Geomorphological evolution of Montserrat (West Indies): importance of flank collapse and erosional processes

Journal of the Geological Society, London, Vol. 161, 2004, pp. 147–160. Printed in Great Britain.

147

Geomorphological evolution of Montserrat (West Indies): importance of flank

collapse and erosional processes

A. LE FRIANT 1,2, C . L. HARFORD 2, C. DEPLUS 1, G. BOUDON 1, R.S .J. SPARKS 2, R.A. HERD 3,4 &

J.C. KOMOROWSKI 1

1Institut de Physique du Globe de Paris & CNRS, Case 89, 4 Place Jussieu, 75252 Paris Cedex 05, France2Department of Earth Sciences, Wills Memorial Building, University of Bristol, Queen’s Road, Bristol BS81RJ, UK

(e-mail: [email protected])3Montserrat Volcano Observatory, Fleming, Montserrat

4British Geological Survey, Keyworth, Nottingham NG12 5GG, UK

Abstract: Analysis of topography and new swath bathymetry as well as geophysical data provides information

about aerial and submarine morphological features and mass transfer processes on Montserrat. The island has

a characteristic shallow (,100 m) submarine shelf, interpreted as having been formed through erosion with a

depth controlled by glacio-eustatic sea-level variation. Several debris avalanche deposits are identified on the

lower submarine flanks of Soufriere Hills volcano, and there is evidence of lateral collapses at the older

volcanic centres. The morphological evolution of Montserrat is interpreted in terms of three stages. The first

stage comprises submarine growth. The second stage, subaerial growth, is represented by the active South

Soufriere Hills–Soufriere Hills volcanic centre. During the current eruption of Soufriere Hills volcano (1995–

2002) more than half of the lava erupted was transported into the sea. Flank collapses occurred several times

during this stage, such as the English’s Crater event (c. 4000 years ago) or the Boxing Day event during the

current eruption (26 December 1997). Montserrat’s older volcanic centres, the Centre Hills and Silver Hills,

illustrate the third stage of evolution, extinction and erosion. Magma production, long-term erosion and total

sedimentation rates on Montserrat have been estimated as 0.17 km3 ka�1, 0.0125 km3 ka�1 and 0.11 km3 ka�1

(i.e. 1.1 cm ka�1), respectively.

Keywords: Montserrat, geomorphology, submarine shelf, debris avalanches, flank collapse.

Volcanic islands and their submarine flanks result from the

interplay of volcanic, tectonic, sediment transport and erosion

processes. The long-term evolution of a volcanic island can be

understood by analysis of morphology, volcanic activity, erosion

processes and sediment transport. In Montserrat, there is an age

progression of volcanic centres, with the oldest in the north (c.

2.6 Ma) and the youngest active volcano of Soufriere Hills in the

south (Harford et al. 2002). The volcano has experienced major

prehistoric flank collapses (Deplus et al. 2001; Le Friant 2001).

English’s Crater was interpreted as the scar of a flank-collapse

event by Wadge & Isaacs (1988). Associated debris avalanche

deposits were identified in the Tar River valley and in coastal

cliffs (Boudon et al. 1996, 1998). Pyroclastic deposits at the base

of the collapse structure give a minimum age of 3950 � 70 a bp

for this event (Roobol & Smith 1998).

The current eruption has produced nearly 0.5 km3 of andesitic

magma in 8 years. This eruption began in July 1995 with

phreatic explosions following 3 years of precursory seismic

activity. Growth of an andesitic lava dome began in November

1995 inside English’s Crater. The eruption has been characterized

by lava dome extrusion, dome-collapse pyroclastic flows and

explosive activity (Robertson et al. 2000). Pyroclastic flows have

often entered the sea, with a total estimated volume of

190 3 106 m3 up to 12 July 2003. On 26 December 1997

(Boxing Day), sector collapse of the southern rim of the crater

and part of the lava dome produced a debris avalanche, with an

approximate volume of (40–50) 3 106 m3. The flank failure was

attributed to pervasive hydrothermal alteration of the south-

western flank where the fumarolic field of Galway’s Soufriere is

located, as well as to internal deformation and overload asso-

ciated with the actively growing lava dome (Voight et al. 2002).

The flank collapse was immediately followed by an energetic

pyroclastic density current that devastated 10 km2 on the south-

western flank and transported most ejecta into the sea (Sparks et

al. 2002).

This paper evaluates the morphological evolution of Montser-

rat. We present geophysical marine data around Montserrat

collected during the Aguadomar (December 1998–January 1999)

and Caraval (February 2002) cruises of R.V. L’Atalante. We

combine the interpretation of the marine data (swath bathymetry

and backscatter data, seismic reflection and 3.5 kHz profiles)

with a morphological analysis of the island to study processes of

gravitational instability and redistribution of volcanic products.

We estimate the magma production, erosion and marine deposi-

tional rates of volcanic products. We propose three stages of

evolution for andesite volcanoes in a marine setting.

Geological setting

Montserrat (168459N, 628109W) is located at the northern end of

the Lesser Antilles island arc (Fig. 1). The 800 km long arc

results from subduction of the North American plate beneath the

Caribbean plate. Arc volcanism initiated at 40 Ma (Martin-Kaye

1969; Briden et al. 1979; Bouysse et al. 1990). To the north of

Dominica, the arc is divided into two groups of islands. The

outer group is older, with thick carbonate platforms covering a

volcanic basement. The inner arc consists of volcanic rocks

younger than 20 Ma and includes all the active volcanoes

(Bouysse et al. 1990).

The subaerial part of Montserrat is 16 km long (north–south),

10 km wide (east–west) and is made up almost exclusively of

volcanic rocks. There are three volcanic massifs (Fig. 1): Silver

Hills (c. 2600 to c. 1200 ka); Centre Hills (at least c. 950 to c.

550 ka); and South Soufriere Hills–Soufriere Hills (at least c.

170 ka to present) (Harford et al. 2002). Andesitic rocks

produced by dome-forming eruptions dominate the geology.

Major deposits are remnants of lava domes, dome talus breccias,

dome-collapse pyroclastic flow deposits, lahar and debris ava-

lanche deposits, and usually thin and subordinate tephra-fall

deposits (Harford et al. 2002). Each massif consists of a core of

andesite lava domes surrounded by aprons of pyroclastic deposits

and volcanogenic sediments. Basaltic to basaltic–andesitic vol-

canism constructed the South Soufriere Hills centre around

130 ka (Harford et al. 2002).

