sedimentary geology 179
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SedimentaryTRANSCRIPT
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Sedimentary Geology
Editorial
Sediment gravity flows: Recent advances in process and field
analysis—introduction
A popular axiom reads: bWe know more about the
surface of the moon than we do about the ocean floorQ.Although the legitimacy of this saying is debatable, as
it compares apples and oranges (Are the bdeadQ moon
and the hydrodynamically and biologically active
ocean floor fair opponents?), there is certainly truth
in the connotation that the deep ocean is one of the
last frontiers of human exploration on Earth. In the
last three decades significant progress has been made
in our understanding of sedimentary processes and
environments in the deep sea. Progress has accelerated
in recent years, driven by social and economic
motives in conjunction with continuing scientific
curiosity. Deep-marine sedimentary rocks store the
world’s largest economic reserves of oil and gas, and
near-surface gas hydrates are a potential future source
of energy. The risk of marine geohazards, made all too
clear by the devastating tsunami in the Indian Ocean
(December 26th 2004), is another incentive for
conducting oceanic research. Typically catastrophic
sediment gravity flows, such as turbidity currents,
debris flows, slides and slumps, are a high risk factor
for man-made engineering works in deep water. These
include communication cables and drilling rigs whose
number, on the continental slope, has multiplied with
the shift of hydrocarbon exploration targets to deeper
water. The deep sea is also one of the most important
storage reservoirs of (palaeo)climate signals, e.g.
through stable isotopes and microfossils. Last but
not least, the desire to understand the fundamental
dynamics of the dispersal of sediment from the
continent via the shelf onto the continental slope
0037-0738/$ - see front matter D 2005 Elsevier B.V. All rights reserved.
doi:10.1016/j.sedgeo.2005.05.003
and into the abyss has grasped academia ever since
the pioneering work of Philip Kuenen, Gerry Mid-
dleton, Emiliano Mutti, Franco Ricci Lucchi, Roger
Walker and others.
Despite the progress made in recent years, our
knowledge of deep-marine sedimentary environments
is still sketchy compared to other environments, and it
is largely based on conceptual models for which
validation is pending. However, new exciting techni-
ques that hold promise for filling these knowledge
gaps have been developed in the various strands of
deep-marine sedimentary geology. Sea-going research
has the benefit of exploration tools with ever
increasing resolution (e.g. giant coring, multibeam
sonar and 3D seismic systems), thus starting to bridge
the gap to scales of observation typical of outcrops of
deep-marine sedimentary rocks. In turn, outcrop
studies have won ground by focussing on well-
exposed seismic-scale sedimentary sequences, such
as in the Karoo Basin (South Africa), East Greenland
and Pakistan. Digital elevation mapping is a promis-
ing new technique in outcrop work, particularly when
considered as an add-on to more classic field work. In
laboratory-scale research of sediment gravity flows
and their deposits, it is now possible to quantify
physical parameters within sediment gravity flows of
high density, using, for example, Ultrasonic Doppler
Velocity methods to measure flow velocity and Ultra
High Concentration Meters to measure concentration
of suspended particles. Before, such data were
obtainable only at the free boundaries of such flows
and in flows of very low density. Moreover, labo-
179 (2005) 1–3
Editorial2
ratory experiments have potential to act as intermedi-
ate between advanced numerical models of deep-
marine sediment transport and real processes in
nature, because, at the moment, validation of numer-
ical models is difficult due to the lack of suitable
natural datasets.
This special issue is a collection of papers
presented at the annual general meeting of the British
Sedimentologists Research Group (BSRG) in Decem-
ber 2003 at the School of Earth and Environment,
University of Leeds, United Kingdom. The papers
investigate sediment gravity flows and deposits, but
from different perspectives and with different meth-
ods. As such, they provide insides into recent
advances in deep-marine sedimentology, following
some of the trends mentioned above. The special issue
starts with three papers that approach the flow
dynamics of sediment gravity flows from an exper-
imental perspective, thus studying scale models of
these flows in laboratory flumes. Amy, Peakall et al.
studied the temporal and spatial evolution of density
currents with imposed bipartite vertical stratification
of flow density and viscosity. It was found that the
lower flow layer outruns or lags behind the upper flow
layer, depending on the density and viscosity contrast
between these layers. This may have important
implications for the depositional signature of stratified
sediment gravity flows in nature, as shown by field
examples. Choux et al. analyse the spatio-temporal
evolution of downstream velocity, turbulence inten-
sity, median grain size and suspended sediment
concentration of two laboratory-scale turbidity cur-
rents of different initial density. Links between
internal fluid and sediment dynamics are proposed,
and dimensionless variables are developed with the
aim to predict the dynamic behaviour of particulate
gravity flows across a wider range of concentrations.
The flows investigated by Choux et al. were low
density, non-cohesive and depositional. This is in
contrast to the experimental particle-laden gravity
currents studied by Felix et al., which were low- and
high-density, cohesive and non-cohesive, and non-
depositional. One of their conclusions is that vertical
flow regions of high turbulence intensity in low-
density, non-cohesive turbidity currents are coupled
through turbulent mixing, while similar regions in
high-density flows are decoupled due to turbulence
suppression. In contrast, however, Choux et al. found
that coherent flow structures generated at the lower
and upper boundaries of their turbidity currents were
discontinuous across the interjacent level of maximum
velocity. Further research is required to explain these
apparently conflicting observations. Detailed analysis
of differences in methodology and flow type (e.g.
depositional versus non-depositional flow) may well
be a proper basis for such work. None of the three
experimental papers provide information on deposits
formed by the laboratory flows. Yet, Amy, Peakall et
al. and Felix et al. included a discussion on the
implications for the geometry and internal organisa-
tion of natural deposits. At present these implications
are rather speculative. Future work should therefore
concentrate on the properties of laboratory deposits as
much-needed intermediate step between scale models
of flow dynamics and prototypes of natural deposi-
tional products.
The two papers that follow the experimental papers
present marine geological data of recent sediment
gravity flow deposits using state-of-the-art geophys-
ical instrumentation. Cronin et al. investigated the
morphology and fill history of Donegal Bay sub-
marine canyon in the Rockall Trough (Irish Sea) using
side scan sonar of different frequencies, 3.5 kHz echo
sounding, multibeam seismic profiling and gravity
coring. Changes in the fill history of the canyon are
closely related to sea level rise since the last
glaciation, with active transport of coarse-grained
sediment during low sea level and infrequent transport
of finer-grained sediment during high sea level. In a
well-illustrated paper, Cunningham et al. evaluate
along- and down-slope sedimentary processes along
the Celtic Margin outer shelf and upper slope of the
Bay of Biscay between Goban Spur and Brenot Spur,
using multi-beam bathymetry and backscatter, 3.5
kHz echo sounding, side-scan sonar and seabed
samples. They reconstruct sediment transport path-
ways from the spatial orientation of giant sandwaves,
and demonstrate that faulting has played an important
role in canyon development and in the initiation of
slope failure events as turbidity currents.
The final three papers of this special issue take the
classic, well-tested approach of outcrop sedimentol-
ogy to reconstruct flow dynamics and sediment
dispersal patterns from deep-marine sedimentary
facies. Edwards et al. focussed their attention on the
Turonian Mancos Shale in the Western Interior
Editorial 3
foreland basin of Utah, USA, where channellised
turbidity current deposits are interpreted as the
product of hyperpycnal flows. Sequence-stratigraphic
concepts are used to explain the occurrence of the
hyperpycnal flows. Wynn et al. introduce a new
technique, called digiscoping, to obtain sedimento-
logical data from inaccessible outcrops. Digiscoping
combines the high resolution of modern digital
cameras with the high magnification of field tele-
scopes to resolve features on the scale of centimetres
from distances of several hundred metres, as exem-
plified by a study of the Fontanelice Channels in the
Marnoso Arenacea Formation, Italian Apennines. By
using extensive correlations of deep-marine deposits
in the same formation, Amy, Talling et al. were able to
distinguish between thick-bedded sandstones of tur-
biditic and debris-flow origin. Principal recognition
criteria are based on kilometre-scale changes in the
geometry of individual beds and submetre-scale
structural and textural properties. Amy, Talling et
al.’s work nicely illustrates the above-mentioned trend
towards studies of large-scale, well-exposed deep-
marine sedimentary successions, yet the authors also
fully realise the value of smaller-scale observations
for process-based interpretations. This approach may
well contribute to bridging the gap between typical
field-scale and seismic-scale observations.
Acknowledgments
This special issue would not have been possible
without the help of more than 20 reviewers. On behalf
of the authors, I thank all of them for their time and
effort devoted to improving the quality of the papers.
The editorial support of Maarten Felix is greatly
appreciated. Tirza Van Daalen and Tonny Smit (both
Elsevier) provided welcome logistical advice.
Jaco H. Baas
Earth Sciences, School of Earth and Environment,
University of Leeds, Leeds LS2 9JT, United Kingdom
E-mail address: [email protected].
Tel.: +44 113 3436624; fax: +44 113 3435259.
www.elsevier.com/locate/sedgeo
Sedimentary Geology
Density- and viscosity-stratified gravity currents: Insight from
laboratory experiments and implications for submarine
flow deposits
L.A. Amy a,b,*, J. Peakall b, P.J. Talling a
aCentre for Environmental and Geophysical flows, Department of Earth Sciences, University of Bristol, Bristol, BS8 1RJ, UKbEarth and Biosphere Institute School of Earth and Environment, University of Leeds, Leeds, LS2 9JT, UK
Received 22 April 2004; accepted 6 April 2005
Abstract
Vertical stratification of particle concentration is a common if not ubiquitous feature of submarine particulate gravity flows.
To investigate the control of stratification on current behaviour, analogue stratified flows were studied using laboratory
experiments. Stratified density currents were generated by releasing two-layer glycerol solutions into a tank of water. Flows
were sustained for periods of tens of seconds and their velocity and concentration measured. In a set of experiments the strength
of the initial density and viscosity stratification was increased by progressively varying the lower-layer concentration, CL. Two
types of current were observed indicating two regimes of behaviour. Currents with a faster-moving high-concentration basal
region that outran the upper layer were produced if CLb75%. Above this critical value of CL, currents were formed with a
relatively slow, high-concentration base that lagged behind the flow front. The observed transition in behaviour is interpreted to
indicate a change from inertia- to viscosity-dominated flow with increasing concentration. The reduction in lower-layer velocity
at high concentrations is explained by enhanced drag at low Reynolds numbers. Results show that vertical stratification
produces longitudinal stratification in the currents. Furthermore, different vertical and temporal velocity and concentration
profiles characterise the observed flow types. Implications for the deposit character of particle-laden currents are discussed and
illustrated using examples from ancient turbidite systems.
D 2005 Elsevier B.V. All rights reserved.
Keywords: Sediment gravity flows; Turbidity currents; Subaqueous debris flows; Flow stratification; Analogue experiments
0037-0738/$ - see front matter D 2005 Elsevier B.V. All rights reserved.
doi:10.1016/j.sedgeo.2005.04.009
* Corresponding author. Centre for Environmental and Geophy-
sical flows, Department of Earth Sciences, University of Bristol,
Bristol, BS8 1RJ, UK. Tel.: +44 117 954 5235; fax: +44 117 925
3385.
E-mail address: [email protected] (L.A. Amy).
1. Introduction
Sediment-laden density flows with a wide range of
sediment concentrations occur in subaqueous environ-
ments (Mulder and Alexander, 2001). These density
flows include turbidity currents and debris flows.
179 (2005) 5–29
L.A. Amy et al. / Sedimentary Geology 179 (2005) 5–296
Collectively they dominate sediment flux from shal-
low to deep marine environments in many locations
(Kneller and Buckee, 2000), and form some of the
most voluminous sediment accumulations on Earth
(Bouma et al., 1985). Individual events may transport
tens, or even hundreds, of cubic kilometres of sedi-
ment (Piper et al., 1999; Wynn et al., 2002). Ancient
deposits of sedimentary density flows form many of
the world’s largest petroleum reservoirs (Weimer and
Link, 1991), whilst modern flow events are a signif-
icant hazard to seafloor structures (Barley, 1999).
Almost all particulate gravity flows contain verti-
cal gradients in suspended-sediment concentration
where particle concentration decreases upwards
away from the bed (Fig. 1). Particle stratification is
a property of laboratory currents (e.g., Middleton,
1966; Postma et al., 1988), numerical simulations
(e.g., Stacey and Bowen, 1988; Felix, 2002), and
is recorded in the few natural flows that have been
instrumented (e.g., Normark, 1989; Chikita, 1990).
This characteristic develops within flows carrying a
range of grain sizes since it takes more energy to
suspend relatively dense and large particles at a point
above the bed. Stratification may become more pro-
nounced due to variable rates of particle settling,
BED
Mixing at upperflow boundary
Particle settling
Particleentrainment
Differentialsupport ofgrains withvarying mass
BED
Fig. 1. Schematic diagram showing particle stratification that occurs
within sediment gravity flows and its principal causes.
mixing with ambient fluid and entrainment of sub-
strate material (Fisher, 1995; Peakall et al., 2000;
Gladstone et al., 2004). Since both the density and
viscosity of particle–fluid mixtures are governed by
particle concentration, flows are also vertically strati-
fied in terms of these properties.
There is strong evidence for stratification in par-
ticulate density currents, however, few studies have
investigated in a systematic manner how density and
viscosity stratification influence flow behaviour.
Gladstone et al. (2004) investigated the behaviour of
two-layer, lock-exchange, stratified density flows
using laboratory experiments. They demonstrated
that layer density and volume have a marked effect
on the current’s evolution and the resulting flow
structure. Their results are applicable to inertial,
surge-type density-stratified currents. In this study
experiments were run to investigate currents that
were viscosity-stratified and density-stratified with
relatively long durations. These experimental currents
should be more representative of natural particle-laden
currents with relatively large volumes and high parti-
cle concentration (Peakall et al., 2001). A series of
experiments is presented in which the initial density
and viscosity stratification of solute-driven currents
was systematically varied. In these experiments the
velocity and concentration structure of flows were
recorded using instrumentation. The interaction be-
tween layers was also analysed from recorded video
footage. These experiments allow the role of density
and viscosity stratification on the behaviour of the flow
to be assessed. Implications for sediment deposition
from particle-laden currents and resulting deposit char-
acter are discussed.
2. Flow stratification
This paper presents new data on the concentration
distributions of density currents. Existing data on the
stratification of sediment gravity flows has been
reviewed by Peakall et al. (2000) and Kneller and
Buckee (2000). Laboratory studies have shown that
different types of flow stratification can occur
depending on flow concentration. Relatively low
concentration (b10% by volume) and fully turbulent,
depositional currents display broadly continuous pro-
files with concentration decreasing gradually up-
L.A. Amy et al. / Sedimentary Geology 179 (2005) 5–29 7
wards (Fig. 2A–C). A gradual concentration profile
occurs under both subcritical and supercritical flow
conditions. Grain-size classes are distributed differ-
ently throughout the flow depth (Garcıa, 1994). Finer
particles tend to be more evenly distributed through-
out the flow depth compared to coarser grains. Nu-
merical models (Stacey and Bowen, 1988; Felix,
2002) and measurements of natural turbidity currents
(Normark, 1989; Chikita, 1990) also suggest contin-
A.
D.
C.
E. F.
Coarse
Fine
Hei
ght a
bove
bed
Concentrationor velocity
VelocityConcentration
B.
Fig. 2. Measured concentration and velocity profiles of laboratory
sediment gravity flows. The vertical dimension is normalised with
respect to the height of the velocity maximum and the horizontal
scale is normalised using the velocity and concentration maxima.
(A) Continuous concentration profile; strongly depositional subcrit-
ical turbidity current (Garcıa, 1994). The distributions of fine (5 Am)
and coarse (32 Am) grain size fractions are also shown. (B) Nearly
continuous concentration profile (slight inflexion above the velocity
maximum); weakly depositional subcritical turbidity current on a
low-angle slope (Altinakar et al., 1996). (C) Nearly continuous
concentration profile; low-concentration (1055 kg m�3) fluid mud-
flow (van Kessel and Kranenburg, 1996). (D) Two-layer model with
a stepped concentration profile above the velocity maximum, based
on visual observations of strongly depositional lock release turbidity
currents (Middleton, 1966, 1993). (E) Stepped concentration and
velocity profile; high-concentration (1200 kg m�3) fluid mudflow
(van Kessel and Kranenburg, 1996). (F) Multi-stepped concentra-
tion profile inferred from video and modified velocity profile mea-
sured using trajectories of moving particles; high-concentration
turbidity current with starting concentration of 35–40% volume
fraction (Postma et al., 1988).
uous profiles for relatively low-concentration cur-
rents that are weakly depositional.
Few measurements of the concentration profiles
of high-concentration, particulate laboratory currents
exist (N20% by volume). Based on visual observa-
tions of strongly depositional currents, Middleton
(1966, 1993) proposed a two-layer model. This
model suggests a stepped profile with a high-con-
centration lower layer overlain by a more dilute and
relatively turbulent upper layer, also observed in
other experiments (e.g., Hampton, 1972; Mohrig et
al., 1998; Hallworth and Huppert, 1988; Marr et al.,
2001). A study by van Kessel and Kranenburg
(1996) measured the concentration profiles of fluid
mudflows and demonstrated a stepped profile for
those with relatively high concentrations. More
importantly, in a set of experiments, they were able
to document a change from a gradual (Fig. 2C) to a
stepped profile (Fig. 2E) with increasing mud con-
centration and were able to show that this occurred
in conjunction with a transition from turbulent to
laminar flow. In another experimental study on a
high-concentration flow, using cohesionless sedi-
ment, a laminar-moving high-concentration basal
layer was observed to form (Postma et al., 1988).
This took the form of a wedge propagating behind
the head of the current and below an overriding
turbulent, but strongly stratified, upper region. The
inferred velocity and concentration profile in these
currents is slightly different to those of high-concen-
tration fluid mudflows (Fig. 2F). The velocity profile
has an doverhanging noseT with a discrete reduction
in values below the velocity maximum, whilst the
concentration profile has two steps one below and
one above the velocity maximum. However, these
measurements were estimated somewhat crudely
compared to the other reported studies, because
flow velocity was derived from the motion of parti-
cles adjacent to the tank wall captured by film, and
thus do not exclude wall affects, whilst concentration
was estimated visually.
3. Experimental method
The experiments were run in a glass-walled, grav-
ity-current tank located in the School of Earth and
Environment, University of Leeds. The tank was 6 m
L.A. Amy et al. / Sedimentary Geology 179 (2005) 5–298
in length, 0.5 m in width and was filled with water
to a depth of ~1.5 m (Fig. 3). Stratified flows were
generated using two aqueous solutions with different
concentrations of glycerol. Each solution was first
mixed in an external reservoir tank and circulated
from this tank into a header box and back again via
an overflow system. Solutions were dyed different
colours and seeded with small amounts (b1% by
volume) of millimetre-sized neutrally buoyant parti-
cles to aid flow visualisation. A small amount of
silica flour (grain size b50 Am diameter) was also
added to help flow velocity measurement (see
below). Experiments were started by opening a
valve on the header boxes and allowing the solutions
to gravity drain into the main tank. The discharge
rate from each header box was kept constant by
maintaining a constant fluid head in the header
box. It was not possible, however, to keep the dis-
charge rate the same for solutions of different gly-
cerol concentrations. The discharge rate varied for
fluids of different glycerol concentrations (i.e., be-
tween different layers and experiments) from 2–4 l/s
owing to their different fluid densities and viscosi-
ties. The solutions passed through an inlet box which
was laid flat on the tank floor. The inlet box parti-
tioned the two glycerol solutions into a vertically
stratified release, with a lower relatively dense
layer and an upper less-dense layer. It also dampened
the initial turbulence by passing the fluid through a
Fig. 3. The experimental set-up used. The apparatus consists of a gravity cu
shown for clarity and is not drawn to scale.
section filled with polystyrene chips. Currents flowed
down a 3.5-m-long, smooth, floor inclined at 38.Fluid at the end of the tank was collected in a
sump and pumped out of the tank to minimise the
effects of flow reflection and changes in the ambient
fluid depth.
3.1. Solute properties
In these experiments solutions were used instead of
particle–water slurries. This approach was chosen
since it was technically difficult to generate high-
concentration currents with long durations and with
controlled initial stratifications using sediment. Solu-
tions of aqueous glycerol were chosen since they have
similarities in their density and viscosity with sedi-
ment–water mixtures. Glycerol (C3H8O3) has a den-
sity of 1260 kg m�3 and viscosity of ~1.5 kg m�1 s�1
at 20 8C (CRC Handbook of Chemistry and Physics).
For aqueous glycerol solutions viscosity increases by
several orders of magnitude with increasing glycerol
concentrations from ~10�3 to 1.5 kg m�1 s�1 (Fig.
4). The viscosity of sediment–water mixtures also
increases strongly by several orders of magnitude at
particle concentrations greater than 40–50% for mix-
tures composed of cohesionless particles (Richardson
and Zaki, 1954; Kreiger and Dougherty, 1959), and at
relatively low particle concentrations for those con-
taining cohesive particles (Major and Pierson, 1992;
rrent tank and two external reservoir tanks. Only one reservoir tank is
0 20 40 60 80 100
Glycerol concentration, % by weight
0.001
0.01
0.1
1
10
Vis
cosi
ty,k
gm
-1s-
1
0 10 20 30 40 50 60
Particle concentration, % by volume
Fluid mixtureAqueous glycerol solution (CRC Handbook)Kaolinite / china clay (Dewit, 1992)Cohesionless particles (Krieger and Dougherty, 1959)Silicon carbide (Ferreira and Diz, 1999)
Fig. 4. The relationship between concentration and viscosity for
glycerol solutions and several types of particle–water mixtures. The
data shown for aqueous glycerol mixtures and china clay bearing
mixtures are based on empirical data from the CRC Handbook of
Chemistry and Physics and from De Wit (1992), respectively. The
relationship for cohesionless mixtures is taken from the theoretical
model proposed by Kreiger and Dougherty (1959) for hard spheres
suspended in water. The relationship for silicon carbide mixtures is
a modified Kreiger and Dougherty model fitted to experimental data
of slurries containing particles with a mean particle size of 13 Amand measured at shear rates of ~100 s�1 (Ferreira and Diz, 1999).
L.A. Amy et al. / Sedimentary Geology 179 (2005) 5–29 9
Coussot, 1997). Density increases linearly with con-
centration for both aqueous glycerol solutions and
particle–water mixtures.
Although aqueous glycerol solutions are excellent
analogues for sediment–water mixtures they do not
reproduce all aspects that are important to flow be-
haviour. Aqueous glycerol solutions do not reproduce
the influence of non-Newtonian rheology on flow
behaviour. In particular, sedimentQwater mixtures
with relatively high concentrations of non-cohesive
particles or significant mud content may possess a
yield strength and display shear-thickening or shear-
thinning behaviour (Barnes, 1989; Coussot, 1997;
Major and Pierson, 1992). Of course, the settling of
sediment particles and erosion of the substrate is not
accounted in experiments using solutions.
3.2. Flow measurements
The flow velocity and concentration were mea-
sured at different heights within the flow at a position
2.5 m downstream of the inlet, and the current was
filmed at this location. Flow velocity and concentra-
tion measuring apparatus were held within a machined
holder that ensured individual probes were set parallel
both to the bed and to the tank walls.
3.2.1. Flow velocity
The downstream component of flow velocity was
measured using ultrasonic Doppler velocity profiling
(UDVP). This method derives velocity using the
Doppler shift in ultrasound frequency recorded from
small particles passing through the measurement vol-
ume (Takeda, 1991; Best et al., 2001). The velocity of
a particle is given by:
U ¼ cfD=2f0; ð1Þ
where c is the speed of sound in the fluid being
investigated, fD is the Doppler shift and f0 is the
ultrasound frequency. Ultrasonic probes simultaneous-
ly measure the velocity in 128 measurement volumes
positioned along the length of the ultrasound beam.
Flow velocity was recorded upstream of the probes,
thus instrumentation placed in the flow did not affect
measurements. A vertical array of eight probes was
used to measure the downstream velocity at heights of
0.1, 1.9, 2.8, 3.7, 5.7, 7.1, 8.5 and 11.4 cm above the
bed. The maximum temporal resolution of the velocity
data was 5.8 Hz; other parameters are listed in Table 1.
Velocity data was post-processed to account for the
effect of flow concentration on velocity measure-
ments; measurements were affected by variations in
glycerol concentration since the speed of ultrasound is
a function of concentration. In order to correct data the
sound velocity of glycerol solutions was empirically
derived using an ultrasonic thickness gauge. The
depth was initially measured in water and then again
in a solution. The velocity of sound in the solution,
csol, was calculated from
csol ¼xsol
xwat
�cwat;
�ð2Þ
Table 1
Starting parameters of the ultrasonic Doppler velocity profilers
Number of probes 8
Height above floor, cm 0.1, 1.9, 2.8, 3.7, 5.7, 7.1,
8.5, 11.4
Ultrasound frequency, MHz 2
Transducer and probe
diameter, mm
Two probe sizes used: 5, 8
and 10, 13
Measurement window, mm 12.4–107.1
Measurement bin length, mm 0.74
Velocity resolution, mm s�1 2.3–3.4
Height of nearest measurement
bin, mm
6.1 and 11.1
Height of furthest measurement
bin, mm
14.35 and 19.4
Ultrasound velocity, m s�1 1480
Sampling frequency/probe, Hz 5.8
The term bbinQ refers to a volume in which velocity measurements
are recorded.
L.A. Amy et al. / Sedimentary Geology 179 (2005) 5–2910
where cwat is the sound of velocity in water and xwatand xsol is the depth measured in water and the solu-
tion, respectively. The velocity of sound of solutions
was found to be proportion to glycerol concentration,
C, so that csol=0.0034C +0.99. The results show that
changes in flow concentration had a relatively small
influence (b5%) on flow velocity measurements.
3.2.2. Flow concentration
Flow concentration was measured using an array of
five vertically-stacked siphons (internal diameter of
0.6 cm) to extract fluid samples at 0.5, 1.7, 3.0, 5.0
and 8.0 cm above the bed. Fluid samples were collect-
ed in beakers positioned on a table on a movable track.
During experiments the table was moved to collect
samples at 5-s intervals. The concentration of glycerol
of each sample was determined by measuring their
refractive index using a temperature controlled refrac-
tometer. The refractive index, RI, of glycerol solutions
is proportional to glycerol concentration, C, and at a
temperature of 20 8C, RI=0.0014C +1.3304.
The temporal record of concentration may become
distorted if the siphon flow rate changes during the
experiment, for example due to fluctuation in flow
velocity, density, and viscosity. This problem occurred
in experiments using relatively high concentrations,
z80% weight glycerol. In these experiments the dis-
charge per unit time from siphons measuring close to
the bed varied by up to five times. Considering changes
in the discharge rate, a cumulative offset of several tens
of seconds or more over the duration of the time of the
flow can be crudely estimated. However, since in these
flows an abrupt change in concentration could be clear-
ly identified based on colouration from video record-
ings, a temporal correction was made to calibrate the
measured concentration of the affected siphon. Smaller
fluctuations of siphon discharge were observed in other
experiments and at measuring positions higher above
the bed; however, these were not corrected for.
3.2.3. Reproducibility
The reproducibility of flow velocity and concen-
tration measurements was tested by repeating one
experiment using the same starting conditions (Fig.
5). The results indicate standard deviations of 10–25
mm s�1 for velocity measurements. These values are
calculated by taking a temporal mean over the time
period of quasi-steady flow, 0–25 s after the arrival of
the flow front. Differences in velocity are relatively
large in the probes positioned at 5.7 and 7.1 cm above
the bed and at times during the waning flow phase
later than 25 s. Trends recorded by the lowest three
UDVP probes are broadly similar for the first 20 s.
Concentration measurements on average have stan-
dard deviations of b3% weight glycerol. A somewhat
larger deviation is recorded by the lowest siphon
during the initial 10 s (Fig. 5A).
3.3. Experimental runs
A set of experiments were run with the initial start-
ing conditions shown in Table 2. The initial basal-layer
concentration (CL) was increased from ~20% to 90%
glycerol whilst the initial upper-layer concentration
(CU) was kept at a low value between 6.5% and
12.3%. Values of glycerol concentration can be com-
pared to the volume fraction, t, of sediment–water
mixtures based on a comparison of values of kinematic
viscosity. On this basis solutions of 10% glycerol are
equivalent to volume fraction of cohesionless particles
of ~0.2t and 20% and 90% glycerol is equivalent to
volume fractions of ~0.3t and ~0.6t, respectively.
3.3.1. Starting stratification
The strength of the stratification of two-layer flows
may be assessed using dimensionless ratios whose
value if small indicate a strong stratification whilst a
value of unity indicates a homogeneous mixture. The
0
20
40
60
C,a
gs%
0.5 1.7 3 5 8
0
100
200
300
u,m
m/s
Run 5 Run 5a
0
100
200
300
u,m
m/s
-5 0 5 10 15 20 25 30 35 40
t, s
-50
50
150
250
u,m
m/s
0
100
200
300
u,m
m/s
0
100
200
300
u,m
m/s
0
100
200
300
u,m
m/s
z = 1 cmstd = 23.0
z = 1.9 cmstd = 11.3
z = 2.8 cmstd = 13.2
z = 3.7 cmstd = 16.6
z = 5.7 cmstd = 17.7
z = 7.1 cmstd = 25.0
z, cm =
A.
B.
Fig. 5. Concentration (A) and velocity (B) data for multiple runs using the same starting conditions (experiments 5 and 5a). The lower and upper
layer had initial glycerol concentrations of ~60% and ~10%, respectively (Table 2). Data show that the reproducibility of data is within F3%
weight glycerol and velocity data is within F30 mm s�1. Standard deviations were calculated for individual measurement probes using the
average standard deviation of measurements taken between 0 and 25 s, during steady input of fluid into the tank. The abbreviations bagsQ andbstdQ are aqueous glycerol solution and standard deviation in mm s�1, respectively.
L.A. Amy et al. / Sedimentary Geology 179 (2005) 5–29 11
dimensionless density ratio between layers was de-
fined by Gladstone et al. (2004):
q4 ¼ qU � qa
qL � qa
¼ gUV
gLV; ð3Þ
where qL and qU are the densities of the lower and
upper layers, respectively, qa is the density of the
ambient fluid and the reduced gravity gV=g(qf� qa /
qa) where g is the acceleration due to gravity, qf is the
density of the flow and qa the density of the ambient
fluid. We introduce a dimensionless viscosity ratio
defined as
l4 ¼ lU � la
lL � la
; ð4Þ
where lL and lU are the viscosities of the lower and
upper layers, respectively, and la is the viscosity of the
ambient fluid. Gladstone et al. (2004) also defined a
Table 2
Starting parameters of the experiments
Experiment CL CU qL qU q* lL lU l* QL QU B* Ta TL TU
1 18.3 9.8 1033.6 1010.9 0.50 0.0020 0.0014 0.446 0.0057 0.005 0.30 11.2 14.4 14.0
2 28.3 10.2 1060.2 1011.9 0.33 0.0026 0.0015 0.331 0.0057 0.005 0.22 10.4 16.6 14.2
3 37.2 6.9 1084.0 1003.1 0.16 0.0038 0.0013 0.116 0.0054 0.005 0.12 11.2 17.6 14.7
4 44.7 8.5 1104.0 1007.4 0.17 0.0054 0.0014 0.096 0.0047 0.005 0.15 10.1 19.3 14.9
5 58.5 9.8 1140.8 1010.9 0.15 0.0120 0.0014 0.038 0.0053 0.005 0.12 10.9 17.6 16.2
5a 59.5 9.4 1143.5 1009.8 0.14 0.0130 0.0014 0.033 0.0054 0.005 0.11 10.4 20.5 14.2
6 67.9 8.5 1165.9 1007.4 0.11 0.0207 0.0014 0.021 0.0054 0.005 0.09 11.6 20.2 15.9
7 79.4 12.3 1196.5 1017.6 0.14 0.0624 0.0015 0.009 0.0030 0.005 0.19 10.5 21.5 15.7
8 87.3 6.5 1217.6 1002.1 0.06 0.1402 0.0013 0.002 0.0020 0.005 0.13 10.7 19.6 15.5
9 88.3 7.7 1220.3 1005.3 0.07 0.1350 0.0014 0.003 0.002 0.005 0.16 11.1 21.1 15.1
The variables are glycerol concentration, C, in percentage by weight; density, q, in kg m�3; dimensionless density ratio, q*; viscosity, l, in kg
m�1 s�1; dimensionless viscosity ratio, l*; discharge, Q, in m3; and dimensionless buoyancy ratio, B*. Subscripts a, L and U indicate values
for the ambient fluid and lower and upper layers of the current, respectively. Viscosity values are corrected for temperature using the empirical
relationship found by Chen and Pearlstein (1987).
L.A. Amy et al. / Sedimentary Geology 179 (2005) 5–2912
ratio for the difference in the driving buoyancy of each
layer, B*, proportional to the layer density and volume.
For continuous input flows B* can be defined as
B4 ¼ QUgUV
ðQUgUVþ QLgLVÞ; ð5Þ
where Q is the discharge per unit width. Values of
0bB*b0.5 indicate a greater driving buoyancy in the
lower layer, whilst those of 1NB*N0.5 indicate a
greater driving buoyancy in the upper layer. The start-
ing conditions were chosen to explore a distinct area of
the bparameter spaceQ of stratified flows where q*, l*and B* are b0.5 (Table 2). The results from these
experiments thus document currents with density
ratios of 0.06bq*b0.50 and viscosity ratios of
0.002bl*b0.48 and those with a greater driving buoy-ancy in their lower layer, 0bB*b0.3. In order to keep
input conditions similar, the upper layer was released
slightly, up to 5 s, before the lower layer in experiments
1–8. In two experiments, 9 and 10, the lower layer was
released first in order to see how this affected flow
behaviour (Table 2).
4. Experimental results
4.1. Visual observations
Gravity currents with a characteristic head and
body structure were formed after releasing the glyc-
erol solutions. The two constituent layers were ob-
served to have variable downstream velocities. Thus,
two distinct types of current developed with either a
faster lower layer or a faster upper layer. In each case
the faster layer ran ahead to form the flow front. In
some experiments, overtaking of one layer by the
other was observed. This occurred within 1 m of the
input point and upstream of where flow measurements
were recorded. The basal layer, when faster, overtook
by pushing lighter fluid of the slower layer upwards
and out of the way (Fig. 6A). The upper layer, when
faster, overtook by propagating over the relatively
slow-moving lower layer (Fig. 6D). The upper layer
became progressively thinner and increasingly
stretched-out as it moved. On approaching the flow
front it intruded into the back of the head along a
density interface between the denser fluid of the
lower layer and the lighter fluid of the wake. In other
experiments overtaking was not observed since the
faster layer was released first (Fig. 6B and C). After
the flow front had reached the end of the tank floor, a
longitudinally uniform current developed, whose two
layer stratification was preserved along the length of
the tank.
4.1.1. Flows with a fast lower layer
A current with a relatively fast basal layer was
formed in experiments run with lower-layer concentra-
tions less than 75% glycerol (experiments 1–6). In
these flows fluid from the lower layer formed the
head of the current (Fig. 7A). Video recordings show
A. Upper layer released first but overtakenby faster lower layer. Experiments 1-6.
B. Upper layer released first and remainingahead of slower lower layer. Experiment 7 & 8.
C. Lower layer released first and remainingahead of slower upper layer. Experiment 10.
D. Lower layer released first but overtakenby faster upper layer. Experiment 9.
Basal fluid Top fluid Mixed fluid
Fig. 6. Schematic diagram showing the evolution of two-layer, stratified gravity currents based on experimental observations. Four types of
evolution were observed depending on which layer had a faster velocity and which was released first.
L.A. Amy et al. / Sedimentary Geology 179 (2005) 5–29 13
that only a small amount of fluid from the upper layer
was able to intrude into the head, instead most of this
fluid was swept back into the wake before reaching the
flow front. Behind the head, the lower layer had an
average thickness of 2–3 cm. The upper layer was
thicker, on average being between 6 and 8 cm thick.
The interface varied in character between experiments
becoming progressively distinct and sharply defined
with increasing lower-layer concentration. In all
experiments interfacial waves developed and thus the
height of the interface varied temporally. In experi-
ments 6 and 7 wave heights were of a similar scale to
the thickness of the lower layer. Tracer particles within
all flows moved in a turbulent fashion, changing both
height above the bed and speed. Mixing was clearly
visible between layers. In relatively high-concentra-
tion flows (e.g., experiments 5 and 6) mixing occurred
in periodic bursts whereby fluid from the denser lower
layer was injected upwards into the layer above.
4.1.2. Flows with a fast upper layer
The lower layer was relatively slow compared to
the upper layer in experiments run with lower-layer
concentrations greater than 75% glycerol (experi-
ments 7 and 8). In these flows the lower layer formed
a slow-moving region that lagged behind the current’s
head. The current’s head was formed by fluid fed from
the upper layer (Fig. 7B). In experiments 7 and 8 it
took the lower layer 8 and 11 s to arrive at the
measurement station after the passage of the head,
respectively. This slow-moving region had a wedge-
shaped front inclined downstream (Fig. 7B). With
time, the lower layer achieved a constant thickness
of several centimetres. In experiment 7 the interface
was noticeably wavy whilst in experiment 8, with the
highest glycerol concentration (CL=90%), the inter-
face between the two layers was flat. Many of the
tracer particles were observed to become concentrated
at the interfacial boundary. For experiment 8, those
Fig. 7. Successive photos taken at a position of 2.5 m downstream of the inlet point. (A) A current with a relatively fast lower layer (experiment 3). Photographs were taken at 2-s
intervals. (B) A current with a relatively fast upper layer and a relatively slow, laminar-moving lower layer that lags behind the flow front (experiment 8). Photographs were taken
every 4 s.
L.A.Amyet
al./Sedimentary
Geology179(2005)5–29
14
L.A. Amy et al. / Sedimentary Geology 179 (2005) 5–29 15
tracer particles on the boundary and within the lower
layer were observed to move in a laminar fashion
whereby they maintained a constant speed and
moved in a straight line at a constant height above
the bed. Mixing between the two layers, especially in
experiment 8, appeared to be suppressed. However,
fluid of a colour indicative of mixing and moving
relatively fast was observed preceding the arrival of
the lower layer.
-10 0 10 20
t, s
Run 8, lower layer ~ 90
02468
10
z, c
m
-10 0 10 20
Run 3, lower layer ~ 40 %
02468
10
z, c
m
-10 0 10 20
Run 7, lower layer ~ 80 %
02468
10
z, c
m
-10 0 10 20
Run 5, lower layer ~ 60 %
02468
10
z, c
m
-10 0 10 20
Run 1, lower layer ~ 20 %
02468
10
z, c
m
-10 0 10 20
Run 6, lower layer ~ 70 %
02468
10
z, c
m
A.
Fig. 8. (A) Maps of flow velocity constructed from temporal measurement
downstream of the inlet point for selected experiments. Data shown is the
flow concentration constructed from temporal measurements taken at five h
inlet point for selected experiments. The time, t, is measured in seconds afte
conditions for limited periods of time some 5 to 20 s after the arrival of t
concentration lower layer. Experiments 7 and 8 had a relatively slow-mov
4.2. Velocity and concentration profiles
Maps of flow velocity and concentration through
time and for different heights above the bed are shown
in Fig. 8. These plots show that currents achieved
quasi-steady conditions for both flow velocity and
concentration for a period of 10 s or more. The initial
recorded flow velocity and concentration, however,
are unsteady and typically waxing for the first 5 to
30 40 50 60
% glycerol
0
40
80
120
160
200
240
280
30 40 50 60
glycerol
0
40
80
120
160
200
240
280
30 40 50 60
glycerol
0
40
80
120
160
200
240
280
30 40 50 60
glycerol
0
40
80
120
160
200
240
280
30 40 50 60
glycerol
0
40
80
120
160
200
240
280
mm/s
30 40 50 60
glycerol
0
40
80
120
160
200
240
280
s taken at eight heights (z) above the bed and at a position of 2.5 m
mean velocity of 60 bins and three successive cycles. (B) Maps of
eights (z) above the bed and at a position of 2.5 m downstream of the
r the arrival of the flow front. Measurements show quasi-steady flow
he flow front. Experiments 1, 3, 5 and 6 had a relatively fast, high-
ing, high-concentration, lower layer.
-10 0 10 20 30 40 50 60
Run 1, lower layer ~ 20 % glycerol
02468
10
z, c
m
0.0
4.0
8.0
12.0
16.0
20.0
- 10 0 10 20 30 40 50 60
Run 5, lower layer ~ 60 % glycerol
02468
10
z, c
m
0.0
12.0
24.0
36.0
48.0
60.0
-10 0 10 20 30 40 50 60
Run 7, lower layer ~ 80 % glycerol
02468
10
z, c
m
0.0
16.0
32.0
48.0
64.0
80.0
-10 0 10 20 30 40 50 60
Run 3, lower layer ~ 40 % glycerol
02468
10
z, c
m
0.0
8.0
16.0
24.0
32.0
40.0
-10 0 10 20 30 40 50 60
t, s
Run 8, lower layer ~ 90 % glycerol
02468
10
z, c
m
-10.0
10.0
30.0
50.0
70.0
90.0
Conc. % ags
-10 0 10 20 30 40 50 60
Run 6, lower layer ~ 70 % glycerol
02468
10
z, c
m
0.0
14.0
28.0
42.0
56.0
70.0
B.
Fig. 8 (continued).
L.A. Amy et al. / Sedimentary Geology 179 (2005) 5–2916
15 s (Fig. 8). This is most apparent in the lower part of
the flow immediately above the bed. Flow unsteadi-
ness at these times is related to the passage of the
current’s head and subsequent large-scale eddies. Cur-
rents began to wane after about 20 s marking the time
at which the input supply was turned off (Fig. 8).
The waning of flows becomes noticeably stronger
with increasing lower-layer concentration.
4.2.1. Flows with a fast lower layer
Experimental flows 1–6 with fast moving bases are
characterised by temporal concentration and velocity
profiles that mirror one another in that maximum and
minimum values occur at similar times (Fig. 8). Near-
bed velocities and concentrations tend to increase with
time to a quasi-steady state before waning (Fig. 9).
Maximum measured velocities of ~300 mm s�1 oc-
curred in experiments 6 and 7 with lower-layer gly-
cerol concentrations of 60% and 70%, respectively.
Vertical profiles in the body (Fig. 10) are similar to
those described for other low-concentration currents
(Fig. 2A–B); concentration displays continuous or
nearly continuous profiles whilst velocity has a con-
cave upward-facing shape above the velocity maxi-
mum (Fig. 10A–D). The velocity maximum occurs
between 1 to 2 cm above the bed and at a fraction of
0.1–0.15 of the total flow depth. The velocity maxi-
mum in experiments 1 to 6 occurs at a similar height
to the interfacial boundary between layers. Root-
mean-squared (RMS) values of the temporal velocity
-10 0 10 20 30 40 50 60
t, s
0
100
200
300
020406080100
020406080100
C, %
ags
020406080100
0
100
200
300
020406080100
C, %
ags
020406080100
0
100
200
300
020406080100
C, %
ags
020406080100
0
100
200
300
020406080100
C, %
ags
020406080100
0
100
200
300
020406080100
C, %
ags
020406080100
C, %
ags
0
100
200
300
B. Run 3 (40%)
A. Run 1 (20%)
C. Run 5 (60%)
D. Run 6 (70%)
E. Run 7 (80%)
F. Run 8 (90%)
u, m
m/s
u, m
m/s
u, m
m/s
u, m
m/s
u, m
m/s
u, m
m/s
Fig. 9. (A–F) Temporal profiles of velocity and concentration measured at 1 cm and 0.5 cm above the bed, respectively, and at a position of 2.5
m downstream of the inlet point, for selected experiments. The time, t, is measured in seconds after the arrival of the flow front. The dashed line
of flow concentration for experiment 8 shows data that have been corrected for temporal displacement resulting from variable siphon flow rates.
See text for further explanation.
L.A. Amy et al. / Sedimentary Geology 179 (2005) 5–29 17
time-series, a proxy for flow turbulence, have maxi-
mum values close to the bed and show a decrease
upwards away from the bed. This type of distribution
has been observed in other quasi-steady turbulent
density currents (e.g., Buckee et al., 2001).
4.2.2. Flows with a fast upper layer
The temporal and vertical profiles of experimen-
tal currents 7 and 8 with slower moving lower
layers are markedly different to those with faster
ones. In these currents the flow velocity and con-
0 100 200 300
u, mm/s
0
2
4
6
8
10
12
z,cm
0 20 40 60 80 100
C, % ags
0 100 200 300
u, mm/s
0 10 20 30 40 50
Urms, mm/s
u, mm/s u, mm/s C, % ags Urms, mm/s
0
2
4
6
8
10
12
0 100 200 300
u, mm/s
0 20 40 60 80 100
C, % ags
0 10 20 30 40 50
Urms, mm/s
0
2
4
6
8
10
12
0 100 200 300
u, mm/s
0 20 40 60 80 100
C, % ags
0 10 20 30 40 500 10 20 30 40 50
Urms, mm/s
0
2
4
6
8
10
12
0 100 200 300
u, mm/s
0 20 40 60 80 100
C, % ags
0 10 20 30 40 50
Urms, mm/s
0
2
4
6
8
10
12
0 100 200 300
u, mm/s
0 20 40 60 80 100
C, % ags
0 10 20 30 40 50
Urms, mm/s
B. Run 3 (40%) E. Run 7 (80%) F. Run 8 (90%)C. Run 5 (60%)
0 20 40 60 80 1000
2
4
6
8
10
12
z,cm
0
2
4
6
8
10
12
0 20 40 60 80 1000 20 40 60 80 1000 20 40 60 80 1000
2
4
6
8
10
12
0 20 40 60 80 1000 20 40 60 80 1000 20 40 60 80 1000
2
4
6
8
10
12
0 20 40 60 80 1000 20 40 60 80 1000 20 40 60 80 1000
2
4
6
8
10
12
0 20 40 60 80 1000 20 40 60 80 1000 20 40 60 80 100
G. Run 1 (20%) H. Run 3 (40%) K. Run 7 (80%) L. Run 8 (90%)I. Run 5 (60%)
0
2
4
6
8
10
12
0 100 200 300
u, mm/s
0 20 40 60 80 100
C, % ags
0 10 20 30 40 50
Urms, mm/s
D. Run 6 (70%)
0
2
4
6
8
10
12
0 20 40 60 80 1000 20 40 60 80 1000 20 40 60 80 100
J. Run 6 (70%)
Dimensional values
Normalised values
A. Run 1 (20%)
υ, σ , χ υ, σ , χ υ, σ , χ υ, σ , χ υ, σ , χ υ, σ , χ
υ συ χ
L.A.Amyet
al./Sedimentary
Geology179(2005)5–29
18
L.A. Amy et al. / Sedimentary Geology 179 (2005) 5–29 19
centration do not mimic one another (Fig. 8). Near
bed velocities increase with time and then decrease
whilst concentration is initially low before increas-
ing (Fig. 9). In experiment 8 an abrupt increase in
flow concentration is seen at about 15 s, marking
the arrival of the slow-moving lower layer (Fig.
