, both conditions might have operated simultaneously.
Robinson's model was applied also to the entrances to Poole and Pagham Harbours, where more recent research has revealed more plausible alternative explanations. However, in the case of the "twin" offset spits of Christchurch Harbour, it continues to be an attractive hypothesis. The more appropriate context is that of mid to late Holocene landward barrier migration, for which this part of Christchurch Bay might be a suitable setting in terms of hydrodynamics and sediment supply. Given that a formerly much larger Hengistbury Head salient would have been an impediment to sediment supply from Poole Bay, onshore-directed transport is a plausible alternative. Further research on the sediment stratigraphy of Christchurch Harbour, and of the spits themselves, is needed to resolve this problem. The "roll over" behaviour of segments of the extended Mudeford spit between 1880 and 1935, and its subsequent "welding" onto Avon beach, is consistent with barrier morphodynamics.
5.3 Hurst Spit - Introduction
This internationally known multi-recurved barrier spit has a main axis approximately 2 km in length, orientated at 130oN, whose proximal end is attached to the shoreline at Milford-on-Sea. At Hurst Point, its 800 m distal sector recurves very sharply to a north/north-west orientation of 100oN, and terminates with an active recurve aligned westwards. Three former, now inactive recurves are present to the west of the modern feature, increasing in size and morphological complexity from west to east. Prior to its substantial modification during the 1980s, particularly along its proximal and central sectors due to a series of major storm surges, it had a crest elevation of between 2 and 4 m above mean sea-level. This has been raised to 7 m along its proximal sector, tapering to 5 m at Hurst Point as a result of a comprehensive programme of stabilisation completed in 1996. The overall planform of Hurst Spit appears to have been stable since at least the mid eighteenth century; Hurst Castle, close to the point of distal recurvature, was built in 1544 and has apparently only been subject to potential threat from beach recession over the most recent 50 to 80 years. Nonetheless, the spit is a dynamic landform that has adjusted to the impacts of historically infrequent major storms by steadily receding landwards. Its behaviour as a barrier structure has been evaluated fully in recent years (Nicholls, 1984 and 1985; Bradbury, 1998 and 2000).
5.3.1 Geomorphological development
As Christchurch Bay was opened out during the mid Holocene due to sea-level transgression (Section 1), dominant south-westerly waves drove sediment both onshore and alongshore in a west to east direction. A substantial proportion was gravel, derived from the erosion of Pleistocene river terrace deposits originally deposited by the Solent River (West, 1980; Nicholls, 1987). An ancestral form of Hurst Spit developed following the creation of the entrance to the Western Solent circa 7000 to 6,500 years BP (Nicholls and Webber, 1987a; Nicholls, 1987; Velegrakis, et al, 1999). With continuing sea-level rise both updrift cliff recession and offshore sea bed erosion released large quantities of coarse sediment that created the Shingles Bank. This provided a large store of material which, together with a rate of littoral drift up to five to seven times what it is at present (Nicholls, 1985) created a substantial barrier spit. Low wave energy conditions to its lee promoted mudflat and saltmarsh accretion. Nicholls and Clarke (1986) have described truncated sequence of estuarine muds and peat deposits that outcrops seawards of the modern beach face, indicating that spit recession was in progress at least by approximately 4,500 years BP. This process is presumed to have been continuous (in a time-averaged sense) since then.
The progressive south-eastwards growth of Hurst Spit appears to have been episodic, or phased, as indicated by the three main "fossil" recurves. None have yet been precisely dated, but each must represent a stage of temporary equilibrium between sediment supply and loss. As now, loss of sediment at the distal end would have been due to a combination of wave action and tidal currents; waves propagating into the Western Solent, and refracted by the Shingles Bank, would have caused distal curvature and powered south to north littoral drift. Tidal currents would have increased in velocity and capacity to transport sand and fine to medium gravel, due to narrowing of the entrance channel between the Isle of Wight coastline and the Spit terminus. The dominant pathway of gravel transport then, as now, would have been towards the Shingles Bank, thereby establishing a sediment circulation that sustained growth. (King and McCullough, 1971; Nicholls and Webber, 1987a; Velegrakis, 1994). The ultimate position of the distal point - Hurst Point - may not have been achieved until the historical period and it was determined by the presence of a steep, tidally-scoured slope at the beach toe that prevented any further accretion. This stability is evident form the fact that the modern distal recurve is substantially larger than its predecessors, and that the latter increase in size with decreasing age. There may, however, be additional explanations for earlier phases of distal recurvature, such as short-term sea-level stillstands; "pulses" of gravel supply from submerged sources or differences in wave climate. Alternatively, major storms might have caused barrier breakdown over the proximal sector, introducing large quantities of sediment into both the longshore and onshore transport pathways feeding the distal end. This would imply that forcing conditions and morphodynamic response were similar to those prevailing today (Bradbury, 1998).