Data collection and processing

Marine data were gathered during cruises of the R.V. L’Atalante: the

Aguadomar cruise (December 1998–January 1999) and the Caraval

cruise (March 2002). During both cruises we collected Simrad EM12D

swath bathymetry and backscatter data, and 3.5 kHz echosounder, gravity,

magnetic and six-channel seismic reflection data (Deplus et al. 2001).

During the Caraval cruise we also collected seismic profiles with higher

resolution using a 24-channel streamer and towing streamer and guns at

lower immersion (Deplus et al. 2002), and collected sediment piston

cores and dredge samples. Navigation used Starfix differential global

positioning system (GPS) during the Aguadomar cruise and GPS with no

degradation during the Caraval cruise; a positioning accuracy of a few

metres was achieved. Swath bathymetry data were processed using the

caraibes software developed by Ifremer. Digital terrain models (DTM)

were constructed with a cell size of 50 m according to the depth of the

sea floor surrounding the island (,1000 m). The multi-channel seismic

reflection data were filtered, stacked and migrated using seawater velocity

with Seismic Unix software (Cohen & Stockwell 1996).

Fig. 1. Geodynamic setting of Montserrat

(West Indies). Predicted bathymetry from

Smith & Sandwell (1997). Contour interval

is 250 m, and 2000 m isolines are in bold.

Volcanic islands are black and subaerial

coral reef platforms dark grey. The 100 m

depth submarine shelf is light grey. Inset 1:

the four major massifs of Montserrat

showing the evolution of volcanism from

north to south. Inset 2: shaded image of

bathymetry collected during the Aguadomar

cruise (illuminated from N 308E) with the

Bouillante–Montserrat graben between

Montserrat and Guadeloupe.

A. LE FRIANT ET AL .148

Morphological analysis of the island was performed on a 10 m

resolution digital elevation model (DEM) digitized from air-photo derived

topographic maps provided by G. Wadge. A 50 m grid DEM covering an

area of 71.25 km by 74.5 km centred on the island was constructed by

combining the swath bathymetry data of the Aguadomar cruise with on-

land data. The land part of the grid represents the period before the

current eruption and the swath bathymetry was that gathered in January

1999. Data for the coastlines along the Tar and White River fans were

obtained from the Montserrat Volcano Observatory (MVO) differential

GPS surveys in 1998, and for shallow water (,100 m) areas from a

single-beam survey carried out by the University of Rhode Island in July

1998 (Shufeldt et al. 2003) and data of the British Admiralty 1:50 000

Montserrat sea chart (1986 with corrections of 1996).

Observations

Regional-scale bathymetry

Montserrat lies at the western edge of a shallow arc platform

between the Limestone and Volcanic Caribbean islands (Fig. 1).

The island is located at the end of a morphological graben

(Bouillante–Montserrat) between Guadeloupe and southern Mon-

tserrat with scarps up to 100 m high (Fig. 1). This structural

trend reappears on the west side of Montserrat. The graben

indicates perpendicular extension along the length of the arc.

Southeast of Montserrat several isolated large submarine volcanic

hills rise to c. 500 m above the surrounding sea floor (Fig. 1).

Feuillet et al. (2001) suggested that the southern rim of the

graben accommodated a component of sinistral motion with

several en echelon oblique faults. Several graben are also

observed at a high angle to the arc (east–west and NE–SW

trends) east of the volcanic islands. Feuillet et al. (2001)

proposed that the Soufriere volcano of Guadeloupe and the

Soufriere Hills volcano of Montserrat lie at the western termina-

tion of the Marie-Galante and the Eperon Bertrand–Falmouth

graben, respectively.

Morphology of Montserrat

Montserrat is mountainous with four distinct volcanic massifs

(Silver Hills (c. 2600 to c. 1200 ka), Centre Hills (at least c. 950

to c. 550 ka), Soufriere Hills and South Soufriere Hills (at least

c. 170 ka to present). Erosional effects on the massifs decrease

with age from the north to the south (Fig. 2a and b). We analysed

the island using raw and shaded digital topography and slope

gradient map, paying particular attention to the directions of the

valleys related to the hydrographic system.

Erosion of old lava dome complexes has resulted in deep

valleys that dissect pyroclastic and volcanoclastic fans. Silver

Hills are deeply eroded with a highest point at 403 m. Some

debris avalanche deposits have been recognized at Little Bay

(Fig. 2b) (Harford et al. 2002). The Centre Hills volcanic

complex culminates at 741 m (Katy Hill) and is bordered by high

coastal cliffs, 140 m high to the east and 75 m high to the west.

Away from the coasts, the slopes of the Centre Hills massif are

generally less than 108. Two main topographic discontinuities are

identified on the eastern flank of Centre Hills. The first

discontinuity (Sa in Fig. 2b) is a SE-facing scarp on the NW

bank of Bottomless Ghaut that extends from the summit to the

coastline, and is clearly defined by the hydrographic system. On

either side of the scarp, the hydrographic system has developed

in divergent directions: south–north in the northwestern area and

N308 to N458 in the southeastern part. Hanging valleys on the

northwestern side indicate that their upper part has been cut by

the discontinuity. The second morphological discontinuity (Sb) is

formed by a cliff facing SE on the NW bank of Pelican Ghaut.

To the SE of the crest, valleys have a different direction (N808)

than to the NW (N208). We infer that these rims have a structural

origin that has modified the hydrographic system. Debris ava-

lanche deposits outcrop along the eastern coastline of Centre Hill

(see outcrop location in Fig. 2b), and so we suggest that these

discontinuities have a flank-collapse origin. A morphological

discontinuity on the southeastern margins of Centre Hills may be

either a fault or a flank-collapse scar (Fig. 2b).

The Soufriere Hills and South Soufriere Hills are only slightly

modified by erosion. Before the current eruption, Soufriere Hills

consisted of a central mountainous lava dome complex (Gages

Mountain, Chances Peak, Galway’s Mountain, Perches Mountain)

with associated volcanoclastic aprons, the highest point of

Montserrat being Chances Peak (921 m). Garibaldi Hill and St.

George’s Hill form two smaller, isolated massifs in the west.

They are composed of andesitic lavas, laharic breccias, pumice

fall and pyroclastic flow deposits. A pyroclastic flow deposit of

Garibaldi Hill has an 39Ar/40Ar age of 282 � 8 ka (Harford et al.