9F). An area of enhanced flow velocity also occurs
several centimetres above the bed before the arrival
of the slow-moving lower layer at 7–12 s (Fig. 8).
Initial flow unsteadiness is related to the passage of
the head but also to their longitudinal structure.
Vertical concentration profiles of experiments 7
and 8 display a less gradual and more step-like
stratification than experiments with lower concen-
trations, e.g. experiments 1 and 3 (Fig. 10E–F).
Currents with a slow moving lower layer have a
convex shaped velocity distribution above the max-
imum being quite different to those recorded for
experiments 1–6. Also in experiment 8, the velocity
maximum is relatively high in the flow and it is
situated in the upper-layer at a fractional depth of
between 0.3 and 0.4 (Fig. 10L). This velocity
distribution is similar to those of high-concentration
suspension flows reported by Postma et al. (1988).
RMS values of the temporal velocity time-series
show profiles and values similar to currents with
a faster lower layer; maximum values occur close
to the lower flow boundary and values decrease
upwards. The high RMS values recorded near the
bed are surprising, especially for experiment 8,
since visual observations suggest that the lower
layer moved in a laminar fashion. Since the veloc-
ity measurements taken by the lowest position
probe were taken from an area close to the inter-
facial boundary between layers, we suggest that the
high RMS values are caused by turbulence related
to the interfacial boundary. Alternatively, conditions
close to the tank wall may have been different to
those in the centre of the tank where the data were
recorded; the interfacial boundary may have been
lower or the lower layer more turbulent in the
centre of the tank.
Fig. 10. (A–F) Vertical profiles of the downstream velocity (u), the root-m
for selected experiments. Profiles are taken from the body of currents at 10
was taken at 20 s. Velocities indicated by crosses show values corrected
uncorrected values. (G–L) Non-dimensional values given as a percentage o
mean-square of downstream velocity (r) and concentration (v).
5. Discussion
The experiments show that the behaviour of con-
tinuously-fed gravity currents is strongly controlled
by their stratification. Initial flow unsteadiness is re-
lated to the passage of the head, but also to the
current’s longitudinal structure. In the ranges of q*,l* and B* (all b0.5) investigated, two distinct types
of behaviour were observed. A summary of the char-
acteristics of each current type is shown in Fig. 11.
Those with low to moderate maximum concentra-
tions, b75% glycerol, have fast-moving, high-concen-
tration, basal regions (Fig. 11A). In contrast, currents
with relatively high maximum concentrations, N75%
glycerol, have a slow moving basal region that lags
behind the flow front (Fig. 11B). Particle-laden labo-
ratory currents with similar flow structures to these
two types have been observed previously. Flows with
high-concentration fast-moving bases have been noted
by Hampton (1972), Mohrig et al. (1998) and Marr et
al. (2001), whilst currents with slow moving bases
have been observed by Postma et al. (1988).
5.1. Interpretation of flow behaviour
We interpret the observed change in behaviour to
correspond to a transition in the flow dynamics of the
lower layer from being inertia-driven to viscosity-
controlled. A reduction in the lower layer velocity at
concentrations exceeding 75% glycerol can be ex-
plained by the enhanced drag at the lower flow bound-
ary, a characteristic of high viscosity flows. The dimen-
sionless Reynolds number (Re) is a measure of the ratio
of inertial to viscous forces and may be used to assess
the transition between turbulent and laminar flow re-
gimes. The Reynolds number is evaluated here using
Re ¼ uhql
ð6Þ
where u is a velocity scale, h is a length scale and qand l are the density and viscosity of the fluid,
respectively. The drag force experienced by flows
ean-square of downstream velocity (Urms) and the concentration (C)
s after the arrival of the flow front except the profile for Run 8 which
for fluid density (see text) whilst those indicated by squares show
f the maximum value in each profile; downstream velocity (t), root-
Time
u,c
Velocity
Concentration
iii
Time
u,c
Velocity
Concentration
A. Current with a relatively fast-moving high-concentration phaseE.g., experiments 1-6
X Z
B. Current with a relatively slow-moving high-concentrationE.g. Experiments 7 and 8
Vertical bed structure
Vertical bed structure
High-concentration Low-concentration
Y
Fig. 11. Summary diagram of experimental data showing the two different current types observed. (A) A current with a relatively fast-moving,
high-concentration lower layer that also forms the flow front. (B) A current with a relatively fast-moving, high-concentration upper layer. The
lower-layer lags behind the flow front. For each current type the near-bed temporal trends of velocity and concentration and corresponding
inferred depositional sequence at a single location is shown.
L.A. Amy et al. / Sedimentary Geology 179 (2005) 5–2920
L.A. Amy et al. / Sedimentary Geology 179 (2005) 5–29 21
varies with Reynolds number. The amount of drag is
similar in turbulent flows with relatively high Rey-
nolds numbers but increases significantly at smaller
Reynolds numbers in the range of transitional and
laminar flow conditions. This relationship between
drag and Reynolds number is known to apply to
pipe flow (Chadwick and Morfett, 1992). van Kessel
and Kranenburg (1996) showed that the drag coeffi-
cient varies strongly at low Reynolds numbers,
Re b103, for gravity currents of fluidQmud (Fig. 12).
The Reynolds number of gravity currents is usually
deduced using the bulk current properties to yield a
single value for the whole current. Since viscosity-
stratified currents may exhibit both laminar and tur-
bulent flow at different heights above the bed, the
application of a single Reynolds number may not be
appropriate. One way to assess the flow character of
strongly stratified currents is to calculate separate
Reynolds numbers for regions near the bed and higher
up in the flow (Table 3A). Reynolds numbers were
calculated for below (RebUmax) and above (ReNUmax
)
the velocity maximum using measured flow properties
Re ¼ UhP
M; ð7Þ
where U, P, and M are the layer depth-average veloc-
ity, concentration, and viscosity respectively, and h is
10 100 1000 10000 100000
Ree
0.001
0.01
0.1
1
CD
Fig. 12. Drag coefficient, CD, as a function of the effective Rey-
nolds number, Ree, for fluid mud gravity currents presented by van
Kessel and Kranenburg (1996). Their original data is approximated
by the curve CD~(12+0.1Ree)/Ree, where Ree is an effective
Reynolds number (see Eq. (11) in van Kessel and Kranenburg,
1996). The graph shows that drag increases significantly at low
Reynolds numbers and in the range below turbulent flow conditions,
Reeb3000, using the criterion proposed by Liu and Mei (1990).
the layer depth (Table 3A). Various values of the
Reynolds number have been proposed for the thresh-
old between laminar to turbulent flow ranging over
one order of magnitude. However, laboratory studies
often take the transitional number as 500–2000 (Simp-
son and Britter, 1979; Allen, 1985). This threshold is
also used in this study and given the range of values
calculated for the experimental currents, would appear
to be a reasonable approximation. Values above the
velocity maximum fall into the turbulent regime,
3000bRe b12000, corresponding to observations of
turbulent particle movement in the upper part of the
flow. Smaller values of the Reynolds number charac-
terise the region below the velocity maximum. Impor-
tantly, they show a decrease from 2500 to 900 Re with
increasing glycerol concentration. These Reynolds
numbers fall into the range where values of drag are
expected to vary strongly and support our interpreta-
tion of a varying drag influence at the lower flow
boundary. The flow Reynolds numbers calculated are
subject to some measurement error in flow concentra-
tion. The value of the Reynolds number below the
velocity maximum for experiment 7 (with the second
highest glycerol concentration) is spuriously high at
1600 compared to values for flows of similar concen-
tration. This corresponds to relatively low values of
glycerol concentration recorded for this flow and is
likely to be an artefact of selective siphoning of fluid
with a lower concentration and viscosity. Reynolds
numbers calculated using the initial values of fluid
density and viscosity for the lower layer (ReL), may
therefore give a better estimate for high-concentration
flows in which mixing was unimportant (Table 3B).
These values indicate Reynolds numbers for the high-
concentration flows of b100, consistent with visual
observations of laminar particle movement in these
currents.
5.2. Stratified flow regimes
Based on experimental data, Gladstone et al.
(2004) constructed a flow regime diagram for the
behaviour of two layer, stratified currents of various
density ratios, q* and buoyancy ratios, B* (Fig. 13A).
In their scheme, the layer buoyancy determines which
layer runs ahead to form the leading edge of the flow;
the lower and upper layers run ahead when B*b0.5
and B*N0.5, respectively. All the currents studied
Table 3
Calculated flow characteristics for laboratory experiments
A. Parameters based on depth-averaged values and Umax subdivision
Experiment UC hC PC MC ReC Fr RiB UbUmaxhbUmax
PbUmaxMbUmax
RebUmaxUNUmax
hNUmaxPNUmax
MNUmaxReNUmax
1 0.10 0.13 994.1 0.0011 11,768 1.1 0.9 0.154 0.02 1004.9 0.0012 2513 0.08 0.11 992.2 0.0011 7768
2 0.10 0.13 1000.6 0.0012 10,510 0.7 1.3 0.136 0.02 1017.5 0.0014 2034 0.09 0.11 997.5 0.0011 8486
3 0.11 0.13 997.4 0.0011 12,069 1.0 1.0 0.116 0.02 1025.5 0.0015 1553 0.09 0.11 992.3 0.0011 9000
4 0.10 0.13 995.1 0.0011 11,598 1.0 1.0 0.130 0.02 1027.3 0.0016 1635 0.08 0.11 989.3 0.0011 8360
5 0.11 0.13 1003.5 0.0012 11,732 0.8 1.3 0.135 0.02 1025.5 0.0015 1800 0.09 0.11 999.5 0.0012 8388
6 0.10 0.13 1003.6 0.0012 10,990 0.7 1.4 0.105 0.02 1064.2 0.0026 847 0.09 0.11 992.6 0.0011 8873
7 0.13 0.13 1001.9 0.0012 14,539 1.0 1.0 0.135 0.02 1044.9 0.0017 1623 0.11 0.11 994.1 0.0011 11,266
8 0.07 0.10 1011.0 0.0013 5522 0.5 2.1 0.095 0.03 1094.1 0.0037 855 0.06 0.07 991.1 0.0011 3664
B. Parameters based on initial starting values and layer subdivision
Experiment UC hC qC lC ReC Fr RiB UbUmaxhL qL lL ReL UNUmax
hU qU lU ReU
1 0.10 0.13 1014.4 0.0015 8804 0.5 1.8 0.154 0.02 1033.6 0.0020 1622 0.08 0.11 1010.9 0.0014 6012
2 0.10 0.13 1019.4 0.0017 7470 0.5 2.1 0.136 0.02 1060.2 0.0026 1114 0.09 0.11 1011.9 0.0015 6454
3 0.11 0.13 1015.6 0.0017 8305 0.6 1.8 0.116 0.02 1084.0 0.0038 669 0.09 0.11 1003.1 0.0013 7497
4 0.10 0.13 1022.3 0.0020 6559 0.5 2.1 0.130 0.02 1104.0 0.0054 531 0.08 0.11 1007.4 0.0014 6303
5 0.11 0.13 1030.9 0.0031 4796 0.5 2.2 0.135 0.02 1140.8 0.0120 255 0.09 0.11 1010.9 0.0014 6994
6 0.10 0.13 1031.8 0.0044 3129 0.4 2.3 0.097 0.02 1165.9 0.0207 109 0.09 0.11 1007.4 0.0014 6925
7 0.13 0.13 1045.1 0.0109 1666 0.5 2.0 0.135 0.02 1196.5 0.0624 52 0.11 0.11 1017.6 0.0015 8368
8 0.07 0.10 1066.7 0.0430 176 0.3 3.9 0.095 0.03 1217.6 0.1402 25 0.06 0.07 1002.1 0.0013 3022
(A) Flow characteristics calculated using depth-averaged values for the whole current and for portions of the current above and below the
velocity maximum. Variables are: depth-averaged velocity, U, in m s�1; height, h, in m; depth-averaged density, P, in kg m�3; viscosity, M, in
kg m�1 s�1; dimensionless Reynolds number, Re =UCPh/M; dimensionless Froude number, Fr =UC/(hC(gP�q0/q0)1/2 (where g is the
acceleration due to gravity and q0 is the density of the ambient fluid); and dimensionless Richardson number, RiB=1/Fr. Subscripts C, NUmax
and bUmax indicate values for the whole current, and the portion of the current below and above the velocity maximum, respectively. Current
height, hC, was estimated from photos. The height of the velocity maximum was estimated using velocity measurements (Fig. 10). (B) Flow
characteristics calculated using initial values for the whole current and for upper and lower layers. Variables are: depth-averaged velocity, U, in
m s�1; height, h, in m; layer density, q, in kg m�3; viscosity, l, in kg m�1 s�1; dimensionless Reynolds number, Re =UCqh/l; dimensionless
Froude number, Fr =UC/(hC(gqC�q0/q0)1/2; and dimensionless Richardson number, RiB=1/Fr. Subscripts L, U and C indicate values for
lower and upper layers and whole current, respectively. Note that the velocity values used are the same in both A and B.
L.A. Amy et al. / Sedimentary Geology 179 (2005) 5–2922
here had a greater buoyancy in their lower layer with
values of B*b0.5. Experiments 1 to 6 with a relative-
ly fast-moving lower layer displayed behaviour con-
sistent with the findings of Gladstone et al. (2004) for
inertial, lock-release currents. In experiments 7 and 8,
the lower layer was relatively slow-moving compared
to the upper layer, despite it having a much greater
driving force, B*b0.2. This indicates that the flow
buoyancy does not control flow behaviour in currents
with strong viscosity stratification l* (of the order of
1�10�3 in the present experiments). Hence, a third
axis to Gladstone et al.’s proposed regime diagram
may be added to describe currents with varied viscos-
ity stratification (Fig. 13B).
The second axis of the regime diagram constructed
by Gladstone et al. (2004) indicates relative amounts
of mixing between layers based on q*; for currents
with q*b0.4 the initial stratification is maintained
whilst for those with q*N0.4 stratification is quickly
destroyed by mixing. Comparison between the present
experiments and those of Gladstone et al. (2004), in
terms of q*, however, is not straightforward for sev-
eral reasons. Firstly, the degree of mixing in continu-
ous-flux flows will be significantly different to surge-
type flows since the latter is dominated by the
dynamics of the current’s head. Secondly, the entire
duration of flow in Gladstone et al.’s study was ob-
served from the point of initiation to their arrest where-
as in the present experiments, currents could not be
observed to their natural stopping point. Bearing this in
mind, it was observed that in all flows with density
ratios b0.35, the initial two layer stratification was
maintained, this being consistent with Gladstone et
al. (2004) results.
Fig. 13. (A) Regime diagram constructed by Gladstone et al. (2004) summarising the behaviour of two-layer density-stratified, surge-type
currents. The graph describes currents in terms of the dimensionless parameters density ratio q* and distribution of buoyancy B*. Modified
from Gladstone et al. (2004). (B) Diagram showing the parameter space varied in the present set of experiments in terms of the three
dimensionless parameters, density ratio q*, distribution of buoyancy B* and viscosity ratio l*. The range in these parameters explored is
shaded. Schematic cartoons for current behaviour are shown for different regimes of flow behaviour.
L.A. Amy et al. / Sedimentary Geology 179 (2005) 5–29 23
L.A. Amy et al. / Sedimentary Geology 179 (2005) 5–2924
5.3. Implications for deposit character
The vertical characteristics of sediment gravity
flow deposits are controlled by temporal variations
in flow velocity and concentration at a point, and
thus, by the stream-wise structure of the current
(Branney and Kokelaar, 2002; Kneller and McCaf-
frey, 2003; McCaffrey et al., 2003; Choux et al.,
2005—this issue). The experimental data show that
stratified flows with relatively fast and slow-moving
lower layers have markedly different near-bed tem-
poral trends in flow properties (Figs. 9 and 10).
Assuming natural sediment-laden flows also display
these current structures, it follows that several dis-
tinct bed types should be deposited. Here we spec-
ulate on the characteristics of these beds, and
propose two simple depositional models for stratified
sediment-laden currents (Fig. 11). These models as-
sume deposition occurs throughout the passage of
the head, body and tail of the flow and at a single
location. The models developed should be consid-
ered as didealT deposit types. As for other models,
such as the Bouma sequence, variations should be
expected given the range of controlling factors on
the final deposit character.
5.3.1. Deposit interpretation
In order to identify the deposits of stratified
submarine currents, sedimentary features that reli-
ably record deposition by flows of low and high
sediment concentration need to be defined. This has
been a controversial subject, especially with regard
to the interpretation of massive sandstones (see dis-
cussions in Kneller and Buckee, 2000; Mulder and
Alexander, 2001; Amy et al., 2005—this issue).
Gradual deposition of particles from a relatively
low sediment-concentration phase (turbidity current)
may be recognised by the presence of tractional
bedforms, vertical normal grading under waning
flow conditions and a high degree of grain-size
sorting within the deposit. However, structureless
intervals such as the Bouma Ta division may be
produced under relatively high sediment-load fall-
out rates leading to the suppression of traction
(Arnott and Hand, 1989). In circumstances of steady
flow the Ta division may also lack grading (Kneller,
1995). Deposition from a high-concentration phase
(debris flow) occurring by en masse settling will
also tend to produce ungraded beds. However,
given a wide grain-size distribution (mud to centi-
metre clasts), debris flow deposits will be distinctive
from Bouma Ta divisions on account of their poor
sorting, relatively high matrix mud contents and
randomly distributed outsized clasts. In addition
debrites may display a distinct shear fabric. These
criteria follow those used by others (e.g., Lowe,
1982; Ghibaudo, 1992; Mulder and Alexander,
2001).
5.3.2. Deposits from flows with fast-moving high-
concentration phases
Particle-laden currents with relatively fast moving,
high-concentration phases will have an initial phase
of deposition from flow with relatively high sediment
concentrations followed later by deposition from
flow with lower sediment concentrations. The result-
ing depositional sequence will comprise a high-
concentration flow deposit overlain by a low-concen-
tration flow deposit (Fig. 11A). Beds displaying this
type of vertical character and ranging from several
metres to over tens of metres in thickness are com-
monly exposed in the Eocene/Oligocene sedimentary
sequence of the Gres de Peıra Cava Formation, SE
France (Fig. 14A). These sediments were deposited
in a relatively small (tens of kilometres long and
wide) deep-water basin in which flows were confined
by the local basin bathymetry (Hilton, 1994; Amy,
2000; McCaffrey and Kneller, 2001; Amy et al.,
2004). In relatively proximal sections, many beds
contain a coarse-grained (small-pebble to very coarse
sand grade), very poorly sorted, clast-rich basal in-
terval. The basal portion of these beds is interpreted
as having been deposited from a high sediment-con-
centration flow phase. The mud content varies later-
ally in the basal interval, implying that the cohesive
strength of the flow may have varied locally. The
upper part of the beds is finer grained and better
sorted and displays normal grading and usually cur-
rent lamination. The upper parts of beds are inter-
preted to record deposition from a relatively low-
concentration portion of the current deposited after
the high-concentration phase had passed. Correla-
tions indicate that these are the deposits of single
flow events and not an amalgamated sequence pro-
duced by multiple flows of different particle concen-
trations (Amy, 2000).
L.A. Amy et al. / Sedimentary Geology 179 (2005) 5–29 25
5.3.3. Deposits from flows with slow-moving high-
concentration phases
Particle-laden currents with relatively slow-mov-
ing, high-concentration phases will deposit initially
Fig. 14. Examples of beds preserved in ancient turbidite successions. (A)
France. This bed is interpreted as the deposit of a stratified sediment grav
Sandstone bed with a tripartite bed structure from the Marnoso-arenacea,
been deposited by a stratified sediment gravity flow with a slower-moving
from flow with low to intermediate sediment concen-
trations, followed by deposition from flow with high
sediment concentrations and finally from the trailing
flow with relatively low sediment concentrations. At a
Sandstone bed from the Gres de Peıra Cava, Maritime Alps of SE
ity flow with a faster-moving high-concentration lower region. (B)
northern Apennines of Italy. This type of bed is interpreted to have
, high-concentration, lower region. See text for further explanation.
L.A. Amy et al. / Sedimentary Geology 179 (2005) 5–2926
single location the deposit will show a tripartite struc-
ture with the deposit of the high-concentration phase
encased between the deposits of the relatively low-
concentration flow phases (Fig. 11B). A significant
proportion of sediment gravity flow deposits of the
Miocene, Marnoso-arenacea Formation located in the
Italian Apennines display this type of tripartite bed
structure. These beds have been interpreted as debris
flow deposits sandwiched between turbidites (Ricci
Lucchi and Valmori, 1980; Talling et al., 2004; Amy
et al., 2005—this issue) formed in an open basin-plain
environment of the Apennine foredeep (Ricci Lucchi
and Valmori, 1980; Argnani and Ricci Lucchi, 2001).
The basal interval is composed of b20 cm, mud-poor
(b15% in thin section; Talling et al., 2004), coarse- to
fine-grained sandstone (Fig. 14B). The middle debrite
sandstone interval is usually slightly finer and thicker
(~20–90 cm) than the basal interval and relatively
mud-rich (15–22% in thin section; Talling et al.,
2004). It contains floating out-sized clasts several
millimetres to tens of centimetres in diameter. The
upper division is usually relatively thin (b20 cm),
fine- to very fine-grained sandstone with millimetre-
scale cross-lamination or parallel lamination.
Evidence that these deposits record a single flow
event rather than several amalgamated event beds are
(a) that the clast-rich debris flow units always occur
within this tripartite vertical bed sequence and (b)
long-distance correlations show that tripartite beds
do not dbreak-apartT into individual beds moving lat-
erally (Talling et al., 2004; Amy et al., 2005—this
issue). Correlations also show that the middle clast-
rich interval pinches out rapidly (over b5 km) down-
stream (Talling et al., 2004). This geometry suggests
that this portion of the bed was deposited en masse by
a high-concentration flow. In comparison, the lower
interval extends downstream of the pinch-out position
of the middle interval (Talling et al., 2004). Beds with
a similar tripartite vertical bed profile have been de-
scribed from the Pennsylvanian Jackfork Group in
Kansas (Hickson, 1999), Jurassic fans in the North
Sea (Haughton et al., 2003), and the Miocene and
lower Pliocene Laga Formation, Italy (Mutti et al.,
1978), demonstrating that these bed types commonly
occur in deep-water systems. Alternative explanations
for the generation of these dsandwichT beds are dis-
cussed by Haughton et al. (2003) and Talling et al.
(2004).
6. Conclusions
The behaviour of stratified gravity currents was
investigated using two-layer, laboratory flows com-
posed of aqueous glycerol solutions. In a set of
experiments the initial density and viscosity stratifi-
cation was systematically changed in a manner that
might occur in particle-laden currents with relatively
low to high sediment concentrations. It has been
shown previously that the vertical distribution of den-
sity and buoyancy profoundly affects the behaviour of
laboratory currents (Gladstone et al., 2004). Results
from this study show that the viscosity stratification
also has an important effect on flow behaviour. In
currents with relatively weak viscosity stratification
the high-concentration basal layer is driven by inertia
and propagates to the nose of the current, provided it
has a larger buoyancy than the upper layer. On the
other hand, in currents with relatively strong viscosity
stratification the high-concentration lower layer is
controlled by viscous forces and lags behind the
flow front regardless of its relative buoyancy. These
two flow types, with a relatively fast- and slow-mov-
ing lower layer, correspond to those with relatively
high and low Reynolds numbers, respectively. We
suggest that a transition in flow type occurs with the
onset of transitional and laminar flow conditions be-
cause of enhanced drag at the lower flow boundary. In
the present experiments this transition was observed at
concentrations of between 70% and 80% glycerol for
the lower layer.
The recorded temporal profiles of velocity and
concentration of currents with relatively weak and
strong viscosity stratification are different. Conse-
quently, stratified currents carrying particles are likely
to show different depositional histories and produce
deposits with varied characteristics. The experimental
results allow some speculation about the character of
stratified flow deposits. Currents with a relatively
fast-moving, high-concentration phase should deposit
beds with high-concentration flow deposits overlain
by those of more dilute flow. A current with a
relatively slow moving, high-concentration phase
should produce a bed with a tripartite structure with
a high-concentration flow deposit sandwiched be-
tween the deposits of more dilute flow. Deposits
with these bed structures are commonly observed in
ancient turbidite successions. Experiments on strati-
L.A. Amy et al. / Sedimentary Geology 179 (2005) 5–29 27
fied flows using particles is suggested as a fruitful
area of future work.
Acknowledgments
This research was funded by the United Kingdom
Natural Environmental Research Council and Con-
oco (now ConocoPhilips) through the Ocean Mar-
gins LINK scheme (Grant number NER/T/S/2000/
0106). Experiments were conducted in the Universi-
ty of Leeds, School of Earth Sciences, Fluid Dy-
namics Laboratory. Mark Franklin, Bob Bows, Gary
Keech, David Forgerty (School of Chemistry), Phil
Fields, Tony Windross are thanked for technical
support and members of Sedimentology Group for
assistance in running experiments. Jaco Baas,
Suzanne Leclair and an anonymous reviewer helped
to improve the original manuscript. Andy Hogg and
Charlotte Gladstone provided useful discussion.
Funding for the laboratory facilities used was pro-
vided by EPSRC (GR/R60843/01) and by a consor-
tium of oil companies including BG, BHP, Chevron
Texaco, Total, Exxon Mobil, ConocoPhillips, Ame-
rada Hess and Shell. We also acknowledge the
award of UK Natural Environment Research Council
grant GR3/10015 that funded development of the
UDVP system.
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www.elsevier.com/locate/sedgeo
Sedimentary Geology 1
Combined measurements of velocity and concentration in
experimental turbidity currents
M. Felix*, S. Sturton, J. Peakall
School of Earth Sciences, University of Leeds, Leeds LS2 9JT, United Kingdom
Abstract
Three different sets of experimental turbidity currents were run in which velocity and concentration were measured
simultaneously, for several different heights above the bed. One set with cohesive sediment had an initial volumetric
concentration of 16% kaolinite, and the other two sets with non-cohesive sediment had concentrations of 28% and 4% silica
flour. Velocity was measured at 104–122 Hz using an Ultrasonic Doppler Velocimetry Profiler and concentration was measured
at 10 Hz using an Ultrasonic High Concentration Meter. The similarity of changes in velocity and concentration at the same
measurement heights are described and it is shown that the similarity depends on flow concentration and position in the flow.
The measurements are analysed using cross-correlations and wavelet analysis. Velocity measurements are compared with
analytical solutions for flow around a semisphere and flow around a half body. Measurements and analyses indicate that
turbulence is diminished by stratification, decoupling of regions where turbulence is generated and by reduction of vertical flow
in the turbidity currents.
D 2005 Elsevier B.V. All rights reserved.
Keywords: Turbidity currents; Laboratory experiments; Velocity and concentration measurements; Cohesive sediment; Non-cohesive sediment
1. Introduction
The temporal and spatial flow structure of turbid-
ity currents is the result of the interaction between
sediment and water (Kuenen, 1951; Lowe, 1982;
Altinakar et al., 1996; Felix, 2002). Sediment is the
driving force of these flows and it is transported as a
result of different mechanisms, such as particle–parti-
cle interactions, matrix strength and turbulence (Lowe,
0037-0738/$ - see front matter D 2005 Elsevier B.V. All rights reserved.
doi:10.1016/j.sedgeo.2005.04.008
* Corresponding author.
E-mail address: [email protected] (M. Felix).
1982; Normark and Piper, 1991). Turbulence is gener-
ated by shear in the flow and can be damped by high
concentration (Kundu, 1990).
Details of the flow structure can influence deposits
and run-out distances of turbidity currents (Kneller and
Branney, 1995; Kneller and McCaffrey, 2003). Such
details include the location of turbulence generation
and the extent of similarity of spatial and temporal
changes in velocity and concentration, which are the
topics of the work presented here. At present, under-
standing of these details is only partial due to a lack of
suitable observations, both in the laboratory and in
nature.
79 (2005) 31–47
M. Felix et al. / Sedimentary Geology 179 (2005) 31–4732
Numerical modelling routinely includes both veloc-
ity and concentration (e.g., Hinze, 1960; van Andel and
Komar, 1969; Parker et al., 1986; Stacey and Bowen,
1988; Eidsvik and Brørs, 1989; Felix, 2001) and results
provide a good insight into flow structure and run-out
length. However, the link between velocity and con-
centration is generally prescribed by assuming that all
sediment has the same velocity as the water except for
the vertical direction where settling is included. Be-
cause of this assumption, laboratory experiments and
natural-scale observations can be useful.
Measurements in nature are generally for flows
with low sediment concentration as these are more
common and less destructive than high-concentration
flows. Combined velocity and turbidity observations
have been made in a number of studies, such as
Tesaker (1975), Hebbert et al. (1979), Fan (1986),
Chikita (1989), Chikita et al. (1991), Mitsuzawa et
al. (1993) and Khripounoff et al. (2003). These obser-
vations are limited to single heights in a flow or
present only single vertical profiles. Extensive tempo-
ral observations were made by Samolyubov (1986)
and Samolyubov and Bystrova (1994), but again,
these flows were of low concentration.
Combined measurements of velocity and concen-
tration in laboratory turbidity currents are presented
by Bonnefile and Goddet (1959), Tesaker (1969,
1975), Parker et al. (1987), Garcia and Parker
(1993), Garcia (1993, 1994), Altinakar et al. (1996),
Lee and Yu (1997) and Yu et al. (2000). These
studies present vertical profiles but do not show
changes with time. Many laboratory experiments
were small scale and flows were barely able to
keep sediment in suspension so the results are valid
for depositional flow stages only.
To address some of these limitations, combined
velocity and concentration measurements are pre-
sented here to look at both the temporal and spatial
flow structure of turbidity currents in short-lived,
rapid, lock-exchange flows. Results for flows of dif-
ferent concentration and sediment type (cohesive and
non-cohesive) are presented.
2. Methodology
Three different sets of lock-exchange experiments
were run where velocity and volumetric concentration
were measured simultaneously at the same height in
each run, with changing heights between runs in each
set. One set of five runs with cohesive sediment
(kaolinite) had initial volumetric concentration of
16% and a second set of five runs with non-cohesive
sediment (silica flour) had initial volumetric concen-
tration of 28%. Measurement heights for these two
sets were 23, 41, 94, 150 and 211 mm above the bed.
The third set of runs had non-cohesive sediment with
initial volumetric concentration of 4% and was mea-
sured at 11 heights: 4, 14, 23, 32, 41, 94, 150, 211,
261, 311 and 361 mm above the bed. The grain size of
the silica flour was d10=2 Am (10th percentile),
d50=9 Am (median grain size) and d90=27 Am(90th percentile) and that of the kaolinite was d10=1
Am, d50=6 Am and d90=35 Am.
The experiments were run in a 4.5-m-long-by-0.2-
m-wide-by-0.5-m-high perspex channel which was
open at the top and the downstream end. This chan-
nel was inserted in a larger (6 m by 0.5 m by 1.5 m)
glass-walled flume (Fig. 1). The insert channel rested
on a false floor, tilted at an angle of 58, positioned0.5 m above the bottom of the flume to create a
sump and to minimise end-wall reflections. The
flume was filled with tap water before each experi-
ment. A lock box was filled with a total volume of
120 l of water and sediment mixture as input for
each experiment and stirred using a mechanical
mixer. Stirring was stopped just before the opening
of the lock gate.
Downstream velocities were measured using a 2-
MHz Ultrasonic Doppler Velocimetry Profiler
(UDVP) in each run, giving a measuring frequency
of 104–122 Hz. For a further description see Best et
al. (2001), who also used this technique in turbidity
currents. The velocity probe was positioned 3.42 m
downstream of the lock gate in the centre of the
insert channel at the measurement heights mentioned
above. The water was 0.8 m deep at the measuring
location.
Concentration was measured using an Ultrasonic
High Concentration Meter (UHCM), which has a
measuring frequency of 10 Hz. The UHCM measures
obscuration between a sender and a receiver 10 mm
apart, outputting a voltage between 0 and 10 V. This
output signal is calibrated by measuring several sed-
iment water mixtures of known concentration, using
the same type of sediment as used in the experiments.
0.5
m
0.2
m
5o
688
mm
471
mm
568 mm
UHCM
3.4 m
120 L (not to scale)
mixer
Lock gate
4.5 m
UDVP rack for all measuring heights
Fig. 1. View of insert channel and lock box. The insert channel is open at the top and downstream ends and is placed in a larger flume filled with
tap water. The insert channel is placed 0.5 m above the base of the flume, creating a sump underneath. Note that velocity was recorded using a
single UDVP probe at a different height for each run. The lock box was not filled to capacity to prevent spilling during mixing.
M. Felix et al. / Sedimentary Geology 179 (2005) 31–47 33
The UHCM was positioned next to the UDVP probe
at the same height.
The results are analysed in several ways. Cross-
correlation of the velocity and concentration signals at
the same heights in each flow compares the shapes of
the normalised signals. Values of the cross-correlation
coefficient range from �1 (signals have the same
shape but have a phase shift of 1808) to +1 (signals
have the same shape and are in phase), with a value of
zero indicating no correlation at all. Low-amplitude
variations have little influence on the cross-correlation
coefficient which therefore measures how comparable
the waveforms and shape of the entire signals are
(e.g., Gubbins, 2004).
A more detailed comparison of the fluctuations is
provided by wavelet analysis, which shows the tempo-
ral variations of different scales for the concentration
and velocity signals (e.g., Farge, 1992; Kumar and
Foufoula-Georgiou, 1997). This type of analysis is
well suited for short-lived flows with a large range of
scales and allows comparison of the velocity and con-
centration signals with the same method. Large scales
represent large-scale motions while small scales repre-
sent rapid fluctuations caused by turbulence (Farge,
1992; Brunet and Collineau, 1997; Howell and
Mahrt, 1997). Paul, Haar and complex-valued Morlet
wavelets were used and showed the same features.
Because the Morlet wavelet results show the least
artefacts, only these are shown here. For long-duration
signals of constant frequency content, the relation be-
tween Morlet scale s and frequency f tends towards
s=2.5/f. The results are presented in scalograms where
scale is plotted against time and the amplitude of the
contribution of each scale to the signal is contoured.
Finally, velocities are compared with analytical
solutions for flow around different body shapes (semi-
sphere and half body, see, e.g., Hampton, 1972;
Kundu, 1990; McElwaine and Nishimura, 2001), to
show that velocity and concentration do not only
influence each other locally, but that the influence of
the near-bed flow extends high up into the flow.
Results are only shown for a measurement height of
M. Felix et al. / Sedimentary Geology 179 (2005) 31–4734
211 mm, but other heights where concentration was
significantly lower than the concentration at a height
of 23 mm showed the same results.
Results from these analyses are combined with
previous theoretical work and observations to discuss
locations where turbulence is generated and the im-
portance for flow structure and deposit formation.
3. Experimental results
The results for the runs with 16% kaolinite are
shown in Fig. 2, for the 28% silica flour runs in
0 5 10 150
0.1
0.2
0 5 10 15
0 5 10 150
0.1
0.2
0 5 10 15
0 5 10 150
0.1
0.2
conc
entr
atio
n
0 5 10 15
0 5 10 150
0.1
0.2
0 5 10 15
0 5 10 150
0.1
0.2
time (sec0 5 10 15
Fig. 2. Concentration and downstream velocity profiles for flows of 16% k
of the plots, velocity profiles are thinner grey lines. All velocity profiles ar
bed while the bottom profile is closest to the bed; z is the height above
upstream. Time t =0 is the start of the measurement period when the lock
Fig. 3 and for the 4% silica flour runs in Figs. 4 and
5. Fig. 4 shows profiles at the same five heights as in
Figs. 2 and 3 while in Fig. 5 measurements at all 11
heights are used to construct a contour plot of con-
centration. The velocity profiles in Figs. 2–4 are
moving averages over 10 points, so they are of a
similar frequency as the concentration profiles and
allow easier comparison. Cross-correlation coeffi-
cients are shown in Table 1. All flows were fast
enough to keep sediment in suspension and no depo-
sition took place until the final part of the tail.
After the opening of the lock gate (t =0), all
velocity profiles show a high value followed by a
20 25 3020 25 30-500
0
500
1000z=211 mm
20 25 3020 25 30-500
0
500
1000z=150 mm
20 25 3020 25 30-500
0
500
1000u
(mm
/sec
)z=94 mm
20 25 3020 25 30-500
0
500
1000z=41 mm
20 25 30)
20 25 30-500
0
500
1000z=23 mm
aolinite. Concentration profiles are thick black lines near the bottom
e moving averages of 10 points. The top profile is furthest from the
the bed. Positive velocity is downstream, and negative velocity is
gate is lifted.
0 5 10 15 20 25 300
0.1
0.2
0.3
0 5 10 15 20 25 30
0
500
1000z=211 mm
0 5 10 15 20 25 300
0.1
0.2
0.3
0 5 10 15 20 25 30
0
500
1000z=150 mm
0 5 10 15 20 25 300
0.1
0.2
0.3
conc
entr
atio
n
0 5 10 15 20 25 30
0
500
1000
u (m
m/s
ec)
z=94 mm
0 5 10 15 20 25 300
0.1
0.2
0.3
0 5 10 15 20 25 30
0
500
1000z=41 mm
0 5 10 15 20 25 300
0.1
0.2
0.3
time (sec)0 5 10 15 20 25 30
0
500
1000z=23 mm
Fig. 3. Concentration and downstream velocity profiles for flows of 28% silica flour. Concentration profiles are thick black lines near the bottom
of the plots, and velocity profiles are the thinner grey lines. All velocity profiles are moving averages of 10 points. The top profile is furthest
from the bed while the bottom profile is closest to the bed; z is the height above the bed. Positive velocity is downstream, and negative velocity
is upstream. Time t =0 is the start of the measurement period and corresponds to the opening of the lock gate.
M. Felix et al. / Sedimentary Geology 179 (2005) 31–47 35
period of decrease. In some profiles, negative (up-
stream) velocity is present, especially at the higher
measuring positions. Velocity increases before the
concentration ramp-up, which indicates the water
that is displaced just before the arrival of the
turbidity current at the measurement location. Con-
centration always increases abruptly from zero. All
velocity and concentration profiles in the turbidity
currents show a number of different asymmetrical
peaks consisting of a rapid increase followed by a
more gradual decrease. The main flow duration
decreases from 10–20 s down to 2–5 s at the
highest measurement positions, and the number of
peaks decreases from about 4 to 1.
The cross-correlation coefficients are calculated for
the period of high concentration. Values are high
below 94 mm but somewhat lower for the two top-
most profiles, indicating a high degree of similarity
between the shapes of the two profiles without a shift
in phase. The cross-correlation coefficients indicate a
high degree of similarity, but in detail there are diffe-
rences (Figs. 2–4), with the greatest similarity for
(parts of) profiles with relatively low concentration,
for example, in the tails. Especially at a height of 94
mm, the profiles are very similar; the shapes, number
of peaks and duration of peaks are the same for the
velocity and concentration profiles in all three flows
(Figs. 2–4). Nearer to the bed, where concentration is
0 10 20 30 40 50 600
0.02
0.04
0 10 20 30 40 50 60
-200
0
200
400z=211 mm
0 10 20 30 40 50 600
0.02
0.04
0 10 20 30 40 50 60
-200
0
200
400z=150 mm
0 10 20 30 40 50 600
0.02
0.04
conc
entr
atio
n
0 10 20 30 40 50 60
-200
0
200
400
u (m
m/s
ec)
z=94 mm
0 10 20 30 40 50 600
0.02
0.04
0 10 20 30 40 50 60
-200
0
200
400z=41 mm
0 10 20 30 40 50 600
0.02
0.04
time (sec)0 10 20 30 40 50 60
-200
0
200
400z=23 mm
Fig. 4. Concentration and downstream velocity profiles for flows of 4% silica flour. Concentration profiles are thick black lines near the bottom
of the plots, and velocity profiles are thinner grey lines. All velocity profiles are moving averages of 10 points. The top profile is furthest from
the bed while the bottom profile is closest to the bed; z is the height above the bed. Positive velocity is downstream, and negative velocity is
upstream. Time t =0 is the start of the measurement period, when the lock gate is lifted.
M. Felix et al. / Sedimentary Geology 179 (2005) 31–4736
highest, similarity is not as good (e.g., the 16% flow at
a measurement height of 23 mm; Fig. 2).
For all three flows at all measurement heights,
velocity decreases more gradually with time than
concentration (Figs. 2–4). For example, the temporal
decrease of the 16% flow concentration at 23 mm is
much more abrupt than that of velocity (Fig. 2). The
maximum velocity measured during the entire flow
period occurs in the second period of high velocity
values. The velocity maximum at any given time in
the 16% flow is measured at a height of 41 mm above
the bed. For the 28% and 4% flows, the maximum
velocity measured at a given time at the five heights is
either at 41 mm or at 23 mm above the bed. The
velocity maximum of all flows is therefore around
these heights. The velocity decreases gradually from
the bed towards the highest measurement positions
and only near the top does the maximum velocity
differ significantly from the maximum near-bed ve-
locity (Figs. 2–4). Near the bed, the concentration is
comparable to the input concentration but it decreases
towards the higher measurement positions.
In the 4% flows (Fig. 4), the concentration in the
first period of high values varies little at the different
measurement heights, but the vertical variation
increases away from the head. The highest concentra-
tion is in the second period of high values. The
maximum velocity is also behind the head and gets
closer to the bed with time. The interpretation of this
flow structure is that the amount of mixing is high in
0 5 10 15 20 25 30 350
0.05
0.1
0.15
0.2
0.25
0.3
0.35
time (sec)
z (m
)
0.04 0.04
0.03
0.02
0.01
0.005
0.010.020.03
0.04
0.01
0.01
0.0050.005
Fig. 5. Volumetric concentration contour plot using measurements at 11 different heights z above the bed in 11 different flows of 4% silica flour,
assuming that flow structure of each flow is comparable to the others. Contours are in volume fraction. Time t =0 corresponds to the opening of
the lock gate.
M. Felix et al. / Sedimentary Geology 179 (2005) 31–47 37
the head but decreases in the body, causing the sed-
iment to accumulate nearer the bed and provide a
larger driving force leading to higher velocity. In the
28% flows (Fig. 3), the concentration at the lowest
measurement height varies less than for the 4% flows
and concentration in the head decreases more towards
the top, causing these flows to be more vertically
stratified than the 4% flows. In the 16% flows (Fig.
2), the lowest profiles show an abrupt termination of
high concentration with time, but higher up, the pro-
files are more comparable to the 28% and 4% flows.
The vertical concentration decrease is largest in the
16% flows, the most stratified of the three.
In Fig. 5, all 11 measurement heights of the 4%
silica flour flows are included in contour plots of
concentration that show the temporal and vertical
spatial development. The head is about twice as
thick as the body and flow thickness is about 0.36
Table 1
Cross-correlation coefficients for the velocity and concentration
profiles in the turbidity currents at different measurement heights
23 mm 41 mm 94 mm 150 mm 211 mm
4% 0.9184 0.9343 0.8932 0.8117 0.8452
28% 0.8830 0.8102 0.8955 0.7852 0.6480
16% 0.9075 0.8503 0.8903 0.7670 0.6655
The profiles were normalised by the same maximum velocity and
concentration in each set of runs.
m (flow just reaches the highest measurement posi-
tion). The same set of peaks and troughs as seen in
Fig. 4 can also be seen in Fig. 5. The variations
extend downwards to the lowest measurement posi-
tion but the height towards which they extend up-
wards decreases with time. In the head, concentration
starts decreasing only above 0.15 m, but in later
periods of high concentration, the flow is more
stratified vertically.
4. Wavelet analysis
Results of the wavelet analysis are shown in Figs.
6–8 for the 16%, 28% and 4% flows, respectively. In
both velocity and concentration plots for all three
flows, the large scales normally have the highest
amplitude and these are also most persistent tempo-
rally. Small scales tend to vary more temporally and
have lower amplitude than large scales but this varies
depending on height in the flow and on concentration.
The small scales with the highest amplitudes are
present at the 94 mm height. There, both large and
small scales are of comparable amplitude and are
similar for velocity and concentration. Towards the
bed, small scales for concentration have lower ampli-
tude for flows of high concentration (Figs. 6 and 7)
but are still of relatively high amplitude for velocity.
Fig. 6. Scalograms for the 16% kaolinite runs using a Morlet wavelet. Left side figures are for velocity, and right side figures are for
concentration. The measurement heights are, from top to bottom, z =211, 150, 94, 41 and 23 mm, the same as shown in Figs. 2–4. Dark colours
indicate scales with high amplitude, and light colours indicate low amplitudes.
M. Felix et al. / Sedimentary Geology 179 (2005) 31–4738
The scale with the highest amplitude in the velocity
profile is largest in the displaced water just in front of
the turbidity current and decreases towards the time of
onset of high concentration. This onset of high con-
centration is also associated with small-scale velocity
fluctuations. Near the top of the three flows, where
concentration is relatively low, the decrease with time
of the scale with the largest amplitude is smaller than
for lower measurement heights and not always notice-
able. In the concentration plots, the onset of high
concentration is seen by the pattern characteristic of
Morlet wavelet analysis of a box-shaped signal. The
abrupt decrease in near-bed concentration of the 16%
flow is similarly shown (Fig. 6).