5.3.2 Hydrodynamic regime
Hurst Spit has three distinct sectors of alignment (Lewis, 1938; King and McCullough, 1971, Bradbury, 1998), each of which is in response to spatial variation in wave climate. This is due to complex interrelations between wave approach, nearshore and offshore bathymetry and wave energy that declines from west to east (Hydraulics Research, 1989a; HR Wallingford, 1992; 1993; Wimpey, 1994; Bradbury, 1998). Details are set out in section 1, in the context of Christchurch Bay as a whole. The presence of the Shingles Bank causes shoaling, refracting and diffracting effects, which - together with the proximity of the Isle of Wight - results in mean significant wave height being one metre lower along the sector 700 m west of Hurst Point in comparison to the proximal segment.
Tidal current velocities, both ebb and flood, are high, especially along the distal sector where tidal flow is constricted at Hurst Narrows. Here, where ebb current speeds can exceed 2.5 m s-1, significant quantities of coarse sediment are transported away from the spit and are moved offshore. Full details are given in sections 1 and 4.1.
5.3.3 Littoral Sediment Transport
As stated in Section 3, there is uncertainty concerning contemporary rates of throughput by longshore transport. Because each of the three main segments are swash-aligned, rates are less than they are on confined beaches further west. The rock armour over the 800 m sector between the neck of Hurst Spit and the detached breakwater installed in 1996 has reduced the natural rate of drift, which accelerates immediately downdrift. An average rate of drift between 11 and 15,000 m3a-1 along the central, unprotected, sector has been determined from various studies (Nicholls, 1985; Halcrow, 1999). As well as input from updrift beaches, and probably from offshore, a proportion of throughput derives from beach erosion. Using evidence of recession rates (see Section 4.3.5); and borehole data for the average thickness of gravel beneath the main axis of the spit and relict recurves, Nicholls (1985) calculated a potential supply of 1,600-3,900 m3a-1 for 1979-1982. This has diminished from 8,500-9,700 m3a-1 (1938-1968) and 1,600-5,600 m3a-1 (1969-1978) due to protection measures introduced in 1968/9 and a steady steepening of the beach face. Several episodes of severe beach erosion during the 1980s and early 1990s would have increased annual rates, which will also have been enhanced by several renourishments since 1980 (culminating in the introduction of 300,000 m3 in 1996, nearly doubling the previous volume). There has been no analysis of post-stabilisation monitoring data to determine contemporary drift rates, but the new configuration of the spit should ensure that they are relatively close to the long-term average.
Between Hurst Point and North Point, the mean longshore drift rate is lower than along the more exposed westerly sector. Most of the sediment moved is coarse sand and gravel. It has been estimated that approximately 10,000 m3a-1 is removed into Hurst Narrows due to wave focusing caused by a deep scour hole off Hurst Point (Nicholls, 1985). This reduces the downdrift transport rate, as does the reduction in wave height along this northward-orientated sector. Prevailing rates may be between 3-5,000 m3a-1 (New Forest District Council, 1992; 1996), a quantity apparently dominated by fine gravel.