2002). The rocks of Garibaldi Hill are strongly deformed with

local dips up to 508 and many faults. St. George’s Hill and

Garibaldi Hill are interpreted as tectonically uplifted volcanic

sequences from early eruptive activity of Soufriere Hills volcano.

Roche’s Bluff on the SE coast of South Soufriere Hills volcano

is an uplifted, faulted and tilted (dips up to 408) block of

submarine volcanoclastic rocks interlayered with thin limestones.

An 39Ar/40Ar age of a lava block of 1021 � 20 ka and fossil ages

of ,1.8 Ma (Harford et al. 2002) suggest that these deposits are

related to the early development of the Centre Hills volcano.

A NW-trending volcanotectonic structure defined by these

aligned lava domes and uplifted regions can be traced through the

Soufriere Hills volcano (see Fig. 2b). The central mountainous

dome complex of Soufriere Hills is truncated by the English’s

Crater depression, which is 1 km wide, 1.6 km long and opens to

the east. The English’s Crater depression is in alignment with the

Tar River valley (0.7 km wide, 1.5 km long). Perches and Gal-

way’s domes, which are cut by the English’s Crater, are located in

a depression. Harford et al. (2002) suggested that the Mefraimie

Ghaut is bounded by a large scar facing to the north, which cuts

the South Soufriere Hills massif at the boundary between

Soufriere Hills and South Soufriere Hills volcanoes (Sc in Fig.

2b). This morphological discontinuity is also apparent from

stratigraphic profiles along the east coast (Fig. 3). Using morpho-

logical analysis of topographic discontinuities and valley orienta-

tions a third horseshoe-shaped structure has been identified, open

to the south, on the southern flank of Soufriere Hills between

Chances Peak and Galway’s Mountain (Sd in Fig. 2b). The

summit area of South Soufriere Hills consists of at least three

imbricated horseshoe-shaped structures that open to the ENE.

The southern rim of the main morphological structure is north

facing as delineated by the hydrographic system. Hanging valleys

are cut by this discontinuity. Harford et al. (2002) interpreted

these imbricated structures as flank-collapse features. Debris

avalanche deposits (.20 000 a bp) are exposed along the south-

western coastline of Soufriere Hills (Boudon et al. 2002), but

associated morphological structures have not been identified.

The current eruption has caused significant morphological

changes. The new lava dome temporarily became the highest

point of Montserrat, reaching nearly 1100 m above sea level in

May 2003 (Fig. 2a). Major valleys cutting the volcano’s flanks

have been partly or completely filled in by pyroclastic flow

deposits and lahars. Mosquito Ghaut was a major valley on the

northern flank and guided many pyroclastic flows in 1997 (Cole

et al. 2002). The ghaut was completely infilled by pyroclastic

GEOMORPHOLOGICAL EVOLUTION OF MONTSERRAT 149

flow deposits to become a slightly positive feature. The White

River valley was largely infilled by both pyroclastic flow deposits

and the debris avalanche deposit of 26 December 1997 (Voight

et al. 2002). During the last few months of 2002 Tuitt’s Ghaut

was also largely infilled. New coastal fans have been built (Cole

et al. 2002) to the east (Tar River) and SW (White River). Since

1997, rain-induced erosion on the upper flanks of unconsolidated

deposits has resulted in lahar deposition on the lower flanks,

notably around Plymouth and in the Belham River valley, as well

as in shallow marine areas off Plymouth and the White River.

Fig. 2. (a) Topographic map from the 10 m resolution DEM; contour interval is 25 m. Place names referred to in the text are marked and ages are from

Harford et al. (2002). (b) Shaded image of topography of Montserrat illuminated from N 608W and produced using the 10 m DEM of Montserrat. Some

geological and morphological features are superimposed.

Fig. 3. Simplified Soufriere Hills stratigraphic section on the east coast from Roobol & Smith (1998) and Harford et al. (2002). Location of profile A–B

is shown in Figure 2b. RBSV, Roches’s Bluff submarine volcanoclastic rocks; SH-I, SH-II, SH-III, SH-G, units of Soufriere Hills pyroclastic stratigraphy

from oldest to youngest; SSH, volcaniclastic deposits of South Soufriere Hills. The extent of the Sc structure shown by the dashed lines in Figure 2b and

discussed in the text is indicated.

A. LE FRIANT ET AL .150

Morphology of the sea bottom around Montserrat

Montserrat has a well-developed shallow submarine shelf, mostly

at depths of 20–60 m with a width from 5 km in the north to

0.5 km in the south. The shelf-to-slope transition occurs at depth

of c. 60 to 100 m (Fig. 4a). In deeper water, off the west coast,

many gullies and canyons up to 100 m deep (Fig. 4a and b) form

tributaries of a large valley, which can be traced southwestwards

to the Grenada Basin (Figs. 1 and 4a, b). No deep gullies occur

on the other submarine flanks. The sea floor off eastern

Montserrat is flat-lying (,48) and slopes gently from a depth of

c. 750 m to the north of Montserrat to 1000 m to the south (Fig.

4a). Several areas with rough topography were observed to the

north of Montserrat and on the lower submarine flanks of the

Soufriere Hills volcano (Fig. 4a). They are interpreted as the

hummocky morphology associated with debris avalanche depos-

its (Deplus et al. 2001).

Topographic profiles around Montserrat change slope abruptly

at the shelf-break, from shallow slopes on the shelf (,48) to

flanks with maximum slopes of c. 228 to c. 368 (Fig. 5). The

slope then declines away from the shelf-break, towards the more

subdued topography of the surrounding basins. Slopes are

typically ,48 at 4 km from the shelf-break. Profile shape

depends partly on the depth to the surrounding sea floor.

Topographic–bathymetric profiles of the flanks of the extinct

volcanic centres, the Silver Hills and Centre Hills, are markedly

concave upwards. Profiles of the active Soufriere Hills show a

more linear profile at shallow depths, becoming concave-up only

at depths greater than c. 400 m. These profiles have maximum

slope angles of c. 258, smaller than maximum slope angles of the

extinct volcanic centres of c. 368. Profiles of the relatively young

South Soufriere Hills (c. 130 ka) are intermediate between those

of the older volcanic centres and those of the Soufriere Hills.

Acoustic facies

Four acoustic facies (1–4) are identified according to the echo

character and the reflector geometry (Fig. 6a–e). Facies 1 (Fig.