5. Influence of near-bed flow on ambient water and
upper part of the flow
The two previous methods of analysis (cross-cor-
relation and wavelets) focussed on a comparison of
signals at the same height, but the mutual influence
of velocity and concentration extends throughout the
entire flow. The results for the 4% flows (Fig. 4)
Fig. 7. Scalograms for the 28% silica flour runs using a Morlet wavelet. Left side figures are for velocity, and right side figures are for
concentration. The measurement heights are, from top to bottom, z =211, 150, 94, 41 and 23 mm, the same as shown in Figs. 2–4. Dark colours
indicate scales with high amplitude, and light colours indicate low amplitudes.
M. Felix et al. / Sedimentary Geology 179 (2005) 31–47 39
show several short periods of varying concentration,
while the results for the lowermost measurement
heights of the 16% flows (Fig. 2) show much
more constant concentration with only one dip pres-
ent a few seconds after the onset of high concentra-
tion. These different patterns of concentration lead to
different flow patterns in the upper parts of the
currents, as can be shown by comparing the present
velocity results to analytical solutions for flow
around two different body shapes (Fig. 9). The first
body shape used in the analysis is a semisphere,
which is half a sphere with its flat side on the bed
and a body of limited spatial extent. The second is a
half body which is a blunt body whose thickness is
zero at the bed and increases with distance from the
zero point. A half body is of much larger spatial
extent than the semisphere (e.g., Hampton, 1972;
Kundu, 1990; McElwaine and Nishimura, 2001).
The 4% flow shows several short periods of high
concentration with periods of low concentration in
between (Figs. 4 and 5), while the 16% flow shows a
long period of more consistently high concentration
Fig. 8. Scalograms for the 4% silica flour runs using a Morlet wavelet. Left side figures are for velocity, and right side figures are for
concentration. The measurement heights are, from top to bottom, z =211, 150, 94, 41 and 23 mm, the same as shown in Figs. 2–4. Dark colours
indicate scales with high amplitude, and light colours indicate low amplitudes.
M. Felix et al. / Sedimentary Geology 179 (2005) 31–4740
near the bed (Fig. 2), so the flow pattern in the 4%
turbidity current is expected to be more like that
around a semisphere while for the 16% flow the
pattern is expected to be more like that for a half
body.
The analytical solutions are strictly valid only
for ambient flow. However, the persistence of
wavelet scales between the displaced water in
front of the turbidity current and in the current
itself indicates a smooth transition between the
two. Additionally, the wavelet pattern at the highest
measurement position for all flows shows that the
influence of concentration on velocity is small at
these heights. For these reasons, the present exper-
imental results can be compared with the analytical
solutions.
-4 -3 -2 -1 0 1 2 30
1
2
3
4
z
flow lines around a half-body
0 1 2 3-1
-0.5
0
1
u
horizontal velocity above a half-body
-3 2 -1 0 1 2 30
0.5
1
1.5
2
z
flow lines around a semisphere
-3 -2 -1 0 1 2 30
0.5
1
x
u
horizontal velocity above a semisphere
-4 -3 -2 -1
0.5
Fig. 9. Analytical solutions for flow around a half body and flow
around a semisphere. Horizontal velocities at one height (see text
for equations) above the bodies are shown below the 2D vertical
cross sections through the bodies. Non-dimensional axis units.
M. Felix et al. / Sedimentary Geology 179 (2005) 31–47 41
For two dimensional flow around a half body, the
downstream velocity u is given (in polar coordinates)
by
u ¼ U þ mcos hð Þ=r;
where U [m/s] is the velocity of the ambient, m [m2/s]
is the material flux, r [m] is the radial coordinate and
h is the angular coordinate. For two-dimensional flow
around a semisphere, downstream velocity is given by
u ¼ U þ a2U=r2;
where a [m] is the radius of the hemisphere (see Fig.
9). These two analytical solutions are compared to the
velocity measurements in a reference frame moving
with the turbidity currents (Fig. 10). The 16% kaolin-
ite flow fits the pattern of flow around a half body,
consisting of initial velocity decrease, rapid increase
at the onset of the high concentration and final de-
crease. The first second of observed data does not fit
the trend, which is interpreted to be the result of the
initial disturbance due to lock gate opening and the
slumping phase of the flow (Simpson, 1997). The 4%
silica flour flow fits the pattern of symmetrical veloc-
ity increase and decrease of flow around a semisphere.
The 28% silica flour flow does not fit either pattern,
but is transitional between the two, with an initial
velocity decrease characteristic of the half body but
symmetrical flow as for the semisphere.
6. Discussion: implications for turbulence
generation and sedimentation
Turbulence is generated in turbidity currents by
shear throughout the entire flow, but three main loca-
tions have been recognised in previous studies. As in
all wall-bounded flows, turbulence is generated near
the bed, leading to high values close to the bed but
diminishing away from it. Turbulence is also generat-
ed near the top of the flow due to shear with the
ambient fluid, although this is less important for
flow in deep ambient than for laboratory experiments.
A third-generation region has been recognised in a
middle region of the flow, at a height just above the
height of the velocity maximum (Dallimore et al.,
2001; Buckee, 2000; Felix, 2002) but the cause of
turbulence at this location has not been well explained
yet. An explanation will be given below using obser-
vations from previous experimental and theoretical
work. The relative importance of turbulence generated
at the different heights varies. For example, a rough
bed will lead to high turbulence near the bed, while a
smooth bed will lead to less turbulence near the bed,
so that turbulence generated just above the height of
the velocity maximum may be more important.
Turbulence is highest near to where it is generated
and diminishes away from it. If turbulence generated
in different regions interacts, sediment can be distrib-
uted vertically throughout the flow more easily. A
lack of interaction has been described in the turbu-
lence minimum just below the height of the velocity
0 1 2 3 4 5 6 7 8 9 10-500
0
500
u (m
m/s
ec) 16 % kaol
0 1 2 3 4 5 6 7 8 9 10-500
0
500
1000
u (m
m/s
ec) 28 % sf
0 2 4 6 8 10 12 14 16 18 20-500
0
500
u (m
m/s
ec) 4 % sf
time (sec)
Fig. 10. Comparison of analytical solutions for flow around bodies with the measured velocity values for the three different flows. All
measurement heights are at 211 mm above the bed. Grey lines are the velocity measurements, continuous black lines are the analytical solutions
for flow around a half body, dashed black lines are the analytical solutions for flow around a semisphere, and the vertical dotted lines indicate
time of onset of high concentration.
M. Felix et al. / Sedimentary Geology 179 (2005) 31–4742
maximum (e.g., Peakall et al., 2000), which is where
turbulence generated near the bed and turbulence
generated in the middle of the flow overlap at the
outer reaches of their influence.
In all three flows in the present experiments, the
wavelet analysis shows that the middle region (see
results for the 94 mm height), above the velocity
maximum, has high-amplitude velocity fluctuations
which will lead to high shear and therefore high
turbulence. For all three flows, the middle region, at
the transition between underflow and dragged upper
flow part, consistently generates turbulence, despite
the differences in concentration and stratification.
Stratification dampens turbulence because it is more
difficult for the denser fluid to be transported upwards
and for lighter fluid to move downwards. As a result,
generation of turbulence in the middle of the flow may
be expected to decrease for increasing stratification, in
contrast to the results shown. To understand sediment
distribution in the flows, it is important to understand
why waviness is generated in the middle of the flow.
The height above the velocity maximum with high
turbulence is the interface between the underflow and
the upper part of the flow which is dragged along.
Lock (1951) derived an analytical solution for the
velocity profile between an underflow and dragged
ambient and found a kink at this height. A similar
kink was found by Ippen and Harleman (1952) in
experiments with clay flows, at the upper surface of
the clay underflow. This kink increases shear and
generates additional turbulence, which was seen in
the numerical experiments of Felix (2002). Although
increased mixing due to turbulence may make this
transition smoother, the division remains clear even
after long flow duration as shown by the numerical
experiments of Felix (2004), where most of the flow
momentum remains concentrated in the underflow
below z1/2, where z1/2 is the height above the velocity
maximum where the velocity equals half the maxi-
mum velocity. Mixing will transport sediment up-
wards in the flow so that this transition height is not
obvious by visual inspection of flows only.
The interface above the height of the velocity max-
imum is not only the location for increased shear, but
also where internal waves will develop. If the transition
height between the lower and upper flow parts is also a
density interface, Kelvin Helmholtz waves may devel-
op (Kundu, 1990) and breaking of these waves leads to
turbulence. However, even if there is no or a very weak
density interface, waves will still develop due to shear
M. Felix et al. / Sedimentary Geology 179 (2005) 31–47 43
instability between two initially different fluid bodies
of the same density (Scorer, 1978) but distinguished by
a non-smooth change in the velocity profile at the
transition between the two flow parts. Samolyubov
and Bystrova (1994) applied the analytical solution of
Scorer describing this mechanism to their observations
of an underflow in a water storage reservoir. This
solution for flow of uniform density describes the
displacement pattern caused by wave formation. The
maximum displacement occurs at the transition be-
tween underflow and upper flow part, and, as for a
density interface, wave breaking may lead to turbu-
lence generation. For decaying waves, such as can be
expected for flows which are more vigorous at the front
than at the rear, this analytical solution shows displace-
ment as in Fig. 11, which is similar to that observed in
the 4% flows (Fig. 5). So for both high-concentration
flows with a density interface and for flows without a
density interface, waves will form. In most flows, both
mechanisms will occur simultaneously.
The results presented for a measurement height of
94 mm show very similar velocity and concentration
profiles (Figs. 2, 3 and 4) and comparable velocity
and concentration wavelet scales with both large and
small scales of high amplitude (Figs. 6, 7 and 8).
-5 0 50
0.5
1
1.5
2
2.5
3
3.5
dista
heig
ht
Fig. 11. Non-dimensional schematic plot of streamlines for the wavy moti
decaying waves. Solid line is the interface line. Method from Scorer (197
These results can be explained with the theoretical
models described above. Although turbulence is al-
ways generated at the middle height, irrespective of
stratification, concentration does have an influence on
the turbulence. For high-concentration flows, turbu-
lence does not have much influence on concentration
for the lowermost measuring positions (Figs. 2 and 3:
absence of small fluctuations, Figs. 6 and 7: absence
of small scales), where concentration is more constant
than for the low-concentration flows (Figs. 4 and 5).
Neither turbulence generated near the bed nor turbu-
lence generated near the middle influences the lower-
most measurement positions (which are, at 23 mm,
between these two heights where turbulence is gener-
ated), so their influence does not extend this far and
the two generation regions can be considered to be
decoupled, diminishing the vertical exchange of sed-
iment within the flow. Such decoupling, of course,
depends on the actual velocity, with high velocity
leading to high turbulence and more overlap between
different regions where turbulence is generated. Low
velocity leads to less turbulence and a decrease in
overlap of the regions. This means that decoupling
may result in a change from the front to the back of
the flow.
10 15 20nce
on formed at the interface of two different displacement regimes for
8).
M. Felix et al. / Sedimentary Geology 179 (2005) 31–4744
The near-bed concentration varies less in higher
concentration flows and the flow pattern caused by
the concentration patterns in the lower part of the flow
is also different, as shown by the comparison of flow
around the different bodies (Fig. 10). For low-concen-
tration flows, the rapid alternation between periods of
(relatively) high and low concentration (see, e.g., Figs.
4 and 5) leads to many periods of upwards and down-
wards flow, as indicated by flow in front of and behind
the semisphere (Fig. 10). For high-concentration flow,
vertical motion is more restricted and flow is more like
flow around and parallel to a half body (Fig. 10). The
imposed forcing up and down of the water for low-
concentration flows leads to high shear and generation
of turbulence. For high-concentration flows, the im-
posed up and down flow decreases as the flow is more
parallel to the long lower part of the flow, leading to a
decrease in shear caused by vertical motion and result-
ing in a decrease of turbulence. It will therefore be
more difficult to mix and suspend sediment in the
higher regions of the flow.
The results of wavelet analysis and flow around a
body therefore show that in higher concentration
flows, turbulence is not only damped by stratification,
but also by decoupling of turbulence generation
regions which inhibits vertical mixing of sediment,
and by a change in the flow patterns in the dragged
upper part of the flow, again inhibiting mixing. These
flow patterns are shown schematically in Fig. 12. The
results now allow a short discussion of the implication
of these flow patterns for deposition, but this discus-
sion necessarily remains speculative as the flows in
the present experiments were non-depositional, so the
implications cannot be confirmed from them. Results
from numerical modelling or observations in deposi-
tional natural-scale flows would be most suitable for
this, being of the right spatial and temporal scales.
The experimental results show the similarity be-
tween the velocity and concentration profiles, with
variations depending on height in the flow and on
flow concentration. When profiles are similar, a
change in either velocity or concentration will be
followed rapidly by a change in the other. Deposition
takes place from sediment transported near the bed
and rapid or slow changes in concentration will result
in rapid or slow changes in deposits and bedforms.
Prediction of deposit changes based on velocity
changes (e.g., Kneller and Branney, 1995; Kneller
and McCaffrey, 2003) depends on the reaction of
concentration to velocity changes.
For low-concentration flows, different turbulence
generation regions interact (see Fig. 12) and sediment
will be transported vertically throughout the flow
relatively easily. The vertical transport of sediment
near the bed is influenced both by turbulence gener-
ated near the bed and by turbulence generated just
above the height of the velocity maximum. A change
in velocity will only be slowly followed by a change
in concentration as near-bed concentration will be
affected relatively little by a change in turbulence
generated near the bed which will be compensated
by turbulence generated just above the velocity max-
imum. Bed thickness will change gradually and
changes in bedforms will also be gradual in the down-
stream direction.
For increasing concentration, the near-bed region
will become decoupled from the middle region (see
Fig. 12) and sediment is transported less easily
vertically. Near-bed concentration is mostly influ-
enced by turbulence generated at the bed with little
influence from turbulence generated above the ve-
locity maximum. A change in velocity and bed
shear will accordingly have a relatively larger
change in concentration than for low-concentration
flows. The resulting deposits will therefore record
smaller changes in velocity more accurately than
low-concentration flows. This is the case for both
deposition and erosion. Deposit thickness can vary
rapidly if the underlying topography causes the
flow velocity to increase or decrease suddenly and
bedforms can also change rapidly from one location
to another.
The two wave generation mechanisms described
above show how turbidity currents may be wavy and
the resulting velocity variations lead to local waxing
and waning behaviour. Although in the present lab-
oratory flows the velocity varies on a scale of sec-
onds, such behaviour has been observed in natural-
scale flows (Samolyubov and Bystrova, 1994) to be
on the order of minutes to hours. This might be
recorded in the deposit through thick laminae.
Lowe (1982) described such lamination (traction
carpets) in deposits of high-concentration flows
being formed as the result of rapid freezing of the
near-bed sediment, but this will only happen for
near-bed concentration close to packing concentra-
High density with stiff layer near the bed that can deform: pluglike flow.
Intermediate density with high density layer near the bed.
TKE
ULow density overall.
LD
dense, mixed
LD
LD
coherent
Fig. 12. Schematic diagram for the three different flows. Left hand column shows a 2D side view of the currents with the different concentration
regimes and the ambient flow around the current (thick black arrow). Right hand column show regions of turbulence generation and their
interaction as well as vertical turbulent kinetic energy (TKE) profiles, linked to height in vertical velocity profiles. LD=low-density part of the
flow. The figure is based on present results and previous theoretical and observational work. See text for discussion.
M. Felix et al. / Sedimentary Geology 179 (2005) 31–47 45
tion. Lower concentration flows can lead to similar
laminations as a result of temporal variations in flow
structure. In such waxing and waning flows, vertical
grain size trends such as inverse grading are more
likely to be preserved in deposits from high-concen-
tration flows if it is assumed that the increase or
decrease in velocity is accompanied by equivalent
changes in transported grain size. In high-concentra-
tion flows, deposition can be so rapid that smaller
grain sizes already deposited are not eroded in a
waxing flow because relatively little turbulence is
available to transport it away from the bed.
7. Conclusions
Combined measurements of velocity and concen-
tration in experimental turbidity currents show that the
similarity of temporal and vertical spatial changes in
velocity and concentration in turbidity currents
depends on bulk concentration of the flow and on
position in the flow. Similarity is high near the middle
of the flow at the transition height between underflow
and ambient. For high concentration, the similarity
decreases, especially near the bed. Despite these dif-
ferences, cross-correlation coefficients of the two sig-
M. Felix et al. / Sedimentary Geology 179 (2005) 31–4746
nals are high, indicating the similarity of the overall
signals and the absence of a phase shift.
Wavelet analysis shows that all profiles have one
large scale of high amplitude which is comparable for
velocity and concentration, while smaller scales gener-
ally have lower amplitude. Small and large scales at the
transition height between underflow and upper part of
the flow are of comparable amplitude as a result of
turbulence generation at this height. For high-concen-
tration flows, turbulence generation regions near the
bed and at the transition height become decoupled
which inhibits vertical mixing of sediment.
Temporal changes in the low-concentration flow
are rapid with several periods of alternating low and
high concentration. This pattern results in flow com-
parable to flow around a semisphere. For high-con-
centration flows, the concentration near the bed varies
more slowly, resulting in a flow pattern comparable to
flow around a half body. Mixing of sediment is re-
duced for high-concentration flows by the decreased
amount of imposed vertical motion as the near-bed
sediment concentration is more constant than for low-
concentration flows.
Acknowledgements
This research was funded by the UK Engineering
and Physical Sciences Research Council, Grant GR/
R60843/01. The UDVPs were funded by UK Natural
Environment Research Council Grant GR3/10015.
Mark Franklin and Gareth Keevil are thanked for
help in the laboratory. Reviews by Yu’suke Kubo,
Jeff Parsons and Jaco Baas helped to significantly
improve the clarity of the paper.
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www.elsevier.com/locate/sedgeo
Sedimentary Geology 1
Comparison of spatio–temporal evolution of experimental
particulate gravity flows at two different initial concentrations,
based on velocity, grain size and density data
C.M.A. Choux a,*, J.H. Baas a, W.D. McCaffrey a, P.D.W. Haughton b
aEarth Sciences, School of Earth and Environment, University of Leeds, Leeds, LS2 9JT, UKbDepartment of Geology, University College Dublin, Belfield, Dublin 4, Ireland
Received 1 May 2004; accepted 6 April 2005
Abstract
Flume experiments were conducted to investigate the spatio–temporal structure of subaqueous particulate gravity flows with
an initial concentration of 14% by volume. Time series of downstream flow velocity and its calculated degree of turbulence,
median grain size and sediment concentration at different positions along the path of nominally identical flows are analysed and
combined to constrain the spatio–temporal evolution of a single idealised flow. Comparison of the 14% flow with a flow of 5%
initial concentration reveals similarities in the basic spatio–temporal structure of velocity, turbulence, grain size and concen-
tration. Both flow types exhibit a velocity maximum at about 1 /3 of the flow height above the flume floor. At that level,
velocity decreases slowly in the flows’ body and more rapidly in their tails. Moreover, turbulence intensity is highest in the head
and at the base of the flows, whereas the level of maximum velocity and the tail of the flows typically are weakly turbulent. The
zones of high turbulence are associated with shear at the front and base of the gravity flows. The flow of 5% and 14% initial
concentration also agree in stratification patterns of median grain size and concentration. Grain populations are relatively well
mixed in the head, show normal grading in the main part of the body and normal to inverse grading in the rear of the body and
tail. The inverse grading is thought to originate from particles transported from the head upward and backward into the body of
the flows, where they subsequently settle. The main difference between the flow of 5% and 14% initial concentration is that the
higher-density flows appear to develop from a jet into a turbidity current closer to the inception point than the lower-density
flow. This difference is interpreted from dimensionless vertical profiles of the flow parameters: horizontal velocity, concen-
tration and grain size distribution. In the turbidity current phase of both flows, the dimensionless variables collapse well. This
indicates that the flows behave in a dynamically similar manner and inspires confidence that the dimensionless variables can be
used to predict the dynamic behaviour of particulate gravity flows across the measured concentration range in the flume, which
due to dilution/sedimentation effects, was from ~7 to b1 vol.% concentration.
D 2005 Elsevier B.V. All rights reserved.
Keywords: Turbidity current; Flume experiments; Horizontal velocity; Root-Mean-Square velocity; Concentration; Grain size
* Corresponding author.
0037-0738/$ - s
doi:10.1016/j.se
E-mail addre
79 (2005) 49–69
ee front matter D 2005 Elsevier B.V. All rights reserved.
dgeo.2005.04.010
ss: [email protected] (C.M.A. Choux).
C.M.A. Choux et al. / Sedimentary Geology 179 (2005) 49–6950
1. Introduction
Particulate gravity currents, both subaerial and
subaqueous, exhibit a wide range of density. Turbidity
currents are an intrinsic part of the spectrum of sub-
aqueous sediment gravity flows. Although they too
encompass a wide range of concentration, for several
decades now, turbidity currents have been consistent-
ly described as sediment-laden gravity-driven flows in
which the sediment is supported principally by fluid
turbulence (Sanders, 1965; Middleton and Hampton,
1973; Middleton, 1993; Simpson, 1997; Shanmugam,
1997). Nevertheless, grain support mechanisms other
than fluid turbulence may co-occur in turbidity cur-
rents, such as hindered settling, particle–particle inter-
actions and buoyancy enhancement (Hiscott, 1994;
Mulder and Alexander, 2001; also see reviews by
Stow et al., 1996 and Kneller and Buckee, 2000).
The contribution of these mechanisms to grain sup-
port is highly dependent on the local concentration of
suspended sediment in the flow, and thus the relative
importance of these mechanisms may change if flows
change their concentration structure as they develop.
In most cases, turbidity currents do indeed evolve in
concentration as they flow, either through sediment
erosion and entrainment (e.g., Pantin, 1979), or
through deposition and entrainment or detrainment
of ambient water (Simpson, 1997 and references
therein). This triggers the question of how flows of
differing initial concentration compare in terms of
internal grain size distribution, concentration and ve-
locity structure as they develop. Do initially dense,
depositional flows, propagating for sufficient time to
become dilute, show the same basic dynamical be-
haviour as initially dilute flows? Are high concentra-
tion turbulent flows viable as a long range transport
mechanism, or are high concentrations only devel-
oped transiently, during sediment entrainment and/or
deposition?
A large volume of experimental work has been
undertaken in order to analyse the role of particle
concentration in turbidity current behaviour and struc-
ture, as well as the geometry and internal structure of
their deposits (e.g., Kuenen, 1966; Middleton, 1967;
Britter and Simpson, 1978; Luthi, 1980; Laval et al.,
1988; Middleton and Neal, 1989; Altinakar et al.,
1990; Bonnecaze et al., 1993; Garcia and Parsons,
1996; Gladstone et al., 1998; Hallworth and Huppert,
1998; Kneller et al., 1999; Stix, 2001; Choux and
Druitt, 2002; McCaffrey et al., 2003; Baas et al.,
2004; Al-Ja’Aidi et al., 2004; see also reviews by
Edwards, 1993; Middleton, 1993; Kneller and
Buckee, 2000, and Shanmugam, 2000). Despite the
value of these experimental works, the results cannot
be used directly to answer questions regarding the
spatio–temporal evolution of natural particulate grav-
ity currents. This requires a detailed characterisation
of both the spatial and temporal evolution of the
properties of experimental particle-driven flows in
terms of velocity, granulometric and concentration
structure, across a range of concentrations. Until re-
cently, however, only temporal (time series) data were
collected (see discussion in Peakall et al., 2001), with
experiments therefore focussing on flow unsteadiness
rather than flow non-uniformity.
Best et al. (2001) used 4-MHz ultrasonic Doppler
velocity profiling (UDVP) to quantify, for the first
time, the spatial and temporal evolution of mean
flow and turbulence structure of sediment-laden par-
ticulate flows. However, concentration and grain size
data were not collected, and the length over which
flow evolution was characterised (up to 85 mm) was
relatively small compared to the scale of the flows.
By coupling the UDVP method with a siphoning
technique and sampling several identical depositing
flows at different locations, McCaffrey et al. (2003)
were the first to produce a detailed description of the
spatio–temporal evolution of a particulate current of
5% initial concentration in terms of instantaneous
velocity, grain size and concentration. Their study
was limited to flows of a single initial starting
concentration.
Computer modellers also endeavour to shed new
lights on describing turbidity current structure (e.g.,
Stacey and Bowen, 1988; Zeng and Lowe, 1997;
Mulder et al., 1998). With the help of a non-depth
averaged model, using a multiphase flow approach,
and particle–particle interaction, incorporating the tur-
bulence model of Mellor and Yamada (1982), Felix
(2001, 2002) produced a 2-D, vertical plane, numer-
ical model that simulates unsteady flow behaviour in
terms of velocity, turbulence, grain size and concen-
tration distribution. However, to date his results have
been tested only against a small number of incomplete
historical flow data, because of the lack of suitable
experimental data (Felix, 2002).
C.M.A. Choux et al. / Sedimentary Geology 179 (2005) 49–69 51
The experimental data of McCaffrey et al. (2003)
can only be used to validate numerical models of
flows in which particle–particle interactions can be
ignored, as the initial particle concentration (5% by
volume) was below the 9 vol.% threshold proposed by
Bagnold (1966) above which these effects become
significant. In the present paper, a new set of experi-
ments is presented, using the same experimental set-
ting as McCaffrey et al. (2003), but for a higher initial
concentration (14%), at which particle–particle inter-
actions should influence flow behaviour, via moderate
turbulence suppression (Middleton and Hampton,
1976; Lowe, 1982). Thus the aim of this work is to
enable the spatio–temporal evolution of the internal
structure of relatively high concentration turbidity
currents (14%, hereafter referred to as high-density
flow) to be compared with that of more dilute flows
(5%, hereafter referred to as low-density flow) in
terms of the vertical gradients in instantaneous hori-
zontal velocity, grain size and concentration distribu-
tion, as well as turbulence structure.
Accordingly, the experimental set up of McCaffrey
et al. (2003) are described, and the key results are
summarised. The flow structure of a turbidity current
with 14% initial sediment concentration is then de-
tailed and compared with the 5% flows results of
McCaffrey et al. (2003). Subsequently, original analy-
sis of the turbulence structure of flows of both 5% and
14% initial concentration (as expressed by root-mean-
square–RMS–velocities) is presented. Dimensionless
parameters are established with the aim of comparing
the low- and high-density flows in more detail and
expanding the results to a wider range of initial sed-
iment concentrations.
2. Previous related work
Based on experimental data, the approach of
McCaffrey et al. (2003) allowed for the first time,
the structure of a turbidity current to be reconstructed
at any position and time, for the parameters stream-
wise velocity, grain size and suspended sediment
concentration. In a series of flume experiments,
McCaffrey et al. (2003) generated subaqueous partic-
ulate gravity flows through release of a 30 l suspension
of non-cohesive material (silica flour) at an initial
concentration of 5% by volume. They measured si-
multaneously the temporal evolution of the vertical
stratification in streamwise velocity, flow concentra-
tion and grain size distribution as the entire flow
passed a measurement location. A series of five nom-
inally identical flows were run, with measurements
repeated at five different locations along the flow
path. The results were then combined to constrain
the spatio–temporal evolution of a single idealised
flow. The inbound jet transformed into a gravity-
driven current at a distance between 1.32 and 2.64
m from the reservoir, and thereafter developed under
its own internal action.
The experimental depositional particulate gravity
currents of McCaffrey et al. (2003) were non-uniform,
i.e., their structure varied spatially (see Allen, 1985
and Kneller and Branney, 1995), indicating that it
would be erroneous to interpret time series data of
such flows in terms of longitudinal flow structure, as
commonly done in the existing literature. For exam-
ple, the velocity data showed that the flow duration
increased downstream, as the flow stretched out. The
concentration data showed that in proximal locations,
the rate of decrease of concentration was high, indi-
cating rapid sedimentation, whereas in distal locations
the rate of decrease was more gradual.
The transition between the head and the body of
each nominally identical turbidity current was de-
scribed by a sharp decrease in the maximum velocity
and median grain size, whereas the transition between
the body and the tail was well defined by a decrease in
the concentration. The velocity maximum was located
at approximately one third of the flow’s height from
its base. An interesting normal to inverse vertical
pattern in grading observed in the grain size distribu-
tion of material suspended in the flow’s body was
linked to the presence of coarse sediment inferred to
have been swept upwards and backwards over the
head then falling passively into the upper part of the
flow.
3. Experimental set up
The experimental set up in this study was the same
as that used by McCaffrey et al. (2003), with the sole
difference that the concentration of the initial suspen-
sion was increased from 5% to 14% by volume (i.e.,
with an initial suspension density of 1231 kg m�3
C.M.A. Choux et al. / Sedimentary Geology 179 (2005) 49–6952
compared to 1082 kg m�3). The flume in which the
five nominally identical flows were run was 10 m
long, 0.3 m wide and 0.3 m deep (Best et al., 2001),
with an overhead reservoir containing 30 l of suspen-
sion (Fig. 1a). A homogeneous suspension in water of
silica flour particles (density: 2650 kg m�3) was
created in the reservoir, and kept well mixed by a
mechanical stirrer. The particle size ranged from
b0.01 to c60 microns, with a median grain size
(D50) of about 8 Am (Fig. 1b). At the start of each
experiment, the stirrer was stopped and a sealing
stopper at the bottom of the reservoir was removed
swiftly. Time series of the low-density experiments of
McCaffrey et al. (2003) taken 0.04 m downstream
from the reservoir outlet (their Flows 1 and 2, illus-
trated in their Fig. 4) showed that the inbound flows
were steady in terms of velocity, grain size and con-
centration for 21.5 s before swiftly decelerating. From
this it may be inferred that the suspension in the
reservoir remained essentially uniform as it drained.
This is probably due to the inherited turbulence from
the mixer, plus any turbulence generated by shear
against the reservoir walls as the suspension flowed
out and into the flume. Although similar outlet time
series were not collected for the high-density experi-
ments reported here, it is inferred that the input to the
flume was essentially steady in this case too. The
suspension drained into the water-filled flume through
a circular pipe of 0.063 m diameter, emptying the
reservoir in about 21.5 s at constant discharge (cf.,
McCaffrey et al., 2003) and forming a particulate
gravity current that propagated along the length of
the flume. The flow was sampled by an array of
instruments all positioned at the same location (Figs.
1 and 2). Due to the intrusiveness of the data acqui-
sition method, each flow could be measured at one
location only. Thus the reservoir was shifted upstream
Fig. 1. a) Experimental set up and
by an interval distance of 1.32 m between each of the
five nominally identical flows, increasing the distance
between the entry point of the flow and the measure-
ment point (Fig. 1a). The height of the flow’s head
was fairly constant at about 0.08–0.085 m, when
reaching the sampling devices, at all locations. Up-
ward flow motion was observed in the head of the
flow as the flow propagated; this upward moving fluid
was then forced back to horizontal by the ambient
water swept over the front of the head (Fig. 2). At
Location 1, the Reynolds number, calculated using
average values within the head, was 2�104, and so
was well within the turbulent regime. Data were ac-
quired from the time of flow inception until upstream-
propagating, solitary waves (e.g., Pantin and Leeder,
1987; Edwards, 1993), generated by the reflection of
the inbound flow from a distal overflow weir (Fig. 1)
passed the array of instruments.
Two sets of instruments were used at each mea-
surement location. A vertical array of six 4-MHz
UDVP (Ultrasonic Doppler Velocity Profiler) probes,
positioned at 6, 16, 26, 36, 46 and 76 mm above the
bottom of the flume, recorded the streamwise compo-
nent of flow velocity upstream of the probes, follow-
ing the technique described by Best et al. (2001). The
sampling rate of each UDVP probe was 4.5 Hz.
Located at the same height (except for 76 mm) and
adjacent to the UDVP probes were 5 siphoning tubes
of 6 mm diameter, which continuously sampled the
flow as it passed by. The suspension samples were
collected continuously in 5 aligned rows of 20 sample
containers, one row for each of the 5 siphoning tubes.
The 100 container array was set up on a sliding
trolley, which was rapidly advanced below the out-
flow ends of the siphon tubes at 4 s intervals, allowing
each successive column of 5 beakers to fill synchro-
nously over ~4 s. The tubes were 1.2 m long, and
b) particle size distribution.
Fig. 2. Frame captured from the video recording of the turbidity current of 14% initial concentration at Location 4, i.e., at 5.28 m from the inlet.
1, 2, 3, 4, 5 refers to siphon and UDVP probe positions, located at respectively 6, 16, 26, 36, 46 mm heights. 6 refers to UDVP probe only,
located at 76 mm height. The arrow points to the turbulent eddy, which generates a rapid deceleration event in the velocity time series. The field
of view is approximately 0.55 m long and 0.35 m high.
C.M.A. Choux et al. / Sedimentary Geology 179 (2005) 49–69 53
allowance was made for the measured transit times
when registering the timing of sample collection with
that of UDVP data collection The content of each
container was later analysed using a Malvern Master-
sizer Plus laser diffraction grain sizer, yielding grain
size distribution and suspended sediment concentra-
tion information at a rate of 0.25 Hz. We will show
below (Section 5.1) that the duration of the head,
delimitated by the horizontal velocity data, is 4 s for
the low-density experiments, i.e. the sampling dura-
tion, and about 2.5 s for the high-density experiments.
Care was taken that as little ambient fluid as possible
was collected prior to arrival of sediment-bearing fluid
at the siphon outlet. This implies that for the high-
density flows, the sampling beaker of the head will
have incorporated some body material for up to 1.5 s.
However, we will show, in Section 6.3, that the di-
mensionless concentration and grain size data in the
head and body are roughly similar; hence we assume
that any effect of dilution from the body material into
the head were insignificant.
A detailed study of the spatio–temporal evolution of
the downstream velocity, grain size distribution and
concentration of the flow was feasible. Moreover,
additional information on the spatio–temporal evolu-
tion of the downstream component of turbulence of the
flow was obtained from the calculation of root-mean-
square (RMS) values of downstream velocity. RMS
velocity is equal to the standard deviation of velocity
averaged over a certain time period (Kneller et al.,
1997; Buckee et al., 2001; Baas and Best, 2002).
Time series of RMS velocity were calculated by aver-
aging the instantaneous velocity data over a 2-s long
period along the whole duration of the velocity time
series and for each measurement height. Tests carried
out to verify the effect of other lengths of averaging
periods on the time series showed no significant differ-
ences. In order to remove the unwanted effect on RMS
velocity of long-term flow deceleration, the velocity
profile in each time window was de-trended using
standard linear regression analysis prior to calculating
the RMS velocities. In analysing velocity signals from
UDVP probes in still water, the long-term average
deviation from the mean was found to be between 2
and 3 mm s�1. These values are therefore considered a
threshold value for the UDVP instrument noise. Areas
of the experimental flows with RMS values below 3
mm s�1 need not be caused by turbulence.
C.M.A. Choux et al. / Sedimentary Geology 179 (2005) 49–6954
4. Description of 14% experimental data
The time series of downstream velocity, grain size,
and concentration are presented for each sampling
location in Fig. 3. This figure also illustrates the
calculated root-mean-square (RMS) values of the
downstream velocity data (Fig. 3b). The reference
time is taken as the time from the removal of the
reservoir stopper. The flow arrival time ranges from
7 s at Location 1 to 35 s at Location 6.
4.1. Downstream horizontal velocity data
The six UDVP probes acquired time series of the
downstream horizontal velocity as the flow travelled
by. The temporal evolution of the streamwise velocity
field for each location is shown in Fig. 3a. A coherent
structure that evolves slightly between the different
locations is observed.
A zone in which the velocity values are consis-
tently high (up to 265 mm s�1) through the flow
depth is recorded for 2–3 s after the arrival of the
flow front (Fig. 3a). Thereafter the velocity drops at
each measurement height. The flow deceleration is
very rapid for the upper two probes (at heights 46
and 76 mm), forming a zone with velocities as low
as 40 mm s�1. The velocity decrease is reduced for
the lowermost probe (6 mm high), with velocities
reduced by a third of the maximum value at that
level. The probes at heights of 16, 26 and 36 mm
reveal a zone of high velocity, with the maximum
velocity occurring close to 20 mm. The height of
this interpolated maximum does not vary temporal-
ly, and varies only slightly spatially (+/�2 mm; Fig.
3a). However, the internal structure of this zone
does evolve with time and distance. At Location
1, the time series exhibits rapidly fluctuating flow
velocity for about 30 s after passage of the head.
These fluctuations become less conspicuous at in-
termediate locations and disappear distally (e.g.,
Location 5 in Fig. 3a). The flow duration, measured
as the time taken for the downstream velocity to
fall below a predefined reference value (cf., McCaf-
frey et al., 2003), progressively increases as the
flow propagates downstream. For example, using a
reference velocity of 100 mm s�1, duration
increases from 28 s at Location 1 to 32 s at
Location 5.
4.2. RMS downstream velocity data
The time series of the RMS of downstream
velocity is shown in Fig. 3b. All graphs exhibit
short-term changes in RMS velocity, which take
the form of small concentric structures in the con-
toured plots presented. The highest RMS values
were found immediately after flow arrival and for
a couple of seconds only, for the highest UDVP
probes. Rapid fluctuations in RMS values are also
found close to the bed at 6 mm. The periodicity of
these fluctuations is ~3 s. The time span between
the passage of successive zones of high RMS ve-
locity remains quasi-constant from Location 1 to
Location 5 whilst the RMS values decrease.
Above the velocity maximum, between 25 and 50
mm, and after the zone of highest RMS values, a
zone of intermediate RMS velocity exists (Fig. 3b).
RMS values are lowest at the level of the maximum
velocity and particularly during the last 10–15 s of
the time series. The most variable RMS time series
is observed at Location 1.
4.3. Grain size data
At all locations along the flow path, median
grain size evolves in a temporally and spatially
consistent way (Fig. 3c). At each location, and
during the first 5 s after the flow arrival, the
grain size data exhibit enrichment in coarse grains
relative to the initial particle distribution. Thereaf-
ter, the vertical grain size profile is characterised
by an upward decrease of the median grain size for
~25 s at Location 1, ~20 s at Location 2 and ~15
s at Location 3. Subsequently, the normal grading
changes into a characteristic vertical pattern of
normal to inverse grading for these locations. The
normal to inverse grading is particularly well-de-
veloped at Locations 4 and 5. At each location, the
zone of grain size reversal moves closer to the
base of the flow with time. A period during
which the flow carries relatively coarse grains is
seen near the base of the flow at about 10 s after
flow arrival at Location 1 and at 18–20 s after
flow arrival at the other locations (Fig. 3c). During
these periods, the flow contains the coarsest sedi-
ment measured, with D50-values of up to 7.7 Am(Location 1).
Fig. 3. Time series of a) downstream velocity (millimeter per second), b) calculated root-mean-square (RMS) values of downstream velocity (millimeter per second), c) median grain
size (micron), and d) concentration (volume percent), at six different flow heights 6, 16, 26, 36, 46 and 76 mm, for five different measurement locations, for the flows with 14% initial
concentration. Note that the 76 mm time series is only collected for downstream velocity. The time, in seconds, is expressed from the removal of the stopper from the bottom of the
reservoir, i.e. the time of inception of the flows. For each graph, the two dashed lines mark the position of the uppermost and lowermost probes, above which and below which no
more data are acquired. The scales and grey shades are the same as in Fig. 4, for easier comparison.
C.M
.A.Chouxet
al./Sedimentary
Geology179(2005)49–69
55
Fig. 4. Time series of a) downstream velocity (millimeter per second), b) calculated root-mean-square (RMS) values of downstream velocity (millimeter per second), c) median grain
size (micron), and d) concentration (volume percent), at six different flow heights 6, 16, 26, 36, 46 and 76 mm, for five different measurement locations, and for the flows with 5%
initial concentration. All data, except RMS velocities, were presented in McCaffrey et al. (2003), but have been redrawn at the same scale and grey shades as in Fig. 3. The time, in
seconds, is expressed from the removal of the stopper from the bottom of the reservoir, i.e. the time of inception of the flows. For each graph, the two dashed lines mark the position of
the upper and lower probes, above which and below which data extrapolation is likely to be biased.
C.M
.A.Chouxet
al./Sedimentary
Geology179(2005)49–69
56
C.M.A. Choux et al. / Sedimentary Geology 179 (2005) 49–69 57
4.4. Concentration data
The time series of sediment concentration as a
function of flow height are given in Fig. 3d. Concen-
tration consistently decreases upwards. At each mea-
surement location, the concentration values are
initially very low (down to 2–3 vol.%) compared
with the initial concentration of 14 vol.% within the
reservoir. The maximum concentration, located at the
base of the flows, occurs at each location some 15–20
s after passage of the flow front. The maximum
concentration at Location 1, which is closest to the
inlet, equals 6.4%. At other locations, the maximum
measured concentration ranges from 5.9% at Location
2 to 7% at Location 5, i.e., furthest away from the
inlet. The interpolated near-bed concentrations values
are 6.4% at Location 2 and 7.9% at Location 5. The
maximum heights reached by concentration contours
V4.5% gradually decrease in a downstream direction.
5. Interpretation of high-density flow data and
comparison with low-density flow data
In this section, the data of the high-density experi-
ments are interpreted and compared with the low-
density experiments of McCaffrey et al. (2003). The
spatio–temporal graphs of the low-density experi-
ments are reproduced in Fig. 4 to facilitate the com-
parison. Fig. 4 also includes time series for RMS
values of downstream velocity, which have not been
published before. Due to the fact that the low-density
flows were slower than the high-density flows, the
low-density flows arrived later at each measurement
location than the high-density flows, explaining the
different initial times on the graphs in Figs. 3 and 4.
Particulate gravity currents are commonly divided
into three flow regions: head, body and tail (see
review by Kneller and Buckee, 2000). The head,
with its overhanging nose due to no-slip condition at
the lower boundary and frictional resistance at the
upper boundary, is the area where mixing of the
current with ambient fluid occurs, essentially by
detraining of dense fluid out of the back of the head
in a series of transverse vortices (Allen, 1971; Britter
and Simpson, 1978; Simpson and Britter, 1979). The
body is the area which has a thin, relatively dense
layer of fluid near the base of the current, and which is
overlain by a mixing zone at its upper boundary
displaying succession of large eddies (Ellison and
Turner, 1959; Middleton, 1966). The tail is the termi-
nal part of the flow, where velocity is low and grad-
ually decreases to zero; here slow settling from
suspension is the dominant depositional process.
The same subdivision is applied below, because it
was possible to confidently delimit head, body and
tail by trends in the experimental data (see also
McCaffrey et al., 2003).
5.1. Downstream velocity
The head of the flow with 14% initial concentra-
tion is delimited by the rapid increase in velocity at
the flow front (Fig. 3a) and the midpoint of a slightly
longer period of relatively strong flow deceleration
present at heights of 46 and 76 mm in all locations.
The head passes the measurement locations in about
2–3 s. The strong deceleration may relate to the
presence of an eddy at the back of the head; analysis
of video recordings of the experiments confirms the
existence of such a structure, located at the back of,
and defining the extension of the head (Fig. 4). A
more gradual flow deceleration event at 20–25 s,
prominent in particular at proximal locations, defines
the transition between the body and tail of the flow. It
corresponds to the beginning of the waning tail of the
flow, probably because the reservoir empties and thus
no longer supplies the flow (the reservoir was seen to
empty at ~21.5 s after the start of the experiments).
The velocity structure of the high-density flow
evolves in time and with distance along the flume.
As a result of the absence of a bed slope and progres-
sive deposition of sediment, flow velocity gradually
decreases at all levels in the flow and at all locations
along the flume. Also, the short-term fluctuations in
flow velocity, which were particularly clear at Loca-
tion 1, gradually disappear. The increasing duration of
the flows away from the inlet is primarily caused by
extension of the tail away from the inlet, because the
duration of the passage of the head and body is almost
constant at the studied locations.
In general terms, the high- and low-density flows
have similar time-dependent velocity structure. In
detail, however, there are important differences. The
dense flow is thinner and travels faster than the dilute
flow, with a shorter head duration (about 2.5 s instead
C.M.A. Choux et al. / Sedimentary Geology 179 (2005) 49–6958
of 4 s for the low-density flows). The height of
maximum velocity is lower for the high-density
flow (c20 mm) than for the low-density flow
(c25 mm). However, the velocity maximum is
located roughly at about 0.3 of the height of the
head in each case, which is in agreement with
previous experiments (Altinakar et al., 1996; Kneller
et al., 1997, 1999; Best et al., 2001).
5.2. RMS downstream velocity
The time series of RMS downstream velocity,
shown in Figs. 3b and 4b, indicates that, overall, the
flows of 14% initial concentration have higher RMS
values than the flows of 5% initial concentration,
although the distribution of RMS values is similar.
Below, the RMS velocity structure of the low-density
flows is described first, then the RMS velocities of the
low- and high-density flows are compared.
All the graphs of RMS velocity for the low-density
flows exhibit numerous rapid fluctuations in the head
of the flows as well as close to their base (Fig. 4b).
The highest RMS values are obtained from the upper
front of the head at Location 2. Locations 3–5 have
their maximum RMS velocities at the same position,
but absolute values decrease downstream. Location 1
is characterised by strong fluctuations in RMS veloc-
ity in the entire head and body. The rear part of the
flows and the upper part of their bodies are charac-
terised by low RMS velocities (Fig. 4b). Particularly
striking are the regular fluctuations in RMS values
near the base of the flow. As for the high-density
flows, a zone of relatively low RMS values exists at
the level of the velocity maximum, above which
intermediate fluctuating RMS values are observed.
In both the low- and high-density flows, the zone
of maximum RMS velocity within the upper part of
the head is interpreted to result from high turbulence
levels linked to the friction between the propagating
flow and the ambient fluid, leading to the formation of
Kelvin–Helmholtz waves and mixing at the back of
the head (Best et al., 2001). The concentric RMS
structures near the base of the flows, whose periodic-
ity slightly increases with time, are interpreted as
coherent flow structures (Baas et al., in press),
corresponding to turbulent eddies generated by fric-
tion at the lower flow boundary. The relatively low
RMS velocities at the level of maximum velocity,
supporting previous measurement by Kneller et al.
(1999), Best et al. (2001) and Buckee et al. (2001)
as well as in the tail of the flows indicate that these
areas are less turbulent than other areas. It thus
appears that the propagation distance of turbulent
eddies generated by shear at the lower, upper and
frontal flow boundaries is relatively small (cf. Felix
et al., 2005). The degree of turbulence decreases
downstream along the flume in both flows, which
correlates with decreasing downstream velocity and
thus decreasing shear.