5.3.4 Sediment Composition
Most information is restricted to sampling of clasts from the beach face and crest (Bradbury, 1998). The dominant constituent is sub-angular to sub-rounded flint gravel, with a mean diameter of 15 mm. Median clast size diminishes slightly from west to east in the direction of wave energy reduction (Nicholls, 1985). Contemporary sediment character has been modified by several replenishments in the 1980s using gravel from inland sources that had different size and shape characteristics than the indigenous material. The 1996 stabilisation scheme introduced some 280,000 m3 of gravel dredged from the Shingles Bank. As this is a store for sediment moving away from, and then back to, Hurst Spit, its textural indices are similar to the "natural" population. Limited evidence (Bradbury, 1998) suggests that the proportion of sand to gravel increases slightly with depth, but does not prevent infiltration and cross-barrier percolation (Nicholls, 1985). Both cross-shore and longshore grading are relatively poorly developed, the reasons for which have not been specifically investigated.
5.3.5 Morphodynamic Behaviour: Historical Trends to 1980
The main processes that have controlled the cross-profile form, and steady landward recession, over this period are berm formation, over-washing and overtopping (Bradbury and Powell, 1992; Bradbury, 1998; 2000). The last two are associated with surge conditions and other factors creating high water levels and high energy breaking waves at or above crest level. Overwashing involves swash passing over the crest and then running down the backslope towards the back barrier saltmarsh (and incised channels). Crestal breaching occurs under conditions of overwash sluicing, which creates low crestal points that facilitate further overwash. It is almost certain that there have been numerous overtopping and overwashing events over the past 3-4,000 years of the history of Hurst Spit, but the first to be fully documented occurred in 1954. Analysis of successive editions of large-scale maps and plans; air photographs and hydrographic charts reveals that these processes, including crest cut-back, have resulted in net landward recession since the mid nineteenth century. For the sector between Milford-on-Sea and Hurst Point, the rate was 1.0-1.5 ma-1, 1867-1968 (May, 1966; Halcrow, 1982; Nicholls, 1985; Hooke and Riley, 1987). It is probable that recession occurred intermittently during this period, with phases of stability (eg 1890-1910) alternating with short-term retreat of several meters under occasional high magnitude wave conditions. Spatial variations in the mean retreat rate can be deduced from the historical record, eg maximum recession of 1.8 ma-1 occurred at the "neck" of the spit between 1931 and 1965, with a minimum retreat of only 0.2 ma-1 over the central corridor during the same period (Hooke and Riley, 1987). Retreat rates everywhere appear to have accelerated (though variably) after about 1940, and Nicholls (1985) concluded that the spit as a whole moved landwards at 3.5 ma-1, 1968-1980. If no major overwashing events occurred during the period before 1950, most of the displaced material may have moved offshore. This would have resulted in progressive beach narrowing and steepening, a response that is implied from plotting the movement of the positions of mean low and mean high water (Hooke and Riley, 1987). Steady volume loss is therefore a feature of Hurst Spit over at least the past 140 years; from a programme of direct field measurements in 1980-1982, Nicholls (1985) calculated this to be 14,000 m3a-1 (but only 7-8,000 m3, 1970-1979).
Accelerating recession in the 1950s and 1960s posed particular protection problems at the far proximal end, where breaching might - potentially - result in spit detachment. Thus, in 1969, massive rock armour was placed along a 600 m frontage. Although effective as a protection measure, it reduced the rate of longshore drift and provoked erosional outflaking (Mackintosh and Rainbow, 1995).
The position of Mean High Water along the distal sector also moved landwards during the period 1870 to 1980, at a rate of 0.9 ma-1 (Halcrow, 1999). Volume loss at Hurst Beach was calculated by Nicholls (1985) to have been 1-2,000 m3a-1, 1965-1982. However, North Point, at the recurve tip, advanced some 60 m during this period; it deflected the mouth of the Keyhaven River across adjacent saltmarsh and mudflats, causing erosion. Previous to about 1920, possibly as far back as the mid seventeenth century, North Point was the site of aggregate removal. As it was more or less stable during the period 1860 to 1920, the rate of gravel extraction may have balanced potential accretion supplied by littoral transport. The 'Point of the Deep' was also stable in depth from 1870 to 1908, but experienced shallowing thereafter (Halcrow, 1999). Progressive beach retreat had converted the site of Hurst Castle into a salient by the early 1960s (Halcrow, 1982). Zig-zag timber breastwork and groynes were installed in front of the Castle in the mid/late 1960s, but were only partially effective as a conservation measure. Furthermore, they promoted downdrift sediment starvation, thus increasing the rate of recession of Mean Low Water (Hydraulics Research, 1982; Halcrow, 1982; Nicholls, 1985).