6a) is characterized by parallel reflectors and penetration of

several tens of metres. It indicates the presence of layered

sedimentary deposits. Large areas of the flat-lying sea floor

surrounding Montserrat show this facies. Facies 2 (Fig. 6a)

displays a strong echo without penetration. It is observed on the

submarine flanks where the slope decreases and may correspond

to coarse-grained deposits (Pratson & Laine 1989). Facies 3 is

characterized by hyperbolic reflections (Fig. 6a, c and d) and has

three variants. Facies 3a is found on steep submarine slopes in

areas where numerous canyons lead into a rough topography

(Fig. 6b). Facies 3b occurs in flat-lying areas, indicating rough

and irregular sea-floor surfaces (Fig. 6a, 6c and d). Such

hyperbolic signatures on a flat-lying area have been interpreted

as submarine debris avalanche deposits (Lipman et al. 1988;

Watts & Masson 1995; Urgeles et al. 1997; Deplus et al. 2001;

Le Friant et al. 2003). Facies 3c is characterized by a thin

sedimentary layer (Fig. 6c and d). An intermediate facies 3d with

weakly defined hyperbolic reflections below subhorizontal paral-

lel reflectors is also observed (Fig. 6a). Facies 4 displays small

blurred hyperbolic reflectors and corresponds to recent pyroclas-

tic flow deposits (Fig. 6e).

Two acoustic units were identified from seismic reflection

profiles (Fig. 7a and b). One unit consists of continuous

subhorizontal reflectors corresponding to parallel-bedded sedi-

mentary layers. The other unit displays chaotic reflectors with

large energy diffraction. In previous papers (Deplus et al. 2001;

Le Friant et al. 2003), we showed that such a chaotic unit

indicates the presence of debris avalanche deposits.

Submarine deposits and gravitational instabilities

Pyroclastic flows deposits

The eruption of Soufriere Hills has produced over 500 3 106 m3

dense rock equivalent volume of magma. By May 2003 the lava

dome had a volume of about 200 3 106 m3. About 30 3 106 m3

had been erupted as tephra fall (Bonnadona et al. 2002). Nearly

280 3 106 m3, i.e. 60% of the erupted products, are pyroclastic

flow deposits. Most of the dome collapsed on 12 July 2003 to

generate pyroclastic flows that were emplaced into the sea. In the

following discussion we use volume estimates of different

products before this collapse.

Pyroclastic flow deposits are divided into those deposited on

land and submarine deposits formed by pyroclastic flows entering

directly into the sea, principally down the Tar River and the

White River (Cole et al. 2002). Before 12 July 2003 the volumes

of submarine pyroclastic flows were about 120 3 106 m3 for Tar

River and 60 3 106 m3 for the White River. These volumes were

estimated from measurements of collapse scar dimensions and

have uncertainties of about 15% (Cole et al. 2002; Sparks et al.

2002). Significant volumes of pyroclastic flow deposits have also

been reworked by lahars on land, with some of this material

being transported into the sea. We estimate that before 12 July

2003 at least 80% of the products erupted as pyroclastic flows

had been deposited into the sea (i.e. 48% of the total erupted

material).

The entrance of pyroclastic flows into the sea has created

about 1 km2 of new land principally at the mouths of the Tar

River and the White River valleys (Cole et al. 2002). These

coastal fans extend to the shelf-break. The pre-eruption bathyme-

try shows a submarine extension of the Tar River valley. The

new bathymetry of July 1998 (Shufeldt et al. 2003) and January

1999 (Aguadomar cruise) shows that these submarine valleys

have been infilled by pyroclastic flow deposits to form positive

relief on slopes less that 158 (Fig. 8). The new submarine

pyroclastic flow fan forms a widening and tapering lobate wedge

with a maximum thickness of 50 m that can be traced 5 km

offshore (Shufeldt et al. 2003). Shufeldt et al. estimated a

minimum deposit volume of 55 3 106 m3 up to July 1998,

consistent with on-land volume estimates. The deposits are

identified at greater depths on 3.5 kHz profiles by small blurred

hyperbolic reflectors of facies 4 (Fig. 6e).

Debris avalanche deposits

Seven debris avalanche deposits with characteristic hummocky

topography (named Deposits 1–7, Fig. 8) have been identified on

the lower submarine flanks around Montserrat at depths of 700–

1000 m. Only blocks larger than 50 m across can be detected

with our multibeam data. Off the Tar River valley, two main

deposits were identified inside the Bouillante–Montserrat graben

(Deposits 1 and 2 in Fig. 8a). Deposit 1 has a typical hummocky

morphology (Fig. 8a) and displays characteristic hyperbolic

reflections (facies 3b) on 3.5 kHz data (Fig. 6a). The deposit,

oriented east–west near the island, has a NW–SE direction in its

distal part. Megablocks are 100–400 m in diameter and up to

40 m high (20 m on average). Between them, smaller blocks

(,10 m high) are also present. Megablocks are found at depths

ranging from 900 m to 1000 m. They are prominent around the

1000 m isobath, where slope changes between the submarine

GEOMORPHOLOGICAL EVOLUTION OF MONTSERRAT 151

Fig. 4. (a) Contour map of Montserrat land

and sea DEM. Contour interval is 100 m

down to 500 m below sea level and 20 m

below that. Key morphological features are

labelled. Montserrat land and sea DEM has

a 50 m resolution and has been constructed

with different kind of data: (1) on land, we

used 10 m resolution DEM provided by G.

Wadge and data from MVO GPS surveys

for the coastlines along the Tar and White

River fans; (2) close to the coastlines (depth

,100 m), we used isolines and soundings

from the British Admiralty 1:50 000

Montserrat sea chart (1986 with corrections

of 1996), and data from a single-beam

survey carried out in July 1998 by the

University of Rhode Island (Shufeldt et al.

2003); (3) for the main marine area (depth

.100 m), we used swath bathymetry from

the Aguadomar cruise (N/0 L’Atalante,

December 1998–January 1999). (b) Shaded

image of topography–bathymetry of

Montserrat illuminated from N 158W. Key

morphological features are labelled. Dashed

lines indicate the location of topographic

profiles shown in Figure 5. Continuous lines

indicate the location of 3.5 kHz profiles

shown in Figure 6. Locations of core En 16

from Reid et al. (1996) and core CAR 4

from the Caraval cruise are shown.