5.3. Median grain size
The median grain size structures of the high-den-
sity flows (Fig. 3c) and the low-density flows (Fig. 4c)
evolve in a similar way. Yet, the basal zone of maxi-
mum grain size is less well developed in the low-
density flows. At each location, the grain size helps to
define the transition between the head and the body of
the flow. A drop in median grain size by up to 1 Ammarks this transition.
The inverse grading in the body of the flows is
interpreted to result from the movement of coarse
sediment from the upper part of the head, enriched
in coarse particles (cf., Section 4.3), upwards and
backwards by turbulent motion (McCaffrey et al.,
2003). During the backward motion, the coarse sedi-
ment is probably located above the measurement area,
i.e. above the highest sampling tube located at 46 mm.
They then fall passively back into the body and tail of
the flow. Video data (Fig. 2) and velocity (Figs. 3a and
4a) support the interpretation that an eddy is present at
the back of the head, which may be responsible for
this redistribution of coarse sediment towards the rear
of the flows.
A major difference between the low- and high-
density flows is that a basal zone enriched in coarse
grains is present at 18–20 s after the passage of the
head at Locations 2–5 in the high-density experiments
(Fig. 3c). This phenomenon was observed only at
Locations 3 and 5 in the low-density experiments.
Their position below the velocity maximum and far
behind the head classifies these coarse-grained zones
as coarse tail lags sensu Hand (1997). The fastest
settling grains are thus concentrated towards the
base of the flow, while slower settling grains are
distributed more evenly throughout the flow depth
Fig. 5. Spatio–temporal evolution of a single idealised flow, created by using the data acquired at the five measurement locations, for a) downstream horizontal velocity (millimeter
per second), b) RMS velocity (millimeter per second), c) median grain size (micron), and d) concentration (volume percent), at 21.5, 29, 36.5, 44 and 53.5 s after inception of the
flows. The time, in seconds, is expressed from the removal of the stopper from the bottom of the reservoir, i.e. the time of inception of the flows. The grey dashed lines show the
position of the uppermost and lowermost probes respectively above which and below which no more data are acquired. A small contouring artefact is noticeable before the flow
arrival, mainly visible for the downstream horizontal velocity and RMS velocity graphs, at 21.5 and 29 s.
C.M
.A.Chouxet
al./Sedimentary
Geology179(2005)49–69
59
C.M.A. Choux et al. / Sedimentary Geology 179 (2005) 49–6960
(Middleton and Southard, 1984; Hand, 1997 and
references therein).
5.4. Concentration
The initial sediment concentration in the reservoir
was 14%. The fact that a maximum concentration of
6.4% was observed at Location 1 (Fig. 3d) indicates
that strong flow dilution occurred due to flow expan-
sion and entrainment of ambient water and/or sedi-
mentation. In the flows of 5% initial concentration,
McCaffrey et al. (2003) also noticed an abrupt change
in concentration between Locations 1 and 2 (Fig. 4d).
They explained the reduction of nearly 50% by in-
voking high rates of sedimentation from suspension
between those locations. In the high-density experi-
ments (Fig. 3d), no such drastic change is observed.
This point will be discussed in more detail in the
section on dimensionless analysis below. The progres-
sive spatio–temporal decrease in the height of con-
centration contours, from Locations 1 to 5, attests to
ongoing sedimentation and dilution as the flow pro-
pagates, although these processes are less marked than
for the low-density flows. At Location 5, the maxi-
mum measured concentration of 7% by volume, seen
at the base of the flow at around 55 s, is higher than
the corresponding maximum seen in Location 4,
immediately upstream, at around 45 s, which is
equal to 6% by volume, interrupting the overall pat-
tern of downstream-decreasing concentration. A pos-
sible explanation is that the flow undergoes a slight
increase in the rate of fallout of sediment from sus-
pension between Locations 4 and 5 due to decreasing
velocity and turbulence intensity.
5.5. Spatial flow evolution
A series of instantaneous snapshots of the high-
density flow was constructed at five selected times
(i.e., 21.5, 29, 36.5, 44 and 53.5 s), for downstream
velocity (Fig. 5a), RMS velocity (Fig. 5b), median
grain size (Fig. 5c) and concentration (Fig. 5d). This
was done by extracting the respective data for each
point in time and for each measurement location from
the time series, and then plotting the data as a function
of distance along the flume for each point in time.
These spatial plots permit the visualisation of the
internal structure of a single flow and its temporal
evolution. They would also allow for a direct com-
parison with numerical modelling results, as sug-
gested by Felix (2002).
The snapshots of downstream velocity (Fig. 5a)
show maximum values at the front of the head and
progressively slower flow with increasing distance
behind it. The decrease in velocity in the tail part is
particularly evident around the height of maximum
velocity. The corresponding snapshots of RMS veloc-
ity data (Fig. 5b) do not show a steady evolution.
Generally, the head of the flow has the highest RMS
values, and therefore is the most turbulent part of the
flow. In the body, RMS velocities are clearly less than
in the head, except for the basal part of the flow,
where turbulence remains strong even at large dis-
tances behind the flow front. The spatial plots of
median grain size (Fig. 5c) show that the zone of
minimum grain size becomes more pronounced as
the flow evolves temporally and that coarse tail lag-
ging occurs behind the head of the flow (e.g., at 44s).
The spatial plots of concentration reveal a higher rate
of sedimentation as time goes by. Indeed at 29 s, the
height between the 2% and 6% concentration contours
is about 40 mm whereas it is only 20 mm at 53.5 s. A
zone of maximum concentration is visible close to the
base of the flow (Fig. 5d). It moves progressively
down the flume as the flow evolves. The zone is
centred at ~1 m after 29 s, at 2.5 m after 36.5 s, at
6.5 m after 44 s and beyond 7 m after 53.5 s.
The temporal evolution of the internal structure of
the high-density flow seen in Fig. 5 for all measured
flow and sediment parameters, indicates that the flow
is non-uniform. Decreasing velocities and overall con-
centrations indicate that this is probably caused by the
flow’s depositional character. This interpretation rein-
forces the conclusions drawn by McCaffrey et al.
(2003) and extends them to flows of higher concen-
tration. The fact that the low- and high-density flows
are non-uniform implies that it is impossible to deduce
the structure of flows from studies of time series data
obtained at only one location, at least over the concen-
tration range encompassed by the flows as they evolve.
6. Dimensionless analysis
Dimensionless analysis was carried out in order to
compare the flow structure of the low- and high-
C.M.A. Choux et al. / Sedimentary Geology 179 (2005) 49–69 61
density flows in detail, and to investigate if data
collapse could be achieved. Here, the flows are com-
pared by means of dimensionless vertical profiles of
normalised downstream velocity (Fig. 6), RMS veloc-
ity (Fig. 7), median grain size (Fig. 8) and concentra-
tion (Fig. 9). A careful choice of normalisation
parameters is essential in order to ensure that flow
and sediment parameters are compared in analogous
zones of the flows (cf., Felix, 2004). Therefore, the
dimensionless parameters used in this study are de-
fined first. Thereafter, the location of vertical profiles
in the head, body and tail of the turbidity currents are
selected. The dimensionless profiles of flow and sedi-
ment parameters are presented in the last part of this
section.
6.1. Normalisation parameters
Selection of reference heights, velocities, median
grain sizes and concentrations was necessary in order
Fig. 6. Dimensionless flow velocity as a function of dimensionless height a
low- and high-density (5% and 14% initial concentration, respectively) tu
to normalise and compare the data from the two sets
of experiments. The height of the maximum down-
stream velocity was used as reference for the calcula-
tion of normalised height (cf., Altinakar et al., 1996;
Kneller et al., 1999). This reference height, which was
shown to be independent of measurement location, is
~25 mm for the low-density flows (Fig. 4a) and ~20
mm for the high-density flows (Fig. 3a). Other refer-
ence heights, such as the height at which the down-
stream velocity in the upper part of the flow is half the
maximum velocity (e.g., Kneller and Buckee, 2000),
could not be reliably applied because there were too
few sampling heights above the velocity maximum to
accurately determine them.
In the vertical profiles for the head (Fig. 6a) and
body (Fig. 6b), dimensionless downstream flow ve-
locity was defined as the ratio between average ve-
locity over a time span of 1.25 s (flow of 14% initial
concentration) or 2 s (flow of 5% initial concentra-
tion) and average head velocity for all measurement
nd measurement location for the head (a), body (b) and tail (c) of the
rbidity currents.
Fig. 7. Dimensionless RMS velocity as a function of dimensionless height and measurement location for the head (a), body (b) and tail (c) of the
low- and high-density (5% and 14% initial concentration, respectively) turbidity currents.
C.M.A. Choux et al. / Sedimentary Geology 179 (2005) 49–6962
locations. The time spans of 1.25 and 2 s were chosen
to obtain velocity values over approximately equal
lengths in both flows. Average downstream head ve-
locities, at the height of the velocity maximum, were
133.5 and 189 mm s�1 for the low- and high-density
flows, respectively. The use of a single head velocity
for each flow concentration is warranted, because
head velocity changes between the most proximal
and most distal measurement locations were insignifi-
cant. The calculation of the dimensionless flow veloc-
ity for the tail zone (Fig. 6c) was carried out using the
maximum velocity found at the body–tail transition
because the head and tail are sufficiently far apart that
processes in the head may not be relevant to tail
dynamics.
The dimensionless RMS velocity for the head (Fig.
7a) and body (Fig. 7b) was calculated by normalising
the RMS velocity values to the average head velocity.
For the tail (Fig. 7c), the same method was used as for
the normalised downstream velocity, i.e., the maxi-
mum downstream value measured at the body–tail
transition was selected.
Median grain size (Fig. 8) was normalised to the
initial median grain size (8 Am) in the overhead
reservoir for the low- and high-density flows. Dimen-
sionless concentrations (Fig. 9) were calculated by
dividing the local concentration values by the initial
concentration in the overhead reservoir.
6.2. Division of flows into head, body and tail
segments
Vertical profiles of flow parameters were outlined
through predefined segments in the experimental
flows. The underlying methodology relies on the de-
termination of equivalent zones in the flows in each of
the two sets of experiments, and subsequent selection
of equivalent locations within these zones where ver-
tical profiles are to be compared. Below, the zones are
defined in terms of the flows’ head, body and tail.
Fig. 8. Dimensionless median grain size as a function of dimensionless height and measurement location for the head (a), body (b) and tail (c) of
the low- and high-density (5% and 14% initial concentration, respectively) turbidity currents.
C.M.A. Choux et al. / Sedimentary Geology 179 (2005) 49–69 63
6.2.1. Head length
In the time series, the passage of the head is
defined as the period immediately following the ar-
rival of the flow, during which the velocities are high
along the entire vertical profile (Fig. 3a). The up-
stream boundary of the head, i.e. the transition be-
tween head and body, is defined by the rapid
decelerating event in the velocity time series, recorded
by the probes at 46 and 76 mm height. The average
period in which the head passed a measurement loca-
tion was 4 s in the low-density flow and 2.5 s in the
high-density flow. In length, this corresponds to about
0.38F0.005 m for both flows, using the average head
velocity as reference. It was then decided arbitrarily to
select a relative distance of 25% of the head length
behind the front of the head to locate the vertical
profiles. Thus, at 0.095 m from the front of the
head, the original data located along a vertical profile
were chosen for normalisation. Because of the large
fluctuations of the downstream and RMS velocity,
instead of presenting an isolated profile from this
location, average velocity values were calculated for
all the values between 0 and 0.19 m (giving the
average velocity value at 0.095 m).
6.2.2. Body length
The body stretches from the upstream limit of the
head to the point marked by a sudden change from
high to low velocity (at proximal locations, Fig. 3a)
and a change from relatively high to low RMS veloc-
ity (predominantly at distal locations, Fig. 3b). This is
interpreted to represent the time when the overhead
reservoir emptied. The reservoir emptied in 22 s for
the low-density flow and 21.5 s for the high-density
flow. Allowing for the head duration, the duration of
the passage of the flow bodies in the low- and high-
density flows was thus 18 and 19 s, respectively, with
corresponding respective body lengths of 1.73 and
Fig. 9. Dimensionless concentration as a function of dimensionless height and measurement location for the head (a), body (b) and tail (c) of the
low- and high-density (5% and 14% initial concentration, respectively) turbidity currents.
C.M.A. Choux et al. / Sedimentary Geology 179 (2005) 49–6964
2.85 m (once again using average head velocity as the
reference velocity). The vertical profiles for dimen-
sionless downstream velocity and its RMS values,
median grain size and concentration were drawn at
an arbitrary relative distance of 30% of the length of
the body from its front, hence at 9.4 s and 0.90 m for
the low-density flows, and 8.2 s and 1.23 m for the
high-density flows. Here, 9.4 and 8.2 s refer to the
time since the arrival of the flow.
6.2.3. Tail length
The tail is the most distal part of the flows bounded
by the body at one end. In the experiments, the tail
was disrupted by the arrival of flow reflections before
it had come to rest. Because the tail is the section
where most of the flow stretching takes place, a
different method is required to select an equivalent
location for the vertical profiles. First, a power func-
tion was fitted to the velocity time series at a height of
2.6 cm (i.e., close to the level of maximum velocity)
in the tail of each flow. Subsequently, the best fit
power function was used to calculate the time period
from the time of first arrival of the tail to the time at
which velocity reached 10 mm s�1. Finally, the time
for the tail velocity to decrease by an arbitrary 40% of
the range between its value at the body–tail boundary
and the 10 mm s�1 boundary was calculated for each
measurement location. At these times vertical profiles
for downstream flow velocity and RMS velocity,
concentration and median grain size were determined.
As mentioned above, dimensionless velocities were
calculated by dividing the values at the 40% boundary
by the velocity at the body–tail boundary rather than
by the average head velocity, because the head and tail
are so far apart that head processes should not affect
the tail.
6.3. Dimensionless vertical profiles
The normalised downstream velocity (Fig. 6),
RMS velocity (Fig. 7), median grain size (Fig. 8)
and concentration (Fig. 9) are plotted versus the di-
mensionless height for each of the three zones of the
flow (i.e., body, head and tail).
C.M.A. Choux et al. / Sedimentary Geology 179 (2005) 49–69 65
6.3.1. Downstream velocity profiles
For the head and body (Fig. 6a and b), normalised
downstream flow velocities greater than 1 indicate
flow towards the front of the turbidity current, while
normalised velocities smaller than 1 signify flow that
moves away from the head for an observer moving
with the flow (i.e., within a Lagrangian reference
frame). For the tail (Fig. 6c), velocities are relative
to that of the body–tail transition. The normalised
downstream velocities for the low- and high-density
flows at Locations 2–5 are similar for the head (Fig.
6a), body (Fig. 6b) and tail (Fig. 6c), hence the data
collapse in a satisfactory manner. At Location 1,
however, the normalised velocity of the high-density
flow is significantly higher than that of the low-den-
sity flow (Fig. 6a), particularly within the head. Be-
tween Locations 1 and 2 the dimensionless values for
the low-density experiment increase drastically, thus
supporting the inferred occurrence of an episode of
high sedimentation rate and change from jet to tur-
bidity current (McCaffrey et al., 2003) between these
locations. In contrast, the dimensionless velocity
remains quasi-constant between Locations 1 and 5
in the high-density flow. This suggests that no regime
change occurred along this transect. The flow may
therefore have developed into a turbidity current by
the time that it reached Location 1. In turn, this
implies that any episode of high sedimentation rate
must have occurred upstream of Location 1, and thus
within 1.32 m of the inlet.
6.3.2. RMS downstream velocity
The vertical profiles of dimensionless RMS ve-
locity versus height (Fig. 7) in the head region of
the flows (Fig. 7a) are irregular, with little similarity
between the low- and high-density experiments.
However, a broad trend with higher RMS velocity
values at the top of the flow compared with the rest
of the profile and a slight increase of RMS close to
the base of the flow, exists. In the body (Fig. 7b),
normalised RMS values are more uniform than in
the head, and maximum RMS velocities are almost
exclusively found at the base of the flows. A weak
zone of relatively low RMS velocities is discernable
at or around the height of the velocity maximum,
particularly at distal locations. The RMS velocities
in the tail of the low- and high-density flows col-
lapse well (Fig. 7c), displaying a quasi-uniform
pattern of RMS velocities along the entire flow
depth, except for a slight increase at the lowest
data point.
At Location 1, the vertical profile of the flow of 14
initial concentration is broadly concave to the right (as
are the profiles at all other Locations), whereas the
profile of the flow of 55 initial concentration is con-
cave to the left. This difference is interpreted to arise
because the high-density flow has undergone the jet to
turbidity current transition at this Location, whereas
the low-density flow still has elements of jet structure.
The higher dimensionless RMS values observed near
the base of the body and tail profiles represent the
turbulent eddies generated by friction with the base of
the flume. This pattern is also seen close to the base of
the flow for the head, yet the basal non-dimensional
RMS values are less than those in the upper part of the
flow, confirming that the head is more turbulent at its
top than at its base. The profiles also show the general
loss of turbulence as the flow propagates. The profiles
become quasi-vertical straight lines with RMS values
close to zero in the tail areas, which indicate that the
flow approaches a laminar regime over its entire
height.
6.3.3. Median grain size profiles
The vertical profiles of median grain size (Fig. 8)
essentially redisplay the data in Figs. 3c and 4c, but
now allow a more direct comparison between the low-
and high-density flows. The vertical profiles of the
low- and high-density flows generally exhibit a simi-
lar trend for the head (Fig. 8a), body (Fig. 8b) and tail
(Fig. 8c). At most locations, the sediment is relatively
coarse near the base of the flow. In the central part of
the flows, the sediment is relatively fine, while its size
increases slightly in the uppermost part of most flows.
The height of minimum grain size decreases from
body to tail and from Locations 2 to 5, the grain
size minimum being more prominent distally than
proximally. Location 1 once again differs from the
other locations in that the normal-to-inverse grading
in the low- and high-density flows cover different
dimensionless size ranges, and normally graded pro-
files (high-density flow) versus weakly graded to non-
graded profiles (low-density flow) prevail in the body
and tail. Once again, these discrepancies suggest that
the flows may not have developed to the same state at
Location 1.
C.M.A. Choux et al. / Sedimentary Geology 179 (2005) 49–6966
6.3.4. Concentration profiles
The vertical profiles of dimensionless concentra-
tion are presented in Fig. 9. The vertical profiles show
a good collapse of concentration data, especially in
the body (Fig. 9b) and tail (Fig. 9c) of the flows. The
exception is again Location 1, where dimensionless
concentration for the low-density flow is consistently
higher than for the high-density flow. In all the graphs
of Fig. 9, the highest normalised concentration is
observed at the base of the flow.
Fig. 9 reinforces the above-mentioned conclusion
that the low- and high-density flows had not evolved
to the same state when they reached Location 1. The
flows show similar nondimensional behaviour from
Locations 2 to 5. The high-density flow is compara-
tively more dilute than the low-density flow when it
reaches Location 1, having lost most of its particulate
load by sedimentation while the flow was sedimenting
rapidly. We infer that the high-density flow underwent
a period of strong sedimentation and change from
inbound jet to turbidity current, before reaching the
most proximal measurement location, at which point
the low-density flow was still propagating with a
relatively high concentration.
7. Discussion and conclusions
Detailed experiments with flows of 14% initial
concentration were undertaken to gain insights into
the spatio–temporal development of flow structure in
terms of horizontal velocity, RMS velocity, concen-
tration and grain size, and to make a comparison
with the evolution of flows of 5% initial concentra-
tion, as described by McCaffrey et al. (2003). The
high-density flows were shown to be non-uniform.
The current duration extends via the stretching of the
tail only, so that the length and duration of the head
and body remain constant throughout the flow length
studied. Behind the head of the flow and below the
maximum velocity zone, a coarse tail lag in sus-
pended sediment has been observed at all locations.
These experimental data therefore strongly support
the model of Hand (1997), in which coarser grains
lag behind the head of the current by virtue of being
concentrated at a level below that of the velocity
maximum, therefore advecting forward more slowly
than the relatively finer-grained material carried
higher up. Although the flows of 14% initial con-
centration are non-uniform, the dimensionless analy-
sis shows a generally good collapse of the data from
the low- and high-density flows for the different
parameters studied. It follows that the fundamental
flow structure is essentially the same for flows of
either initial starting concentration at any stage in
their development after the jet to turbidity current
transition. The implications of this conclusion are
considered below.
The common evolution of the flows of 5% and
14% initial concentration might suggest that the par-
tial turbulence suppression predicted to develop at
concentrations above 9% (Bagnold, 1962; Shanmu-
gam, 2000; Gani, 2004) does not appear to influence
the development of the high-density flow. It should be
borne in mind, however, that flows of both initial
concentrations were considerably more dilute at Lo-
cation 1, i.e. 1.32 m from the flow inlet, than when
generated. Such a dilution could be produced by
turbulent entrainment of water and/or by sedimenta-
tion. McCaffrey et al. (2003) sampled two flows (their
Flows 1 and 2) with one siphon tube/UDVP probe
pair located 0.04 m from the outlet of the reservoir,
and found that concentration had fallen from 5% to
2.5% over this short distance. In the absence of a
significant deposit at this location, it was concluded
that rapid water entrainment had occurred. Thus, even
in the moderate Reynolds regime of the inbound jet
flow, high concentrations could only be maintained
transiently. This might imply that high-concentration
turbidity currents may not be viable as a long-range
transport mechanism. However, entrainment-related
dilution might also be an artefact of the flow genera-
tion mechanism. In these experiments the flows en-
tered the flume as a jet, necessarily incorporating
ambient water. Improved experimental design, or gen-
erating flows with yet higher initial concentrations,
might overcome this problem. Thus, the viability of
high density turbidity currents as a long-range sedi-
ment transport mechanism remains open. It seems
likely that no partial turbulence suppression was no-
ticeable because the turbidity currents produced did
not exceed the 9% concentration threshold of Bagnold
(1962). It follows that any dimensional analysis re-
garding flow evolution should not be applied to con-
centrations higher than those observed in these flows,
i.e. about 7%.
C.M.A. Choux et al. / Sedimentary Geology 179 (2005) 49–69 67
Even if the initial decrease of concentration is
dilution-related, it seems likely that sedimentation-
related concentration decreases may also have played
a role by Location 1. Certainly, such changes can be
inferred to have occurred between Locations 1 and 2.
Thus both low- and high-density flows may have
undergone deposition-related decreases in concentra-
tion, with the low-density flows reaching a lower
concentration than the high-density flows. This sug-
gests that the flows may have an upper concentration
threshold or capacity, indirectly controlled by the
inherited concentration, above which they cannot sus-
pend and transport the whole load of sediment. This is
presumably linked via the gravitational driving force
to its effect upon turbulence generation.
Finally, it is evident that after the jet to turbidity
current transition, the two data sets can be combined
to illustrate the path of an idealised flow of 14% initial
concentration (~7% concentration in the flume) until it
has developed the structure of a flow of 5% initial
concentration measured at the most distal location.
Such a flow would therefore develop over a length
scale greater than that provided by the experimental
facility. In addition, for new flows generated in similar
initial starting conditions but with initial concentra-
tions less than or equal to 14% by volume, dimen-
sionless concentrations could be used to calculate
local concentrations at various positions and times.
For these new flows, and as a first approximation,
head velocities could be computed by linear interpo-
lation, and then used (via the dimensionless values) to
estimate downstream velocities, concentrations and
grain size distributions of suspended material at var-
ious spatio–temporal positions.
Acknowledgements
We are grateful to T. Sakai, P. Julien and M. Felix
whose thorough reviews greatly improved and clari-
fied earlier versions of this manuscript. This research
was funded in part by the Turbidites Research Group
(TGR) Phase 4 programme, sponsored by BG, BHP-
Billiton, BP, ConocoPhillips, Shell and Norsk Hydro.
Caroline Choux acknowledges funding from Marie
Curie Individual Fellowship grant HPMF-CT-2001-
01298. The UDVP system was purchased under Natu-
ral Environment Research Council (NERC) Grant
GR3/10015 to Jim Best and colleagues. The overhead
reservoir flow generation mechanism and bslidingtrolleyQ siphon sample collection equipment were ini-
tially developed by J. Peakall. Complete data sets of
the experimental flows are available from the follow-
ing web site: http://trg.earth.leeds.ac.uk/.
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www.elsevier.com/locate/sedgeo
Sedimentary Geology 1
Morphology, evolution and fill: Implications for sand and
mud distribution in filling deep-water canyons and slope
channel complexes
Bryan T. Cronin a,*, Andrey M. Akhmetzhanov b, Adriano Mazzini a,
Grigorii Akhmanov c, Michael Ivanov c, Neil H. Kenyon b
TTR-10 Shipboard Scientists
aUniversity of Aberdeen, Department of Geology and Petroleum Geology, King’s College, Aberdeen AB24 2UE, United KingdombChallenger Division, Southampton Oceanography Centre, European Way, Southampton, SO14 3ZH, England, United Kingdom
cUNESCO Centre for Marine Geology and Geophysics, Science Park, Moscow State University, Russia
Received 21 April 2004; accepted 6 April 2005
Abstract
A survey of the northeastern margin of the Rockall Trough on the Irish margin examined the transition from shelf edge to
basin floor, in morphology and sedimentary activity, of a deeply incised submarine canyon system, the Donegal Bay submarine
canyon. The survey produced superb 3D profiling of the canyon along its entire length, marking a transition from dcauliflowerTshaped head region with numerous tributary gullies feeding into one main canyon, to a single trunk canyon. This canyon, with
an initial combined width and depth of N17 km and N800 m in the dcauliflowerT head area, decreases rapidly to N4.5 km wide
and N450 m deep after the zone of tributary confluence. Eighteen kilometers further down dip, the canyon loses topographic
expression as it approaches the lower rise and floor of the Rockall Trough.
Degrees of recent sedimentary activity are evaluated by comparing side scan sonar systems of different frequency, and thus
of different penetration sub sea, and by ground-truthing using drop (gravity) cores. The canyon was a very active system,
dominated by sand transportation towards the floor of the Rockall Trough, along the slope as coarse-grained contourite, or as
sand spillover from the shelf. Sand was also deposited as overbank deposits outside the main head region of the canyon,
presumably by large volume turbidity currents and more active lateral gullies. The head area of the canyon system has been
progressively cut off from sand source by progressive sea level rise since the last glaciation. Sand was locally deposited on
terraces but not in the overbank area. Less frequent, lower volume and finer grained turbidity currents have become more
common in the system. The initial sand and bypass-dominated system with small sediment waves, which may be gravels, has
become dominated by muddy debrites in the lower reaches and by slumps in the upper reaches. Slumping in those upper reaches
leads to ponding of sand in the head and upper reach areas, with only very occasional turbidity currents transporting sand
further down the system in small channels.
0037-0738/$ - s
doi:10.1016/j.se
* Correspondi
E-mail addre
79 (2005) 71–97
ee front matter D 2005 Published by Elsevier B.V.
dgeo.2005.04.013
ng author.
ss: [email protected] (B.T. Cronin).
B.T. Cronin et al. / Sedimentary Geology 179 (2005) 71–9772
A model is produced to explain the mechanism and expression of backfilling in a large deep-water canyon system whose
hinterland has been flooded back since glacial drawdown of sea level in the eastern Rockall Trough area. This model explains
how sand may be trapped in large volumes in the upper reaches of a canyon system, due to slumping from the canyon margins
and nearby upper slope regions. The focusing of sand deposition in areas where this is not usually expected will have important
implications for hydrocarbon explorationists who wish to map the distribution of potential reservoir sand bodies within large,
confined deep-water canyon systems.
D 2005 Published by Elsevier B.V.
Keywords: Deep-water canyon; Rockall margin; Debrites; Backfilling; Slumps; Sand ponding
1. Introduction
1.1. Canyon morphology
There are few documented examples of canyons
from their head regions down to where they usually
lose topographic expression on the lower slope,
though some recent papers illustrate seafloor render-
ings from 3D seismic datasets (e.g. Barrufini et al.,
2000). Industrial seismic datasets are rarely collected
on lower slope areas where working petroleum sys-
tems are assumed not to work due to the lack of burial
of potential source rocks, so most of these renderings,
though usually spectacular, do not extend the full
length of the canyon system. There are some exam-
ples of sidescan sonar datasets where canyons have
been followed for long distances (Cronin et al., 1995).
For the moment, we still rely on conventional models
for deep-water canyons (e.g. Shepard, 1977; Scruton
and Talwani, 1982).
As a general rule, canyons typically have sinu-
ous courses with straight sections; floors which
deepen seawards; a V-shaped profile which is lost
as the adjacent continental slope reduces in gradi-
ent; tributaries, which are normally gullies in the
canyon head region; and currents which move up
and down their axes even at very great depth,
typically with tidal periodicities, or episodic turbid-
ity currents. They are normally associated with
large rivers (e.g. Amazon Canyon: Heezen and
Tharp, 1961; Damuth et al., 1983); rare on gentle
slopes, and densely spaced on steeper slopes; ero-
sive into any substrate; and older features, where
the canyon head is more deeply recessed into the
slope, or indeed, back across the shelf, indicating
that canyons develop shorewards by headward ero-
sion (Shepard and Marshall, 1978). Younger can-
yons are typically incised several hundred meters
below the shelf break. The morphologies of can-
yons are usually divided into four geomorphological
categories (Cronin, 1994):
(i) Canyon heads: Canyon head regions have been
described as amphitheatre-shaped areas with
steep rims (Belderson and Stride, 1969; Kenyon
et al., 1978). The oldest canyon on any one
slope is usually the canyon whose head is high-
est on the continental slope. Though tidal peri-
odicities are recorded in the lower reaches of
canyons using current meters, turbidity currents
are thought to be the most important process in
the head region (Shepard and Marshall, 1978;
Shepard, 1982). In summary, conventional mod-
els for canyon heads indicate that they are active
features which erode upslope and will ulti-
mately erode across the shelf; they entrain
shelf sediments, and funnel them basinwards;
they usually have currents moving through them
most of the time, both up and down canyon
axis; and turbidity currents account for most
of the intermittent flow through canyon heads,
and in combination with slumps, slides and
debris flows, are thought to be the main cause
of headward erosion.
(ii) Canyon axes: Canyon axes may be straight for
some sections but are locally sinuous, and are
commonly deflected by fault line escarpments
(e.g. Cap Ferret, Cap Breton and Guinivec Can-
yons, Brittany: Kenyon et al., 1978; Baltimore
and Wilmington Canyons, northeastern US con-
tinental slope: Twichell and Roberts, 1982;
Almeria Canyon, SE Spain: Cronin, 1995, Cro-
nin et al., 1995), or other seafloor topography
such as slumps (e.g. Lagos Canyon, Gulf of
B.T. Cronin et al. / Sedimentary Geology 179 (2005) 71–97 73
Cadiz: Gardner and Kidd, 1983; Monterey Fan
Canyon: Gardner et al., 1991) or salt domes
(Shepard and Emery, 1973). Otherwise, the
chief control on planform geometry of deep-
water canyons appears to be slope gradient
(Belderson and Kenyon, 1976). In summary,
deep-water canyon axis orientation is usually
driven by underlying structure, courses are
strongly affected by a range of seafloor topog-
raphy, and spacing appears to be controlled by
slope gradient.
(iii) Canyon walls: Canyon walls are usually V-
shaped, and this profile changes downslope,
in contrast to the U-shaped gullies which feed
the canyons in the head region, and of deep-
sea channels, which are usually flat-bottomed
(Carter, 1988). Most canyon walls are steep
or vertical, with local terracing giving most
canyons a dsteerheadT profile (e.g. Gollum
Channel System: Wheeler et al., 2003). A
downslope change in canyon profile from
V-shaped initially to U-shaped until the can-
yon eventually loses topographic expression,
has been observed by many workers (e.g.
Malahoff et al., 1982; Twichell and Roberts,
1982).
(iv) Distal canyon reaches: Distal parts of canyons
have not been imaged routinely. Most documen-
ted observations of lower canyon reaches date
back to the 80s. Canyons develop thalwegs in
their lower reaches and these thalwegs channels
may have significant dimensions (e.g. Stoe-
chades Canyon thalwegs, 300–500 m wide
and 25–75 m deep: Le Pichon and Renard,
1982; St Tropez Canyon, 200 m wide and 700
m deep: Le Pichon and Renard, 1982). These
thalwegs are also made remarkable by the fact
they rarely have a talus, which suggests that
lateral slumping of material from the canyon
and thalwegs walls is transferred into axial
transport within the thalwegs. These thalwegs,
found in most canyon lower reaches, pass into
constructional channels once the canyon has lost
topographic expression. There may be an inter-
vening transitional zone, called a canyon-fan
transition, which is characterized by a rise area
dissected by channels with an erosional aspect
(Akhmetzhanov et al., 2003).
1.2. Canyon evolution and filling
Current understanding of deep-water canyons, par-
ticularly those fed by subaerial drainage systems, is
that they form initially by retrogressive slumping on
the upper continental slope (e.g. Bouma et al., 1985:
Mississippi), and are maintained by funnelled density
flows. In studies of the Pliocene–Pleistocene eustatic
cycles in the Gulf of Mexico, the activity within the
Mississippi Canyon began as sea level fell (where
mass-transport complexes were deposited as the can-
yon was excavated). At the lowest point of sea level
sand was transported through the canyon into the deep
basin as channel levee systems, and as sea level began
to rise 30,000 years ago, deposition on the turbidite
fan ceased, and in the canyon area continued well into
the Holocene (12,000–11,000 years BP). As this pe-
riod corresponded to deglaciation, large volumes of
sediment continued to be funnelled through the Mis-
sissippi Canyon during transgression (Bouma et al.,
1989; Weimer, 1990). Sequence stratigraphic models
show cessation in turbidite deposition in shorter sub-
marine canyons, or those not fed by major rivers,
suggesting that submarine canyon activity stops just
prior to transgression (e.g. Posamentier and Vail,
1988).
Outcrop studies of canyon fills typically recognize
canyon fills that include chaotic deposits and thin-
bedded turbidites, with locally developed coarse-
grained, usually lenticular bodies, particularly in the
lower part of the fill (e.g. Doheny Channel, Piper and
Normark, 1971; San Carlos submarine canyon, Mor-
ris and Busby-Spera, 1988; Charo Canyon, Mutti et
al., 1988; Point Lobos submarine canyon, Clifton,
1981, 1984; Cronin, 1994; Cronin and Kidd, 1998).
In all of these examples, the submarine canyons are
interpreted to have been excavated initially due to
major external factors such as sea level fall or recon-
figuration of slope morphology by thrusting. Their
fills are then made up of one or more broadly fining-
upwards sequences, typically with clast-supported
conglomerates on an erosive surface (always inter-
preted as residual lags) and usually capped with thick
mud-prone intervals. In short, ancient submarine can-
yon fills are interpreted to reflect (a) erosion by
excavation of the upper slope by mass wasting; (b)
sediment bypass to the deep basin, (c) a major filling
phase; and (d) thin-bedded turbidites, probably con-
B.T. Cronin et al. / Sedimentary Geology 179 (2005) 71–9774
centrated by the local topography (Piper and Nor-
mark, 1971).
Clearly, a lot of what we know about the down-
slope evolution and interpreted behaviour of subma-
rine canyons on deep-water slopes is still rooted in
work from the 70s and 80s, with some excellent case
studies appearing in recent years. As more examples
of high resolution 3D seismic data showing renderings
of the seafloor, multibeam bathymetry surveys, and
deep-towed side scan sonar surveys of slopes are
acquired, we will develop a better understanding of
these systems. However at present, little is known
about the longer-term behaviour of canyons, and
their mechanisms of filling. This is important for oil
and gas exploration and production from slope can-
yons. Current thinking on the distribution of potential
reservoir facies, particularly sand bodies, has canyons
as zones of bypass as they develop, which backfill
with sand as their hinterland floods back, and then get
plugged with mud. This paper is an example of a
modern deep-water canyon that was surveyed from
the head area on the upper continental slope, down to
the upper rise area at the base of slope, with two
different side scan sonar tools. Ground-truthing of
the acoustic facies allowed the mapping of sand on
the present sea floor and in the shallow subsurface, to
test some of the models for sand distribution in these
features. It is hoped that the models produced will be
used as a proxy for predicting sand distribution in the
types of scenarios described.
The objective of this study was to examine the
planform expression of the Donegal Bay deep-water
canyon system on the NE Rockall Margin, and to
design a sampling strategy to recover sediment from
different parts of the canyon, to test the recent activity
of the system. Two side scan sonar systems were used,
the medium range OKEAN and the deep-towed high
frequency O.R.A.Tech systems, which operate at dif-
ferent frequencies and thus produce sonar mosaics
that have different depths of penetration. These two
systems were compared with parts of the TOBI
(Towed Ocean Bottom Instrument) side scan sonar
data collected on the Irish Margin by the Southampton
Oceanography Centre (S.O.C.) as part of the ENAM
II project (see Stoker and Hitchen, 2003 for ENAM II
objectives and results summary). Sediment was recov-
ered using a 6 m gravity corer. Usually, areas of high
backscattering on the side scan sonar mosaics were
targeted, to examine sand distribution at or near the
sea floor in the canyon area. Such a survey over much
of the length of a deep-water canyon system would
highlight that the mechanisms of sediment transport,
chiefly from sediment gravity flows, have changed
over time as the canyon has progressively shut down.
This would thus allow testing of the simple assump-
tions that are routinely made about canyon backfilling
and shutdown in response to relative sea level rise.
2. Geological setting
The Rockall Trough is situated to the west of
the Hebridean and Irish Shelves on the NE Atlantic
European Margin; the deeper parts of the basin are
mostly within Irish territorial waters (Fig. 1). It is a
major NE–SW trending sedimentary basin that
extends as far to the north as the Wyville–Thomson
Ridge, NW of the Faeroes, where it is 1000 m
deep. It increases in depth to the SW to 4000 m
as it opens out into the Porcupine Trough. The
Rockall Trough is between 200 and 250 km wide,
and is flanked to the east by the Hebridean and Irish
shelves, and to the west by the Rockall Plateau and
adjacent smaller banks to the north (Stoker and
Hitchen, 2003). The deeper part of the Rockall Trough,
below the 2000 m isobath, extends across approxi-
mately half of the basin, from the late Cretaceous–
early Palaeogene volcanic edifices of the Anton Dohrn
and Hebrides Terrace Seamounts, to the SW. The
complex configuration of plateaus, troughs and sea
mounts on this highly topographically irregular
ocean margin is a reflection of Mesozoic–Cenozoic
rifting that led to the opening of the NE Atlantic Ocean
in the early Eocene. This phase has been overprinted
by later, post-Eocene subsidence, Oligo-Miocene com-
pression associated with the Alpine orogeny and late
Neogene uplift (Dore et al., 1999; Naylor et al., 1999).
The study area is located on the northeastern mar-
gin of the Rockall Trough on the Irish Margin to the
west of the Malin Shelf, offshore Eire. This region is
on the dividing line of a major change in the mor-
phology of the deep-water slopes on that basin margin
flank. To the north of the study area the Hebridean and
east Rockall Bank shelves have seen considerable
uplift since the early Pliocene, with the shelf margins
having prograded by up to 50 km (Stoker and
Fig. 1. Location of the northeastern Rockall Trough margin and study area. Bathymmetric data source: GEBCO. SSF: Sula Sgeir Fan; BF: Barra
Fan; DF: Donegal Fan; RKB: Rockall Bank; N.I.: Northern Ireland; Smt: Seamount.
B.T. Cronin et al. / Sedimentary Geology 179 (2005) 71–97 75
Hitchen, 2003). This uplift was related to the regional
Miocene uplift that also strongly affected the northern
North Sea. The Sula Sgeir Fan off the NW Hebridean
shelf, the Barra Fan southwest of the outer Hebrides
(north of the Hebrides Terrace seamount) and the
Donegal Fan (immediately to the north of the study
area) are the most prominent of these progradational
wedges, and record extensive shelf glaciation since
the mid-Pleistocene (Stoker and Hitchen, 2003). To
the south of the study area towards the western flanks
of the Porcupine Bank, the slope has a destructive
profile and is characterized by mass-wasting that has
persisted since the late Palaeogene (Stoker et al.,
2001). The complex topography that resulted from
this structural configuration of elements implies that
the entire margin has had a profound influence on
palaeoceanographic circulation and deep-water sedi-
mentation. The eastern margin of the Rockall Trough
in particular has been constructed by a combination of
along slope and downslope processes, and is referred
to as a sediment loaded (in contrast to a sediment-
starved) slope, with turbidite fans (e.g. Sula Sgeir Fan,
Barra Fan, Donegal Fan; Fig. 1) overlying and inter-
digitating with older and contemporary along slope
sediment drift deposits (Stoker and Hitchen, 2003).
Along-slope transport of sand is known from the
upper slope in this area (Kenyon, 1986).
The general morphology of the southern Rockall
Trough on the Irish margin is known from previous
surveys. The slope is characterized by a series of
canyons which cut back almost to shelf break, and
another series of associated canyons which appear to
initiate part or halfway down the continental slope.
These canyons have areas of high backscatter associ-
ated with their floors. Apart from the rare but often
dramatic canyon and channel features, the slope is
characterized by a series of features indicative of
mass wasting. These include slumps, slides, slump
scars, failure terraces (oriented parallel to slope), and
other more enigmatic features that could be inter-
B.T. Cronin et al. / Sedimentary Geology 179 (2005) 71–9776
preted as sedimentary bodies beyond the canyon
mouths, or bodies scattered across the slope, of un-
known origin or composition.
The Donegal Bay Canyon (first named here) is the
focus of examination within the study area, and is the
most prominent submarine canyon surveyed in detail
to date in the region. The head region is located at
approximately N55845VW10840V, below the shelf
break to the west of Donegal Bay in the Republic of
Ireland. Its existence has been known for some time,
and it is shown as a marked wide embayment into the
shelf edge profile in many atlases (e.g. Times Books,
1999).
3. OKEAN side scan sonar and 3.5 kHz profile
data
The study area was surveyed using OKEAN, a
Russian medium-range sidescan sonar tool, a hull-
mounted 3.5 kHz profiling system, and simultaneous-
ly with multichannel seismic. Fig. 2 shows the profile
tracks over the study area from the upper to lower
slope. Fig. 3 shows the OKEAN side scan sonar
mosaic, with high backscattering areas shown in
light shades, and lower backscattering areas shown
in darker shades. Cores locations are also highlighted
and will be discussed below in the ground-truthing
section. Fig. 4 shows an interpretation of the
OKEAN mosaic, which clearly shows the Donegal
Bay submarine canyon. The head area is impressive,
with a large dcauliflowerT shaped amphitheatre up to
20 km in diameter, fed into by a network of incised
tributary gullies just below the shelf break. The
OKEAN shows this area to be broadly highly back-
scattering in character, interpreted to correspond to
mud-poor and sand-rich sediment at or near the
canyon floor. The middle and lower parts of the
OKEAN mosaic are medium backscattering in char-
acter, and this is interpreted to reflect depositional
topography. This part of the slope shows one main
trunk canyon running downslope from the complex
head region.
The 3.5 kHz hull-mounted profiling survey was
carried out simultaneously with the seismic and
OKEAN data acquisition, at a speed of 6 knots.
The data were of superb quality, and were used
initially for the siting of gravity core locations. Sub-
sequently the profiles were used more systematically
to describe the shallow subsurface and the sea floor
topography. Fig. 5 shows the 3.5 kHz lines from the
entire survey area (Fig. 2) from just below the shelf
break to the lower slope area. This figure is thus a 3-
dimensional figure down the complete Donegal Bay
Canyon. A series of 15 acoustic facies were ob-
served and a scheme of their characteristics is pre-
sented (Fig. 6). In the section below, the main
features are described and interpreted. A generalized
map of the acoustic facies distribution is shown in
Fig. 7.
The survey area has three different characteristic
geographic regions of sea floor with associated acous-
tic facies, specific features and sediment recovery
from core. These three areas are (i) erosional area;
(ii) erosional and depositional area; and (iii) deposi-
tional area (Fig. 4).
3.1. Erosional area (lines PSAT 160, 159, 163, 162)
3.1.1. Erosional canyon
Figs. 4 and 5 show that the upper headward parts
of the Donegal Canyon are characterized by evidence
for erosion-dominated processes. Fig. 7 shows that
from 1000 to 2000 m water depths, the acoustic facies
are dominated by 3A, 4A, 5 and 6, which are facies
indicative as hard bottom or outcrop, separated by
areas of parallel acoustic bedding, which increase in
width down system. From 1000 to 1500 m the sea
floor is dominated by closely spaced, V-shaped gullies
separated by ridges. The gullies become tributary to
one another within a broader deep-water valley at
1500 m. The valley width expands rapidly to almost
20 km with a maximum depth of 625 m at this zone of
confluence of the V-shaped gullies. At 1800 m water
depth the sea floor is dominated by one trunk valley
with a width of just over 5 km, and with associated
levees, overbank area, terraces and valley thalwegs.
The valley has the typical characteristics of a
dsteerheadT channel profile. It is ca. 450 m deep
with walls of poor acoustic response.
3.1.2. Intervening ridges
Between the gullies and trunk valley area are in-
tervening areas that are characterized by a variety of
acoustic facies. In the erosional head region the ridges
have internal reflections (e.g. Fig. 5, PSAT 160, 159).