5.3.6 Morphodynamic Behaviour, 1980 to 1996
A sudden increase in the frequency and magnitude of overtopping, overwashing, breaching and breakdown events occurred in the early 1980s, preceded by one major overwash in 1962. Crest lowering, breaching and rollback occurred in 1981/2; 1984/5; 1989/90 and 1994 (Dobbie and Partners, 1984; Wright and Bradbury, 1994; Wright, in Bray and Hooke, 1998; Mackintosh and Rainbow, 1995; Bradbury, 1998).
Rollback (landward recession) of 10-25 m took place over a length of 2,300 m on 29 October 1989. During a few hours, a 1 in 100 year storm on the 16/17 December 1989 caused overwashing to lower the crest of the western sector by 2.5 m and displace it landwards between 60-80 m. Some 50,000 tonnes of gravel were moved across the crest and backslope onto the marsh and infilled Mount Lake channel via large overwash fans. A further 50,000 tonnes was moved seawards, resulting in a total volume loss 20-30 times the annual average for the previous decade. This recession exposed some 600 m2 of foreshore to erosion. Emergency reconstruction used 25,000 tonnes of gravel from the site and another 20,000 tonnes imported from external sources. However, erosional scour of the foreshore during the storm; shearing and settlement in the weakly consolidated basement sediments below the new site of the barrier and the scale of displacement meant that the stabilised spit was 12 m landwards of its previous position. These events were carefully observed and measured (Bradbury, 1998) demonstrating also the important role of landward seepage at the interface between basement materials and overlying barrier sediments.
The distal sector experienced relatively much less dramatic morphodynamic change over this period. Beach width east and west of Hurst Castle marginally increased between 1984 and 1989 (New Forest District Council, 1990; Halcrow, 1999), and there was also net accretion over some 200 m updrift of Hurst Castle. However, after about 1990, it appears that defence measures at Hurst Castle caused local erosion due to outflanking immediately downdrift, with overall advance of Mean High water further north. Net accretion between 1991 and 1995 was 12,000 m3a-1 for this sector as a whole.
A major storm was experienced on 1 April 1994, causing overwashing, elevation lowering, crest cut-back and fans extending up to 26 m across the backbarrier slope (Bradbury, 1998). This event emphasised the urgent need for a comprehensive scheme of barrier spit stabilisation. Various mathematical and physical modelling investigations into forcing conditions, and evaluation of the relative merits and probable performance of alternative protection measures, had been initiated in 1990 to this end (New Forest District Council 1990; 1992; 1996; Mackintosh and Rainbow, 1995; Posford Duvivier, 1992; HR Wallingford, 1992 and 1993; Wimpey, 1994; Bradbury, 1998).
5.3.7 The 1996 Stabilisation Scheme; Post-Scheme Morphodynamics
The main measures that were implemented (New Forest District Council, 1996; Bradbury and Kidd, 1998; Bradbury, 1998) were:
In the seven years since scheme completion, Hurst Spit has successfully resisted over 20 storms that would otherwise have caused overtopping or overwashing. However, crest cut-back has occurred, with crest cliffing also a feature before fine sediment in recharge material was winnowed out or moved down-profile (Bradbury, 1998). Between 1996 and 2000, there was a trend towards steepening of the lower foreshore slope, and net retreat of the upper foreshore (New Forest District Council, 1997-2001). Accretion behind the breakwater has occurred as predicted, though the rate of accumulation is not known. Overall, the morphodynamic behaviour of the spit has been close to model predictions, with no reduction of crest height due to profile adjustment, "winnowing" of fines and settlement. Over the distal sector, regular topographic surveys at Castle Point have revealed an erosional trend of just over 1,000 m3a-1, 1996-2000 (New Forest District Council, 2000). This has been concentrated along the southern beach face, with a small accretion "fillet" downdrift of the protection works. Since 1997 it has been moving northwards towards North Point (New Forest District Council, 2000; 1997-2001).