A. LE FRIANT ET AL .152

flanks and the graben floor (Fig. 8a). Absence of sedimentary

cover on 3.5 kHz profiles (Fig. 6a) indicates a recent age for the

deposit. Deposit 1 covers 25 km2 and its maximum volume is

estimated as c. 0.5 km3 using a 20 m mean height of the blocks.

Deposit 2 has a larger extent (Fig. 8a), with a less hummocky

morphology and shows hyperbolic reflection on the 3.5 kHz

profiles interpreted as blocks (facies 3d) (Fig. 6a). Hyperbolae

are covered by parallel-bedded reflectors characteristic of fine-

grained layered sediments with an estimated thickness less than

3 m. Deposit 2 is significantly older than Deposit 1 and in

proximal areas it lies below Deposit 1. In a seismic reflection

profile, Deposit 2 appears chaotic (Fig. 7a), a characteristic of

debris avalanche deposits. In distal areas, a layered sedimentary

unit appears under the chaotic unit. The thickness of the chaotic

unit reaches a maximum of 0.05 s two-way travel time (twt) i.e.

40–50 m using a velocity ranging from 1600 to 2000 m s�1

(Urgeles et al. (1997) and Collot et al. (2001) used 1800 m s�1

for debris avalanche deposits) and thins southwards. The south-

ern limit of Deposit 2 is identified by a thinner chaotic layer and

a morphological front. Several other morphological fronts are

observed along line 52, which may correspond to different flow

units in the debris avalanche. Disturbed sedimentary reflectors

reappear at the base of the chaotic unit. Deposit 2 extends to

30 km from the coastline but its lateral extent is not well

constrained because of the lack of sufficient transverse high-

resolution seismic profiles in the distal part. Deposit 2 covers at

least 200 km2 and has probably a volume of the order of a few

km3.

Deposits 1 and 2 extend from submarine embayments in the

flanks of Soufriere Hills and South Soufriere Hills between the

coastline and the 1000 m isobath (Fig. 8a). Deposit 2 extends

from a large embayment (C2), 3 km long and 3.5 km wide,

striking roughly west–east. Its northern rim is 10–60 m high

whereas its southern rim is up to 100 m high in its distal part.

Inside C2, an elongated chute (3.5 km long and 1.5 km wide)

extends from the coastline (C1). Its northern rim is the same as

that of C2. Its southern rim is up to 20 m high. Deposit 1 extends

from this smaller chute. Pyroclastic deposits of the current

eruption have been emplaced inside these two structures.

Deposits 3 and 4 are identified on the southern submarine

flanks. For deposit 3, swath bathymetry data show prominent

hummocky morphology at depths of 800–900 m. This area

shows hyperbolic reflections on the 3.5 kHz data at depths of

700–1000 m (Fig. 8a). Megablocks are 100–300 m in diameter

and 10–40 m in height. Deposit 3 covers an area of about

45 km2 and its maximum volume is estimated to be about

0.9 km3 (using a mean thickness of 20 m deduced from the mean

height of megablocks). On 3.5 kHz records, hyperbolic reflectors

are not covered by parallel reflectors. A slight submarine

embayment (C3, Fig. 8a) was identified on the southern sub-

marine flank between the coastline and the 500 m isobath. It is

1 km long and 1.5 km wide, and is bounded by two rims 10–

20 m high. This structure is partly buried by pyroclastic deposits

of the current eruption.

Deposit 4 extends further offshore than Deposit 3 (Fig. 8a).

On 3.5 kHz profiles, hyperbolic reflectors are commonly covered

by parallel-bedded reflectors and on seismic reflection profiles

Deposit 4 is covered by sedimentary layers, indicating a greater

age than Deposit 3 (Fig. 7b). The area of Deposit 4 is difficult to

estimate because it is covered in its proximal part by Deposit 3.

Deposit 5 is located on the southwestern submarine flanks. It

covers an area of 27 km2 and abuts against a small submarine

hill (Fig. 8a). It displays hummocky morphology with mega-

blocks 10 m high. On 3.5 kHz data, hyperbolic reflectors are not

covered by parallel reflectors. Deposit 5 is linked to a prominent

scar (C4).

The submarine flanks of the islands of Redonda and Antigua

display typical morphology of debris avalanche deposits (Fig.

8b). Deposits 6 and 7 are located on the southwestern

submarine flank of Antigua and the eastern submarine flank of

the small island of Redonda, respectively. Each deposit covers

an area of about 50 km2. Megablocks are 100–400 m in

diameter and 10–40 m high. On 3.5 kHz data, hyperbolic

reflectors are not covered by parallel reflectors, which suggests

relatively young events.

Flank-collapse events

Hereafter, we use the term ‘flank collapse’ when the subaerial

part of the volcano is affected by collapse and ‘submarine slope

failure’ when collapse is below sea level.

English’s Crater event. The submarine rims of C1 are aligned

with the rims of the horseshoe-shaped English’s Crater and Tar

River valley. The debris avalanche Deposit 1 is located at the

mouth of C1, which either could represent the submarine

continuation of English’s Crater and Tar River valley or, more

likely, is the result of flow erosion. Le Friant et al. (2003)

documented similar features at Montagne Pelee, Martinique,

where subaerial horseshoe-shaped structures can be traced to

submarines chutes. Deposit 1, the submarine chute, the elongated

on-land structure and exposures of debris flow deposits on land

(Boudon et al. 1998) are inferred to be formed by the same

flank-collapse event (Fig. 9a). The volume of the submarine

Deposit 1 (c. 0.5 km3) is consistent with the estimated missing

volume in the English’s Crater and the Tar River valley (c.

0.5 km3). The minimum age of 4 ka proposed for the formation

of English’s Crater (Roobol & Smith 1998) is consistent with the

absence of a sedimentary cover over the submarine deposits.

South Soufriere Hills–Soufriere Hills events. Deposit 2 could

have resulted either from submarine flank failure, which formed

C2 (Fig. 9b) or from subaerial flank collapse (Fig. 9b) or both

together. The southern rim of Deposit 2 is aligned with a

submarine scar that extends back to the geological contact

between the South Soufriere Hills and Soufriere Hills volcano

(marked Sc in Figs. 2b and 3). Perches and Galway’s lava domes

occur within a depression abutting a north-facing scarp in the

South Soufriere Hills Massif. A putative northern rim of this

1 km

0.5 km

- 100 m

0 mSea Level

Soufrière Hills

South-Soufrière Hills

Silver Hills

Centre Hills

Fig. 5. Topographic profiles of submarine flanks from the combined land

and sea DEM and illustrated as relative distance from 100 m depth;

vertical exaggeration is 33.