22:34
23:00
23:30
0:00
0:29
1:00
1:30
2:00
2:30
2:583:02
22:00
3:30
4:00
4:234:27
5:00
5:30
6:00
6:18
6:45
7:00
7:30
8:00
8:30
8:40
9:00
9:30
9:36
PSAT-164
23:560:00
0:30
1:00
1:30
PSAT-165
PSAT-161
1:37
2:00
2:30
3:00
3:37
PSA
T-166
4:10
4:30
5:00
5:30
6:00
6:17
PSAT-1
67
PSAT-1
62
PSAT-1
63
PSAT-1
59
PSAT-1
60
PSAT-1
58
6:30
7:00
7:27
PSAT-168
7:33
8:00
8:30
9:00
9:309:32
PSAT-1
69
10:05
10:30
11:00
11:30
12:00
12:14
PSAT-1
70
12:2012:30
13:00
13:29PSAT-171
13:35
14:00
14:30
15:00
15:30
15:44
PSAT-1
72
16:12
16:30
17:00
17:30
18:00
18:30
PSA
T-173
18:35
19:00
19:20PSAT-174
19:26
19:30
20:00
20:30
21:00
21:30
21:44
PSAT-1
75
11:30
12:00
12:30
13:00
13:30
14:00
14:30
15:00
15:30
16:00
16:30
17:00
17:30
18:00
18:30
ORAT-39
ORAT-39
11°30’W 11°20’W 11°0’W 11°00’W 10°50’W 10°40’W 10°30’W
11°30’W 11°20’W 11°0’W 11°00’W 10°50’W 10°40’W 10°30’W
54°40’N
54°50’N
55°00’N
55°10’N
AT282G
AT283G
AT284G
AT285G
AT286G
AT287G
AT288G
AT289G
AT290G
AT291G
AT292G
AT293G AT294G
AT295GAT296G
OKEAN 10 kHZ coverage
OREtech 100 kHZ coverage
AT286G Sampling sites
0 10 20 30
Km 1000
2000
Fig. 2. Map of the Irish Margin with profile (seismic, OKEAN side scan sonar and 3.5 kHz sub-bottom profiler), gravity core and O.R.E.Tech
(deep-towed side scan sonar) locations. Bathymetry from S.O.C. unpublished data.
B.T. Cronin et al. / Sedimentary Geology 179 (2005) 71–97 77
Those ridges bordering the main Donegal Canyon
show typically strong acoustic response with moder-
ate to deep penetration (facies 1D). The ridges thus
have a constructional appearance that contrasts with
the erosive nature of the gullies and canyons. The
ridges are interpreted to be a combination of erosional
remnants of slope sediments into which the gullies
and canyons have incised, and local evidence of levee
build-up.
3.1.3. Transparent lenses
There is evidence for transparent lenses on the
acoustic profiles in the erosional head area, though
these are restricted to PSAT 159, within the main valley
profile, where they have a hummocky appearance and
laterally restricted occurrence (facies 6). The lenses
have a transparent seismic character and are associated
with local slope changes. They aremuchmore common
in lower reaches of the system. They are interpreted as
11°30’W 11°20’W 11°10’W 11°00’W 10°50’W 10°40’W
10 °40’W10°50’W11°10’W 11°00’W11°20’W11°30’W
10°30’W
10°30’W
54°40’N
54°50’N 54 °50’N
55°00’N
55°10’N 55°10’N
54°40’N
55°00’N
0 10 20 30
Km
AT282GAT282G
AT283GAT283G
AT284GAT284GAT285GAT285G
AT286GAT286G
AT287GAT287G
AT288GAT288G
AT289GAT289G
AT290GAT290G
AT291GAT291G
AT292GAT292G
AT293GAT293GAT294GAT294G
AT295GAT295G
AT296GAT296G
1000
2000
Fig. 13C
Fig. 10C
Fig. 10A
Fig. 3. OKEAN mosaic from the Irish Margin, showing the location of lines and cores. Light shades are high backscatter, dark shades are low
backscatter. Gravity core locations are shown in white. Location of O.R.E.Tech profile 39 shown in white. Interpretation shown in Fig. 4.
B.T. Cronin et al. / Sedimentary Geology 179 (2005) 71–9778
either slumps from the valley walls or as muddy debris
flow deposits filling an older, wider valley, that have
subsequently been eroded near the thalwegs.
3.1.4. Mounded features
A fourth feature observed on the erosional area of
the slope was that of transparent mounds on the
terrace of the Donegal Canyon. The best of these is
seen on the northern side of PSAT-162 at a water
depth of about 2100 m (Fig. 6, facies 6). This
mound is internally transparent on the acoustic profile.
Its origin is unknown.
3.2. Mixed erosional-depositional area (lines PSAT-
167, 166, 170)
The second geographic area is one with evidence
for combined erosion and deposition (Figs. 4 and 5).
284
283
285
292 294
296295293
291
288289290
282
ORAT-39
286
287
sand prone mud pronedepositional
Erosional
Mixederosional
depositional
Depositional
5 kmS H E L F B R E A K
'Cau
liflo
wer
' - s
hape
d he
ad r
egio
n
Fig. 4. Interpretation of OKEAN mosaic shown in Fig. 3.
B.T. Cronin et al. / Sedimentary Geology 179 (2005) 71–97 79
The main Donegal Canyon is one distinct trunk feeder
channel between water depths of 2200 and 2500 m,
with smaller channels feeding in from the south and
north. There is evidence for hard sea floor bottom
locally, particularly on the floor of the canyon and the
smaller channel. The canyon loses topographic ex-
pression rapidly, down to less than 100 m deep.
3.2.1. Canyon
The Donegal Canyon has reduced in depth, has a
flat bottom, with rough sea floor topography including
remnants of the ridges seen in the erosional part of the
upper slope. The canyon is less than 150 m deep and 4
km wide. Two of the canyons which are seen at the
edge of the survey area (PSAT-166, Fig. 5) are filled
900
1000
1100
1200
1600
1700
2100
2200
2600
2300
2700
2400
2800
2500
1800
1900
2000
2000
1300
1400
1500
1600
1700
2100
2200
2600
2300
2700
2400
2800
2500
1800
1900
2000
2000
1400
1500
2100
2200
2600
2300
2700
2400
2800
2500
2000
2000
2100
2200
2600
2300
2700
2400
2800
2500
2000
2600
2300
2700
2400
2800
2500
2600
2700
2400
2800
2500
2600
2700
28002600
2700
2800
2700
2800
2700
2800
2700
2800
Pr. 160
Pr. 159
Pr. 163
Pr. 162
Pr. 167
Pr. 166
Pr. 170
Pr. 169
Pr. 173
Pr. 172
Pr. 175
N
Erosional
area
Erosional and
depositional area
Depositional area
Canyondirection
Transparent lensmudflow deposit
Main canyondirection
KEY
Dep
th, m
ORAT-39
Fig. 5. Three-dimensional view of the northeastern Rockall Trough slope from shelf to rise (times in milliseconds TWTT). All panels are
interpreted from 3.5 kHz profiles (see Fig. 2).
B.T. Cronin et al. / Sedimentary Geology 179 (2005) 71–9780
1A
1B
1C
1D
2A
2B
Facies Schematic Description Type example
Deep penetration; manysimilar reflectors;
arithmetic deterioration ofreflectors
Deep penetration; strongreflectors in middle;
irregular reflectordeterioration
Deep penetration; one hardreflector; several feint,
continuous reflectors; deeperdispersal/diffusion pattern
Deep penetration; arithmeticdecrease in reflector strength;discontinuous lower reflectors
Transparent lens withindistinct base
Transparent lensunderlain by strong
reflectors
Transparent lens withdistinct base
Shallow penetration 1-2hard reflectors
Shallow penetration withrapid decline in reflector
strength
2C
3B
3A
PSAT 168
PSAT 175
PSAT 172
PSAT 172
PSAT 166
PSAT 166
PSAT 164
PSAT 159
PSAT172
Damuth and Hayes, 1977 DescriptionFacies
IA
IB
IIIC
IIIB
IIIA
IIB
IIA
Continuous, sharp bottom echoeswith no sub-bottom reflectors
Continuous, sharp bottom echoeswith continuous, parallel sub-bottomreflectors
Semi-prolonged bottom echoes withintermittent zones of semi-prolonged,discontinuous, parallel sub-bottomreflectors
Very prolonged bottom echoes withno sub-bottom reflectors
Large, irregular overlapping orsingle hyperbolae with widelyvarying vertex elevations abovethe sea floor
Regular single or slightly overlappinghyperbolae with conformablesub-bottom reflectors
Regular overlapping hyperbolae withvarying vertex elevations above thesea floor
4A Rough sea floor; hardand diffuse returns;
diffractions
4B Rough sea floor, lessdark returns; diffuse PSAT 162
PSAT 166
Schematic Description Type example Damuth and Hayes, 1977 DescriptionFacies
6 PSAT 162
5 PSAT 163
7 PSAT 160
8 PSAT 166
Transparent sea floor;transparent mounds
Small blocks above mainSea floor reflector
Steeper slopes; variousInternal reflectors, some
oblique
Hard returns in channels;multiple diffraction
IIIF
IIIE
IIID
Regular, intense, overlappinghyperbolae with verticesapproximately tangent tosea floor
Zones of irregular, intense,overlapping hyperbolae withvertices tangent to the sea floorwhich are interrupted by zones ofdistinct echoes with parallel sub-bottom reflectors
Broad, single, irregular hyperbolaewith disconformable, migratingsub-bottom reflectors
(A)
(B)
Fig. 6. Acoustic facies scheme for the Irish Margin in the vicinity of the Donegal Bay Canyon (left), with Damuth and Hayes (1977) for
comparison (right).
B.T. Cronin et al. / Sedimentary Geology 179 (2005) 71–97 81
11°30’W 11°20’W 11°10’W 11°00’W 10°50’W 10°40’W 10°30’W
54°40’N
54°50’N
55°00’N
55°10’N
X
X
XXX
XX
XXXXXX
X
X
XX
XXXXXX
XX
X
XXXXXX XX
XXXX
XXXXXXXXX
XX
LEGEND:
PARALLEL ACOUSTIC BEDDING (Facies 1, 3B)
TRANSPARENT LENSES (Facie 2)
HARD BOTTOM (Facies 3A, 4A, 5)
OUTCROP (Facies 3A, 6)
TRANSPARENT HUMMOCKY (Facies 6, 4B)
XXXXX
1000
2000
500
2500
0 10 20 30
Km
Fig. 7. Map showing the generalized distribution of 3.5 kHz acoustic facies.
B.T. Cronin et al. / Sedimentary Geology 179 (2005) 71–9782
with transparent acoustic facies (facies 2B; locally
facies 2A and facies 2C). The main Donegal Canyon
has higher acoustic penetration (facies 1B).
3.2.2. Transparent lenses
The intervening low-relief ridges seen on PSAT-
166 in the middle of the erosional–depositional area
have small transparent lenses, in addition to those
recognized at the bases of the smaller feeder canyons.
These lenses are found predominantly on the slopes of
the ridges. Some of the lenses seen on the canyon
floors can be traced laterally, updip, onto flatter or
steeper areas of slope. These lenses are interpreted
here as muddy debris flow deposits.
3.3. Lower Slope area (lines PSAT 169, 173, 172,
175)
This geographic area includes the lower reaches of
the Donegal Bay Canyon. Depths range from 2595 to
B.T. Cronin et al. / Sedimentary Geology 179 (2005) 71–97 83
2665 m, which is a relatively flat area compared with
the other two areas. A generally strong reflector char-
acter with rough topography is observed, with either
parallel acoustic bedding (facies 1, 3B) or transparent
lenses (facies 2).
3.3.1. Transparent lenses
The lower reaches of the Donegal Bay Canyon are
characterized by laterally extensive transparent
acoustic facies lenses, which are locally sheet-like
in geometry, and are concentrated in topographic
lows. These lenses are seen on the lower profiles in
Fig. 5. Typically the lenses pass laterally or pinchout
onto topographic highs with facies 3A acoustic char-
acter. Lenses are present at depth ranges between
2650 and 2740 m, and are of variable thickness.
The thickest lens (b7 m) is partially on a slope to
the east where it onlaps or passes into diffuse and
then sharp reflectors, though several of the lenses are
seen to abruptly terminate against areas with very
sharp, sometimes even vertical terminations. One of
the wider lenses is 2 km wide. This lens occupies the
much shallower (75 m deep) down dip reach of the
Donegal Bay Canyon.
The transparent lenses are interpreted as canyon
mouth depositional areas occupied by slump or deb-
rite deposits and perhaps coarser material, bordered
on the slopes on either side by hemipelagic material.
The debris flow deposits appear to infill an eroded
canyon topography. There is evidence for erosion
here with a terraced erosive cut that is filled by
debrite material, which is probably muddy. The west-
ern margin of the valley appears to be associated
with faults.
4. O.R.A.Tech side scan sonar data
The O.R.A.Tech side scan sonar is a high-resolu-
tion system that was operated at 100 kHz providing a
swath of 1000 m (500 m each side). The vehicle was
towed at about 50 m above the seabed. The
O.R.A.Tech system also includes an acoustic sub-
bottom profile (SBP) that operates at 7 kHz in order
to give good resolution.
One O.R.A.Tech profile was designed on the
basis of backscatter characteristics of a previous
TOBI (towed ocean bottom instrument) survey in
the area (S.O.C., unpublished data), and information
gathered from the OKEAN mosaic collected during
this leg (Fig. 3). High backscattering features on the
TOBI were thought to be sand or gravel features,
within a broader canyon morphology. An exercise in
ground-truthing and comparison of side scan sonar
tools with different depths of penetration and reso-
lution was undertaken to resolve the enigmatic fea-
tures, and this is discussed in more detail in the
next section.
4.1. ORAT 39 description
The O.R.A.Tech deep-towed 100 kHz side scan
sonar provided detailed (high resolution) imaging of
the area surveyed. Line 39 (Fig. 2) is positioned in the
lower reaches of the down slope-running canyon sys-
tem on the continental rise of the Irish margin. The
line runs from 55804VN–10857VW to 54852VN–11812VW in a NE–SW direction. It is about 24 km
long and located halfway between the 3.5 kHz profile
and OKEAN/seismic lines PSAT 169 and PSAT 170
(Fig. 3). The positioning of ORAT 39 was based upon
the previous TOBI 30 kHz side scan mosaic in which
some of the acoustic facies were difficult to interpret
(Fig. 8).
On the basis of the O.R.A.Tech profile it is noted
that topography is flat, denoting a broadly aggrada-
tional relief. The depths range from 2580 to 2680 m.
The 2580 m depth corresponds to a topographic
elevation, which is a wide, gently sloped ridge run-
ning SE–NW as seen on the bathymetric chart (Fig.
2), on the northeastern section of the line. The 2680
m depth is correlated to the thalweg of a V-shaped
channel partly infilled with sediments of low acous-
tic backscatter that give the channel its current U-
shaped profile. It is located on the mid-section of the
line, together with five other channels of higher
backscatter and reflectivity on the profile. Four of
them are U-shaped, with very low angle walls, and
one is V-shaped with an asymmetric profile. This
system of channels is positioned on a topographic
depression correlated to the main canyon in the
upper reaches of the area surveyed by OKEAN
and 3.5 kHz profile.
Towards the southwestern end the depths gently
decrease to 2650 m, corresponding to a gently sloped
wide ridge, running SE–NW from the upper reaches
5
7
11
3
1a9
1a
3
11
7
6
5
6
NE11:34
SW15:20
Start
SW
NE15:20
18:45
End
2600m
2550m
2600m
2550m
1a
4a
2
1b
3
4a
9
8
10
11
3
9
4b
10
9
700 m
700 m
Fig. 8. Mosaic and facies outline of O.R.E.Tech line ORAT-39—Rockall Trough. Acoustic facies are shown in Fig. 9, and described in the text.
See Figs. 2–4 for location.
B.T. Cronin et al. / Sedimentary Geology 179 (2005) 71–9784
B.T. Cronin et al. / Sedimentary Geology 179 (2005) 71–97 85
of the area. The declivity between the deepest chan-
nels and the NE ridge is 0.18 or 1:150 (the bathym-
etry increases 1m for every 150 m). This is twice
that of the declivity between the same channel thal-
weg and the southwestern ridge, which is 0.178 or
1:300.
(a)
(b)
Parallellineations of mediumbackscatter in a homogenous areaof low backscatter.
as with subtle lineations(a)
(a)
(b)
Slightly rough texture; smallstipples of medium backscatter;low scatter background.
as with parallellineations(a)
low
back
scat
ter
high
back
scat
ter
2
1a-b
3
4a-b
6
7
8
9
10
11
DescriptionExampleFacies
Elongate patches of mediumbackscatter on a smoothbackground of low backscatter
Smooth, homogenous areaof lowbackscatter;little acoustic variation(occasional stippling)
Smooth; homogenous; featureless;medium backscatter
Slightly rough texture; stippledhigher backscatter;mediumbackscatter background
Rough, granulartexture; highbackscatter
Distinct features; varying width;very high backscatter;roughtextures
Irregularzone; high backscatter;occasional parallellineations; somegrooves and ridges
Discontinuous linear features; verynarrow;fine-medium texture
Scattered lineations; highbackscatter;coarse granulartexture; dispersive
5
med
ium
back
scat
ter
line
arfe
atur
es
Fig. 9. Facies scheme for O.R.E.Tech 100 kHz line—ORAT-39
(Fig. 8).
4.2. ORAT 39 acoustic facies description
Eleven acoustic facies have been recognized on
line ORAT 39 (Fig. 9). They have been classified
into four groups according to their backscatter re-
sponse and the presence of linear sea floor shapes.
The low backscatter group comprises facies 1a, 1b, 2,
3, 4a and 4b. The medium backscatter group is made
up of acoustic facies 5 and 6. Facies 7, 8 and 9 are
within the high backscatter group. The final group
comprises acoustic facies 10 and 11 that have different
patterns of linear features. All of the facies are dis-
tinctive and readily recognized. This side scan sonar
line is particularly interesting because of the large
number of contrasting backscatter patterns, whose
origins are rather enigmatic. Ground-truthing of
these acoustic facies is discussed in the next section.
The polarity on the images is positive which
implies high backscatter response is in black–dark
grey colour, and low backscatter is in light grey–
white. Fig. 8 shows the acoustic facies/backscatter
patterns along line ORAT 39. The facies can be
described as wide, linear features of approximately
180 m width. The wide, linear facies may occur
either orientated SE–NW, normal to the line (facies
1a, 1b, 2, 3, 4a, 4b, 5, 6 and 7) or oriented obliquely
(608) to the line (facies 8, 10 and 11). Only facies 9
has two different orientations. It either occurs as a
narrow, linear feature (30–200 m wide on average)
oriented both SE–NW (normal to the line) and E–W
(oblique to the line at 308 and 608) or as an irregular,
U or horseshoe-shaped (in plan view) patch, opening
to the NW.
5. Ground-truthing
A series of gravity cores were collected on the
basis of backscatter acoustic facies from the
O.R.A.Tech profile ORAT 39 and its 7.5 kHz profile,
and the 3.5 kHz profile sections (Figs. 10–13). In this
section we discuss the gravity cores from three differ-
ent geological/physiographic parts of the Irish Margin
deep-water slope, and see how the sedimentary layer-
ing may explain the backscattering on the OKEAN
and O.R.A.Tech sonographs. These three areas are: (i)
Canyon Head region; (ii) Lower Slope region and (iii)
Upper Rise region. In all regions, a pattern-coding of
Fig. 10. (A) Sedimentary core logs (depth in cm) AT-282 to AT-285G with positions on (B) 3.5 kHz hull mounted profiler line PSAT 162 and (C) on OKEAN side scan sonar mosaic.
AT-282 hit buried turbidite sand in the overbank area; AT-283 hit sand on the sea floor on a mound on the terrace; AT-284 and AT-285 both hit dstickyT debris flow that had slumped in
from the side. AT-286 was outside the line of section. See Fig. 2 for location.
B.T.Cronin
etal./Sedimentary
Geology179(2005)71–97
86
Fig. 11. 7.5 kHz O.R.A.Tech profiler line ORAT-39 (see Fig. 3) with core logs (depth in cm) AT-292 to AT-295G. AT-292 and AT-293G
penetrated muddy debrites. AT-294G penetrated into shallow turbidite sands in a narrow channel seen on the profiler. AT-295G penetrated a
wide, empty channel. Core AT-296 was not on the profile.
B.T. Cronin et al. / Sedimentary Geology 179 (2005) 71–97 87
the core numbers at their locations (as shown on the
sonograph details) is used. A stippled pattern is used
for those cores where sand (of both turbidity current
and bottom-reworking/contour current origin) was re-
covered, a grey pattern for reworked mud (both
muddy debris flow and slump origin), a pebble pattern
for prominent layers of clast supported gravel of
enigmatic or compound origin, and a white pattern
for hemipelagic sediment. The same patterns are used
in the core logs (Figs. 10–13).
Of the two side scan systems used on this cruise,
the OKEAN system with a frequency of 10 kHz has a
maximum penetration of 10 m, but lower resolution
than deep-towed systems. The O.R.A.Tech has a dual
frequency but in this ground-truthing exercise we
operated the 100 kHz band, which results in penetra-
tion of up to approximately 0.3 or 0.4 m, and has the
highest resolution of the two. TOBI (not used on this
cruise) may resolve features that are too deeply buried
for O.R.A.Tech to identify.
5.1. Canyon Head region
Six gravity cores were taken in the Canyon Head
region: four in a transect across the upper reaches of
the main deep-water canyon (AT-282G to AT-285G),
one from the floor of the same canyon approximately
8 km nearer the shelf break up the canyon axis (AT-
286G) and one core (AT-287) was taken from a
mound off the canyon flank in the extreme SW of
the surveyed area, approximately 17 km up dip from
the canyon transect (Figs. 2 and 10C).
AT292GAT292G
AT293GAT293G
AT294GAT294GAT295GAT295G
AT296GAT296G
2 km
- mud flow - clast-supported layer - sand - hemipelagic
AT-292GAT-293G AT-294G AT-295G
OKEAN (10kHz) O.R.E.tech(100kHz)
(A)
(B)
N
50m
700 m
Fig. 12. (A) Comparison of two different side scan sonar tools operating at different frequencies. At 10 kHz, OKEAN may penetrate several to
10s of meters below the sea floor—thus buried features are frequently imaged that drop cores may not reach. At 100 kHz, O.R.A.Tech will
penetrate only up to 0.2/0.3 m, and thus only very shallow or sea floor features are imaged. Backscattering penetration is tested by targeting
features with drop cores. See text for details. (B) O.R.A.Tech profile showing core locations. Compare with Fig. 11.
B.T. Cronin et al. / Sedimentary Geology 179 (2005) 71–9788
5.1.1. Canyon transect
Fig. 10C shows the core logs, the OKEAN data in
the canyon transect and PSAT 162 3.5 kHz profile
data in the immediate vicinity of the transect with the
positions of the cores.
Core AT-282G was targeted in an area of multiple
layering on the 3.5 kHz profile data which is outside
the canyon, and is thought to be the overbank area to
that canyon. The backscattering on the OKEAN is
medium to low. A linear feature at the core location
running NW–SE and then doglegging towards the
canyon and is probably a buried channel or canyon
(Fig. 10C). The core is not thought to penetrate this
object. The core recovered bedded hemipelagic sedi-
ment and three sands between 2.4 and 3.5 m that
contain siliciclastic material with foraminifera and
local glacial dropstones.
Core AT-283G was targeted on a topographic high
between the overbank and canyon axis areas, and is
thought to be the highest part of the canyon levee (or
erosional remnant, if not aggradational), with a trans-
parent mound structure clearly visible on the 3.5 kHz
data. The OKEAN sonograph shows medium–high
backscatter, although this is somewhat obscured by
the ship’s-track and by a nearby acquisition anomaly
on the sonograph. The core location is near the in-
ferred buried channel at core station AT-282G. The
core recovered bedded sands and foraminiferal ooze,
though it is dominated throughout by sand. Penetra-
tion was poor (0.8 m), and this was thought to be due
to the core barrel hitting deeper, impenetrable sand
below this level.
Core AT-284G was targeted on the deepest part of
the canyon, i.e. the inferred canyon thalweg, where
the 3.5 kHz profile shows a hard, shallow-penetra-
tion acoustic response with some diffractions. On the
OKEAN, the backscatter is medium and the sedi-
mentary body is clearly part of the bcauliflowerQ-shaped area of the Canyon Head region. The core
recovered hemipelagic sediment and hit the top of a
muddy slump or debris flow interval, though core
penetration was poor (0.6 m). This area corresponds
to the location of two large transparent lenses within
the canyon.
Core AT-285G was targeted on a mounded fea-
ture on the canyon floor with highly irregular sea
floor topography, strong acoustic returns on the side
scan sonar, and diffractions on the 3.5 kHz profile.
Fig. 13. (A) Sedimentary core logs (depths in cm) AT-288 to AT-290G with (B) 3.5 kHz hull mounted profiler line PSAT 175 with core positions, and (C) core positions shown on the
OKEAN side scan sonar mosaic. AT-290 penetrated a debris flow with pebbles, corresponding to the transparent lens on the 3.5 kHz profile; AT-289 hit the feather edge of the debris
flow floored by turbidite sand, and AT-288 pelagic marls.
B.T.Cronin
etal./Sedimentary
Geology179(2005)71–97
89
B.T. Cronin et al. / Sedimentary Geology 179 (2005) 71–9790
The feature was interpreted as a slump mass on the
canyon floor. On the OKEAN the backscattering is
medium, a little higher than at station AT-284G, but
clearly in the axis of the main canyon at the down-
dip end of the bcauliflowerQ structure. The core
recovered a hemipelagic sequence with intercala-
tions of slumped material. The slumped intervals
are dewatered and extremely sticky in texture, mak-
ing it difficult to extract the sediment from the
core-liner.
5.1.2. Core AT-286G: Canyon Head core (Fig. 10)
Core AT-286G was taken approximately 7 km up-
canyon towards the headwall of the cauliflower pat-
tern imaged on the OKEAN mosaic (Figs. 3 and 10).
The core station was not targeted on the profile, but it
does coincide with the sonographs. On the OKEAN,
the location has low backscatter within the central part
of the cauliflower structure. A deep canyon with high
backscatter margins is seen clearly within the cauli-
flower. The core was targeted at the axial part of the
canyon floor, and recovered intercalated thick sandy
layers (the thickest recovered on the margin) and
hemipelagic sediments.
5.1.3. Core AT-287G: marginal mound feature
Core AT-287G was targeted using the 3.5 kHz
profile on a mounded structure on the flank of the
main canyon further up canyon from AT-286G. The
structure, despite its presence at an unusual depth
(~1300 m), was targeted as a possible carbonate
mound. Such carbonate mounds are commonly
found on the Porcupine Trough margin but at shal-
lower depths (~500–1200 m: Kenyon et al., 2003),
though Lophelia pertusa, the dominant coral species
that is found on these mounds, is known in water
depths up to 2000 m in the Atlantic (Le Danois,
1948). The structure is one of several features of the
same size seen on PSAT-161, at water depths between
1240 and 1280 m. On OKEAN, the core station is in a
localized area of high backscatter situated on the flank
or just outside a major SSW–NNE trending canyon
that feeds into the SW part of the bcauliflowerQ struc-ture. There was no recovery in the gravity corer, but
there were sand and clasts (interpreted as dropstones)
in the core catcher, suggesting that the features are
more likely to be sediment waves than carbonate
mounds.
5.1.4. General interpretation of the Canyon Head
region
The bcauliflowerQ-shaped head area seen so spec-
tacularly on the OKEAN mosaic is not seen on the
shallower penetration TOBI mosaic (S.O.C. unpub-
lished data). On OKEAN, stacked linear canyon and
channel features and tributaries with low backscatter
are interpreted to record major phases of erosion and
sand deposition in the head area that do not occur at
present. On 3.5 kHz profiles the dominant sedimen-
tary bodies overprinting this older phase are slumps
and muddy debris flows. Most of the smaller-scale
mudflows and slumps are directed towards the canyon
and channel axes, away from the canyon walls. This
indicates that the system does not transport sand at
present. One slump body seen on TOBI and cored
twice is thought to have blocked the main canyon
axis, and subsequent sand transportation may have
been ponded behind it. In targeting sand with the
gravity corer, it was found that coring the topographic
highs near canyons, coring areas of low backscatter
near to the head of the system, and collecting samples
from the deepest part of a V-shaped thalweg, are the
most reliable ways. Other potentially sand-rich areas
were too deep below the sea floor (seen on OKEAN)
to be sampled by the 6 m gravity corer, largely due to
the thickness of the slumped masses. The overbank
area in particular (e.g. AT-282G) shows that relatively
recent large sandy turbidity currents moved through
the canyon, and spilled into this region. This is par-
ticularly impressive when one considers the width (17
km) and depth (~800 m) of the canyon. The timing of
these flows is thought to be towards the end of the last
glacial period, confirmed by their association with
dropstones.
5.2. Lower Slope region
Five cores were collected in the mixed erosional–
depositional, lower slope area of the Irish Margin
(Figs. 3 and 11). Cores AT-292G to AT-295G were
all selected on the basis of their backscatter character-
istics on O.R.A.Tech, in combination with informa-
tion about seafloor topography and acoustic response
from the 7.5 kHz profile (Fig. 12). Core AT-296G was
selected based on backscatter patterns. The cores were
taken in a strike (along-slope) transect (Fig. 3). Cores
AT-292G and AT-293G were targeted on depositional
B.T. Cronin et al. / Sedimentary Geology 179 (2005) 71–97 91
features on the profile, to test the reason for the
difference in backscatter on OKEAN and O.R.A.Tech.
Cores AT-294G and AT-295G were targeted on ero-
sional features. Core AT-296G was targeted on a very
high backscattering area seen on TOBI, where sea-
floor topography was unknown.
Core AT-292G was taken from an area of medium
backscatter on OKEAN and TOBI, and low backscat-
ter on O.R.A.Tech (Figs. 3 and 8). On the profile the
sub-bottom is characterized by deep penetration, mul-
tiple layering. A transparent layer was observed on the
profile, which can be traced laterally under core sta-
tion AT-293G, pinching out before station AT-292G
(Fig. 11). The core recovered a thick mudflow deposit
capped by Holocene marl.
Core AT-293G was taken from an area of medium–
high backscatter on both OKEAN and O.R.A.Tech
(Fig. 12). On the profile the seafloor reflects strongly
but otherwise is the same as station AT-292G. The
core recovered a mudflow deposit capped by a marl,
but separated by a 0.05 m thick clast-supported gravel
(Fig. 11).
Core AT-294G was taken from an area of medium–
high backscatter on OKEAN, a medium–high back-
scattering, narrow, linear feature on TOBI (with a
NW–SE orientation), and an area of low backscatter
on O.R.A.Tech. On the profile this area is in a small,
V-shaped channel with a thin layer of transparent
acoustic response underlain by stronger returns and
poor penetration (Fig. 11). The core had poor recov-
ery, with a Holocene marl capping sand. The marl is
interpreted to correspond to the transparent acoustic
facies on the 7.5 kHz, and the sand to be the top of the
higher amplitude acoustic facies below the channel
base. The O.R.A.Tech does not show the sand (0.38
m) due to lack of penetration.
Core AT-295G was taken from an area of medium
backscatter on OKEAN and low backscatter on
O.R.A.Tech. On the profile the core is from the mid-
dle of a zone of strong reflectivity within a flat-
bottomed, shallow depression. The core recovered a
thick hemipelagic sequence without sand, draped by
Holocene marl. The depression is interpreted as an
inactive, draped channel (Fig. 11).
Core AT-296G was taken from an area of me-
dium backscatter on OKEAN. The core recovered a
thick hemipelagic interval capped by a Holocene
marl, and separated by a thin clast-supported gravel
(reworked dropstone) layer, identical to that seen in
AT-293G.
5.2.1. General interpretation of the Lower Slope area
The different side scan sonar systems allow 3D
analysis of the upper sedimentary layers, in combina-
tion with 3.5 kHz, 7.5 kHz and seismic profiling
across the same transect (Figs. 11 and 12). The high
backscattering gravel layer seen on TOBI (S.O.C.
unpubl. data) is not as evident on OKEAN, because
it is not thick enough, and not as evident on
O.R.A.Tech, because it is partially buried. The same
can be said for sand-filled features such as the narrow
linear channel seen on TOBI, which is not clear on
OKEAN (Fig. 12), but perhaps shows up as a field of
linear features (acoustic facies 10 on O.R.A.Tech, Fig.
9). The transparent lens on the profile to the SW is
attributed to the (cored) mudflow that is buried by the
gravel layer in the centre of the profile. It is a sheet-
like lens extending across half of the profile, and is
thought to terminate at this locality rather than be
removed by erosion. The complex nature of mudflow
deposition, inferred from the profile and cores, cannot
be resolved from the sonographs, but it is inferred
here that there are many, interfingering mudflows
possibly coming from different local directions. This
is in agreement with conclusions from the Canyon
Head region; the system is now dominated by rework-
ing of intraslope mud. The gravel layers sampled at
AT-293 and AT-296 are thought to be of identical age.
Therefore the central part of the profile is interpreted
as an erosive valley (low relief area of erosion) that
truncated the gravel, and possibly the transparent lens
underneath.
5.3. Upper Rise region
Four cores were taken in the Upper Rise area.
Cores AT-288G to AT-290G were targeted on a
range of acoustic facies on the 3.5 kHz profile and
on OKEAN line PSAT-175. This section is a slope-
parallel transect in the depositional area of the Irish
Margin (Fig. 13).
Core AT-288G was targeted on the edge of a very
thin transparent lens on a topographic high otherwise
characterized by medium penetration, multiple layer-
ing (which deteriorates rapidly) in the sub-bottom.
The area is outside the wide, shallow valley that
B.T. Cronin et al. / Sedimentary Geology 179 (2005) 71–9792
corresponds to the area downdip of the main canyon
mouth. The core recovered a hemipelagic sequence
with two coarse-grained, clast-supported layers at
0.85 and 0.9 m. These coarse-grained layers are over-
lain by silty hemipelagic sediments, which may cor-
relate with the sands in AT-289G, and the succession
is capped by Holocene marl. Two thin, silty–sandy
layers were recovered towards the bottom of the core.
Core AT-289G was targeted on the edge of the
shallow valley, on an undulose low-angle slope. The
3.5 kHz profile is characterized by the feather edge of
a thin transparent lens underlain by a strong reflector
and some deeper strong reflectors. The core recovered
a thick sequence, with the Holocene marl underlain by
a thin mudflow, underlain by several graded silty–
sandy layers (max. thickness 0.2 m), interbedded
with hemipelagic sediments, a thin, clast-supported
gravel at 1.3 m, and a thick package of hemipelagics
without gravel. Several thin silty layers are seen be-
tween 2.5 and 2.7 m.
Core AT-290G was targeted on an area of slightly
higher backscatter on the OKEAN, where the trans-
parent lens on the 3.5 kHz profile is thicker, and
nearer the deepest part of the shallow valley than in
core AT-289G. The core recovered a thick sequence,
with the Holocene marl underlain by a 3.7 m thick
mudflow, with various layers of gravel, some clast-
supported, at its lower parts. Due to their isolation in
mudflows or in situ deep-water muds, and distance
from potential source of such rounded gravels, they
are interpreted as dropstones. In situ Holocene sedi-
ments underlie the mudflow.
6. Discussion
Submarine canyons are a major feature of deep
marine slopes, and act as transfer zones for clastic
sediment to deeper water. Fan models usually place
the submarine canyon as the point source for the fan,
and other areas of the slope are typically mud prone
with mass-wasting being the main process of sediment
remobilisation. The Donegal Bay Canyon is an exam-
ple of such a sediment transfer zone from the north-
western European margin. The morphology of the
canyon was examined using a variety of methods,
including high resolution 3.5 kHz seismic profiling,
low and high-resolution sidescan sonar imaging, and
ground truthing of the acoustic facies using gravity
corers. The canyon was examined from its head area
down to the continental rise. The head area is a classic
example of a dcauliflowerT-shaped, head area, previ-
ously described as an damphitheatreT-like head area byother workers (e.g. Kenyon et al., 1978). This head
region is the widest part of the canyon system, and
comprises a confluence area of a series of V-shaped
gullies, each with dimensions in the region of 100–
500 m wide and 150–300 m deep. The confluence
area is one of highly irregular sea floor, with multiple
V-shaped incisions within the larger canyon fairway.
Downdip the canyon passes rapidly into a single U-
shaped fairway that narrows to 2–3 km, with a depth
of 300 m. Downslope, the canyon is joined by several
other tributary canyons of varying dimensions, before
passing into a broad, shallow, scoured area and then
losing topographic expression on the continental rise.
The margins of the canyon are characterized in the
upper and medial reaches by terracing, giving the
canyon a dsteerheadT cross-sectional profile, and by
multiple slump scars, giving a dscallopedT morpholo-
gy to the edges of the fairway. This terraced character
and the prevalence of slump scars along the margins
has been described by other authors from slope can-
yons and channel complexes (e.g. Cronin et al., 1995;
Wheeler et al., 2003). The slump scars are interpreted
to reflect lateral wasting of canyon margins towards
the canyon axes by slumping. Terraces are interpreted
as reflecting larger scale wasting of the margins to-
wards the axes, by slow creep caused by rotational
slumping along discrete gravity faults. The surface
expression of this type of dslowT margin failure to-
wards the axis, seen in the OKEAN mosaic interpre-
tation in Fig. 4, is compelling. The implications of this
are far-reaching, as this type of canyon or slope
channel margin terracing, and even the position of
the interpreted faults from this model, is interpreted
as the product of dinside leveeT development in some
current models (Deptuck et al., 2003). In these mod-
els, the dipping reflectors that we interpret as parts of
the rotational slumps, are interpreted as the aggrada-
tional elements of a confined channel levee complex
within the main canyon or slope channel. Further-
more, the scalloped pattern, interpreted as axis-direct-
ed slump scars, is interpreted as the erosional margins
of underfit sinuous channel elements in the same
models. The dinside channelT model does not work
B.T. Cronin et al. / Sedimentary Geology 179 (2005) 71–97 93
for the Donegal Bay Canyon, and both end-member
models need to be tested more closely for other slope
channel and canyon complexes.
Submarine canyons are thought to initiate by head-
ward erosion of a submarine slump scar in the upper
slope region. As the feature gets larger due to contin-
ued retreat of the headwall, the canyon is modified by
episodes of mass wasting and the headwall retreats
towards the shelf break, where it becomes inclined to
trap turbidity currents spilling over the break. At this
point the canyon becomes a major sediment transfer
zone and if this process continues, the canyon may
erode back across the shelf towards any fluvial or
longshore drift sediment source. Most sediment is
transported to the end of the canyon system by tur-
bidity currents, and larger flows may even spill over
the walls of the canyons in the upper reaches, though
this is usually more common in the lower reaches,
where the canyon may pass into a leveed valley
system. Canyon bypass is the main phase of turbidite
fan growth, and corresponds to the classic Type I
turbidite system of Mutti (1985).
As relative sea level rises, present models infer that
the lower parts of the canyon will become deposition-
al, and that sand deposition is confined to the lower
slope area, perhaps with the fan backstepping into the
canyon mouth and even into the canyon itself. This
corresponds to a Type II turbidite system (Mutti,
1985). Continued sea level rise will see the sand
trapped on the shelf area, unable to be transported
into the submarine canyon as it is locked up-system.
The submarine canyon and slope area will thus revert
to a mud-starved doverbank wedgeT and the canyon
will be mud-plugged. This corresponds to a Type III
turbidite system (Mutti, 1985). Though this scheme
has been modified to take in subsequent concepts such
as flow efficiency, the main tenets remain the same in
the literature.
Fig. 14 is a model proposed to explain the mechan-
isms of filling observed in the vicinity of the Donegal
Bay Canyon. Three stages of activity within the main
canyon area are recognized during this filling period
(Fig. 14).
Stage 1 corresponds to the stages of sediment
bypass through the canyon, inferred to represent the
end of the last glacial low stand, when sand was
probably transported directly from a shallow Malin
Shelf area, where large volumes of sand were present
on the shelf, and initially bypassed the major Donegal
Bay Canyon into deeper water. The canyon base is
strongly erosive, even in the lower reaches, during this
stage. Occasionally sand was deposited adjacent to the
head area of the main canyon, thus spilling onto the
flat upper slope area in general. This has major impli-
cations for flow volumes as the canyon is 17 km wide
and up to 800 m deep in that area.
Stage 2 corresponds to a phase of mixed erosion and
deposition within the Donegal Bay Canyon. Sand is no
longer found outside the canyon in the upper reaches,
and the largest flows only managed to spill onto the
terraces within the main canyon in that area. Sand was
deposited in the upper and medial reaches of the can-
yon, and flows were smaller in volume and usually
confined to smaller channels within the main canyon.
Stage 3 saw a transition from sand and silt trans-
portation to one of muddy debris flow deposit and
slump dominance in the medial and lower reaches of
the Donegal Bay Canyon. These deposits are clearly
seen to overlie erosive features and coarser sediment
from stages 1 and 2 on the 3.5 kHz records. Slumps
were particularly common in the upper reaches, where
they blocked the main canyon trunk axis. Subsequent
turbidity currents that were transporting sand were
ponded in the canyon in the upper reaches, rather
than depositing sand sheets in the lower reaches of
the canyon. This is also complicated by sand spill over
from the shelfal region and from along slope trans-
portation of sediment as coarse sandy contourites
(Kenyon, 1986).
This paper presents a dataset where mechanisms
for canyon filling on a glaciated margin, which has
progressively drowned since the last glacial low stand,
are investigated. After a sustained phase of coarse
clastic sediment bypass through this major erosive
feature on the Irish Margin, the locus of sand deposi-
tion moved into the upper reaches. A later phase of
mass-wasting plugged the canyon and ponded the
sand in these upper reaches.
There are implications for the understanding of
deep-water slope channel complex filling, particularly
in West Africa where they are known to host major
fields of oil and gas (Mayall and Stewart, 2001).
There are many more potential stratigraphic hydrocar-
bon plays in this new model, particularly during stage
3, than predicted in existing models. Current under-
standing of the architecture of slope canyons is that
1
Stage 1:
All bypass (Mutti type I);Some overspilling of sandin head region
erosion
Overbanksand
Sandy terraces
2
Stage 2:
Occasional Bypass;Decrease in volume of flowsin lower reaches
3
Stage 3:
Plug and pond;Slumps from wallsdominate lower canyon fill and pond in uppercanyon
Continental rise(2700 m)
Shelf edge(1000 m)
5 km
Slope (2500 m)
overbanking
Bypass
Bypass and
Channelling, erosion and deposition (mixed)
Bypass
Slump:ponding of sand
Drape, slump,rare channelised sand
Slump anddebris flow
Fig. 14. Model for canyon filling in response to relative sea level rise.
B.T.Cronin
etal./Sedimentary
Geology179(2005)71–97
94
B.T. Cronin et al. / Sedimentary Geology 179 (2005) 71–97 95
they have complicated multi-phase fills between ini-
tial excavation, bypass and erosion, and eventual
plugging. In all of these models, the dquiescentTphases of activity are always characterized by sinuous
channel elements, which underfill the larger canyons,
and produce sediment fill packages that are highly
heterogeneous and have distinctly lower net:gross
ratios. Sand is found as either erosional remnants of
point bars or as longitudinal, in-channel bars. Slumps
and debris flow deposits are not included in any of
these models. This model implies that longitudinal
slumping and debris flows from canyon margins are
major processes in canyon filling, in distal as well as
proximal canyon reaches. This will produce dmud
podsT of local derivation within canyon fills, and
will also produce sand bodies caused by ponding
like that described in this paper, which are not includ-
ed in current models for slope channel complexes
(Beaubouef and Friedmann, 2000; Campion et al.,
2000; Cronin et al., in press; Mayall and Stewart,
2001; Mayall et al., in press, Wonham et al., 2000).
Acknowledgements
The TREDMAR TTR Program and R/V Logachev
crew are warmly thanked for their support and assis-
tance during data collection on cruise TTR-10. We
wish to thank Jaco Baas, the editor of this special
issue of Sedimentary Geology for this contribution,
for pursuing our revisions of this paper, and pushing it
to completion. BTC wishes to thank the supporting
Oil Companies from the dMesostratigraphy of Deep-
Water SandstonesT consortium: Amerada Hess, BP-
Amoco, Elf, Enterprise, and Conoco; and laterally,
ENI Agip, during the period that the work was carried
out. We wish to thank the following referees for their
considered and thoughtful reviews of earlier versions
of this paper: David Piper, Bill McCaffrey and Jamie
Pringle.
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www.elsevier.com/locate/sedgeo
Sedimentary Geology 1
An evaluation of along- and down-slope sediment transport
processes between Goban Spur and Brenot Spur on
the Celtic Margin of the Bay of Biscay
M.J. Cunningham *, S. Hodgson, D.G. Masson, L.M. Parson
Challenger Division for Seafloor Processes, National Oceanography Centre, Empress Way, Southampton, SO14 3ZH, United Kingdom
Geological Survey of Ireland, Department of Communications, Marine and Natural Resources, Ireland
Petroleum Affairs Division, Department of Communications, Marine and Natural Resources, Ireland
BT, United Kingdom
Flag Telecom, United Kingdom
Gemini, United Kingdom
Tyco Telecommunications, United Kingdom
Global Crossing and Global Marine Systems, United Kingdom
Abstract
Multi-beam bathymetry and backscatter, 3.5 kHz pinger profiles, side-scan sonar and seabed samples have been examined to
evaluate along- and down-slope sedimentary processes along the Celtic Margin shelf and upper slope in water depths of 200 to
1500 m. The continental shelf and slope are indented and dissected by major canyon systems. The shelf is characterised by major
northeast – southwest trending sand banks that are orthogonal to the shelf edge. Along the shelf edge, several fields of asymmetric
sandwaves oriented orthogonal to the canyon axes indicate sediment transport into the canyon heads. Less commonly, sandwaves
with weak asymmetry suggest sediment transport onto the shelf. These may be reworked and are partly overprinted by more recent
sandwaves. Down-slope sediment transport by turbidity currents is the dominant process through the major canyons. Recent
faulting has also played a role in canyon development. Turbidity currents are most likely initiated by faulting, and/or slope failure of
the walls that bound the canyon head drainage basins and sediment migration from the shelf. This leads to deep incision of sinuous
thalwegs in the upper reaches of canyon floors and downcutting and sediment transport on the mid to lower continental slope. The
canyons are V-shaped in the upper reaches and become U-shaped progressively down-slope, suggesting they represent either a
transition from erosive to depositional processes or sediment bypass conduits carrying sediment between the shelf and abyssal
plain. Over-bank spill from canyons leads to deposition of unconfined turbidite deposits (muds) on the intervening canyon spurs.
D 2005 Elsevier B.V. All rights reserved.
Keywords: Sedimentary processes; Morphology; Celtic Margin; Multi-beam bathymetry; Sediment waves; Submarine canyons
0037-0738/$ - s
doi:10.1016/j.se
* Correspondi
Southampton, S
E-mail addre
79 (2005) 99–116
ee front matter D 2005 Elsevier B.V. All rights reserved.
dgeo.2005.04.014
ng author. Navigation Safety Branch, Maritime and Coastguard Agency, Bay 2/30, Spring Place, 105 Commercial Road,
O15 1EG, UK. Tel.: +44 23 8032 9198.
ss: [email protected] (M.J. Cunningham).