Research by Bradbury (1998; 2000) into the relations between gravel barrier morphodynamics and hydraulic conditions was fundamentally based on extensive field and model tests carried out on Hurst Spit. This work has demonstrated the critical importance of antecedent barrier geometry in controlling behaviour under extreme forcing conditions; it has also identified a "dimensionless predictive inertia parameter" for defining threshold conditions leading to barrier breakdown. This advance will be a considerable value in assessing the future response of Hurst Spit following modification of its natural evolution.
5.4 The Shingles Bank - References Map
This extensive shoal, described by Velegrakis (1994) as the West Solent Ebb Tidal Delta has a main axis of 6.5 km that trends approximately north-east to south west. It has two distinct components, the elongate form of The Shingles Bank proper and, in the north, North Head Shoal. There is considerable spatial variation in morphological form, especially crest height, but both components are asymmetrical in cross-section. The steepest slope of The Shingles Bank is along its eastern flank and is clearly delimited by the Needles Channel; the steeper slope of North Head Shoal is along its north-eastwards facing margin.
The dominant sediment type composing both shoals is relatively poorly-sorted gravel. Velegrakis and Collins (1993; 1994) state that vibrocore and surface sampling suggest that flint gravel accounts for a minimum 70%, with various grades and textures of sand contributing 20-25%. Silt and clay, and finely-communited biogenic (shell) debris, together constitute less than 2%. HR Wallingford (1994) propose a composition of 90% gravel, 9% sand and 1% silt derived from grab samples taken in the area up to 1 km seaward of Hurst Spit. Seismic reflectors reveal that much of the Shingles Bank rests on an eroded bedrock surface (Velegrakis, 1994), and has a variable mean thickness of 2-8 m; a maximum thickness of 12 m is achieved over a part of the north-eastern sector. This data has provided an estimation of present volume to be 40 to 60 million m3.
Analysis of hydrographic charts for the period 1880 to 1968 suggested that the Shingles Bank, as a whole, was accumulating approximately 29,000 m3a-1 (Lacey, 1985; Nicholls, 1985). Velegrakis (1994) re-examined this data, and analysed subsequent charts up to 1990. From this, he has identified considerable variation in both planform and morphology. In the last decade of the nineteenth century, the Bank had two areas with distinct crestal orientations, but North Head Shoal was barely apparent. This feature was beginning to take on a separate identity by 1921, but at this stage a single crestal orientation prevailed. Over the next 60 years there were non-periodic variations in the extent of separation of North Bank shoal and the Shingles Bank, but there was a distinct trend for (a) the eastern flank to encroach on the Needles Channel; (b) profile asymmetry to be accentuated. By 1979, the Shingles Bank was morphologically unified, with apparent infilling of the channel separating it from North Head Shoal. However, over the next 10 years, this channel became strongly re-established, although the bank complex continued to move eastwards. The analysis of detailed changes from hydrographic charts is limited by both survey accuracy and revision intervals. New Forest District Council (1992) undertook an independent analysis based on digitised hydrographic charts back to 1882 which revealed complex spatial variations of morphology that were interpreted as the product of unsteady processes of both accretion and erosion. This study revealed that the southern area of the Shingles Bank was relatively more stable, with no clear evidence there of sustained eastward movement. The earlier work of Lacey (1985) and Nicholls (1985) was confirmed, with net accretion of 3,226,000 m3 between 1882 and 1988 (30,400 m3a-1). During this 100 year period, decadal rates of accumulation varied between more than 3% to less than 1% of total volume; between 1972 and 1988, 183,000 m3 were added (0.46% of volume calculated for the preceding 25 years). Annual hydrographic surveys of the area dredged for the replenishment of Hurst Spit in 1996 (New Forest District Council, 1997-2001) revealed accretion to be a continuing trend, with almost exactly 1 million m3 added over the 5 years up to 2000. This high rate of accretion is unlikely to be sustained, as it is in response to the 300,000 m3 of gravel removed in 1996, and related morphological disturbance). Survey results also revealed complex morphological change, with the persistence of the channel separating North Head Shoal from the main body of the Shingles Bank. Overall westwards migration of North Bank shoal continued up to 2001, with a degree of fragmentation in the central area. The width and depth of North Channel fluctuated annually in precise width and depth, indicating a continuation of spatially variable erosion and accretion. The south-eastern extremity of The Shingles Bank was described by Velegrakis (1994) as building out as a "fan", with some possible transfer of sand to Dolphin Bank.