GEOMORPHOLOGICAL EVOLUTION OF MONTSERRAT 153

Fig. 6. Selected 3.5 kHz echosounder profiles (location in Fig. 4b) showing the main acoustic facies described in the text. (a) Line 202 crossing eastern

debris avalanche deposits displays facies 1, 2, 3b and 3d. (b) Line 209 crossing the westward canyons system displays facies 3a with hyperbolic reflectors.

(c) Line 207 crossing southern debris avalanche deposit showing facies 3b. (d) Line 139 crossing the southern rim of the Bouillante–Montserrat graben

and debris avalanche deposit displays facies 3c. (e) Line 200 crossing submarine pyroclastic flow deposit illustrates facies 4.

A. LE FRIANT ET AL .154

depression has not been identified. However, between Mefraimie

Ghaut and Irish Ghaut, all units are younger than 130 ka (see

Fig. 3). We suggest that this depression was produced by a flank-

collapse event associated with Deposit 2. The age of this collapse

is constrained between 130 ka (age of South Soufriere Hills) and

112 ka (age of Galway’s Mountain). Deposit 2 is covered by

about 3 m of sediments (line 202 in Fig. 6a). For an age of

130 ka a calculated sedimentation rate of 3 cm ka�1 is consistent

with regional sedimentation rates of 1–3 cm ka�1 (Reid et al.

1996).

The eastward-facing horseshoe-shaped scarps of South Sou-

friere Hills volcano (Fig. 2b) pose an unresolved problem.

Roches Bluff stands east of the South Soufriere Hills and is

overlain by lavas of this complex (Harford et al. 2002). Sector

collapses of the South Soufriere Hills, however, could not have

occurred with Roches Bluff in the way. These observations can

be reconciled if Roches Bluff was uplifted contemporaneously

with the construction of South Soufriere Hills.

Other flank-collapse events. The horseshoe-shaped Sd appears

too insubstantial to account for both Deposits 3 and 4. Instead,

we propose that these deposits are related either to submarine

flank failure or to older flank-collapse structures that have been

buried by the South Soufriere Hills volcano. We collected a core

over Deposit 4 (location shown in Fig. 8). The base of the core

consists of a dark basaltic scoria deposit at least 60 cm thick

derived probably from the South Soufriere Hills volcano. Thus

Deposit 4 must be at least 130 ka or older. It could have

originated from an older flank-collapse event with the South

Soufriere Hills volcano infilling the collapse depression. The

younger Deposit 3 is here related to a submarine flank failure,

perhaps related to C3.

Fig. 7. Selected high-resolution air-gun seismic profiles gathered during

the Caraval cruise (location shown in Fig. 8). twt, two way travel time.

(a) Line 52 crossing eastern debris avalanche deposits (Deposit 1 and 2,

see Fig. 8), with interpretative section. Debris avalanche deposits are

identified by an incoherent–chaotic unit and contrast with subhorizontal

and well-bedded sedimentary layers corresponding to normal

sedimentation. (b) Line 56 crossing southern deposits (Deposits 3 and 4).

Deposit 4 is covered by a thicker sedimentary layer.

GEOMORPHOLOGICAL EVOLUTION OF MONTSERRAT 155

Fig. 8. Maps showing submarine gravity flow deposits around Montserrat. Contour interval is 100 m down to 600 m below sea level and 20 m below that.

Dark grey area shows extent of submarine pyroclastic flow deposits; mid-grey area shows extent of debris avalanche deposits with hummocky terrain

identified on the bathymetry; light grey area shows debris avalanche deposits identified with 3.5 kHz and seismic data. (a) Deposits off southern part of

Montserrat. (b) Deposits between Redonda and Antigua.

A. LE FRIANT ET AL .156

Submarine slope failures

Submarine embayment C4 is well marked along the 100 and

200 m isobaths and is considered the source area for Deposit 5

(Fig. 9c). Deposits 6 and 7 off Redonda and Antigua probably

also result from submarine slope instability. The gap in high-

resolution marine geophysical data around these islands pre-

vented the identification of associated collapse structures. Off the

southern flanks of Montserrat, isolated megablocks are present,

which are not associated with acoustic or seismic facies typical

of debris avalanche deposits. We propose that they result from

several small-scale, local submarine slope instabilities.

The shallow shelf

Montserrat is surrounded by a well-developed shallow submarine

shelf. Its width is related to the age of each of adjacent volcanic

centres: 5 km wide around Silver Hills (1200 ka); 3 km wide

around Centre Hills (550 ka); 0.5 km wide around the South

Soufriere Hills (130 ka); 0.1–0.3 km wide around active Sou-

friere Hills (Fig. 4a). The northern submarine part of the island

represents the original expanse of the Silver Hills volcano, which

was once of similar dimensions to South Soufriere Hills–

Soufriere Hills. Subaerial erosion by fluvial and mass wasting

processes as well as coastal erosion reduced the size and area of

the extinct volcanic centres and produced the flat-lying shelf

topography.

The submarine shelf extends to c. 100 m depth, with most

depths at c. 20–60 m. A similar shallow shelf is observed around

other islands of the arc (Fig. 1). The islands of the Limestone

Caribbees (Outer Arc) have well-developed shallow shelves, as

have Grenada and the Grenadines. St. Kitts, Nevis and St.

Eustatius are joined by a single shallow shelf area. Redonda is

only a few hundred metres long but has a shallow shelf

approaching the size of that of Montserrat. In the middle

segment of the arc, from St. Vincent to Dominica, the small shelf

of a few kilometres width is less prominent compared with the

large size of these islands, which are composed of several

volcanic centres (Fig. 1).

Fig. 9. Maps showing spatial relationships between submarine deposits, submarine chutes and flank-collapse structures. Contour interval is 100 m down to

100 m below sea level for (a) and (c), and down to 600 m below sea level for (b), and 20 m below that. The 100 m isolines are annotated. (a) English’s

Crater flank-collapse structure; (b) submarine slope instabilities or flank-collapse structures of the South Soufriere Hills–Soufriere Hills volcano (see text

for the discussion of the two models); (c) submarine slope instabilities off SW Montserrat.