Fig. 1. Regional setting of study area. a) Shaded bathymetric image
of SW Eire, SWApproaches and northern Bay of Biscay. Resolution
is 2 km. b) 3D image of Celtic Margin showing an amalgamation of
multi-beam bathymetry sourced from the Geological Survey of
Ireland, the Irish Petroleum Affairs Division of the Department of
Communications, Marine and Natural Resources (Ireland), subma-
rine telecommunications industry data and GEBCO. Horizonta
resolution is 150 m. Note that most prominent canyons along the
Celtic Margin shelf trend north-northeast – south-southwest and
then develop into a series of smaller canyons further west. The
small canyons however coalesce down-slope with the main canyon
system. BS - Brenot Spur, CSSB - Celtic Sea Sand Banks, GC -
Gollum Channel, GS - Goban Spur, GSDA - Great Sole Drainage
Area, KAC - King Arthur Canyon, LCB - La Chapelle Bank
LSDA - Little Sole Drainage Area, MT - Meriadzek Terrace, PB -
Porcupine Bank, SS - Shamrock System, WS - Whittard System
M.J. Cunningham et al. / Sedimentary Geology 179 (2005) 99–116100
1. Introduction
Over the last 10 to 15 years, multi-beam (swath)
bathymetric sonar systems have become the dominant
survey instrumentation at all water depths: deep ocean
(e.g., Exon et al., 1997), continental slope (e.g., Shaw
and Courtney, 1997; Todd et al., 1999; McAdoo et al.,
2000), continental shelf, shallow bays and estuaries
(Courtney and Fader, 1994). When integrated with
other information such as seabed samples and seismic
profiles, it is possible to chart bathymetry, and study
morphology and geology in unprecedented detail in-
cluding the mapping of fault scarps, sediment slides,
debris flows and turbidite deposits, the recognition of
sediment transport pathways to offshore depocentres,
studies of fauna and habitat, and the evaluation of
sediment transport processes.
The study area is located on the Celtic margin of
the Bay of Biscay between 11819VN/48855VW and
9825VN/48830VW and 150 to 4500 m water depths
(Fig. 1). Whilst new multi-beam bathymetric coverage
reveals exceptional details of sea floor morphology
within the study area, our knowledge of sedimentary
processes is poorly constrained. For example, which
canyons contribute to sediment flux between the shelf
and the abyssal plain? Is there a genetic link between
the Celtic Sea Sand Banks located on the continental
shelf as a sediment source (Marsset et al., 1999), and
the sandy sediments of the Celtic Deep Sea Fan
(Zaragosi et al., 2000) at the foot of the continental
rise? What sediment patterns and processes affect
individual channels: which are likely to be experien-
cing erosion and canyon head retreat, which are swept
clean or filling up with sediment? To improve our
knowledge of the Celtic Margin, we need to better
understand processes such as bottom currents and
their temporal variability, the structural and erosional
characteristics of canyon evolution, bedload transport,
and sedimentary depositional patterns by tidal cur-
rents. There is also a significant commercial aspect
(e.g., for the submarine telecommunications industry)
to increasing our understanding of sedimentary pro-
cesses such as the assessment of risk from slope
instability and the prevention of cable breaking.
The main aims of this paper are to gain a better
understanding of sedimentation processes on the Cel-
tic Margin through the interpretation of integrated
multi-beam bathymetry/backscatter, side-scan sonar,
l
,
.
3.5 kHz profiles, bottom sediment grabs and gravity
cores. We first outline the regional setting of the study
area, describe our observations of erosive and depo-
sitional features, and then discuss the significance and
main implications for along- and down-slope sedi-
ment transport processes.
The data presented in this paper are based on 1)
submarine telecommunications industry cable route
surveys, 2) multi-beam bathymetry/backscatter and
3.5 kHz profiles of the Geological Survey of Ireland
M.J. Cunningham et al. / Sedimentary Geology 179 (2005) 99–116 101
(GSI) collected as part of the Irish National Seabed
Survey (Cullen, 2003), and 3) multi-beam bathymetry
of the Irish Petroleum Affairs Division of the De-
partment of Communications, Marine and Natural
Resources, Ireland; acquired in 1996 as part of the
United Nations Convention Law Of the Sea (Part VI,
Article 76).
2. Materials and methods
The bathymetry/backscatter and shallow seismic
data presented in this paper were collected onboard
the RV Bligh in 2001 using Kongsberg Simrad
EM1002S (water depths b1000 m) and EM120
(water depths N1000 m) echo sounders and a hull-
mounted 3.5 kHz Oretech 3010 S transceiver, respec-
tively. Seismic reflectors were mapped from 3.5 kHz
profiles and subsequently transformed into contoured
3.5 kHz echofacies maps of penetration depth. Pene-
tration contours were further constrained using shaded
relief bathymetry by forcing the interpolation algo-
rithms to follow bathymetric contours and avoid areas
of steep slope (N158).Data supplied by the submarine telecommunica-
tions industry were acquired onboard the NO Jean
Charcot in 2000. These data include multi-beam ba-
thymetry/backscatter, shallow seismic and side-scan
sonar and were collected using a Konsberg Simrad
EM120 echo sounder, a hull-mounted 3.5 kHz ORE
4�4 Pinger and a GeoAcoustics Deepwater Model
159 fish, respectively. Bathymetry and backscatter
were processed and gridded using open source MB-
System software (Caress and Chayes, 1995). Shallow
seismic sub-bottom profiles (SBP) and side-scan
sonar (SSS) were supplied as analogue rolls. In addi-
tion to the geophysical/morphological data, a gravity
corer with a 3 m long barrel and a Shipek grab
sampler were used to acquire seabed samples during
survey operations (Fig. 1b).
3. Regional setting
The Celtic Margin is a sediment-starved passive
margin trending west-northwest – east-southeast (Fig.
1; Roberts et al., 1981; de Graciansky et al., 1985).
The margin is characterised by a relatively steep
continental slope with a mean gradient of 118, butlocally attaining very steep gradients to vertical along
canyon walls. The geological evolution of the margin
is a combination of several major phases of tectonic
activity (de Graciansky et al., 1985) associated with
the Palaeozoic formation and subsequent Mesozoic
disintegration of Pangea, and the initiation of North
Atlantic rifting and sea-floor spreading (Ziegler,
1981). The Celtic Margin shelf is characterised by a
series of fault-bounded, rift basins trending west-
southwest – east-northeast (Dingle and Scrutton,
1979; Roberts et al., 1981; Ziegler, 1987). Seawards
of the margin, there is a series of tilted and rotated
fault blocks bounded by north-northwest – south-
southeast trending, listric normal faults (Dingle and
Scrutton, 1979; Roberts et al., 1981; de Graciansky et
al., 1985), which relate to rifting prior to seafloor
spreading. Late Cretaceous eustatic sea-level rise
combined with a cessation of tectono-sedimentary
activity resulted in the progressive overstepping of
earlier basin margins and the deposition of a wide-
spread transgressive sequence (Roberts et al., 1981;
Ziegler, 1987).
Post-Pliocene, there was a widespread marine
transgression due to continuing subsidence (Naylor
and Shannon, 1982) with relatively low sedimentation
rates (Dingle and Scrutton, 1979). It has been sug-
gested that rejuvenated Late Neogene –Holocene can-
yon cutting and slope wasting has removed much of
the earlier Neogene sediments deposited on the West-
ern Approaches slope (Evans et al., 1990).
3.1. Hydrodynamics
The hydrodynamics of the Celtic Margin are typi-
fied by high-energy conditions, in particular related to
storm surges and spring tides, which transport sediment
from the near-shore to the shelf-edge (Reynaud et al.,
1999; Zaragosi et al., 2000). Sediment transport may
also occur in the form of mass flows, either due to
erosion of the Celtic Sea Sand Banks, or from
regressive canyon head erosion (Reynaud et al.,
1999; Zaragosi et al., 2000).
Tidal conditions in the Celtic Sea are typically
vigorous and sediments are transported south-west-
ward (Zaragosi et al., 2000), related to ebb-dominated
tidal flow (Reynaud et al., 1999). These currents
decrease in strength from the southeast (around La
M.J. Cunningham et al. / Sedimentary Geology 179 (2005) 99–116102
Chapelle Bank, velocities close to the seabed are
locally 0.9 m s�1, Heathershaw et al., 1987) to the
northwest (currents of 0.2 m s�1 are more typical of
the shelf southeast of the Goban Spur, Huthnance et
al., 2001). At the shelf-edge, the dynamics of these
large tidal currents generate correspondingly large
internal tides, which may also be important for sedi-
ment transport across the slope area (Reynaud et al.,
1999; Wollast and Chou, 2001; Huthnance et al.,
2001). In addition, internal waves may serve to rein-
force bottom current velocities, which cause sediment
erosion (Huthnance et al., 2001) and can occur at
depths of up to 400 – 500 m on the Celtic Margin
(Pingree and Le Cann, 1989).
Whilst bottom current data in the region are rela-
tively sparse, results from the Ocean Margins EX-
change project (Wollast and Chou, 2001) show
typical along-slope residual current velocities, close
to the shelf-break, of 0.05 m s�1 (Huthnance et al.,
2001). Such currents are responsible for north-west-
ward transport and are associated with a general
northerly flow west of the UK driven by the North
Atlantic Current and at depth, by the Mediterranean
Outflow Water (900 – 1000 m) and the Lower Deep
Water (N2500 m) (Colling, 2001; Huthnance at al.,
2001). Internal tides are probably important for trans-
port and reworking of fine- to very fine-grained sedi-
ments within canyons and on the intervening spurs
(see Section 4.4).
3.2. Submarine canyons
Steep-sided canyons crosscut the continental shelf-
edge and deeply incise the margin (Fig. 1). The
orientation of canyons at the shelf edge is normally
north-northwest – south-southeast and north-north-
east – south-southwest, although numerous smaller
canyon segments also occur and vary towards east –
west trends. The shape and form of the canyon heads
is semi-circular with a concave profile. The canyons
provide conduits for the transport of sediment from
the shelf to the abyssal plain (e.g., Zaragosi et al.,
2000) and for over-bank turbidity currents, which
deposit on the intervening terraces and spurs.
A number of the shelf-break canyon heads may
possibly be the seaward expression of Pleistocene
river mouths, which supplied large volumes of sed-
iment directly to the shelf edge (Hadley, 1964;
Marsset et al., 1999; Droz et al., 2003; Bourillet
and Lericolais, 2003). The present-day deposits of
the area are dominated by marine bioclastic sands
(Bouysse et al., 1979), with fine-grained material in
the form of terriginous silts and clays, which by-
pass the hydrodynamically energetic shelf to settle
on the continental slope below water depths of
500 m (Auffret et al., 1979).
4. Morphology and sedimentary processes —Celtic
Margin
The Celtic Margin shelf/slope transition has a gen-
eral west-northwest – east-southeast trend from the
King Arthur Canyon (eastern margin of the Goban
Spur) to the Brenot Spur (Fig. 1a). The slope break
occurs at water depths of between 170 and 300 m,
landward of which the shelf morphology (at a cell
resolution of 25 m) is relatively smooth (Fig. 1b). The
margin is dissected by canyons with dominant south-
southwest and south-southeast trends. In general, two
types occur along the margin: 1) canyons with rela-
tively long, narrow upper reaches and V-shaped pro-
files that incise the shelf-break; and 2) canyons with
relatively short, broad upper reaches, U-shaped pro-
files and heads deeper than the shelf-break on the
continental slope. This pattern breaks down further
west where morphology consists of smaller canyon
segments in various orientations, but is re-established
towards the King Arthur Canyon (Fig. 1a,b).
Morphological evidence for sediment transport is
seen throughout the canyon systems. Some of the
most important features are described in the following
Sections, beginning at the shelf edge and progressing
down-slope.
4.1. Sediment waves north of Brenot Spur
The Brenot Spur forms an important divide between
two main canyon networks, the Whittard and Sham-
rock systems. The Whittard system is located to the
west of the spur and drains the Great Sole drainage area
(linked to the southern end of the lowstand Irish Sea
river system) along the Celtic Margin. To the east, the
Shamrock system drains the Little Sole drainage, which
is primarily associated with the lowstand English
Channel river system (Bourillet and Lericolais, 2003;
M.J. Cunningham et al. / Sedimentary Geology 179 (2005) 99–116 103
Zaragosi et al., 2003). The western margin of the
Brenot Spur is bounded by a major north – south trend-
ing canyon. This canyon swings to a north-northwest –
south-southeast trend upslope and offsets three north-
northeast – south-southwest canyons (Fig. 1b).
The largest of the canyons occurs in the east. Two
sets of sediment waves occur on either side of the
canyon head, one to the southeast and one to the
northwest (Figs. 1 and 2). Pinger (3.5 kHz) profiles
show zero penetration of the sediments within the
waves, which suggests that the substrate is sandy
and that these bedforms are sandwaves.
The sandwaves southeast of the main canyon head
are oriented northwest – southeast (Set A in Fig. 2)
and cover a minimum area of 150 km2 (sandwaves
extend beyond our bathymetric coverage). They are
Fig. 2. Sandwaves. Shaded bathymetric image of two sets of prominent sed
inferred. West-northwest striping in this and subsequent images represents
consistently asymmetrical with angular crests and lee
slopes facing southwest (Fig. 2) and are mainly par-
allel. They have wavelengths of 450 – 650 m, and
typical relief of 3 – 5 m. These sandwaves occur in
water depths of 170 – 200 m and continue into the
canyon head area (Fig. 2).
The larger set of sandwaves (covering an area
N200 km2 in water depths of 180 to 230 m) lies
some 3 km to the northwest of the canyon head and
is oriented west-northwest – east-southeast (Set B in
Fig. 2). The morphology of these sandwaves is quite
different to that of Set A. They are sub-parallel,
swinging clockwise to a more northwest–southeast
trend towards the west, and show frequent bifurcation
(Fig. 2). They have a lower slope angle and opposite
sense of asymmetry (lee slopes facing northeast) com-
iment waves from which opposing sediment transport directions are
artefacts (ship-tracks).
M.J. Cunningham et al. / Sedimentary Geology 179 (2005) 99–116104
pared to Set A (Fig. 3). Some are also near symmet-
rical and tend to have more rounded crests and
troughs in comparison to Set A. Wavelengths vary
between 350 and 650 m with relief of 2 – 10 m. A
secondary field of east–west trending sandwaves
occurs on the northeast slopes of the main sandwaves.
These secondary features have a wavelength of 250 –
350 m, and relief of 1 – 3 m (Fig. 2). The wavelength
of Set B sandwaves decreases westward to 250 – 300
m. The decreasing relief and subsequent disappear-
ance of the sandwaves are associated with a change of
slope aspect (Fig. 3), which marks the uppermost
extent of another canyon head.
To the west, there is a series of arcuate sandwaves
at the head of the dominant north-northwest – south-
southeast canyon, covering an area of N10 km2 (Fig.
4). The sandwaves become slightly sinuous towards
the northeast and have an overall northwest – southeast
trend. These sandwaves have a relief of 0.5 – 2 m and a
wavelength of 100 – 120 m. They show minor asym-
metry, generally towards the southwest, but occa-
sionally to the northeast.
In summary, the crests of the sandwaves are or-
thogonal to the axes of canyon heads. The strong
asymmetry of Set A shows sediment movement is
towards the canyon heads and illustrates that the
Fig. 3. Slope-aspect of sandwaves. a) Gradient map (averaged over 75 m
head. b) Aspect image showing facing directions of both sets of sandwaves
images represent flat areas. Note the change in aspect related to the cany
canyons have a sediment source. The potential build
up and subsequent failure of sandwaves in the canyon
heads may initiate gravity and/or turbidity currents, an
important agent in canyon incision and erosion. The
sandwaves of Set B show subtle asymmetry, are N1
km distance from the canyon heads, and give an
inferred sediment movement away from the canyon
heads. These features can be explained by considering
the bifurcation and rounded morphology of the sand-
waves. The presence of a secondary set of sandwaves
suggests that Set B originally formed during the last
glacial low-stand, possibly through wave-dominated
processes and the secondary set are forming under
present day current regimes. This is in harmony with
the primary sandwaves now being re-worked and
more symmetrical in form. It is also possible, how-
ever, that these sandwaves are related to temporal
variations in storm events, which reflect the prevailing
southwesterly trade winds.
4.2. Canyon head- faulting, slumping
An important feature of the north-northwest –
south-southeast trending canyon are the well-devel-
oped drainage basins (also known as bamphitheatre
rimsQ e.g., Belderson and Kenyon, 1976) (Fig. 4) on
length scale). Note bathymetric Profile 1 is upslope from a canyon
and canyon heads. Areas of grey on image and subsequent daspectTon head to the south.
Fig. 4. Tectonics and sediment erosion. a) Shaded bathymetric image at head of north-northwest – south-southeast canyon. The eastern margin
of the canyon is fault controlled with a minimum vertical displacement of 120 m. Most recent movement post-dates major canyon scarp retreat
of a tributary system. b) 3D shaded relief showing east-southeast trending hanging wall tributary entering the main canyon system on the
western margin. Also shown are sediment waves, which infer a sediment transport direction towards the southwest.
M.J. Cunningham et al. / Sedimentary Geology 179 (2005) 99–116 105
the eastern margin (marked as AT on Fig. 4). Within
the catchment boundary there is a well-developed
channel network with interfluves bounded to the
west by a scarp forming the main canyon margin.
The headwalls of the canyon occur at water depths
of 200 – 330 m and 200 – 460 m on the west and
eastern margins, respectively (Fig. 4).
Erosion and downcutting of the canyon floor have
resulted in the carving of a thalweg, with the floor
having an average down-slope gradient of 38 com-
M.J. Cunningham et al. / Sedimentary Geology 179 (2005) 99–116106
pared to average gradients of 0.68 on the shelf and 118on the canyon walls (Fig. 5). The main canyon has a
single tributary on the western margin, which trends
west-northwest – east-southeast and is 6.5 km in
length. This canyon tributary has a thalweg with
moderate sinuosity of 1.17 compared to the main
canyon thalweg sinuosity of 1.08 (Schumm and
Khan, 1972). There is also deep incision in the trib-
utary canyon floor of up to 150 m (Fig. 4) and a
channel gradient of 68, which is coupled with strongly
developed incised meanders.
On the northern margin of the tributary, above its
confluence with the main canyon channel, there are
two small plateaux (b2 km2 in total), one of which
dips gently west-northwest, the other dipping gently
east-southeast (Fig. 4). There is no obvious drainage
system crossing either of these platforms and they
may represent rotational slides from over-steepened
canyon walls, or terraces that may have formed by
lateral migration of the thalweg (e.g., see Mulder et
al., 2004).
4.2.1. Upper reaches
The upper reaches of the canyon head are desig-
nated as those above the confluence of the main
channel and its primary tributary (Fig. 4). Here the
canyon is oriented approximately north – south and
the main canyon thalweg shows a low degree of
sinuosity (1.01) with weakly developed meanders
Fig. 5. Slope-aspect of faulted canyon. a) Gradient map (averaged over 60
on the eastern margin of the canyon. This fault trace continues across par
prominent sediment waves to the northeast margin of the canyon head.
and a down-slope gradient of 38. The upper reaches
are characterised by a more distinct incision of tribu-
tary canyon heads on the eastern margin of the main
channel than those on the western margin.
4.2.2. Lower reaches
At the channel confluence there is a subtle change
in the orientation of the canyon axis, towards the
north-northwest – south-southeast. Below this point,
the canyon is essentially linear in character and has
an average down-slope gradient of 28. The east-north-east edge of this part of the canyon is defined by a
steep scarp with an average height of 130 – 140 m and
a gradient in excess of 608 (Fig. 4). The top of the
scarp lies at water depths of 420 – 460 m and is linear
in plan. It can be traced into the upper reaches of the
canyon (though less well defined than in the lower
reaches), over a distance of approximately 7 km and
has a south-southwest facing slope. In general, this
scarp is parallel to the lower reaches of the canyon
floor.
Given the linear nature, steepness and relatively
large relief of the eastern canyon margin scarp, we
interpret this feature as a fault. To the northwest the
scarp disappears, although it may coincide with the
beginning of channel incision in the upper reaches of
the canyon. This suggests that there may be a signi-
ficant tectonic control on down-slope erosion and
subsequent sediment deposition.
m length scale). Note the steep northwest – southeast trending fault
t of the amphitheatre. b) Aspect image of canyon system. Note the
M.J. Cunningham et al. / Sedimentary Geology 179 (2005) 99–116 107
The fault offsets the drainage basins (bamphitheatre
rimsQ) and with sufficient time, displacement along
the fault could result in a steepening of channel
gradient from the drainage basin. The timing of
such changes in fluvial systems is poorly understood;
similarly there are no documented, quantified studies
of the response of submarine systems to such tec-
tonic movements. However, it is clear that insuffi-
cient time has passed since fault movement for these
changes to occur, as the drainage basins do not
appear to have developed a more bmatureQ morphol-
ogy, e.g. rounded interfluves, gentler slopes, up-
stream migration of channel knickpoints (currently
located on the fault scarp), and formation of a par-
abolic thalweg (re-equilibrium with post-fault dis-
placement boundary conditions). Therefore, we
Fig. 6. Canyon incision and slumping. a) Shaded relief image showing thr
displays deep incision of up to 150 m in its upper reaches, with a V-shaped
U-shaped profile (Profile 2). The Eastern canyon is broader in form with
Central canyon forms a broad, flat plateau. The white star represents the lo
60 m length scale). Note very high angle slopes clearly define the incised c
map: this shows that shelf surfaces (interfluves) are relatively flat (shaded g
artefacts. CC - Central canyon, DB - Drainage basin, DIC - Deeply incise
speculate that this fault is an active structure (see
Discussion).
4.3. Down-slope sediment transport
The main canyon tributary of the fault-controlled
canyon is separated by a small ridge from a broad
U-shaped canyon further west. The orientation of this
canyon is northeast – southwest and forms the eastern
component of a set of three sub-parallel canyons
(Fig. 6). The two other canyons, referred to here as
the Central and Western canyons, are much narrower
and more V-shaped in their upper reaches. At the
head of the western canyon, a series of asymmetric
northwest – southeast trending bedforms indicate a
south-southwest sediment transport direction.
ee major north-northeast trending canyons. The canyon to the west
profile (Profile 1). Compare this to further down-slope where it has a
a far lower degree of incision. The interfluve between this and the
cation of a gravity core. b) Image of maximum slope (averaged over
hannel and its terraces, whilst the shelf is relatively flat. c) Gradient
rey). Closely spaced stripes (parallel to depth contours) are gridding
d canyon, EC - Eastern canyon, WC - Western canyon.
M.J. Cunningham et al. / Sedimentary Geology 179 (2005) 99–116108
4.3.1. Eastern canyon
The morphology of the canyon head is less incised
than those further west and occurs on the slope at a
depth of 400 m. This may suggest that it is younger
than the shelf-indenting Central and Western canyons
(Fig. 6) where retrogressive mass wasting of slope
sediments along headwalls (bamphitheatre rimsQ) hasled to headward migration and eventual indention of
the shelf (e.g., Twichell and Roberts, 1982; Farre et
al., 1983). However, it has also been suggested that
canyons are eroded by turbidity currents on the upper
slope/shelf edge (Daly, 1936). Slope failures lead to
gravity flows moving down-slope through pre-exist-
ing bathymetric lows and thus lead to canyon evolu-
tion. Subsequent flows widen the canyons by
retrogressive canyon wall failure, i.e. undercutting of
the canyon walls leading to instability and thus caus-
ing slope failure (Pratson et al., 1994; Pratson and
Coakley, 1996).
A gravity core on the drainage divide between the
Eastern canyon and the canyon tributary to the east
consists of green to brown homogeneous soft, slight-
ly clayey, silty, fine sand. The substrate material is
capable, given the existence of appropriate flow
regimes (possibly internal tides), of producing close-
ly spaced sandwaves. However, the apparent lack of
bedforms and the broad U-shape of this canyon head
might indicate that the system once extended further
upslope and has subsequently been cut by the fault-
controlled canyon system to the east, and, as shown
in Section 4.4, erosive processes are active in this
canyon.
4.3.2. Central canyon
The Central canyon occurs in water depths of
190 – 240 m, with retrogressive canyon wall failure
occurring on both margins. The canyon thalweg
shows moderate sinuosity (1.23) and has a V-shaped
profile (Profile 2, Fig. 6). The canyon has an average
down-slope channel gradient of 2.58 and typical can-
yon wall gradients of 10 – 148 (Fig. 6b). Drainage
basin development is more extensive on the eastern
margin (DB on Fig. 6). Whilst it is clear that slumping
has occurred on the eastern margin, the basin now
appears to be infilling with sediment. In general, this
drainage basin, with an ideal parabolic thalweg, is
more maturely developed than the fault-related drain-
age basin (Fig. 4).
4.3.3. Western canyon
The Western canyon has a typical down-slope
channel gradient of 38, with average canyon wall
gradients in the order of 11 – 158. Canyon margins
occur in water depths of 200 – 300 m and the canyon
has a deeply incised channel with drainage basins on
the margins. The drainage basins are less well devel-
oped than those of the Central canyon. The primary
features of the deeply incised thalweg include two sets
of parallel terraces and steep channel walls of 458(Fig. 6a,b). Consequently the canyon has a stepped,
relatively narrow, V-shaped profile (Profile 1, Fig. 6),
which broadens downstream to a U-shape profile
(Profile 2, Fig. 6). The channel has a relatively linear
morphology and is 100 – 150 m deep. Terraces are
more strongly developed on the eastern margin of the
canyon (Fig. 6b), suggesting an increased rate of
incision and channel migration to the west.
4.3.4. Interfluve
The interfluve between the Western and Central
canyons has a width of 2 km at the canyon heads,
but broadens down-slope to 3.5 km at its widest
point. This plateau lies in water depths of ~200 m
and is typified by a low down-slope gradient (typi-
cally 1 – 28) towards the south-southwest. The inter-
fluve probably represents relict continental shelf
bounded to the east and west by the two canyon
systems (Profile 2, Fig. 6).
4.4. Relict canyons
Down-slope and to the east of the Eastern can-
yon, a broad terrace forms average slope gradients of
1 – 2.58, in water depths of 680 – 1200 m. This ter-
race is crosscut by three sub-parallel canyons (Fig.
7), trending west-northwest – east-southeast in their
upper reaches, and swinging north – south down-
slope. These crosscutting canyons occur in water
depths of 700 – 1100 m and their morphology is
atypical with respect to other Celtic Margin canyon
systems. They are U-shaped in cross-section, with
broad, sub-horizontal floors 1 – 1.5 km across and
300 – 350 m deep. Average down-slope canyon floor
gradients are in the region of 3 – 4.58. There is no
obvious incision of the canyon walls or canyon
floors. This suggests that these canyons represent
areas of sediment deposition.
Fig. 7. Sediment deposition and relict canyons. a) 3D Shaded relief image showing a spur bounded on the west by the Eastern canyon, and
cut by broad U-shaped canyons (Profile 1). West of the flow divide, these canyons are convex upward in longitudinal profile and to the east,
they are concave upward (Profiles 2 and 3). Towards the south, the spur steepens and has a convex upward slope; long wavelength, low
amplitude sediment waves are apparent (Profile 4). b) Bathymetric image of the same area overlain with an interpolated map of sediment
penetration based on 3.5 kHz profiles. Note the areas of high (light) and low (dark) penetration on the terrace surfaces and canyon floors,
respectively. c) 3D shaded multi-beam backscatter image showing areas of high backscatter in the canyon floors and low backscatter on the
terraces. CF - Canyon floor.
M.J. Cunningham et al. / Sedimentary Geology 179 (2005) 99–116 109
At the headward end of the canyons there is a
north – south drainage divide (Fig. 7). East of the
divide, canyon floors slope towards the west-south-
west (with spurs sloping mainly south) and steepen
towards the divide (canyon floor gradient V 68)forming a lip on the eastern margin of the divide
(Fig. 8). West of the divide, canyon floors slope
north-northeast at a more typical gradient of 3 – 4.58
Fig. 8. Sediment erosion and overspill deposition. Shaded relief image of transverse (relict) canyons and surrounding area. Side-scan sonar
images show that erosive processes appear to be dominant in the Eastern canyon and that sediment waves occur in the floors of the canyons.
CF - Canyon floor, D - Divide, GC - Gravity core.
M.J. Cunningham et al. / Sedimentary Geology 179 (2005) 99–116110
until they are captured by the walls of the neigh-
bouring canyon system. Spurs slope predominantly
southwest.
It is unclear if headward erosion of the canyons
has led to a progressive narrowing of the divide, or
whether they represent an older trend that is now
truncated by younger canyons where turbidity over-
flow has resulted in sediment deposition on the spurs,
with partial burial of the canyons. A study of canyons
to the east of New Jersey (western North Atlantic) has
shown similar patterns where abandoned, and partial
to fully buried, canyons are offset by modern canyons
(Pratson et al., 1994; Pratson and Coakley, 1996).
The canyon floors are associated with areas of low
3.5 kHz seismic penetration and high 95 kHz back-
scatter (Fig. 7b,c) which may imply basement subcrop
or relatively coarse sand. A gravity core from the edge
of the most northerly canyon head and a bottom
sediment grab from between the two southern canyons
(Fig. 7) consist of brown-grey, sandy, silty clay with
occasional shell fragments and correspond with high
penetration and low backscatter. In the canyon heads
to the east of the divide, the side-scan sonar shows a
series of relatively closely spaced sediment waves
(Fig. 8). The sediment waves are generally orthogonal
to the canyon axes and show an east – southeast sed-
M.J. Cunningham et al. / Sedimentary Geology 179 (2005) 99–116 111
iment transport direction. However, these sediment
waves are too closely spaced to be resolved by
multi-beam bathymetry (50 m cell resolution). To
the west of the divide side-scan sonar shows that the
Eastern canyon is swept clean of sediment with ex-
tensive areas of bedrock exposure (Fig. 8).
Side-scan sonar data show that the terrace surface
between the active and relict canyons is generally
featureless. However, to the southwest where the
terrace slope gradient increases to form a convex
northeast – southwest profile, (Profile 4, Fig. 7) there
is a series of relatively high relief, long wavelength
sediment waves (Fig. 8). The sediment waves have a
wavelength of 1500 – 1800 m and relief of ~30 m.
These features are approximately parallel to each
other, broadly follow the slope contours and have an
overall northwest – southeast orientation. The sedi-
ment wave crests are generally orthogonal to the
upslope canyon axes and coincide with areas of high
3.5 kHz seismic penetration (~10 to N80 m). In cross-
sectional profile they form a rounded, terraced mor-
phology where the sub-horizontal slopes have a gra-
dient of b18 towards the northeast and the steep
slopes face southwest with a gradient of 8 – 118(Fig. 8). 3.5 kHz seismic profiles reveal that these
sediment waves have a greater depositional thickness
on the sub-horizontal slopes facing the sediment trans-
port direction.
In summary, the high backscatter intensities and
low 3.5 kHz penetration of the canyon floors cou-
pled with the existence of closely spaced sediment
waves is opposite in character to the intervening
canyon spurs with no detectable sediment waves.
This shows that relatively coarse material is accumu-
lating in the canyons with fines being preserved on
the spurs, and may imply that along-slope processes
affect these relict canyons, with only the coarsest
material being preserved.
Toward the south of the relict canyons the terrace
surface is covered by a series of low relief, long
wavelength sediment waves which coincide with
high 3.5 kHz seismic penetration, locally in excess
of 80 m. Internally, these sediment waves have a
greater depositional thickness on the sub-horizontal
lee slopes, which face up-slope. These bedforms have
been detected and described in other areas (e.g., Wynn
and Stow, 2002) and are believed to have been de-
posited by unconfined turbidity currents.
5. Discussion
5.1. Sediment supply and transport
On the Celtic Margin, at the shelf/slope transi-
tion, sandwave fields occur immediately upslope of
major canyon heads (Fig. 9). The sediment wave
crests are mainly perpendicular to the canyon axes,
with lee slopes facing towards the canyon heads.
This suggests that sediment transport is towards the
canyon systems. In general, the sandwaves lie to the
northeast of the canyon heads and this, combined
with the direction of asymmetry of Set A sandwaves
(Fig. 2), ties in with the regional southwesterly active
sediment transport direction (Heathershaw et al.,
1987).
A relatively large field of sandwaves to the
northwest of the Brenot Spur (Set B, Fig. 2) is
more equivocal in its association with present-day
sediment transport into the canyon system. These
sandwaves are more rounded in morphology, and
where asymmetry exists the inferred sediment trans-
port direction is towards the northeast. It is there-
fore suggested that these bedforms represent relict
features that originally developed during the last
sea-level lowstand under a current regime different
to that of the present day, and are now being
reworked to equilibrate with the existing current
regime.
The source of the shelf edge sandwaves is probably
to the north-east where the tidally dominated Celtic
Sea area is characterised by a series of elongate, linear
sand banks with a maximum length of 200 km, width
of 7 km, and height of up to 60 m (Reynaud et al.,
1999). The sand bank crests are oriented broadly
orthogonal to the shelf edge and are at an average
water depth of 150 m (Fig. 1a, Berne et al., 1998;
Marsset et al., 1999). The origin of these sand banks is
uncertain.
Previous workers have suggested that the Celtic
Sea Sand Banks may be tidal, formed during the last
glacial sea-level lowstand when tidal velocities were
high enough to generate them (e.g. Pantin and Evans,
1984; Berne et al., 1998; Zaragosi et al., 2000).
Alternatively, the sand banks may be the remnants
of a shallow marine delta channel-levee system
formed at the mouth of the English Channel River
(Marsset et al., 1999; Droz et al., 2003).
Fig. 9. Inferred sediment transport direction on shelf-break. Northeast shaded bathymetric image from ~150 to 400 m bathymetric contour
showing shelf-break and heads of major north-northeast-south-southwest trending canyons. The data is artificially cut off at 400 m to capture the
canyon heads and associated sandwaves. There is an obvious association of sediment waves at major canyon heads, with inferred direction of
sediment transport being normal to the canyon axes. In most cases, sediment transport is towards the shelf-break (canyon heads), but there is
also a series of sub-symmetrical to weakly asymmetrical sandwaves showing an inferred transport direction to the northeast. Given their
morphology (see Section 4.1), these are probably in the process of being reworked and originally formed during the last glacial low-stand. CC -
Central canyon, EC - Eastern canyon, FCC - Fault controlled canyon, WC - Western canyon.
M.J. Cunningham et al. / Sedimentary Geology 179 (2005) 99–116112
5.2. Structure
The character of the Celtic Margin has been
strongly influenced by seafloor spreading and the
opening of the North Atlantic Ocean (at around 53
Ma), Palaeocene – Eocene Alpine foreland deforma-
tion, and mid-Cenozoic Pyrenean orogenesis (Zieg-
ler, 1987; Naylor and Shannon, 1982; Tucker and
Arter, 1987; Naylor, 2001; Bourillet and Lericolais,
2003). These events have resulted in major fault
blocks bounded by north-northwest – south-southeast
trending, listric normal faults (de Graciansky et al.,
1985).
About 20 km west-northwest of the Brenot Spur,
there is a canyon of north-northwest – south-southeast
orientation (see Section 4.2). On the eastern margin of
the canyon, there is a steep, linear scarp with a relief
of 130 – 140 m. There is also a drainage basin on the
eastern margin, which is clearly offset by the fault and
where insufficient time has passed for the basin to re-
equilibrate with its boundary conditions. These obser-
vations strongly suggest that the canyon is actively
fault controlled (Fig. 4).
The age of the fault is difficult to constrain. The
last regional compressional event associated with
Alpine foreland deformation occurred between the
Oligocene and Miocene (Ziegler, 1987; Cook,
1987) and it is therefore likely that the fault is at
least of similar age. However, the present-day fault
scarp relief of N100 m and the fact that movement
post-dates the morphologically fresh drainage basin
on the eastern canyon margin, suggests a much
younger age. Supporting evidence comes from the
magnitude of downcutting and incision that has oc-
curred down-slope of this canyon (Fig. 10). The
trend of the fault is consistent with other relatively
young, north-northwest – south-southeast trending
faults known from around the British Isles, which
include the Neogene Sticklepath fault of SW Eng-
land (Bristow and Hughes, 1971) and the Codling–
Newrey faults of the Irish Sea (e.g., Geoffroy et al.,
1996).
Fig. 10. 3D shaded relief of the entire Celtic Margin from the shelf to the foot of the continental rise. Arrows show down-slope flow direction
based on slope aspect. BS - Brenot Spur, CC - Central canyon, EC - Eastern canyon, FCC - Fault controlled canyon, RC - Relict canyons, SW -
Sediment waves, WC - Western canyon.
M.J. Cunningham et al. / Sedimentary Geology 179 (2005) 99–116 113
The interaction of sedimentary processes and
structure within the canyon is difficult to quantify,
particularly in the upper reaches where back-wall
retrogressive mass erosion of the eastern margin is
more developed when compared to the western mar-
gin. This greater development may also be, in part, a
function of bottom current movement along the Celt-
ic Margin, which would bring more sediment to the
eastern margin.
5.3. Upper to middle slope sediment deposition
In cross-sectional profile, the canyons of the Celtic
Margin are primarily V-shaped at the shelf edge,
becoming increasingly flat-bottomed to U-shaped
down-slope. The Western canyon (Fig. 6) is a good
example of this change in morphology since it is
strongly V-shaped in its upper reaches close to the
shelf-break, but becomes progressively more U-
shaped with increasing depth. A first-order interpre-
tation of this pattern is erosion dominating the upper
reaches and deposition on the lower reaches. How-
ever, it may also be that canyon morphology and
sediment flows are presently in a state of equilibrium,
and that the canyons act as bypass conduits carrying
sediment to the Celtic Sea Deep Fan which is pres-
ently dominated by low-density, muddy turbidity cur-
rent deposits (Zaragosi et al., 2000, 2003). This would
imply that flows are of insufficient size to overspill the
canyon system and that the deposition of fines on
spurs is probably related to background, terrestrially
derived sediment.
Spurs vary from being sharp-crested to relatively
flat in transverse profile, but are usually convex up-
ward in longitudinal profile and in some cases are
crossed by relict canyons that have been abandoned
and subsequently incised by newer canyon systems.
The Eastern canyon appears to be less incised than
the Central and Western canyons. This could be due to
differences in current velocities, sediment supply and/
or substrate. Alternatively, it could be related to the
development of the fault controlled canyon system to
the east, which has captured the sediment supply that
once fed this canyon.
The terrace east of the Eastern canyon is crossed
by three U-shaped canyons, with a drainage divide
near its western margin (Fig. 7). The multi-beam
bathymetry, backscatter and lack of 3.5 kHz pene-
tration imply relatively coarse material on the canyon
floors, suggesting that they are swept clean of finer
material. The U-shaped profile of these canyons
combined with the occurrence of bedforms on the
canyon floors (close to the drainage divide) supports
this proposal.
The unconfined turbidite deposits forming sedi-
ment waves towards the southern end of the terrace
M.J. Cunningham et al. / Sedimentary Geology 179 (2005) 99–116114
are likely to have originated upslope in the canyon
heads. There is ample evidence from faulting,
drainage basins (bamphitheatre rimsQ), deep incision
and changes of sinuosity (particularly incised mean-
ders) for the retrogressive mass wasting and inden-
tion of the canyon heads, leading to indention of
the shelf-break. Over-steepening of the canyon
headwalls and drainage basins will lead to slope
failure, which will initiate coarse-grained gravity
and turbidity currents that downcut and deepen
the canyon system. The high rates of incision at
canyon headwalls, particularly the Western canyon
(Fig. 6), show that there may also have been a
recent influx of coarse-grained sediment (probably
at the end of the last glaciation), leading to frequent
turbidity currents (although the possibility of slow
rates of incision that have been ongoing over long
periods of time cannot be discounted). Down-slope,
the canyon walls will subsequently over-steepen
and fail, leading to a widening of canyons. There
is little deposition in the main canyons, although
sediment waves, which occur on many interfluve
areas, suggest that down-slope turbidity currents
spill out of the canyon system and form over-
bank deposits on the terrace/spur surfaces (Fig. 10).
Hence, we suggest that recent mud-rich flows have
been of sufficient size to overspill the canyon system
leading to fine sediment deposition on the continental
slope.
Our observations are in agreement with previous
studies (Zaragosi et al., 2000, 2003). For example, the
Celtic Deep Sea Fan, which spreads out at the foot of
the continental rise at water depths of 4200 – 4900 m
(Zaragosi et al., 2003), is a mature, mud-rich system
(Zaragosi et al., 2000; Droz et al., 2003) with a
depositional history commencing in the Miocene
(Droz et al., 1999). Isotope studies show that much
of the fan material has been sourced from high-density
turbidity currents (forming sand-rich deposits) during
glacial low-stands, whereas at present, fan sediments
are derived from the outer shelf by low-density turbi-
dite currents (forming mud-rich deposits) (Zaragosi et
al., 2000).
The Celtic Sea Sand Banks on the continental shelf
are the obvious source for fine muds reaching the fan
system, whilst very fine sands, silts and clays are
deposited further upslope on the spurs as canyon
overspill unconfined turbidite deposits.
6. Conclusions
Multi-beam bathymetry and backscatter, 3.5 kHz
pinger profiles, side-scan sonar and seabed samples
have been integrated to evaluate along- and down-
slope sedimentary processes along the Celtic Margin.
The main conclusions are:
1) Asymmetrical sandwaves occur along the Celtic
Margin shelf-break, are orthogonal to the canyon
axes, and are inferred to demonstrate sediment
transport into the canyon heads.
2) Near symmetrical sandwaves show well-rounded
crests with no obvious association to canyon heads
and occur along the Celtic Margin shelf-break.
Where asymmetry exists, the inferred sediment
transport direction is away from the shelf edge.
3) These dsymmetricalT sandwaves probably formed
prior to the end of last glacial low-stand at ~14 ka
BP. The smaller set of superimposed sandwaves
indicates progressive reworking through present-
day bottom current activity. Alternatively, the sym-
metrical sandwaves may reflect temporal variations
in north-eastward moving storm events.
4) Active down-slope sediment transport in the form
of turbidity currents is the dominant process in the
upper reaches of canyons characterised by deep
incision of sinuous thalwegs into the canyon floors.
5) Drainage basins occur along the margins of the
canyon heads and are interpreted to result from
retrogressive mass wasting and indention of the
Celtic Margin shelf.
6) Active faulting appears to control the development
of a north-northwest – south-southeast canyon. The
canyon shows rapid and deep down-cutting of the
continental shelf down-slope of the fault and off-
sets earlier canyon systems.
7) Earlier west-northwest – east-southeast canyons lo-
cated mid-slope are now areas of sediment depo-
sition. Along-slope sediment transport of fines
occurs in relict canyons that are sub-parallel to
the shelf-break. This process has led to the forma-
tion of closely spaced sediment waves.
8) Active canyons are dominated by sediment trans-
port in the mid to lower slope, but over-bank spill
of turbidity currents has led to the deposition of
muds and clays. Further down-slope, these form
unconfined turbidite deposits.
M.J. Cunningham et al. / Sedimentary Geology 179 (2005) 99–116 115
9) Canyons are V-shaped in the upper reaches and
become U-shaped progressively down-slope. This
may indicate a transition from erosive to deposi-
tional down-slope processes, or that the canyons
act as bypass conduits carrying sediment to the
abyssal plain.
From this work, it appears that the Celtic Sea
Sand Banks of the Celtic Margin have a genetic
link to the Celtic Deep Sea Fan at the foot of the
continental rise. The main agent of sediment trans-
port in the canyon heads is by slope failure and the
seaward migration of sandwaves whose sediment is
sourced from the Celtic Sea Sand Banks. These will
initiate gravity currents, which down-slope, will in-
cise, deepen and widen the canyons. However, over-
bank sediment deposition leads to along-slope depo-
sition in relict canyons and on spurs and terraces,
which further down-slope form unconfined turbidite
deposits.
Acknowledgements
This study was made possible through funds pro-
vided by NERC (grant ner/t/s/2000/01013). We
would also like to thank our project partners, BT,
Flag Telecom, Gemini, Global Crossing Systems,
Global Marine Systems and Tyco Telecommunica-
tions for allowing us access to their submarine cable
data. We thank Fugro Survey and Svitzer for unarch-
iving and supplying us with certain cable route
survey data. We wish to thank Sean Cullen, Archie
Donovan, Mick Geoghagan and Roger Sweetman of
the Geological Survey of Ireland for allowing us to
use data from the Irish National Seabed Survey and
Dr. Peter Croker of the Petroleum Affairs Division of
the Department of Communications, Marine and
Natural Resources, Ireland, for allowing us to use
multi-beam bathymetric data. We benefited from
reviews by R. Brunt and S. H. Chough and fruitful
discussions with Dr Russell Wynn.
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www.elsevier.com/locate/sedgeo
Sedimentary Geology 17
Contrasting styles of shelf sediment transport and deposition in a
ramp margin setting related to relative sea-level change
and basin floor topography, Turonian (Cretaceous)
Western Interior of central Utah, USA
Chris M. Edwards*, David M. Hodgson, Stephen S. Flint, John A. Howell 1
Stratigraphy Group, Department of Earth Sciences, University of Liverpool, 4 Brownlow Street, Liverpool L69 3GP, United Kingdom
Received 9 April 2004; accepted 6 April 2005
Abstract
The Turonian lower Ferron Sandstone and Juana Lopez Members of the Mancos Shale in the Western Interior foreland basin
form a series of regressive shoreface to shelf depositional complexes sourced from the Sevier fold-thrust belt to the west.
Continuous, dip-parallel outcrop exposures around the northern and eastern edges of the Tertiary San Rafael Anticline (SRA) in
central Utah provide an insight into the origins of anomalous, marine mudstone-encased sandstones.