Despite evidence for fluctuations in planform, and inferred transport paths, all investigations conclude that The Shingles Bank has an overall clockwise movement of sediment (Dyer, 1970; Velegrakis and Collins, 1994; Velegrakis, 1994; Brampton et al, 1998; Halcrow, 1999).
If the historically recent rate of accretion is extrapolated retrospectively, the age of the Shingles Bank cannot exceed 2000 years. However, most researchers are agreed that it started to form no later than mid-Holocene times, following the opening of the western entrance to the Solent and the progressive elongation of ancestral forms of Hurst Spit (Nicholls, 1985; Nicholls and Webber, 1987a; Velegrakis, et al, 1999). A substantial part of its original (and contemporary) volume derived from wave reworking of Pleistocene fluvial gravels. The present-day Pot Bank gravel deposit might originally have been a constituent extension of Shingles Bank, but was detached from it by the incision of the Needles Channel (Webber, 1977). This resulted from erosional scour by strong ebb tidal currents exiting Hurst Narrows. A crude estimate of the net rate of accretion since circa 6,500 years BP (the most recent possible date for the initiation of Hurst Narrows) is 7-9000 m3a-1. Apart from supply from the terrace and flood plain deposits of the River Solent, input has come from erosion and recession of the cliffed coastline of Christchurch Bay via the beaches and Hurst Spit. Net onshore transfers and offshore transport out of Hurst Narrows represent potential outputs and inputs. In this sense, the Shingles Bank is a dynamic store; but the fact of long-term increase in volume (net accretion) also justifies its classification as a sink. There remain considerable uncertainties concerning the components of its sediment budget, which can only be clarified by more detailed research on sediment mobility and transport pathways. There would appear to be an input to the southern end of the Shingles Bank via Dolphin Bank (Velegrakis and Collins, 1993; Velegrakis, 1994) but present evidence suggests that it involves sand, rather than gravel. Further north, gravel transport towards the western flank would appear likely.
5.5 Dolphin Bank - References Map
Dolphin Bank is made up of predominantly fine to medium, well-sorted, sand. It is some 7 km in length, 1.4 km broad (at its maximum breadth) and has an elevation of up to 14 m above the adjacent, almost featureless, seabed. Its average thickness is 8 m (Brampton, et al, 1998). Its eastern extremity is separated by only a narrow corridor from the southern end of Shingles Bank; by contrast, its eastern limit is clearly separated from Dolphin Sand, to the west Seismic data suggests that it rests on an erosional surface cut transversely across Eocene bedrock and older, coarser (pre-Holocene?) sands.
Dolphin Bank has an uncertain depositional history, and may pre-date Holocene sea-level transgression. As a site of net accumulation, it must be classified as a sink. However, the evidence for considerable sediment mobility under present-day hydrodynamic conditions gives it some of the characteristics of a sediment store.
6 COASTAL DEFENCE AND HABITAT INTERFACE ISSUES - References Map
Halcrow (1999) provide a comprehensive review of the conservation and habitat attributes of the Christchurch Bay coastline. This has recently been supplemented, for Hurst Spit, by the Solent CHaMP (Posford Duvivier and the Universities of Portsmouth and Newcastle, 2003).
Much of this coastline has high national and international importance for geological (stratigraphical and palaeontological) and geomorphological features. Examples include the Eocene rock succession between Hengistbury Head and Beckton Bunny, and both Hurst and Mudeford Spits. Nonetheless, considerable loss of intrinsic interest resulted from cliff stabilisation schemes in the 1960s and 1970s. Measures to protect and enhance what remains have been put in place at Hengistbury (Turner, in Bray and Hooke, 1998b; Bray, et al, 1996), whilst innovative "soft" control structures have been proposed for Naish Cliffs (Mackintosh and Rainbow, 1995; HR Wallingford, 1995). The Hurst Spit stabilisation scheme (Bradbury, 1998; Bradbury and Kidd, 1998; Wright, in Bray and Hooke, 1998b) has maintained as much as possible of its morphodynamic character as a barrier structure.