GEOMORPHOLOGICAL EVOLUTION OF MONTSERRAT 157

We discount the possibility that the depth of the shelf of

Montserrat is controlled by local tectonic subsidence of the

island, because of both the similarity of shelf depths on the other

islands in the arc and the observation that the shelf depth does

not correlate with the age of the volcanic centre. Additionally, no

subsidence has been documented on any of the other islands.

Instead, we propose that the shelf depth is controlled by glacio-

eustatic sea-level changes. During the Quaternary, sea level has

fluctuated between c. 110–120 m below and 10–20 m above the

present sea level as a result of glacial–interglacial cycles

(Shackleton 1987). For the Montserrat volcanic centres, the

currently submarine parts of the volcano, down to 110–120 m,

have been exposed to subaerial erosion for significant periods of

time. We estimate from the sea-level curve of Haddad et al.

(1993) that since 1.2 Ma, sea level has been at 20–80 m below

current sea level for 60% of the time and deeper than 80 m for

26% of the time. Submarine shelf development can therefore be

explained by subaerial erosion during periods of sea-level low-

stand. Carracedo et al. (1999) pointed out that coastal erosion

appears to be much more rapid during sea-level lowstand.

The formation of Montserrat’s submarine shelf can be related

to two phenomena. First, during lower sea level, subaerial and

coastal erosion remove volcanic material down to c. 100 m below

sea level as explained below. This depth corresponds to the major

change in slope between the shelf and the submarine flanks of

the island (Fig. 5). Second, coral reefs build over the shelf when

sea level rises until the sea-level rise is not too rapid. The depth

of the shelf may be related to only the last sea-level rise (B.

Pelletier, pers. comm.), which would explain why the depth is

approximately the same for all submarine shelves in the Lesser

Antilles arc. High rates of sediment transport away from the

shallow shelf during periods of sea-level lowstand are consistent

with the higher deposition rates observed in adjacent basins

during glacial periods (Reid et al. 1996). The correlation be-

tween shelf width and the age of the most recent activity at a

volcanic centre reflects exposure of the younger centres to fewer

glacial–interglacial cycles.

Well-developed canyons on the western flank ofMontserrat

Prominent canyons occur in many submarine slope environments

and are characterized by highly corrugated topography and large-

scale hyperbolic 3.5 kHz reflectors (facies 4a) as observed off the

western part of Montserrat’s shallow shelf (Fig. 3). These

features are here related to collapse of steep shelf-margin

sediments in response to sediment loading. This process is a

common cause of mass wasting on continental margins; for

example, off the Mississippi River delta (Coleman & Prior

1988). Mass wasting of submarine slope sediments generates

numerous sediment gravity flows, which are responsible for

eroding the canyons and creating the corrugated topography.

Canyons developed only off the western side but not the

northern or eastern sides of Montserrat’s shallow shelf. This

observation could be either related to a tectonic control (such as

the SE–NW tectonic line (Fig. 2b)) or associated with the large-

scale structure of the island arc. A large submarine valley lies to

the west of the island, whereas to the east lies a relatively

shallow flat area formed between Antigua (Limestone Caribees)

and Montserrat (Volcanic Caribbees) (Fig. 1). This flat area is

probably more stable than the outer flanks of the double island

arc. Major canyons are also present on the western flanks of

Guadeloupe and northern Dominica, above the deep Grenada

Basin.

Discussion

Erosion volcanic production and sedimentation rates

The long-term rate of surface erosion is estimated from the

Silver Hills centre. Assuming it had a similar size to the

Soufriere Hills volcano, we estimated the minimum original

volume for the part of the Silver Hills above 100 m depth to have

been c. 17 km3. The difference of 15 km3 between this volume

and the extant volume (c. 2 km3) is attributed to long-term

erosion. Assuming that the last activity of Silver Hills was

c. 1200 ka, the long-term erosion rate is estimated at

0.0125 km3 ka�1.

The time-averaged volcanic production rate for the South

Soufriere Hills–Soufriere Hills volcanic centre has been esti-

mated using available geochronological and volumetric con-

straints (Harford et al. 2002). The age for the oldest subaerial

part of the centre is 174 ka, and the preserved subaerial volume

is c. 12 km3. Pyroclastic flows, tephra fallout and erosion of

subaerial pyroclastic deposits had transported nearly 50% of

eruption products into the sea before 12 July 2003; since then

the amount exceeds 90%. Here we use 50% of erupted products

as a conservative value for the proportion of subaerially erupted

material that is transported into the ocean during eruptions.

Using the previous assumptions and taking into account the

volumes associated with the major collapses and the role of

erosional processes, the total subaerial erupted volume V is

estimated as follows: V ¼ Vp þ Vm þ Ve þ Vc, where Vp is the

actual subaerial volume (12 km3), Vm is the erupted volume that

goes directly to the sea (50% of the original volume), Ve is the

long-term eroded material (0.0125 km3 ka�1, c. 2.2 km3 in

174 ka), and Vc is the collapsed volume (at least 1 km3). So we

obtain c. 30 km3 for the total subaerial erupted volume of the

South Soufriere Hills–Soufriere Hills complex. We thus deduce

the minimum time-averaged magma production rate of the South

Soufriere Hills–Soufriere Hills over the last c. 174 ka as c.

0.17 km3 ka�1. This is nearly 400 times lower than the time-

averaged eruption rate during the current eruption (0.5 km3 in

7 years, i.e 70 km3 ka�1), indicating that periods of magma

extrusion represent only a very small proportion of Montserrat’s

history. Stratigraphic and geochronological data (Wadge & Isaacs

1988; Roobol & Smith 1998; Harford et al. 2002) indicate a

relatively small number of major dome-forming eruptions over

the c. 174 ka history of the volcanic centre, separated by periods

of dormancy of the order of 103 –104 years. The long-term

erosion rate (0.0125 km3 ka�1) is only 7% of the geologically

time-averaged extrusion rate. These results indicate that periods

of eruption supply most of the material into the ocean at rates

much faster than long-term erosion rates.

The total volcanic production rate of the Lesser Antilles arc

over the past 100 ka is estimated at 3 km3 ka�1 based on

estimation of volumes of both subaerial and submarine deposits

(Wadge & Shepherd 1984). Therefore using our calculated time-

averaged production rate for the South Soufriere Hills–Soufriere

Hills, this volcanic centre accounts for c. 6% of the total volcanic

production for the arc.