Lower Ferron strata consist of both coarsening-and-thickening upward and sharp-based shoreface successions, interpreted to
have been deposited as sequential highstand and falling stage systems tracts. Overlying these shallow marine deposits is a
surface that marks sediment bypass that is correlated down-dip into coarse-grained, cross-bedded sandstones, interpreted as the
deposits of channelised turbidity flows. These turbidites are considered to be the products of sustained hyperpycnal flows
exiting rejuvenated lowstand rivers based on sustained flow indicators such as dune-scale cross-bedding, absence of genetically
related slumps and delta front deposits in up-dip areas, surfaces of sediment bypass, coarser grain-sizes than the underlying
shoreface deposits, and regional palaeogeography. The flows are interpreted to have been ignitive in proximal areas down a
tectonically-induced slope in the uplifted area of the Farnham Dome. The change in depositional style from falling-stage to
lowstand systems tracts is attributed to increased rates of sediment supply to the shoreline, promoting hyperpycnal flows. A
flooding surface caps the interval and separates the lower Ferron Sandstone from the overlying Juana Lopez Member. This
interval in contrast, consists only of heterolithic, parallel and ripple-laminated graded sandstone–mudstone couplets interpreted
to have been deposited by turbulent, storm-induced geostrophic flows. The couplets are arranged into bundles and are
interpreted as parasequences, which in turn are arranged into a progradational parasequence set deposited during a period of
sea-level highstand.
0037-0738/$ - s
doi:10.1016/j.se
* Correspondi
E-mail addre1 Present addr
Norway.
9 (2005) 117–152
ee front matter D 2005 Elsevier B.V. All rights reserved.
dgeo.2005.04.011
ng author. Present address: ExxonMobil Exploration Company, 233 Benmar, Houston, Texas 77060, USA.
ss: [email protected] (C.M. Edwards).
ess: Department of Earth Sciences/Centre for Integrated Petroleum Research, University of Bergen, Allegt. 41, N-5007 Bergen,
C.M. Edwards et al. / Sedimentary Geology 179 (2005) 117–152118
Divergent sole mark and bedform palaeocurrents from both members suggest slight eastward deflection of the turbulent
flows off the proto-SRA. Further evidence for this topography is inferred from an absence of turbidites along the western side of
the structure. The inferred influence of local basin floor structures on sedimentation of these units has important implications for
structural models for this basin, for generalized sequence stratigraphic models applied to ramp-type foreland basin sequences
and for elucidating the origin of anomalous marine mudstone-encased sandstone bodies.
D 2005 Elsevier B.V. All rights reserved.
Keywords: Western Interior Basin; Turbidites; Hyperpycnal flows; Geostrophic flows; Sequence stratigraphy
1. Introduction
The Western Interior foreland basin formed as a
flexural response to Late Cretaceous thrust sheet load-
ing in the Sevier thrust belt to the west (Jordan, 1981;
Wiltschko and Dorr, 1983; DeCelles et al., 1995). The
combined effects of subsidence and high Cretaceous
eustatic sea-level (Haq et al., 1988) created an epicon-
tinental seaway that stretched over 5000 km from
Arctic Canada in the north to the Gulf of Mexico in
the south. Shorelines periodically advanced and
retreated into the basin in tandem with changes in
relative sea-level and sediment supply (McGookey et
al., 1972; Williams and Stelck, 1975; Kauffman,
1977). Basin physiography was represented by a shal-
low, eastward-dipping ramp and probable water depths
of no more than a few hundred ms (Kauffman, 1977).
In the central Utah sector of this basin, Late Cretaceous
palaeogeography was dominated by eastward prograd-
ing shorelines fed by rivers that transported sediment
from the uplifted Sevier orogen to eastward-tapering
shorelines in the Castle Valley area. These sediments
are now superbly exposed in the badland and mesa
topography of the Uinta and Piceance basins of central
Utah and western Colorado (Fig. 1).
Within this basin, a number of Late Cretaceous,
marine mudstone-encased sandbodies are found phys-
ically disconnected from their contemporaneous shor-
elines by distances of many tens to hundreds of
kiloms. The depositional origins of these enigmatic
units have been the focus of intense debate in recent
years where the absence of a regional physiographic
shelf-edge precludes the application of conventional
deep-water sediment distribution models (c.f. Posa-
mentier et al., 1991; Kolla and Perlmutter, 1993).
Upper Cretaceous isolated sandbodies of this foreland
basin, such as the Cardium Formation of Alberta (e.g.,
Bergman and Walker, 1987; Plint, 1988; Walker and
Plint, 1992) and the Shannon Sandstone of Wyoming
(e.g., Tillman and Martinsen, 1987; Walker and Berg-
man, 1993; Bergman, 1994), commonly show an
alignment parallel to the palaeoshoreline and consist
of upward-coarsening successions ornamented by a
variety of wave-and storm-generated bedforms. Early
explanations argued their origins as the products of
migrating, linear doffshore barsT. This interpretation
was largely based on observations from modern shelf
environments (e.g., Swift and Rice, 1984; Tillman and
Martinsen, 1987). However it was subsequently con-
sidered that this model could not adequately explain
long transport pathways, the transport of granular
sediment or the coarsening upward motif associated
with these genetic units (Walker, 1984; Plint, 1988).
The resultant explanation relied on changes in relative
sea-level to displace shorefaces long distances into the
basin during relative sea-level fall before isolating the
dislocated shoreface with a sea-level rise (e.g., Berg-
man and Walker, 1987; Plint, 1988; Walker and Plint,
1992; Walker and Bergman, 1993; Bergman, 1994).
In the Utah sector of the Western Interior Basin,
some isolated sandstones have been attributed to rela-
tive sea-level fluctuations (e.g., Taylor and Lovell,
1995; Van Wagoner, 1995; Hampson et al., 1999).
Thin, mud-rich sandstones of the Campanian Prairie
Canyon Member of the Mancos Shale have been inter-
preted as the products of up-dip palaeovalley incision
into sand-rich highstand shorefaces of the Blackhawk
Formation with coeval down-dip forced regression of
muddy shorelines (Hampson et al., 1999). Similar en-
visaged sea-level fluctuations and incised valley devel-
opment have been applied in the interpretation of the
Tocito sandbodies in the northern New Mexico seg-
ment of the basin (Jennette and Jones, 1995). An
alternative hypothesis used to explain the occurrence
of channelised deposits such as the Prairie Canyon
Member adopts an offshore-directed delta plume
Fig. 1. Distribution of the lower Ferron Sandstone and Juana Lopez Members of central Utah, southwestern United States. The interval forms a
westward thickening wedge towards a high-subsidence zone adjacent to the Sevier Orogen. (A) Sediments sourced from the orogen were shed
into the epicontinental Western Interior Seaway that stretched from northern Canada to the Gulf of Mexico (Modified from Kauffman, 1984 and
Eaton and Nations, 1991). (B) To the north lie the Uinta Mountains and, at the time of lower Ferron Sandstone deposition, the bVernal HighQ ofRyer and Lovekin (1986). (C) Middle Turonian strata rim the northern edge of the San Rafael Anticline and extend southwards and eastwards
along the foot of the Book Cliffs. To the west, the lower Ferron is the primary coal bed methane reservoir horizon in the Drunkards Wash and
Helper Fields sited within the Ferron Coal Trend (from Montgomery et al., 2001). Structural elements are taken from Witkind (1988).
C.M. Edwards et al. / Sedimentary Geology 179 (2005) 117–152 119
model associated with periods of sea-level highstand
(Swift et al., 1987; Chan et al., 1991; Cole and Young,
1991; Cole et al., 1997; Stevens and Chaiwongsaen,
2003; Patterson, 2005). This model assumes fluvial
effluent exiting the delta front to have a bulk density
greater than the ambient density of the marine water
into which it enters in order to create a hyperpycnal
underflow (Bates, 1953). A characteristic inverse to
normal grading of individual event beds as a result of
the waxing and waning flood hydrograph (e.g., Mulder
et al., 1998) have been reported in sandstones of the
Prairie Canyon succession (Stevens and Chaiwong-
saen, 2003; Patterson, 2005). The interpretation of
these thin bedded sandstone intervals as hyperpycnal
turbidites provides a further mechanism for the depo-
sition of basinal isolated sandstones.
Fig. 2. Chronostratigraphy of the lower Ferron and Juana Lopez
stratigraphy along Castle Valley. The lower Ferron Sandstone
(dVernal deltaT) is an older and northerly counterpart to the upper
Turonian upper Ferron Sandstone (dLast Chance deltaT) that was
sourced from an area to the southwest of southern Castle Valley
Ammonite biostratigraphy is from Kauffman et al. (1993). Termi-
nology of depositional units defined by Cotter (1975) is given in
small type.
C.M. Edwards et al. / Sedimentary Geology 179 (2005) 117–152120
This paper introduces an unusual succession that
encompasses well-exposed bstrandedQ sandstones be-longing to the Turonian lower Ferron Sandstone, with
contemporaneous shoreline strata located at higher
levels on the shelf. The studied section of the outcrop
extends from a road cut near the town of Wellington,
Utah to the town of Green River on Interstate High-
way 70, a down-dip distance of over 70 km and
permits the correlation of shallow marine strata with
shoreline-detached sandstones. The shallow marine
succession provides a reliable sequence stratigraphic
framework into which the deposition of the mudstone-
encased sandstones may be placed. Overlying depos-
its of the Juana Lopez Member are equally well
exposed, although this interval cannot be traced up-
dip into an equivalent succession of shoreface strata.
Nonetheless, their relative sea-level context can be
sufficiently established using existing regional corre-
lations. Given the reliance upon understanding rela-
tive sea-level as a control on the deposition of isolated
shelf sandstones, this information is critical to their
correct interpretation. In addition to the relative sea-
level understanding, this paper presents data that im-
plicate the role of intra-basinal structures in the depo-
sition of the lower Ferron and Juana Lopez Members.
1.1. Geologic context
The lithostratigraphic term dFerron SandstoneTrefers to Turonian-age fluvio–deltaic units extending
from southern Castle Valley to the westernmost Book
Cliffs in the area of Price, Utah (Lupton, 1916; Spie-
ker and Reeside, 1925; Hale, 1972). These deposi-
tional wedges thicken westwards in the direction of
the thrust belt and taper eastwards into the Tununk
Shale of the Mancos Shale (Fig. 2). Deposits of this
age record the earliest major shoreline regression into
the Western Interior Seaway in central Utah and are
internally characterized by higher frequency sea-level
cycles. One northeastward-prograding depositional
system was sourced from an area southwest of Castle
Valley, and a second, separate system in northern
Castle Valley delivered sediment from an area north-
west of Price to a shoreline located at the northern-
most tip of the San Rafael Anticline (Fig. 2). These
two separate successions constitute the informally
known dLast ChanceT (southern) and dVernalT (north-ern) deltas (Hale, 1972; Ryer, 1991). Detailed sedi-
.
mentological and stratigraphic studies have revealed
that the Vernal and Last Chance deltas were not
contemporaneous. The Vernal delta deposits constitute
an older and entirely separate regressive wedge from
the younger Last Chance delta complex (Cotter, 1975;
Ryer and McPhillips, 1983). This relationship was
confirmed by macrofossil dating using the ammonite
biostratigraphic framework of the Western Interior
Seaway where the lower part of the Ferron Sandstone
was determined to lie within the upper middle Tur-
onian biozone of Prionocyclus hyatti (Molenaar and
Cobban, 1991). The northern succession was infor-
mally renamed the dlower Ferron SandstoneT to estab-
lish a clear distinction from the dupper Ferron
SandstoneT, which refers to the deposits of the Last
Chance delta outcropping in southern Castle Valley
(Ryer and McPhillips, 1983). This paper focuses on
the deposits and stratigraphy of the Vernal delta or
lower Ferron Sandstone, the older and northern coun-
terpart to the well-known upper Ferron Sandstone.
The lower Ferron Sandstone was first described in
detail by Cotter (1975) who subdivided it into the
dClawsonT, dFarnhamT and dWashboard UnitsT. He
noted that sediment was sourced from the west and
northwest and delivered to a NE–SW oriented shore-
line that tapered southwards and eastwards to its
down-dip termination in the Tununk Shale (Fig. 2).
The Clawson Unit was considered to be a shallow
marine shoreface deposit that crops out west of the
San Rafael Anticline (SRA) and the younger Wash-
board and Farnham Units were ascribed to shelf and
tidal inlet depositional environments, respectively
C.M. Edwards et al. / Sedimentary Geology 179 (2005) 117–152 121
(Cotter, 1975). Ryer (1981) subsequently interpreted
the lower Ferron Sandstone as a storm-and wave-
dominated shoreline that was sourced from an area
to the north. Several authors have proposed the exis-
tence of an unconformity spanning early Cenomanian
to late middle Turonian, related to the uplift of an
intra-basinal culmination and possible source area in
the area of the Uinta Mountains (e.g., Weimer, 1962;
the dVernal HighT of Ryer and Lovekin, 1986; Fig. 1).
Merewether and Cobban (1986) and Molenaar and
Wilson (1990) concluded that the area of uplift was
centered in northwestern Colorado, southwestern
Wyoming and parts of northeastern Utah. The pro-
posed culmination was considered to be part of an
array of active mid-Cretaceous basin structures in the
Western Interior Seaway at this time (Merewether and
Cobban, 1986).
Riemersma and Chan (1991) and Gardner (1995)
attempted to place the lower Ferron Sandstone in a
relative sea-level context. Thick sandstones of the
Farnham Unit described by Cotter (1975) were related
to a relative sea-level fall in the Farnham Dome area
(Molenaar and Cobban, 1991; Riemersma and Chan,
1991; Gardner, 1995). The lower Ferron Sandstone of
central Utah constitutes part of a period of a wide-
spread regional regression across the southwestern
part of the Western Interior Basin. It is biostratigra-
phically constrained to be equivalent to shoreface
successions of the dRegressive Coastal SandstoneT ofthe Frontier Formation of the southern Uinta Moun-
tains of northeastern Utah (Hale and Van de Graaff,
1964; Molenaar and Wilson, 1990), an unconformity
at the base of the Dry Hollow Member in the western
Uinta Mountains (Ryer, 1977; Molenaar and Wilson,
1990) and to the dFerron SandstoneT of the Henry
Mountains (Peterson et al., 1980).
Less well understood are the overlying deposits of
the Juana Lopez Member, which crop out along the
base of the western Book Cliffs. Molenaar and Cob-
ban (1991) interpreted the thin mudstone and sand-
stone interbeds of this unit were derived from a source
area to the north and may have been deposited in an
outer shelf environment by storm-generated currents.
In a regional context, deposition of the Juana Lopez
Member was coeval with northeastward progradation
of the upper Ferron Sandstone along the western flank
of the San Rafael Anticline in southern Castle Valley
(Gardner, 1995) and incision into underlying regres-
sive shoreface sandstones of the Frontier Formation
on the south side of the Vernal High (Molenaar and
Wilson, 1990). At present however, no detailed anal-
ysis of depositional mechanisms or stratigraphic ar-
chitecture of this succession exist. Furthermore, its
stratigraphic relationship to the lower Ferron Sand-
stone has not been addressed.
The San Rafael Anticline around which the lower
Ferron and Juana Lopez deposits are exposed (e.g.,
Fig. 3), is traditionally interpreted as a latest Creta-
ceous to early Tertiary feature (Lawton, 1983; Law-
ton, 1986; Witkind, 1988; Franczyk and Pitman,
1991). However, speculations on pre-Tertiary tectonic
activity in this area have been well documented (e.g.,
Peterson, 1986; Eaton et al., 1990; Molenaar and
Cobban, 1991; Gardner, 1995; Martinson et al.,
1998). Peterson (1986) suggested that the SRA is a
remnant late Paleozoic structure that was activated
during the Early Jurassic to early Middle Jurassic,
based on isopach maps and distribution of these age
strata. Eaton et al. (1990) provided evidence of tec-
tonic activity from the identification of reworked
exotic clasts within the Mancos Shale that could
only be derived from older rocks deposited prior to
marine flooding of the Western Interior Seaway pro-
viding supporting evidence to observation by Mole-
naar and Cobban (1991) of marked differential
sediment thicknesses of Turonian strata across the
structure. Gardner (1995) identified a number of de-
positional disconformities within lower and middle
Turonian sediments that underlie the upper Ferron
Sandstone Member that he attributed to periodic tec-
tonic uplift. More recently, Martinson et al. (1998)
proposed the existence of subtle bathymetric relief
during Coniacian and Santonian times based on
cross-structure differential subsidence patterns and
detailed foraminiferal analysis. Elsewhere within the
Western Interior, similar intrabasinal structures are
believed to have had a syn-depositional influence on
sedimentary sequences (Merewether and Cobban,
1986; Schwartz and DeCelles, 1988; Heller et al.,
1993). Typically, the age of these culminations pre-
date or approximate to ages of Sevier-style thrusting
(Merewether and Cobban, 1986; Heller and Paola,
1989; Eaton and Nations, 1991). Structural modelling
studies of Heller et al. (1993) and Stephenson and
Cloetingh (1991) suggests that they may be the man-
ifestations of intra-plate contractional stresses that
Fig. 3. High altitude aerial photograph draped on digital terrain model of the Farnham Dome area at the northern tip of the San Rafael Anticline.
The lower Ferron Sandstone is superbly exposed in three dimensions in canyons and cliffs around the periphery of the anticlinal structure. The
core of the structure exposes Early Cretaceous and older rocks. Sections A and B are shown for reference. Vertical scale=5� horizontal.
C.M. Edwards et al. / Sedimentary Geology 179 (2005) 117–152122
emanated from the fold-thrust belt and became
uplifted along lines of crustal and lithospheric weak-
ness. These zones of differential uplift are either
equant or elongate features that have broad wave-
lengths (10s of km) and amplitudes ranging from
metres to a few 10s of metres (Merewether and Cob-
ban, 1986; Stephenson and Cloetingh, 1991; Heller et
al., 1993). Ultimately, many of these sites of reported
gentle intraforeland uplift eventually became large,
basement-involved Tertiary dLaramideT structures
(Schwartz and DeCelles, 1988).
2. Sedimentology
The facies associations introduced below are sub-
divided into those belonging to the lower Ferron
Sandstone Member and those of the overlying Juana
Lopez Member (Figs. 4 and 5). The lower Ferron
Sandstone Member comprises (A) Teichichnus-bur-
rowed siltstones which encase, or grade into deposits
consisting of (B) Bioturbated fine-grained sandstones,
(C) Hummocky cross-stratified sandstones, (D) Swa-
ley cross-stratified sandstones, (E) Coarse-grained,
cross-bedded sandstones and (F) Interbedded cur-
rent-rippled sandstones, siltstones and mudstones.
The Juana Lopez Member consists of a single facies
association comprising (G) Organic-rich laminated
shales and interbedded rippled sandstones. The reader
is also referred to Cotter (1975) and Riemersma and
Chan (1991) for sedimentological descriptions and
regional correlations of the lower Ferron Sandstone
around the northern San Rafael Anticline. Bioturba-
tion grade is reported using the scheme of Taylor and
Goldring (1993) where a numeric scale from 0 to 6
(lowest to highest) determines burrow intensity, bur-
row overprinting and the level of preservation of
primary sedimentary structures.
2.1. Lower Ferron sandstone member facies
associations
2.1.1. Facies association A: Teichichnus-burrowed
siltstones
The lower Ferron Sandstone is volumetrically
dominated by grey, massively-bedded, intensely bio-
Fig. 4. Example section through the lower Ferron Sandstone at Section A highlighting the principal facies associations and sequence
stratigraphic elements.
C.M. Edwards et al. / Sedimentary Geology 179 (2005) 117–152 123
Fig. 5. Key to symbols in Figs. 4, 7, 8, 9, 12, 13 and 14.
C.M. Edwards et al. / Sedimentary Geology 179 (2005) 117–152124
turbated, silty mudstone and muddy siltstone and
constitutes the bulk of the recessively-weathered
Tununk Shale. This facies association is pervasive
along the length of the outcrop belt. In more basin-
ward (i.e., eastward) settings the unit thins and the
sharp basal contact becomes increasingly diffuse to
gradational with sandier facies of the lower Ferron
Sandstone. Although dominantly siltstone, the asso-
ciation may coarsen to very fine-grained sandstone
or fine, laminated, dark grey or black claystones.
Thin bentonite ashes form cm-thick, white and or-
ange bands within the grey mudstones. The succes-
sion is easily identified by the relatively high
abundance and low diversity of ichnospecies. The
dominant species is Teichichnus and bioturbation
indices of greater than 4 throughout the succession
produce homogeneous siltstones commonly lacking
primary sedimentary structures and lamination.
Where rare lamination is preserved, faint hummocks
and symmetrical ripple laminations are identified as
subtly coarser bands of coarse silts or very fine
sands against the more commonly encountered
grey fine silts.
Fig. 6. Paralic facies of the lower Ferron Sandstone. (A) Intensely bioturb
sediment deformation features such as folds (arrow) and overlying dew
deposited in an offshore transition zone to lower shoreface setting. (B) T
scale) showing large-scale pillows and dewatering on the eastern flank of
truncated (hashed line) by subsequent flow events. (D) Bioturbated bed at t
interpreted as a surface of submarine sediment bypass containing a high d
(arrow). (E) Thick accumulation of swaley cross-stratified (SCS) sandston
SCS sandstones are underlain by an erosive surface, bearing, large, metre
2.1.2. Interpretation
These bioturbated siltstones were deposited from
suspension in a fully marine, offshore shelf environ-
ment when deltaic systems were confined to western
and northern areas adjacent to the thrust front. Distal
deltaic plumes and/or storm-induced agitation of fine-
grained terrigenous material in proximal areas trans-
ported these sediments basinward in suspension.
Thus, these sediments are interpreted to have accu-
mulated below storm wave base in a low-energy
marine setting. Very distal influences of oscillatory
currents and minor sediment reworking by storm
waves higher in the shoreface profile transported
slightly coarser grained clastic sediment offshore, de-
positing it as thin ripple-laminated layers or fine
hummocky cross-stratified bedforms. Well-circulated
waters encouraged bioturbation, giving rise to a
churned substrate and the ubiquitous obliteration of
bedding surfaces. The Teichichnus trace fossil is com-
monly associated with offshore environments and is
thought to be the deposit feeding trace of a worm-like
organism in unconsolidated substrates (Pemberton et
al., 2001). Its occurrence affirms a distal shelf envi-
ronment interpretation.
2.1.3. Facies association B: bioturbated fine-grained
sandstones
Much of the lower Ferron stratigraphy along the
western and eastern sides of the SRA is composed of
a monotonous succession of well-sorted, fine-
grained, intensely bioturbated sandstone with a fria-
ble weathering character. This homogeneous unit
may reach thicknesses exceeding 8 m that grades
southeastwards (i.e., down-dip) into bioturbated silt-
stones of facies association A and northwestwards
(up-dip) into numerous hummocky cross-stratified
(HCS) sandstones (facies association C) and is
encased by bioturbated siltstones of facies association
A. The basal surface of the association is sharp and
ated sandstones of facies association B. Note the abundance of soft-
atering structures. These sandstones are interpreted to have been
hick, hummocky cross-stratified (HCS) bed (1.5 m Jacob’s staff for
the Farnham Dome. (C) Deformed HCS sandstones with top surface
op of lower shoreface sandstones (facies association C) at Section B,
iversity assemblage of ichnofauna, including large Rosellia burrows
es of facies association D (western flank of the Farnham Dome). (F)
-wide gutters at Section A.
C.M. Edwards et al. / Sedimentary Geology 179 (2005) 117–152 125
C.M. Edwards et al. / Sedimentary Geology 179 (2005) 117–152126
flat above underlying siltstones, whilst the upper
surface is flat to gently undulatory. Dominant ichno-
fauna include Thalassinoides, Ophiomorpha, Palaeo-
phycus, Teichichnus and accessory Arenicolites.
Bedding is poorly developed as a consequence of
ubiquitously high bioturbation indices, commonly
attaining index values of 5 or 6. However, several
faint bedding surfaces are present defined by thin,
cm-thick, discontinuous siltstone interbeds. Where
primary lamination is better preserved, the unit dis-
plays widespread and complex internal deformation
with up to metre-wavelength and decimetre ampli-
tude recumbent folds and abundant water escape
structures (Fig. 6A). Elsewhere, plane-parallel lami-
nation or low amplitude HCS is very rarely and only
partially observed.
2.1.4. Interpretation
This facies association is interpreted to have been
deposited in an offshore transition zone setting beneath
fair-weather wave base. The presence of HCS indi-
cates deposition by episodic storm events and rework-
ing and transport of sediment higher to the distal
reaches of the shoreface profile (Walker and Plint,
1992). Rapid deposition of sands above a substrate
of low strength (most likely siltstones of facies asso-
ciation A) promoted soft-sediment deformation and
rapid dewatering. Variations in bioturbation intensity
in lower shoreface bedforms possibly reflects an in-
verse relationship between storm magnitude and/or
duration between successive storm events and the
degree of biological mixing of the substrate (Dott
and Bourgeois, 1982). Intense bioturbation, reflected
also in the underlying Teichichnus-bioturbated silt-
stones, is likely to have obliterated original bedding
surfaces and internal primary stratification. Homoge-
nization of these beds may therefore be a product of
low magnitude, high frequency storm events, deposit-
ing numerous hummocky cross-stratified sandstone
beds, probably no thicker than a few tens of centi-
metres, with sufficient intervening time for complete
biogenic reworking, in a similar manner to those de-
scribed from the Fulmar Formation of the North Sea
(Howell et al., 1996). Further, high diversity ichnofau-
nal assemblages are regarded as an indicator of water
oxygenation (e.g., Bottjer et al., 1986; Savrda et al.,
1991) that may be promoted by storm surges (Brom-
ley, 1996). The homogenization and relatively thick
accumulation of this association distinguishes these
sandstones from more typical heterolithic lower shore-
face to offshore transition zone strata as predicted by
existing facies models (e.g., Walker and Plint, 1992).
These facies are particularly prevalent along the west-
ern side of the SRA and are characterized by numerous
calcite-cemented botryoidal concretions that were de-
termined by McBride et al. (2003) to have precipitated
in the presence of interstitial brackish water pore
fluids, the significance of which is discussed later.
2.1.5. Facies association C: hummocky cross-strati-
fied sandstones
Sandstones of facies association B grade north-
westwards (i.e., up-dip) into interbedded fine-grained
hummocky cross-stratified (HCS) sandstones and bio-
turbated siltstones. The sandstones are sharp-to ero-
sively-based, light brown or beige-coloured and are
broadly sheet-like occurring as beds that pinch and
swell along their lengths. Bed thicknesses range from
15 cm in the troughs of swales thickening to 60 cm at
the crest of the hummocks. Bioclastic material is
contained within these beds and consists of broken,
disarticulated, thick-walled bivalve and occasional
ammonite debris. Bed tops are sharp and may be
characterized by symmetrical and less commonly
asymmetrical ripple laminae. Internal laminae of the
HCS beds are generally well preserved. Soft-sediment
deformation features are pervasive around the periph-
ery of the Farnham Dome in the north of the study
area. These include large pillows of up to several
metres in width in individual beds up to 1 m thick
(Fig. 6B) and beds with convoluted, overturned lam-
inae. In the latter, the top surface of the bed has been
truncated by subsequent flow events (Fig. 6C). The
basal surfaces are erosive and commonly ornamented
with prominent, cm-amplitude gutters up to 5 cm
long. In these units Thalassinoides, Planolites,
Palaeophycus, Rosselia, Asterosoma and Teichichnus
species predominate, displaying a bioturbation grade
up to 4 with bioturbation intensity increasing towards
the top of the unit. In addition the uppermost burrows
may be filled by a red-weathered (?ferric) cemented
sandy fill (Fig. 6D). In sections where bioturbation is
considerably lower, species diversity is restricted to
sparse vertical Skolithos extending down from the top
surfaces and horizontal Planolites burrows. Typical
bioturbation values for these beds range from 1 to 2.
C.M. Edwards et al. / Sedimentary Geology 179 (2005) 117–152 127
2.1.6. Interpretation
Hummocky cross-stratified sandstones are indica-
tive of deposition in the lower part of a wave-or storm-
dominated shoreface (Walker and Plint, 1992). Depo-
sition of continuous sheet sands occurs during storms
when large-scale combined oscillatory flows rework
sediment into hummocky bedforms connected by low-
relief swales with mildly erosive bases and low angle
internal laminae (Dott and Bourgeois, 1982; Harms et
al., 1982; Southard et al., 1990). Wave rippled tops of
the sandstones is a result of reworking from waning
oscillatory currents following storm events (Cheel and
Leckie, 1993). Similarly, asymmetrical ripples pro-
duced by bedload transport of grains under uni-direc-
tional currents are frequently reported in storm
deposits (e.g., Walker, 1984; Brenchley, 1985; Leckie
and Krystinick, 1989). Inter-storm periods are repre-
sented by lower energy, suspension deposition of Tei-
chichnus-burrowed siltstones. The upward thickening
of HCS sandstones and the reduction in thickness of
interbedded siltstones reflects basinward shoreface
progradation. The ichnofaunal assemblage affirms a
shallow marine interpretation of these facies and var-
iations in bioturbation intensity may reflect variations
in the duration of successive storm events, where
longer intervals between storm events leads to a greater
intensity of bioturbation (Dott and Bourgeois, 1982).
2.1.7. Facies association D: Swaley cross-stratified
sandstones
Thick sandstones, up to 10 m thick and consisting
of amalgamated swaley cross-stratified sandstones,
are found only at Section A in the northernmost part
of the study area (Figs. 6E). The base of this arenitic
sandbody is highly undulatory and erosive with
metre-wide, sand-filled gutters (Fig. 6F). Swaley
cross-stratification is characterised by concave up-
ward swales, 10 to 50 cm thick, containing sub-par-
allel to slightly divergent internal laminae. Instead of
passing laterally into convex-up hummocks the struc-
ture is truncated by overlying swales, producing a
thick, continuous succession of fine-grained swaley
cross-stratified (SCS) sandstone. A coarse lag of
abundant bivalve and inoceramid fragments, sharks’
teeth, mudstone intraclasts and scattered chert pebbles
line the basal surfaces of the swales. Ichnospecies are
limited to isolated occurrences of Ophiomorpha and
Skolithos and index values are never greater than 1. A
thick succession of these sandstones at Section A
corresponds to the Farnham Unit as described by
Cotter (1975) and the damalgamated hummocky
cross-stratified sandstone faciesT of Riemersma and
Chan (1991). Intense bioturbation by Palaeophycus,
Ophiomorpha, Skolithos, and Thalassinoides in the
upper 3 m of the sandstone has removed most of the
primary stratification leaving behind a homogenised
sandstone cap. Faint concave-up bedding plane sur-
faces within this bioturbated upper part indicate that
these sandstones were probably deposited as SCS
bedforms. Southeastwards (i.e., down-dip) these sand-
stones pass gradationally into sandstones of facies
association C.
2.1.8. Interpretation
Swaley cross-stratification is associated with storm
events (Leckie and Walker, 1982; Hettinger et al.,
1994) and may be a highly amalgamated form of
hummocky cross-stratification whose occurrence is
commonly associated with storm-dominated shore-
faces (Dott and Bourgeois, 1982; Brenchley et al.,
1986; Walker and Plint, 1992). Consequently this
facies association is interpreted to have been deposit-
ed in a proximal lower shoreface to middle shoreface
position (Walker and Plint, 1992) with sands and shell
debris transported basinward from higher on the
shoreface profile. The abundance of shelly debris
and occasional coarse clasts suggests that the coastline
may have been gravelly or rocky. The erosive basal
contact between underlying offshore siltstones and the
overlying SCS sandstones is interpreted as the effects
of wave ravinement of the sea-floor in relation to an
abrupt shallowing of water depth (e.g., Plint, 1988).
The heavily bioturbated upper 4 m of the unit reveals
intense biogenic reworking by shallow marine ichnos-
pecies towards the end of deposition of this associa-
tion and implies a depositional hiatus towards the top
of the succession.
2.1.9. Facies association E: coarse-grained, cross-
bedded sandstones
Encased within Mancos Shale are coarse-to gran-
ular-grained, poorly to moderately well-sorted, peb-
ble-bearing arenitic sandstones (Fig. 7A). These
sandstones are best exposed west of the Sphinx rail-
way siding (Sections K, L and M) in the southern part
of the study area (Fig. 7B). Mapping of the associa-
C.M. Edwards et al. / Sedimentary Geology 179 (2005) 117–152128
C.M. Edwards et al. / Sedimentary Geology 179 (2005) 117–152 129
tion with integrated palaeocurrent information reveals
a single, low-relief linear body up to 4 km wide
encased within the Mancos Shale, the southern edge
of which is exposed at Section N (Fig. 7C). The lateral
margins of this gently lensing unit show no interdig-
itation with other facies associations indicating gentle
incision into underlying siltstones of facies association
A (Fig. 7C). The basal surface of the unit is flat to
erosional and the uppermost surface is an equally
sharp contact with open marine mudstones. The unit
comprises irregular beds 1 to 40 cm thick that lens,
split and amalgamate and are intercalated with lesser
thicknesses (up to 5 cm) of Teichichnus-bioturbated
siltstone (Fig. 7D).
The association has been removed in much of the
middle part of the study area due to present-day
erosion but reappears in the north where it consists
of thin, highly amalgamated and erosively-based beds
that reach a cumulative thickness of less than 10 cm
(Fig. 7E). In these localities 15 mm-long sharks’ teeth,
mudstone intraclasts, an exotic array of rounded cher-
tiferous, feldspathic and quartzitic granules, pebbles
and cobble-sized clasts show cross-stratification with
southeastward-dipping foresets. Despite the exotic
collection of pebbles present, these clasts are textur-
ally mature in roundness and sphericity. The unit
bears abundant bioclastic material including fragmen-
ted, disarticulated inoceramid shell debris and large
shark teeth in some instances exceeding 1 cm in
length. Bioturbation is generally absent but rare Pla-
nolites and large Thalassinoides networks occur at the
tops of sandstone bedding surfaces. Mean grain size
decreases southwards from granular, locally conglom-
eratic, pebbly sandstone in the northern outcrops
(Sections C, D and E), to coarse-grained sandstone
at Sections K to M. Sorting also improves southwards
from very poorly sorted in the north to moderately and
poorly sorted in the south.
In Section L the association thickens to a maxi-
mum thickness of 2 m. Here individual beds have
Fig. 7. Facies association E of the lower Ferron Sandstone. (A) Interpreted
encased by offshore siltstones of the Mancos Shale and forms a gently-lens
single 1-cm thick bed at Section N and shows no interdigitation with ot
splitting and amalgamation and contain many erosive surfaces (lines) to pr
sandstones are coarser-grained and poorly sorted (Section C) and show
characterised by erosively-based, ungraded, cross-stratified beds bearing co
sustained turbidity flows. (G) Bed amalgamation resulting from successive
bedded lower part beneath line, plane-parallel laminated upper part).
erosive basal surfaces (Fig. 7F), ornamented by mod-
erately intense tool markings. These structures include
spatulate flutes, discontinuous linear grooves and
broad u-shaped gutter casts, commonly exceeding 5
cm in width and revealing southward palaeoflow
directions. Beds exhibit high degrees of scour into
underlying sandstones, forming locally thickened,
amalgamated packages (Fig. 7D). The sandstone
beds form partially-amalgamated lenses that can be
traced over distances of up to 5 m to many tens of
metres. Cross-stratification is common, but some thin-
ner beds can be massive, plane-parallel laminated or
ripple cross-laminated. Highly irregular internal scour
surfaces are common, producing amalgamated sandy
packages of differing juxtaposed stratification (e.g.,
Fig. 7G). South-to southeastward-dipping centimetre-
wide, sand-rich foresets are defined by thin medium to
fine coarse-grained sandstone laminae that dip at
angle of repose and have tangential to asymptotic
bases indicating both planar and trough-cross stratifi-
cation. Coarser material and clasts of up to 7 cm in
diameter can be found at the toes of foresets. Al-
though some beds show gentle normal grading from
very coarse to coarse-grained sandstones, ungraded
beds are most dominant. In nearly all cases, the top
surfaces of the sandstones are always sharp with
overlying burrowed siltstones.
2.1.10. Interpretation
Erosional bases, dominance of tractional structures
and marine ichnofacies reflect deposition by numer-
ous turbidity flows in an offshore to outer shelf set-
ting, forming coarse-grained barforms. These deposits
have many parallels with the cross-stratified sands of
the turbidite classification scheme of Pickering et al.
(1986) such as mean grain size, sedimentary structures
and the unusually coarse grain size relative to bed
thickness. Scoured bases to the sandstone beds reveal
the erosional capacity of a waxing head of a turbulent
cloud prior to the aggrading, traction-dominated de-
graphic log from Section L near the Sphinx siding. (B) This unit is
ing sandbody over 4 km. (C) The lateral margin of this unit thins to a
her facies associations. (D) Individual beds show high degrees of
oduce thickened, sandy packages. (E) In proximal areas, equivalent
faint cross-bedding. (F) In more distal areas, the association is
bbles (encircled) and sharp tops and are interpreted as the products of
erosive flows juxtaposes contrasting sedimentary structures (cross-
C.M. Edwards et al. / Sedimentary Geology 179 (2005) 117–152130
positional phase. The assimilation of sharks’ teeth into
these deposits and the high degrees of scour and
amalgamation in the northern part of the study area
imply cannibalisation of the sea-bed by an erosive
flow prior to deposition (i.e., the flows were
bignitiveQ, sensu Parker, 1982). Up-dip thinning of
the unit is interpreted to reflect greater degrees of
submarine erosion and bypass as a function of greater
flow velocities.
Bedload deposition behind the head of the turbu-
lent cloud records the downstream migration of
straight to sinuous-crested dune forms producing
cross-bedding, and coupled with the absence of either
normal or reverse grading suggests that the flows were
sustained rather than unsteady and waning (sensu
Kneller, 1995; Kneller and Branney, 1995). Gentle
sediment grade changes with down-dip distance
have also been attributed to the effects of flow conti-
nuity (Plink-Bjorklund and Steel, 2004). The presence
of outsized cobble clasts lying many tens of km into
the basin indicates that the velocities of the flows were
also very high (at least 1.7 m s�1 given the modal
grain size of the sandstones and the sedimentary
structures present) and continuous over long dis-
tances. In contrast, there is little evidence for deposi-
tion by waning flows, such as composite upward
successions of progressively finer grain sizes and
lesser stability bedforms or Bouma sequences. How-
ever, interflow periods are represented by passive
suspension settling of thin siltstones that separate the
sandstones. These may become wholly or partially
removed during subsequent flow events.
Development of dune-scale bedforms is a relatively
unusual phenomenon in turbidite systems. Very large-
scale bedforms with wavelengths of tens to hundreds
of metres have been identified on the surfaces of
modern lower slopes and turbidite fans (Prior et al.,
1986; Bornhold and Prior, 1990; Hughes-Clarke et al.,
1990). It is likely that such forms require relatively
continuous, high velocity flows. Channels are an im-
portant morphological element in turbidite systems as
they contribute to enhancing the long-distance trans-
port capacity of coarse sands and pebbles by maximiz-
ing flow velocity (Postma et al., 1988). The lenticular
cross-sectional geometry of this facies association in
the vicinity of the Sphinx siding suggests that deposi-
tion occurred within a channel or chute. However, the
broad, shallow dimensions of this association is some-
what unusual and probably reflects the natural shal-
lowing and widening of submarine channels with
downslope distance (e.g., Clark and Pickering,
1996). At down-dip channel termini, turbidity flows
begin to expand and become net depositional. Exper-
imental analogues have indeed shown that at these
points of expansion, both sinuous and straight crested
bedforms may readily develop (Imran et al., 2002).
Although turbidity currents are known to be capa-
ble of transporting clasts of pebble size and greater,
resting positions at the toes of foresets suggest bed-
load transport. Coarse grain sizes can be maintained in
suspension by turbulence, buoyancy and hindered
setting support mechanisms (Pickering et al., 1986).
Clasts are not found bfloatingQ within beds, implying
that these facies were not deposited by non-turbulent,
laminar inertia-flow (or traction carpets) as simulated
experimentally (Postma et al., 1988). The position of
outsized clasts at the toes of foresets, together with an
absence of other distinguishing features of laminar,
plastic rheology flows (i.e., poor sorting, random
clastic alignment, massive beds displaying reverse
grading, plastic deformation features and so on, e.g.,
Major, 1997; Shanmugam et al., 1995; Shanmugam,
1996) indicates bedload-dominated deposition at the
base of a fluidal, turbulent flow of considerable dura-
tion rather than single pulse-like flow events. Further,
flows that are underladen with respect to their carry-
ing capacity, result in a predominance of tractional
process at the base of the flow and the production of
dune-scale bedforms (Pickering et al., 1986).
Hiscott et al. (1997) listed several criteria to ac-
count for the origin of sharp-topped turbidites, such as
the abrupt termination of deposition of an aggrading
bed, bimodal grain size distributions or longitudinal
grain size partitioning between the head and tail of the
flow. In addition, sharp bed tops might also reflect
flow-lofting processes when deposition at the base of
the flow causes the relatively low density interstitial
fluid to convect upwards, producing a buoyant plume
that lofts off the upper surface of the flow and result-
ing in the enhancement of grain size contrasts between
turbidite sandstones and overlying finer-grained units
(Sparks et al., 1993). Much of the finer-grained frac-
tion carried by the plume may then become dispersed
over wide areas (McLeod et al., 1999). Finally, abrupt
grain size contrasts may reflect the removal of finer-
grained material by flow-stripping over channel mar-
C.M. Edwards et al. / Sedimentary Geology 179 (2005) 117–152 131
gins, resulting in the enhancement of the lower, sand-
rich part of the flow (Piper and Normark, 1983;
Bowen et al., 1984). However, it should be stressed
that channel sinuosity has not been directly observed.
Consequently, whilst flow stripping is an integral part
of submarine channel evolution and fan growth (Hay,
1987), here the process is only inferred.
Riemersma and Chan (1991) and Molenaar and
Cobban (1991) proposed that this coarse sandy inter-
val represent a transgressive lag resulting from a rise
in relative sea-level. However there is no evidence for
lowered wave base and wave impingement on the sea-
floor prior to their deposition and the process of
winnowing subjacent deposits (siltstones) during
transgression would not produce the grain sizes ob-
served. In addition, terrestrial (e.g., single or multi-
storey fluvial sandstones, interfluvial palaeosols and
rooted horizons) and estuarine (e.g., clay-draped tidal
sandstones, bi-directional current indicators, bay and
lagoon mudstones) deposits are absent in these rocks,
precluding their interpretation as the products of val-
ley incision and filling.
2.1.11. Facies association F: interbedded current-rip-
pled sandstones, siltstones and mudstones
This facies association is found at Sections F
through J and again at O. Spatially, it can be found
laterally adjacent to the coarse-grained planar cross-
bedded sandstones. It consists of a 1 to 1.5 m thick
heterolithic succession comprising parallel-and cur-
rent-ripple laminated, very fine-to medium-grained,
normally graded sandstones interbedded with biotur-
bated siltstones (Fig. 8A and B). Sandstone beds
occupy 20 to 50% of the facies association. In sand
poor areas, such as at Section O (Green River) the unit
is populated by numerous concretions of decimetre
diameter. This association also thins northwards, but
rather than becoming coarser grained to the north, the
sandstones maintain a near-uniform grain size across
the entire outcrop belt.
Individual sandstone beds typically range in thick-
ness from 1 to 4 cm although rare larger beds reach
thicknesses of up to 9 cm. Bed geometries vary in
equal measure from discontinuous and lenticular, to
continuous and sheet-like, but bed amalgamation is
rare. Lenticular beds may form small, decimetre-long
ripple-laminated sandstones isolated in bioturbated
siltstones (Fig. 8C). In all cases, bed basal surfaces
are characterised by flat or gently scoured bases with
abundant and varied arrangement of grooves, flutes
and prod marks and indicate southerly flow directions.
These tool marks have mm-scale amplitudes and
widths and axial lengths of less than 5 cm. Typically
beds are ripple laminated, however the thicker beds
may consist entirely of plane-parallel laminations
(Fig. 8D) or upward graded parallel-to ripple-lamina-
tion. Asymmetric ripples are no more than 2 cm in
amplitude and greater than 10 cm in wavelength and
display south-to southeasterly-dipping foresets. Stoss-
side preservation is rare. Mean palaeocurrent values
show a 268 divergence between sole mark and ripple
bedform measurements (Fig 8E). Beds are normally
graded or ungraded. Where grading is observed, the
upward transition from sandstones into overlying silt-
stones is abrupt. In ungraded beds, the top surface of
the bed maintains asymmetrical current ripple bed-
form morphology and has a sharp contact with over-
lying siltstones. Trace fossils are generally rare within
the sandstones, but isolated examples of vertical bur-
rows (possibly Diplocraterion) and small crawling
traces are present.
2.1.12. Interpretation
Scoured bases, normal grading, and composite ar-
rangement of parallel and ripple-laminated sandstones
are all suggestive of transport and deposition by wan-
ing turbidity flows. These facies are comparable to thin
bedded sand–mud couplets facies in the classification
of Pickering et al. (1986). Episodic southward flows
appear to have transported fine-grained sand over large
areas. Planar-and ripple-laminated sandstones there-
fore correspond to Tb and Tc divisions of partial
Bouma sequences (Bouma, 1962) and indicate bed-
load sediment transport. Tbc sequences indicate that
these flows were capable of producing both upper flow
regime plane bedding and sub-critical lower flow re-
gime ripple-drift bedforms. The common presence of
lenticular bedding suggests fluctuations between sed-
iment supply and flow strength, most likely indicating
that flows were underladen with respect to their carry-
ing capacity. Tool marks at the bases of sandstone beds
exhibit the latest, pre-depositional scouring effects of
turbulent uni-directional flows prior to current-ripple
settling and migration from the base of the flow. The
variance in mean palaeocurrent directions taken from
these structures compared to palaeocurrents taken
Fig. 8. Facies association F of the lower Ferron Sandstone. (A) Interpreted graphic log from Section O near Green River. (B) Heterolithic
succession consisting of graded sandstone beds. Sandstones are erosively-based, plane-parallel and ripple laminated and are interpreted as the
products of waning turbidity flows. (C) Beds may be thin, lenticular and ripple-laminated or (D) thicker, more sheet-like and plane-parallel
laminated. (E) Palaeocurrent data show flows were towards the south and southeast. Sole structures show a southerly flow direction, whereas
palaeocurrents from cross-strata show flow directions towards the southeast (incorporates palaeocurrent data from facies association E; mean
values provided).