Most of the cliffline is devoid of biological interest, either because its mobility is inimical to habitat development (eg Naish cliffs, and the cliffline east of Barton-on-Sea), or because of protection structures. Exceptions are provided by (i) the vegetation of the cliffs fronting Highcliffe, which have been deliberately planted with an "engineering sward", to enhance their stability (Tyhurst, in Bray and Hooke, 1998b), and (ii) the vegetated cliffline between Sandhills and Friar's Cliff, Mudeford. The latter is largely artificial, dominated by species that have escaped from cliff top properties, notably the Highcliffe Castle estate. There are, however, plants relict from formerly more natural conditions, including those that colonised previous dune habitats near Sandhills. There are opportunities to introduce a more varied cliff vegetation community at Barton, perhaps drawing upon experience gained at Highcliffe. Between Hordle and Milford-on-Sea, relatively stable sectors of cliffline support a patchy vegetation cover of species tolerant of dry, saline conditions. Its local ecological value should be maintained.
Dune vegetation, dominated by Ammophila arenaria (Marram grass) is locally important on Mudeford Spit, where its survival is encouraged by both present and intended management measures (Christchurch Borough Council, 1999). More attention might be given to the introduction of sub-dominant species. This was, up until the early 1940s, a much more extensive habitat, but has been greatly reduced by the loss of the former extent of the spit and by intensification of recreation pressures in an area where there is no direct control of visitor pressure.
Vegetated shingle is relatively impoverished, except along the backshore between Hordle and Rook cliffs and associated with the 'fossil' recurves of the distal part of Hurst Spit. The recession of the backbarrier slope of the western and central sectors of Hurst beach has inhibited the establishment of vegetation, but this may have some potential following stabilisation in 1996.
The maintenance of the integrity of both Mudeford and Hurst spits is, of course, crucial to the continuing survival of the rich variety of habitats in Christchurch Harbour and the North-West Solent (Keyhaven and Pennington Marshes) respectively. In the latter case, this is one of several objectives justifying the Hurst Spit stabilisation project, with an expected 50 year life. There may, however, be some loss of both intertidal and terrestrial habitats if this barrier structure continues to evolve by transgressing landwards and its far distal point recurves into Keyhaven creek.
Christchurch Harbour is the site of both high and low saltmarsh; wet meadows, dry grassland and both Phragmites and Scirpus reedbeds. Saltmarsh here is unusual for south coast estuaries in that there is a diversity of species, with Spartina anglica failing to gain invasive dominance. Present policy is to maintain the balance between different communities and habitats, an approach that is vitally dependant upon shoreline protection against potential breaching at Double Dykes and Mudeford Spit. Although there are opportunities for managed retreat to allow expansion of inter-tidal mud/sand flat and marsh communities, this would be difficult to put into effect because of loss of landward habitats (eg at Wick). Flood risk along the developed frontage of the northern shores of the harbour would - in theory - benefit from the presence of marginal marsh, perhaps created by recharge from dredge spoil. However, the planform of the Avon channel makes this a difficult option, unlikely to be implemented. Thus, future management is likely to attempt to perpetuate the present-day ecological and habitat diversity of the harbour.
7 KNOWLEDGE LIMITATIONS AND MONITORING REQUIREMENTS - References Map
The coastline of Christchurch Bay has been the subject of a number of research projects, and monitoring programmes, over the past 30 years. Further understanding will be provided by the forthcoming Christchurch Bay Coastal Defence Strategy Study (Halcrow, 2003). As a consequence, there is sound qualitative knowledge of most aspects of the coastal process regime, supported by several quantitative studies. Future research and monitoring (some of which is part of commitments by the district authorities and Environment Agency) should emphasise:
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MMIV © SCOPAC Sediment Transport Study - Christchurch Bay