Incomplete preservation means that the time-averaged produc-

tion rate cannot be accurately constrained for the older parts of

the island. However, a rough estimate of sediment transfer to the

adjacent marine basins can be made at the scale of the island if

it is assumed that the time-averaged eruption rate estimated for

the South Soufriere Hills–Soufriere Hills is representative of the

whole history of the island. The time-averaged rate of South

Soufriere Hills–Soufriere Hills magma production is estimated at

c. 0.17 km3 ka�1. With the assumption that a minimum of 50%

A. LE FRIANT ET AL .158

of material is emplaced directly into the sea during an eruption,

a rate of sediment deposition on the sea bottom of

0.085 km3 ka�1 is estimated. As the long-term erosion rate is

estimated to be 0.0125 km3 ka�1 for Silver Hills, a minimum of

0.1 km3 ka�1 sediment transfer rate can be deduced at the scale

of the island. Episodic emplacement of several debris avalanches

into the sea will also contribute to the large input of volcanoclas-

tic material into the sea and the generation of secondary debris

flows and turbidites. Moreover, the equilibrium profiles of the

drainage system on-land will be modified after a flank collapse,

thus increasing erosion and detrital sedimentation rate. Taking

into account all this information, a minimum rate of volcanoclas-

tic material transfer into the sea of 0.11 km3 ka�1 is derived. If

these sediments were distributed in an area of 100 km by 100 km

around Montserrat, a deposition rate of 1.1 cm ka�1 is calculated,

consistent with sedimentation rates of 1–3 cm ka�1 (Reid et al.

1996).

Applying a rate of 0.11 km3 ka�1 to South Soufriere Hills–

Soufriere Hills for the last 174 ka yields a total volume of

19 km3 of volcanic products deposited on the sea bottom during

the history of the volcano, i.e. about 65% of the total volcanic

material produced. Considering that eruptions on Montserrat

have been low-energy eruptions compared with those elsewhere

in the arc (e.g. Dominica), our inferences are in good agreement

with the work of Sigurdsson et al. (1980), who estimated that

volcanogenic sediments contributed 80% of volcanic products of

the arc over the past 100 ka.

Morphological evolution of the island

The morphological evolution of the island of Montserrat can be

interpreted in terms of three main stages (Fig. 10).

Stage 1: submarine growth. The first stage involves submarine

edifice growth, where long-term erosional processes have little

effect. Kick ’em Jenny, in the southern part of the Lesser Antilles

arc (Fig. 1), is currently at this stage (Sigurdsson & Sparks

1979). The Kick ’em Jenny cone is located inside a horseshoe-

shaped structure, open to the west towards the Grenada basin,

and is interpreted as a flank-collapse structure (Le Friant 2001;

Boudon et al. 2002).

Stage 2: subaerial growth. In the second stage, magma extrusion

builds the subaerial edifice. An active centre supplies material at

a much higher rate than the long-term erosion rate. The South

Soufriere Hills–Soufriere Hills volcanic centre on Montserrat

represents this stage of evolution. At the active Soufriere Hills

volcano, where deposition of material supplied by eruptions is

currently outpacing erosion, there is virtually no shelf. Only a

narrow zone with less steep slopes than the surrounding deeper

submarine flanks had developed since the activity at the Sou-

friere Hills at 17 ka (Harford et al. 2002) and the current

eruption. During the current eruption of the Soufriere Hills

volcano, pyroclastic flows have resulted in a high sediment

supply to the marine environment. Deposits build up large

prograding fans of volcanoclastic material in shallow water

(,100 m), where gradients had been reduced through erosion

since the last eruption (Fig. 4a). In addition, the rugged sea-floor

morphology on the western flank of the volcano, where several

‘ghauts’ (steep-sided valleys) enter the sea, suggests sediment

deposition at the mouths of the ghauts and occasional transporta-

tion to the deep Grenada basin by the system of canyons.

Several large flank collapses and submarine failures have

occurred during this stage on Soufriere Hills volcano. The major

English’s Crater event (c. 4 ka) involved a volume of 0.5 km3

and the South Soufriere Hills–Soufriere Hills event involved a

volume of several km3. On Soufriere Hills volcano, flank

collapses have occurred in different directions (to the east, to the

south). Debris avalanches and sector collapses are a major cause

of transfer of volcanic materials to the ocean during the active

stage. Flank instability can be related to hydrothermal alteration

(Boudon et al. 2002; Komorowski et al. 2002; Voight et al.

2002), tectonic controls and active volcanism.

Stage 3: extinction and erosion. In the final stage, the volcanic

centre becomes extinct. Processes resulting in edifice destruction

become dominant. Material is weathered and eroded on land and

transported into the sea, resulting in development of a shallow

shelf. When the sea level is high, the shallow shelf is covered by

a coral platform, but sediment is swept into the sea at lowstands.

The extinct centres of the Silver Hills and Centre Hills illustrate

this stage in the island’s evolution, with the older Silver Hills

having reached a more advanced stage. The Antigua debris

avalanche shows that submarine slope failure could occur during

this stage, perhaps triggered by regional earthquakes on an

unstable volcanic edifice and submarine flanks.

We thank the captains, officers and crew of the R.V. L’Atalante and the

Fig. 10. Morphological evolution of the volcanic island of Montserrat.

(a) The three key stages are shown schematically: 1, submarine growth

of the volcano; 2, subaerial growth with submarine emplacement of

volcanic products; 3, extinction stage characterized by subaerial erosion.

(b) Radial topographic profiles (vertical exaggeration 6.253) across

Soufriere Hills and Silver Hills illustrate stages 2 and 3, respectively.

Location of Soufriere Hills (SH) and Silver Hills (SvH) profiles is shown

in Figure 4b.

GEOMORPHOLOGICAL EVOLUTION OF MONTSERRAT 159

Aguadomar and Caraval scientific teams for their efficient work at sea.

The marine surveys would not have been possible without help of the

Montserrat Volcano Observatory. We thank B. Pelletier for fruitful

discussions on the origin of the shallow submarine shelf, and G. Wadge

and S. J. Day for their careful reviews. We used the GMT software

developed by P. Wessel and W. Smith for most of the figures. C.L.H. was

supported by an NERC studentship and R.S.J.S. by an NERC Professor-

ship and Royal Society Wolfson award. This study was supported by

CNRS–INSU (Oceans) and PNRN program.

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Received 23 February 2003; revised typescript accepted 12 September 2003.

Scientific editing by Mike Fowler

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