C.M. Edwards et al. / Sedimentary Geology 179 (2005) 117–152132
from ripples indicates subtly different orientation of
incident erosive flows from the subsequent deposition-
al flow. Mudstone interbeds probably reflect deposi-
tion out of suspension of fine-grained material
entrained within the tail of the flow. Since this facies
association can be found adjacent to facies association
E, it is considered to be the result of unconfined over-
bank or interchannel deposition. Collectively the two
associations constitute genetically-related depositional
elements (after Mutti and Normark, 1991) of a chan-
nel-overbank system.
These facies also have many indicative features of
tempestites. Despite their common association with
combined or oscillatory currents, storm-generated
C.M. Edwards et al. / Sedimentary Geology 179 (2005) 117–152 133
tempestite deposits may comprise turbidite-like beds
consisting of tool-marked bases and asymmetrical
ripples (Hamblin and Walker, 1979) with long, high
velocity run-out distances in excess of 100 km (Walk-
er, 1984; Brenchley, 1985; Leckie and Krystinick,
1989; Myrow and Southard, 1996) often parallel to
the shoreline trend (Walker and Plint, 1992). To this
extent, facies association F appears similar to current-
modified tempestites (Myrow and Southard, 1996).
However, the normal-to-shoreline palaeoflows and
synchronous deposition of laterally adjacent coarse-
grained sandstones are not compatible with deposition
by sheet-like, storm-generated currents.
2.2. Juana Lopez member facies association
2.2.1. Facies association G: organic-rich laminated
shales and interbedded rippled sandstones
Dark grey to black, clay-rich mudstones dominate
this facies association and are characterized by an
absence of bioturbation, making it distinct from
deposits of the underlying lower Ferron Sandstone.
Hence the basal surface of the association is distinc-
tive where it forms a textural contrast with underlying
sandstones and Teichichnus-bioturbated siltstones be-
longing to the lower Ferron Sandstone and is easily
identified. The finely laminated shale is extremely
fissile, contains rare, centimetre-thick sandstone
lenses and dominates the lower half of the member
(Figs. 9 and 10A). Southward thickening is observed
in the lower part of the interval from 3 m in north-
ernmost sections to 6 m in the vicinity of Green River.
The upper half of the member also comprises
fissile black shales but is interbedded with numerous
very-fine and fine-grained sandstones (Figs. 9 and
10A). Towards the top of this heterolithic succession,
the association weathers to a lighter brown colour and
becomes increasingly fossiliferous with abundant and
well-preserved ammonite, bivalve and gastropod de-
bris distributed along bedding surfaces. Individual
sandstone beds are typically sheet-like, 1 to 4 cm
thick and separated by shale interbeds of centimetre-
scale thicknesses. The internal fabric of the sandstones
is similar to those of facies association F, comprising
both current ripple-and plane-parallel laminated sand-
stones. Parallel-laminated beds are occasionally found
to grade upwards into current-ripple cross-stratifica-
tion and undulatory ripple-tops. Many of the ripple-
laminated beds occur as discontinuous decimetre-long
lenses. Bed bases are intensely scoured with prod,
groove and flute markings, providing reliable palaeo-
flow information (Fig. 10B). The geometries of the
sole marks in this succession display amplitudes and
widths of mm scale and wavelengths that are limited
to a few cm. Tool mark and ripple lamination palaeo-
current indicators repeatedly show discordance in ori-
entation values within individual bed readings. The
mean palaeoflow orientation attained from sole struc-
tures is 1728S compared to ripple lamination measure-
ments indicating south–southeast flows, averaging
1578S (Fig. 10C). Repetitive sandstone–shale couplets
occur as concentrated packages up to 3 m thick,
separated by papery shale horizons less than 50 cm
thick with widely distributed centimetre-thick, very
fine-grained, ripple-laminated sandstone beds (Fig.
10A). These packages may show coarsening-and-
thickening upward bedset trends, in addition to fin-
ing-and-thinning upward series of bedsets. Packages
become progressively sandier and thicker upwards
although much of the top of the Juana Lopez Member
is poorly exposed due to removal by recent erosion.
2.2.2. Interpretation
These facies are remarkably similar to those of
facies association F and thin bedded sand–mud cou-
plets of Pickering et al. (1986). However their attri-
butes are also remarkably similar to documented
examples of tempestites (Myrow and Southard,
1996). As such this association could be interpreted
either as turbidites or storm-related tempestites for
which an understanding of relative sea-level context
is of critical importance. Nonetheless, the recurring
nature of sand depositional events are interpreted as
the products of repetitive initial erosive flows that
subsequently wane and become depositional. Sole
structures reveal the scouring effects of southward-
directed flows at the waxing head of an erosive tur-
bulent cloud. The sandstone beds record the waning
of uni-directional flows, represented in the transition
from horizontally-laminated to ripple-laminated sand-
stones. These flows appear to have been dilute and
sediment starved as observed by the discontinuous,
ripple-laminated, sandstone lenses. The accumulations
of laminated claystones represent periods of pelagic
deposition between flows with little contamination
from clastic sediment input.
Fig. 9. Example measured section through the Juana Lopez Member at Section N illustrating the principal facies and their interpretations.
C.M. Edwards et al. / Sedimentary Geology 179 (2005) 117–152134
Outflow distances extend many tens of km south
and eastwards beyond the study area, attesting to
widespread, sheet-like sandstone deposition. Similar
to the observed disparity of orientation of directional
data in the lower Ferron deposits, a 158 mean differ-
ence between tool mark and ripple lamination palaeo-
currents reflects subtly differing orientations of the
initial erosive flow and subsequent deposition of
deflected bedload as flow energy dissipated. The bun-
dles of couplets representing numerous flow events are
interpreted as parasequences, punctuated by short-
lived recourses to mudstone deposition, interpreted
as flooding surfaces and parasequence boundaries.
2.3. Lower Ferron sandstone stratigraphic architec-
ture and chronostratigraphy
The sequence architecture of the lower Ferron and
Juana Lopez stratigraphy is summarised below and
illustrated in Figs. 11, 12, 13, 14 and 15. In northern
Castle Valley the lower Ferron Sandstone consists of
the shallowest water facies, namely facies associations
B, C and D (Fig. 11A) which thin southeastwards
until they eventually grade into marine siltstones of
facies association A (Fig. 11B). Southeastwards, these
facies grade into turbidite deposits of facies associa-
tions E (Fig. 11C) and F (Fig. 11D). The lower
Fig. 10. Facies of the Juana Lopez Member. (A) The heterolithic interval consists of numerous sandstone-mudstone bundles separated by mud-
rich sections and is interpreted as a parasequence. Individual sheet-like and lenticular beds are graded and plane-parallel to ripple laminated.
Parasequences thicken and coarsen upward through the succession. (B) The basal surfaces of beds are intensely scoured with abundant sole
structures. (C) Palaeocurrent data from sole marks show southerly flow directions whilst cross strata show flows to the southeast (mean values
provided).
C.M. Edwards et al. / Sedimentary Geology 179 (2005) 117–152 135
Ferron–Juana Lopez succession comprises one com-
plete sequence broken down into component systems
tracts. Initial highstand, falling stage, lowstand and
transgressive systems tracts define the lower Ferron
Sandstone whilst only a highstand systems tract repre-
sents the Juana Lopez Member (Fig. 11A–D).
2.3.1. Highstand systems tract
The lower Ferron shoreface deposits are character-
ized by both coarsening upward intervals of inter-
bedded HCS sandstones and burrowed siltstones
(facies associations A and B) and erosively-based
shorefaces consisting of amalgamated SCS sandstones
(facies association D). Normal progradation is repre-
sented by upward coarsening interbedded HCS sand-
stones and bioturbated siltstones that record the
systematic outbuilding of the shoreline as sediment
flux exceeded accommodation creation (Walker and
Plint, 1992). Termination of the coarsening upward
trend, and replacement by thicknesses of offshore
bioturbated siltstones indicates a rise in relative sea-
level and therefore defines a flooding surface. Each
cycle is interpreted as a storm-dominated parase-
quence (Van Wagoner et al., 1990). The vertical stack-
ing of successively thicker parasequences indicates
net progradation of the shoreline and defines the distal
portion of a highstand systems tract (Van Wagoner et
al., 1990) at the base of the lower Ferron succession
(Fig. 11A and Section A of Fig. 12).
2.3.2. Falling stage systems tract
Sharp-based middle shoreface SCS sandstones
scour into underlying offshore siltstones, and in their
most proximal reaches omit lower shoreface intervals
predicted by normal progradational facies models
(e.g., Walker and Plint, 1992). When traced palaeo-
Fig. 11. Outcrop expression of key stratigraphic surfaces. (A) Upward-thickening and coarsening sandstones of facies association C at Section
A define highstand systems tract parasequences bounded by flooding surfaces. The top of the cliff comprises a thick stack of sandstones of
facies association D. The basal erosive surface and abrupt juxtaposition of facies association D on facies association C is interpreted as a
regressive surface of marine erosion resulting from a lowering of relative sea-level. The overlying shallow marine sandstones are thus
interpreted as the products of a falling stage systems tract. (B) These sandstones grade down-dip into siltstones of the Mancos Shale. Overlying
highstand systems tract deposits of the net-progradational Juana Lopez Member are clearly seen in the top of the cliff. A sequence boundary,
characterised by submarine bypass and erosion, separates the falling stage and highstand systems tract. Photo is taken near Section D. (C)
Down-dip correlative strata to the sequence boundary thicken and comprise turbiditic sandstones of facies association E and (D) laterally
adjacent facies association F.
C.M. Edwards et al. / Sedimentary Geology 179 (2005) 117–152136
seaward, the amalgamated SCS sandstones grade into
HCS sandstones of facies association C (Figs. 11A
and B and Section B, Fig. 12). Shoreface sandstones
belonging to this genetic unit are represented over
several km further basinward than deposits of the
highstand systems tract. The long-distance regression,
omission of lower shoreface facies, erosive base
caused by wave ravinement and down-dip translation
Fig. 12. Stratigraphic correlation panel of the lower Ferron Sandstone in t
represented by deposition of highstand and falling stage systems tract
shoreface units and separate falling stage shoreface deposits form a thin s
transgressive systems tract strata are thus interpreted to downlap onto old
and are interpreted to onlap older shoreface strata in a landward direction.
the Farnham Dome structure (Section B). (B) Map provided for reference
into an apparently conformable succession of shore-
face strata defines an abrupt shallowing of water depth
associated with forced regression (Plint, 1988). Con-
sequently the basal surface is interpreted as a regres-
sive surface of marine erosion and the overlying
sandstones constitute a falling stage systems tract
(Plint, 1988; Hunt and Tucker, 1992; Posamentier et
al., 1992; Plint and Nummedal, 2000).
he northern part of the study area. (A) Initial shoreline regression is
strata. Deposition of lowstand turbidites overlies these regressive
uccession of transgressive systems tract deposits at Section A. The
er deposits. The lowstand turbidites thicken in a down-dip direction
Thinning of falling stage systems tract strata can be observed across
.
C.M. Edwards et al. / Sedimentary Geology 179 (2005) 117–152 137
Fig. 13. Stratigraphic correlation panel of the lower Ferron Sandstone in the southern part of the study area. (A) Lowstand systems tract strata
comprising facies associations E and F thicken basinwards. The lenticular distribution of facies association E and the absence of interdigitation
with facies association F suggest the presence of a broad, low-relief channel. Collectively the two associations constitute genetically-related
depositional elements of a channel-overbank system. (B) Map provided for reference.
C.M. Edwards et al. / Sedimentary Geology 179 (2005) 117–152138
2.3.3. Lowstand systems tract
Sandstones of facies association E can be traced
up-dip where they comprise a thin gravelly cap above
falling stage bioturbated sandstones (Sections C, D
and E, Fig. 12). Because of this relationship, the
turbidites of facies associations E and F are not con-
sidered contemporaneous with deposition of the un-
derlying forced regressive shoreface (Fig. 12). In their
justification of a fourfold systems tract model, Plint
and Nummedal (2000) argued that turbidites lying
basinward of forced regressive shorefaces constitute
lowstand systems tract strata since their deposition
post-dates forced regression. This criterion is adopted
here, and the lower Ferron turbidites are thus consid-
ered to represent a period of lowstand deposition
beyond the basinward limits of the underlying falling
stage strata (Figs. 13 and 15). Consequently the fall-
ing stage and lowstand systems tracts are separated by
an interpreted sequence boundary. In up-dip sectors
this sequence boundary is manifest as a submarine
bypass surface mantled by a thin gravel laggy deposit
of facies association E and intense bioturbation into
underlying falling stage shoreface sandstones (Sec-
tions A to E, Fig. 12). We stress that other features
associated with relative sea-level lowstands (incised
valleys, well-drained palaeosol horizons) are not iden-
tified, indicating that the lowstand shoreline did not
extend basinwards as far as the present day outcrop. A
downslope increase in the thickness of the channel–
overbank depositional elements may be attributed to a
gradient change, analogous to a transition from the
base of slope to the relatively flat basin floor and the
downslope transition from dominantly erosional to
dominantly depositional regimes.
2.3.4. Transgressive systems tract
The top of the lower Ferron succession consists of
a thin heterolithic succession comprising HCS sand-
stone and bioturbated siltstone interbeds. These facies
sit atop bioturbated sandstones capping the forced
regressive strata (Section A, Fig. 12). The return to
lower shoreface deposition at Section A denotes a rise
in relative sea-level and the presence of a flooding
surface at the base of a distal storm-dominated para-
sequence. The flooding surface merges with the un-
derlying sequence boundary and separates falling
stage deposits from the overlying transgressive sys-
tems tract deposits. Sand-poor, laminated mudstones
of the Juana Lopez that overlie the lower shoreface
interval, indicate a landward shift of the shoreline in
Fig. 14. Stratigraphic correlation panel of the Juana Lopez Member along the length of the study area. Progradational stacking of parasequences
indicates deposition under highstand conditions following transgression of the lower Ferron shoreline. (B) Map provided for reference.
C.M. Edwards et al. / Sedimentary Geology 179 (2005) 117–152 139
response to continued sea-level rise and are inter-
preted as the maximum flooding surface of the
lower Ferron succession.
2.4. Juana Lopez member stratigraphic architecture
and chronostratigraphy
2.4.1. Highstand systems tract
Deposition of Juana Lopez turbidites that overlie
the lower Ferron Sandstone can be separated into two,
upper and lower architectural divisions (Fig. 14). The
lower part of the succession is represented by contin-
ued mud-dominated deposition immediately follow-
ing the transgression at the end of lower Ferron
Sandstone time, with sandstones of patchy distribu-
tion confined to central and southern sectors of the
study area. The upper part of the Juana Lopez suc-
cession is characterized by northwest to southeast
progradational stacking of parasequences. This sea-
ward stepping architecture suggests basinward progra-
dation as sediment supply outpaced accommodation
creation associated with conditions of sea-level high-
Fig. 15. Chronostratigraphic diagram summarising the spatial and temporal changes in deposition of the lower Ferron and Juana Lopez
Members. Importantly, the deposition of facies associations E and F that constitute the lowstand systems tract of the lower Ferron Sandstone
correspond to a surface of non-deposition and bypass in up-dip areas. The absence of co-genetic slumped deposits suggests that a direct
connection between lowstand rivers and the deeper basin existed. Consequently these turbidites are interpreted to have been initiated by
hyperpycnal flows exiting flooded river mouths.
C.M. Edwards et al. / Sedimentary Geology 179 (2005) 117–152140
stand. Regional correlations reveal that on the western
side of the SRA the Juana Lopez Member corresponds
to outer shelf and turbidite-free mudstones of the
Washboard Unit in northern Castle Valley (Molenaar
and Cobban, 1991; Riemersma and Chan, 1991; Gard-
ner, 1995) and pronounced seaward stepping of the
upper Ferron delta in southern Castle Valley (Gardner,
1995). This seaward stepping was associated with
paralic and coastal plain accretion and stratigraphic
climb (Gardner, 1995), indicating the progradation of
the shoreline was coupled with rising sea-level typical
of highstand conditions (Van Wagoner et al., 1988;
Van Wagoner et al., 1990).
3. Discussion
The interpretation of sandstones deposited by tur-
bulent flows demands further discussion regarding the
propensity of shallow marine systems to deliver sand-
grade sediment via turbidity currents to offshore
environments over long distances during forced re-
gression, despite the absence of a clearly discernable
shelf break. Any depositional model for the lower
Ferron Sandstone needs to account for the initiation
mechanism for the generation of turbidity currents
during falling and lowstand of relative sea-level in a
basin lacking a shelf edge. Additionally, the signifi-
cance of lower Ferron and Juana Lopez shelf sand-
stone distribution in central Utah, and the possible
cause of variance in palaeocurrent directions between
sole marks and cross-stratification, are discussed.
3.1. Turbidite initiation mechanisms
It is widely regarded that turbidites form as a
consequence of either delta front failure and mass
gravitational sliding and slumping, or by hyperpycnal
underflows resulting from high sediment loads
debouching at flooded river mouths (Normark and
Piper, 1991). The type of flow generated by these
two separate mechanisms, and the types of bedforms
and facies that they produce, will determine flow
continuity and steadiness (Kneller, 1995; Kneller
and Branney, 1995). It is common belief that turbidity
currents are initiated when sea-level falls to, or be-
yond the continental shelf edge (Mutti, 1985; Mutti
and Normark, 1991; Posamentier et al., 1991; Weimer,
1991; Normark et al., 1993). Slides and slumps result-
ing either from earthquake seismicity or depositional
oversteepening of the delta front at the top of the
continental slope transform into debris flows, and
ultimately, surge-type turbidity flows of several
hours duration (e.g., Piper et al., 1988; Normark and
Piper, 1991; Mulder and Syvitski, 1995). Surge-type
flows are typically depletive as current mixing with
C.M. Edwards et al. / Sedimentary Geology 179 (2005) 117–152 141
ambient water results in a loss of density contrast,
sediment concentration and velocity and will produce
a predictable arrangement of facies that records initial
seabed erosion by the waxing head of the turbulent
cloud, followed by deposition of normally graded
sequences in the depletive tail of the flow (e.g.,
Bouma, 1962; Lowe, 1982). The absence of regional
shelf edge physiography inhibits the development of
failure-related turbidites in ramp margins such as the
Western Interior Basin, and as such, turbidites are
relatively rare.
Kneller (1995) and Kneller and Branney (1995)
showed that depletive, unsteady flows typical of
surge-type turbidity currents represent one flow struc-
ture scenario from at least 13 different possible flow
types. Furthermore, a number of recent studies have
suggested that turbidity currents may last for consid-
erably longer durations than the traditionally-held
waning-flow perception (e.g., Kneller, 1995; Kneller
and Branney, 1995; Mulder and Syvitski, 1995;
Mulder et al., 1998; Kneller and Buckee, 2000;
Mulder and Alexander, 2001; Parsons et al., 2001;
Mulder et al., 2003 and many others). This conclusion
is drawn from a wide range of approaches including:
(1) studies of highly sinuous turbidite channels
(Damuth et al., 1988), (2) direct observations and
sampling from lakes, high-latitude fjords and marine
basins (Weirich, 1986; Hay, 1987; Wright et al., 1988;
Prior and Bornhold, 1989; Chikita, 1990; Prior and
Bornhold, 1990; Zeng et al., 1991; Piper and Savoye,
1993; Piper et al., 1999; Kineke et al., 2000; Hicks et
al., 2004), (3) laboratory experiments (Rimoldi et al.,
1996; McLeod et al., 1999; Alexander and Mulder,
2002), (4) depositional and physical stratigraphic rela-
tionships (Plink-Bjorklund and Steel, 2004), (5) bed-
forms not compatible with waning flow processes
(Kneller, 1995; Kneller and Branney, 1995; Nemec,
1995; Mulder et al., 1998; Piper et al., 1999; Mellere
et al., 2002; Mutti et al., 2003; Plink-Bjorklund and
Steel, 2004) and (6) numerical models (Chao, 1998;
Imran et al., 1998; Kassem and Imran, 2001). The
maintenance of sustained flows is most commonly
attributed to the development of hyperpycnal under-
flows at river mouths, that may last for many hours,
days or even weeks (Mulder et al., 1998).
Hyperpycnal flows initiate when low salinity river
discharge enters marine waters with suspended sedi-
ment concentrations in excess of 36 kg m�3 (Mulder
and Syvitski, 1995). The excess density enables sed-
iment-laden flows to become negatively buoyant rel-
ative to the ambient, resulting in a long-lasting
seaward flow. From present day observations, these
conditions may be met following major river floods
carrying large volumes of sediment on a periodic
frequency ranging from b1 to N10 years (e.g., Mulder
and Syvitski, 1995; Johnson et al., 2001; Mulder et
al., 2001; Hicks et al., 2004). Some authors have
proposed that hyperpycnal flows are most common
during periods of sea-level highstand (Mulder and
Alexander, 2001) whilst others have argued that
they may be more frequent during sea-level falls
when large parts of the shelf become excavated
(e.g., Normark and Piper, 1991; Plink-Bjorklund and
Steel, 2004). Regardless, the relatively frequent oc-
currence at which these flows occur during present
day sea-level highstand suggests that ancient hyper-
pycnal flows ought to be more common than presently
documented.
3.1.1. Lower ferron sandstone depositional model
The following observations favour a hyperpycnal
flow origin for the lower Ferron turbidites: (1) the
presence of cross-bedded sandstones, (2) paucity of
genetically related slump and debris flow and delta
front/prodelta deposits, (3) surface of sediment bypass
in up-dip sectors, (4) change in dominant shoreline
depositional process, (5) flow channelisation and later-
al partitioning of sustained and waning flow conditions
and (6) regional palaeogeographic considerations.
3.1.1.1. Dune-scale cross-bedding. The relatively
recent recognition of facies pertaining to deposition
by riverine underflows has thus far revealed only a
limited array of types and arrangement of facies. The
most common of these are thick, ungraded sandstone
bodies that may be massive (e.g., Kneller, 1995;
Kneller and Branney, 1995) or plane-parallel laminat-
ed (e.g., Plink-Bjorklund and Steel, 2004). Mulder et
al. (1998) documented a hyperpycnite from the
Saguenay fjord, Canada that was characterised by a
graded motif from coarsening-upward at the base to
fining-upward towards the top. Waxing and waning of
current strength that gave rise to the coarsening up-
ward-fining upward succession was attributed to in-
creasing and decreasing river outflow during the
course of a flood. Plink-Bjorklund and Steel (2004)
C.M. Edwards et al. / Sedimentary Geology 179 (2005) 117–152142
suggested that hyperpycnal flow beds may be signif-
icantly more complex, and place emphasis on the up-
dip to down-dip variability of grain size trends and
arrangement of sedimentary structures that deviate
from the Saguenay model.
Mulder and Alexander (2001) proposed that climb-
ing ripples might readily form during hyperpycnal
flow events as a result of sustained flow durations
and sediment transport. Little comment has been made
regarding the propensity of these flows to deposit
larger bedforms such as dunes. However, Mulder
and Alexander (2001) suggested that the production
of dunes is likely to be relatively uncommon owing to
the longer durations required for their development.
The cross-bedded lower Ferron examples may howev-
er, represent an unusual example of dune-scale bed-
form development during the passage of turbidity
currents. Such prolonged bedload sediment transport
is consistent with the longer flow timescales associated
with hyperpycnal flows. However, because dune-scale
bedforms may also be associated with surge-type
flows, additional distinguishing criteria are required
to support a hyperpycnal flow interpretation.
3.1.1.2. Absence of slump, debris flow and delta front/
prodelta deposits. Turbidites deposited as a result of
slope failure are commonly associated with surge-type
flows. Under these circumstances, slumped and de-
formed sandstones ought to be found in association
with equivalent, down-dip turbidites. Whilst soft-sed-
iment deformation features are pervasive in the falling
stage shoreface strata of the Ferron Sandstone,
slumped sandstones that are genetically related to
deposition of (lowstand) turbidites are absent. A sim-
ilar paucity of deformed deposits has been noted from
hyperpycnal turbidites in the Spitsbergen Central
Basin (Plink-Bjorklund and Steel, 2004) and on the
seaward slopes of fjord deltas (Bornhold and Prior,
1990; Prior and Bornhold, 1990) and is argued to
reflect deposition by a hyperpycnal mechanism rather
than slope failure. Delta front facies (e.g., mouth bar
sandstones) are also absent in the lower Ferron stra-
tigraphy suggesting a direct connection between low-
stand rivers and the deeper basin.
3.1.1.3. Sediment bypass in proximal zones. The
presence of surfaces of sediment bypass in proximal
settings with down-dip thickening of turbidite facies is
analogous to the btype 1Q clinoforms described from
the Spitsbergen Central Basin (Plink-Bjorklund and
Steel, 2004). These clinoforms are characterised by
bypass from the shelf edge to the lower slope and
thick accumulations of turbidites from the base of
slope out onto the basin floor (Plink-Bjorklund and
Steel, 2004). Turbidite systems associated with these
clinoforms are associated with river systems that
debouched into channels (chutes) that cut into the
shelf edge and slope (Steel et al., 2000). Sand is
interpreted to have been deposited on the slope by
hyperpycnal flows (Plink-Bjorklund and Steel, 2004).
In the lower Ferron deposits, a similar arrangement of
facies is associated with up-dip erosion and bypass to
down-dip depositional thickening of turbidite beds.
Reconstructing palaeogeography (Fig. 16) illustrates
that the zone of sediment bypass, as represented by
thin gravel deposits of facies association E in areas to
the north and extensive bioturbation into underlying
falling stage deposits, occurs across the Farnham
Dome and Mounds Anticline. Both of these structures
are reported to have been present early in the evolu-
tion of the basin (Molenaar and Cobban, 1991; Wit-
kind, 1988). Ignitive flows are assumed to have
accelerated down such a slope. Whilst elevated depo-
sitional slopes are not a prerequisite for the initiation
of hyperpycnal flows, Plink-Bjorklund and Steel
(2004) argue that sediment transport beyond the
slope is more probable if a direct connection to a
river mouth exits. Furthermore, Prior and Bornhold
(1990) proposed that hyperpycnal flows at delta fronts
transporting sands and gravels are likely where the
bottom surface gradient of nearshore areas is greater
than the slope of the river thalweg. To date, channels
equivalent to those described in the Spitsbergen exam-
ples are not observable. The presence of inferred
topography may account for down-dip differential
thickness variability of the falling stage shoreface
strata (e.g., Section B, Fig. 12) and may have promot-
ed slumping and soft-sediment deformation within
these older units.
3.1.1.4. Change in dominant shoreline depositional
process. Highstand and falling stage shorelines of
the lower Ferron were dominated by storm and wave
processes (e.g., facies associations C and D), resulting
in the transfer of fine-grained sand away from river
mouths by longshore processes (e.g., Bhattacharya
Fig. 16. Palaeogeographic reconstructions for the (A) lower Ferron and (B) Juana Lopez Members. (A) The lowstand systems tract is
represented by turbidite deposition far into the basin, long distances from the underlying falling stage shoreface extent (shown). A submarine
channel is extrapolated northwards towards the Farnham Dome and Mounds area. These structures may have contributed a localised slope down
which flow accelerations took place, causing the turbidity flows to become ignitive. (B) Geostrophic flows transported and deposited Juana
Lopez sands to the east of the San Rafael Anticline from a source area to the north. Their absence to the west may reflect interaction with a
basinfloor structure in the position of the San Rafael Anticline. The inferred flow deflections are recorded by divergent palaeocurrent readings
between sole marks and cross-strata.
C.M. Edwards et al. / Sedimentary Geology 179 (2005) 117–152 143
and Giosan, 2003). By contrast, lowstand palaeogeo-
graphy was characterised by offshore sediment trans-
port with minimal coast-parallel sand transport.
Further, the transport of coarser material during the
lowstand interval implies a change in the hydrody-
namic regime, perhaps driven by external factors (e.g.,
hinterland tectonics, rate and magnitude of eustatic
sea-level fall, climate) enabling lowstand rivers to
carry greater concentrations of sediment and to cross
the density threshold required for the generation of
riverine underflows. The inception of riverine under-
flows restricted the lateral supply of sediment to
alongshore areas, resulting in a period of sediment
condensation in areas away from the sediment source
area. The interpreted change in depositional process is
consistent with fluid composition histories for
cements in the lower Ferron shoreface sandstones.
Dilution of saline pore fluids during the precipitation
of calcite-cemented concretions in the uppermost part
of the lower Ferron succession (McBride et al., 2003)
therefore probably reflects greater freshwater runoff
into nearshore areas of a relatively constricted low-
stand sea.
3.1.1.5. Flow channelisation. Models of turbidite
channel inception have centred upon progressive ero-
sion by several events (Clark and Pickering, 1996;
Imran et al., 1998) or by the transport and basal scour
by large glide blocks (Elliott, 2000). However, recent
laboratory experiments and numerical models have
shown that the very nature of continuous flows may
result in the self-generation of channels (Imran et al.,
1998; Alexander and Mulder, 2002). These models
illustrate the lateral distribution of net erosion and
C.M. Edwards et al. / Sedimentary Geology 179 (2005) 117–152144
deposition as underflows exit from a confined point
source (e.g., a submarine canyon). As the flow exits
from its confined pathway, it expands resulting in the
lateral diminution of erosion away from the centre of
the flow towards the transverse edges. Hence the
inception of these channels could conceivably be a
result of self-channelisation by hyperpycnal flows
exiting a rivermouth point source. Additionally, chan-
nel construction would have produced marked varia-
tions in cross-sectional flow velocities inside and
outside of the channel, producing a variety of over-
bank processes and transporting fine-grained material
over the channel margins.
3.1.1.6. Palaeogeographic considerations. It is well
established that the textural properties of turbidites are
primarily controlled by the nature of the sourcing
fluvial feeder system (Reading, 1991). The unusually
coarse grain-size of facies association E must be a
product of a coarse-grained, possibly braided, river
system transporting sediment to a braid or fan delta.
This supply of sediment cannot have been derived by
local erosion of underlying fine-grained sands of the
falling stage systems tract shoreface deposits.
Schwans (1995) correlated lower Ferron shoreface
strata to braided fluvial deposits in the subsurface of
the Wasatch Plateau, the occurrence of which was
attributed to tectonism and source area uplift along
central Utah salients of the Sevier thrust belt. It is
likely that these coarse-grained fluvial sandstones
belong to the highstand systems tract, deposited dur-
ing accommodation creation in alluvial plain settings.
Subsequent lowering of base-level would have
resulted in steeper fluvial profiles and delivery of
even coarser grained sediment to the shoreline with
a low preservation potential (Posamentier and Morris,
2000). The enhanced carrying capacity of these coars-
er-grained feeder systems would have promoted
hyperpycnal underflows at the shoreline. Submarine
imaging of coarse-grained deltas by sidescan sonar
methods shows them to be characterised by a spec-
trum of delta front chutes and channels, over which a
wide variety of frequently-occurring, long-distance
gravity flows are expected to occur, often initiated
by riverine underflows from the transport of high
sediment loads (Prior and Bornhold, 1989, 1990;
Bornhold and Prior, 1990). Unlike the envisaged
palaeogeography of the lower Ferron, many Late
Cretaceous river systems of other stratigraphic inter-
vals in central Utah were significantly finer-grained
and fed lobate, wave dominated deltas (e.g., Balsley,
1980; Van Wagoner et al., 1990; Kamola and Van
Wagoner, 1995; O’Byrne and Flint, 1995; Van Wag-
oner, 1995). As such they were likely to have been of
significantly lower gradient and sediment yield, and
therefore, less likely to turn hyperpycnal.
3.1.2. Juana Lopez member depositional model
Although the sandstones of the Juana Lopez Mem-
ber are interpreted to have been transported and de-
posited by waning turbulent flows, interpreting the
flow initiation mechanism is somewhat more difficult,
primarily because of the shortage of information on
contemporaneous shoreface environments and archi-
tecture. Secondly, in contrast to the lower Ferron
Sandstone, coarse-grained sandstones such as those
of facies association E have not been observed and
hence there is no evidence for channelisation of the
flows. Thus the sheet-like geometry of these sand-
stones can be feasibly attributed to either turbidity or
geostrophic (wind-driven) currents. Our interpretation
of these deposits is largely dependent on parasequence
stacking patterns, regional stratigraphic considerations
and sea-level context from existing studies.
Turbidity currents resulting from either slope col-
lapse or hyperpycnal flows is considered unlikely for
the Juana Lopez Member because of the relatively
high accommodation setting following transgression
at the end of lower Ferron deposition. The invoked
landward shift in the shoreline would have resulted in
a shallowing of fluvial gradients and sediment being
trapped within up-dip shorefaces and coastal plain
settings. An alternative and favoured hypothesis is
that the Juana Lopez sandstones may be considered
as the deposits of shoreline-oblique geostrophic cur-
rents, generated when winds blow at high angles to
the coast, setting up a pressure gradient between
surface waters moving landward and seaward-return-
ing bottom currents (Duke, 1990). The southerly
palaeoflow directions of Juana Lopez flows is sub-
parallel to oblique to the regional coastline trend in
central Utah (~north–south e.g., Roberts and Kirsch-
baum, 1995). South to southeastward returning geo-
strophic flows are plausible since such processes are
likely to be favoured during times of increased basin
width and wave fetch following transgression and
C.M. Edwards et al. / Sedimentary Geology 179 (2005) 117–152 145
subsequent highstand. Importantly, a counter-clock-
wise geostrophic circulation pattern is consistent
with computed Turonian oceanographic models for
the Western Interior (Ericksen and Slingerland,
1990; Slingerland et al., 1996). Thus these sandstones
are interpreted as the products of storm-related, shelf
transport processes during sea-level highstand and
coincident with highstand deposition and pronounced
seaward stepping of the upper Ferron deltaic complex
in southern Castle Valley (Gardner, 1995).
3.1.3. Facies distributions
Sandstone facies of the Juana Lopez Member de-
scribed here from outcrops on the eastern side of the
SRA are lacking along its western flank (Molenaar
and Cobban, 1991). The angular variance between
sole marks at the bases of beds and the overlying
current-ripples is conceivably a consequence of inter-
action between incident flows and a gently-dipping
lateral sea-floor structure in the position of the present
day SRA. Subaqueous flow deflections have been
commonly recognised in the ancient record (e.g.,
Fig. 17. Synoptic model based on the lower Ferron Sandstone illustrating th
architecture. Falling stage strata reaching an inclined sea-floor gradien
physiographic continental slope. In such cases, the elevated gradient and
deposition. Hence lowstands in basins lacking a shelf edge may be repres
and coastal plain strata as otherwise predicted by existing sequence stratig
Kelling, 1964; van Andel and Komar, 1969; Ricci
Lucchi and Valmori, 1980; Pickering and Hiscott,
1985) and have been replicated in laboratory flume-
tank experiments (Pantin and Leeder, 1987; Kneller et
al., 1991; Edwards et al., 1994). Distribution of Juana
Lopez sandstones and palaeocurrent variance implies
that the proto-SRA may have exerted a local positive
palaeobathymetric expression that was of sufficient
relief to prevent south-flowing turbidity currents
from reaching the western flank (Fig. 16B).
The proposed physiography exerted by the Farn-
ham Dome northern extension of the SRA during
deposition of the lower Ferron Sandstone, combined
with palaeocurrent variance parallel to the main axis
of the SRA and the partitioning of Juana Lopez
turbidite facies across structure reinforce models that
propose the presence of intra-basinal antecedent phys-
iography inherited prior to foreland basin develop-
ment and/or pre-Tertiary tectonic uplift of the SRA.
In summary, the sandstones belonging to the lower
Ferron Sandstone that are encased by Mancos Shale in
central Utah are considered the products of turbidity
e potential effect of basin floor topography on lowstand depositional
t may operate akin to a shelf edge delta in basins possessing a
raised sediment loads carried by lowstand rivers promotes turbidite
ented by turbidite fans rather than aggradational stacks of shoreface
raphic models (e.g. Plint and Nummedal, 2000).
C.M. Edwards et al. / Sedimentary Geology 179 (2005) 117–152146
currents that most likely originated as hyperpycnal
underflows exiting a coarse-grained delta front. The
inception of these turbidites is proposed to be a func-
tion of both high sediment loads reaching a braid delta
(in the subsurface) creating negatively buoyant under-
flows as a result of delta progradation to a structural-
ly-generated, transient shelf edge (Fig. 17). Flow
enhancement through scours or channels enabled the
long-distance transport and eventual deposition of
coarse sands over large distances.
4. Conclusions
The Turonian of central Utah was a site of signif-
icant cross-shelf sediment transport. The mechanisms
by which these transport processes were initiated are
inherently related to their relative sea-level context.
Deposits of the lower Ferron Sandstone Member of
central Utah represent a period of shoreline regression
into the Western Interior Seaway under highstand and
falling sea-level conditions. Lowstand deposits are
represented by turbiditic sandstones that are inter-
preted to be the products of sustained, quasi-steady
flows. The proposed hyperpycnal origin is based upon
the following criteria:
1. Turbidite facies consisting of coarse-grained, cross-
bedded sandstones require sustained flow for their
development.
2. An absence of related slump, debris flow and delta
front facies suggesting a direct connection between
lowstand rivers and the deeper basin.
3. A surface of sediment bypass above falling stage
systems tract strata.
4. A change in dominant depositional shoreline
process from oscillatory during falling sea-level
to offshore sediment transport during sea-level
lowstand.
5. Correlations with coarse-grained fluvial deposits in
the subsurface.
Hyperpycnal flows are considered to have been pro-
moted by enhanced sediment supply during sea-level
lowstand and by structurally-generated sea-floor
topography.
In contrast, sandstones of the overlying Juana
Lopez Member are interpreted to have been deposited
by long-distance, southward-flowing geostrophic cur-
rents following sea-level rise and inundation of the
lower Ferron Sandstone clastic wedge. Progradational
stacking of parasequences of this succession is inter-
preted as deposition under sea-level highstand condi-
tions that is coeval with progradation of a major
deltaic complex in southern Castle Valley. Palaeocur-
rent information and facies mapping indicates the
presence of active sea-floor structures in the area of
the San Rafael Anticline.
The propensity of depositional systems to deliver
sand-grade material over large distances across the
shelf has important ramifications for sequence strati-
graphic models. Given suitable conditions, sea-level
falls in basins lacking a regional shelf edge may
produce lowstand turbidite successions provided
river systems are capable of transporting sufficient
concentrations of sediment to exceed the density
threshold of the ambient (marine) water. Sea-floor
roughness exerted by intra-basinal structures may en-
courage the generation of hyperpycnal flows, contrib-
uting an important modification to generalized models
of forced regression.
Acknowledgements
This paper forms a part of a PhD thesis carried out
by the primary author under the guidance of S. Flint
(University of Liverpool) and J. Howell (University of
Bergen) and funded by a studentship provided jointly
by the University of Liverpool and ENI-LTE. In ad-
dition, the British Sedimentological Research Group
is acknowledged for providing additional financial
support for fieldwork via the Farrell Funding Scheme.
The manuscript benefited from critical reviews pro-
vided by E. Cotter (Bucknell University), R. Fitzsim-
mons (ConocoPhillips, Norway), P. Plink-Bjorklund
(University of Gothenburg) and D. Tabet (Utah Geo-
logical Survey). T. Elliott (University of Liverpool),
R. Gawthorpe (University of Manchester), and W.
Nemec (University of Bergen) are also thanked for
their in-depth and thought-provoking discussions.
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www.elsevier.com/locate/sedgeo
Sedimentary Geology 17
Imaging bed geometry and architecture of massive sandstones
in the Fontanelice Channels, Italian Apennines,
using new digiscoping techniques
R.B. Wynn a,*, P.J. Talling b, L. Amy b
aNational Oceanography Centre, European Way, Southampton, SO14 3ZH, UKbDepartment of Earth Sciences, University of Bristol, Queens Road, Bristol, BS8 1RJ, UK
Abstract
In this study we present digital images and sedimentological data from a channel fill succession in the Italian Apennines that
is dominated by massive sandstones. Although the studied outcrop is largely inaccessible, valuable data have now been
obtained using the new technique of ddigiscopingT, which allows features of b10 cm to be resolved from a distance of several
hundred metres.
About 75–80% of the channel fill is composed of massive sandstone beds N1 m thick, with overall sandstone : shale ratios of
~9 :1. Massive sandstones are poorly sorted and overall show little or no normal grading. They are commonly amalgamated and
always have sharp bed tops. Massive sandstone beds show abrupt pinch-outs at the channel margin, whereas overlying thin-
bedded siltstone/mudstone layers taper gradually and drape up the margin more extensively. This suggests that the depositing
flows were stratified into a lower, thin, (hyper)concentrated density flow and an upper, more dilute, turbidity current. In
summary, the digiscoping technique is shown to be a cheap and efficient method for imaging distant and/or inaccessible
outcrops and providing information on bed geometry and architecture.
D 2005 Elsevier B.V. All rights reserved.
Keywords: Channel; Marnoso Arenacea; Massive sandstone; Digiscoping
1. Introduction
Deep-water gravity flow processes and deposits
have been the focus of intensive research over the
0037-0738/$ - see front matter D 2005 Elsevier B.V. All rights reserved.
doi:10.1016/j.sedgeo.2005.04.012
* Corresponding author. Tel.: +44 2380 596553; fax: +44 2380
596554.
E-mail address: [email protected] (R.B. Wynn).
last three or four decades (see recent overviews by
Bouma and Stone, 2000; Shanmugam, 2000; Knel-
ler and Buckee, 2000; Mulder and Alexander,
2001). However, our inability to directly observe
and instrument such flows means that knowledge
of flow processes, and their impact on deposit
character, is restricted to interpretation of preserved
deposits and experimental/numerical modelling.
9 (2005) 153–162
R.B. Wynn et al. / Sedimentary Geology 179 (2005) 153–162154
Such interpretations are inevitably ambiguous and
frequently controversial; this is exemplified by the
continuing debate over the processes responsible for
depositing thick massive sandstones in deep-water
environments (e.g. Lowe, 1982; Kneller and Bran-
ney, 1995; Shanmugam et al., 1995, 1997; Shanmu-
gam, 1996, 2000, 2002; Hiscott et al., 1997; Stow
and Johansson, 2000; Marr et al., 2001). As a
consequence, there is a growing requirement for
new case studies of massive sandstones in both
modern and ancient environments.
In this study we use a new technique called
digiscoping to image the geometry and architecture
of thick massive sandstone beds in a channel fill
succession in the Italian Apennines. Digital images
are then combined with sedimentological analysis
of channel fill deposits to provide some brief
insights into depositional process. The general char-
acter of the studied outcrop has previously been
described by Ricci Lucchi (1969, 1975, 1981) and
Mutti et al. (2002), however, only now can we
examine detailed bed geometries and architecture
by digiscoping. The results will highlight the appli-
cability of the digiscoping technique to outcrop-
based fieldwork, especially in areas where key
exposures are inaccessible.
Fig. 1. Location map of the study area near Fontanelice in the
2. The Fontanelice Channels
The Fontanelice Channels are exposed in an im-
pressive 2D cliff section adjacent to the Santerno
River, near the town of Fontanelice in the Italian
Apennines (Fig. 1). The channels occur within the
uppermost Marnoso Arenacea Formation, which is
Miocene in age (Tortonian). The sediment source for
flows passing through the channels was the southern
Alps and material was ultimately deposited in a basin
environment, although poor outcrop quality has pro-
hibited correlation with the time-equivalent down-
stream sections (Ricci Lucchi, 1969, 1975, 1981;
Mutti et al., 2002).
Two offset-stacked channels are visible in the cliff
section on the northwest side of the Santerno Valley
(Figs. 2 and 3), incising into adjacent and underlying
parallel-bedded sandstone-rich sheets (note that the
lower channel base and underlying beds have recently
been covered by vegetation and talus, see Ricci Luc-
chi (1981) for an older image displaying this section).
Mutti et al. (2002) interpreted the underlying sand-
stone sheets as lobe deposits forming part of the
Castel del Rio dmixedT system. The orientation of
the outcrop relative to palaeoslope (NNW–SSE) sug-
gests that the cliff section is roughly perpendicular to
Italian Apennines. Modified from Ricci Lucchi (1981).
Fig. 2. Digital photo showing the main studied outcrop in the Santerno Valley. Channel bases are shown by dashed lines. Note that the base of
Channel A is no longer visible due to erosion/vegetation cover but is roughly reproduced here (white dashed line) after study of an older photo
in Ricci Lucchi (1981). Large white rectangle shows location of Fig. 3. Small black rectangles show location of Figs. 5 and 6. F=minor fault.
R.B. Wynn et al. / Sedimentary Geology 179 (2005) 153–162 155
channel axes (Mutti et al., 2002). The lower channel is
several hundred metres wide and about 80 m deep,
and is dramatically reincised by the upper channel
which is of similar dimensions (Figs. 2 and 3). For
the purposes of this study, the lower channel will
hereafter be referred to as Channel A and the upper
channel as Channel B. Both channels contain an
aggradational fill comprising thick massive sand-
stones that are commonly amalgamated. Muddy over-
bank/slope deposits are preserved above Channel B,
and are incised by a younger channel complex (Ricci
Lucchi, 1981; Mutti et al., 2002).
Mutti et al. (2002) considered the location and
morphology of the Fontanelice Channels to be struc-
turally controlled, and suggested that the erosional
base of Channel B may represent a slump scar that
was generated in response to increased slope insta-
bility during tectonic uplift. An equally viable hy-
pothesis would be that the channel was initially cut
by high-energy bypassing flows. The channel fill
itself was interpreted to represent deposition from
sand-bearing gravity flows, possibly related to remo-
bilisation of unconsolidated sand from upslope
(Mutti et al., 2002).
3. Digiscoping equipment and technique
Digiscoping combines the improved resolution and
adaptability of modern digital cameras with the high
magnification of field telescopes. The digital camera
is attached to the telescope allowing the photographer
to take ultra-high magnification images of distant
subjects. This method is now widely used in the
field of wildlife photography, but can be applied to
any situation where the subject is a long way from the
observer. The technique is ideally suited to photogra-
phy of distant inaccessible outcrops which, until now,
have been too distant for conventional single-lens
reflex (SLR) photography with zoom lenses.
Fig. 3. Digital photo (a) and line drawing interpretation (b) showing the erosional base and channel fill of Channel B. For location see Fig. 2.
Black lines denote thin-bedded heterolithic intervals. Note the terraced erosional profile of the channel, the tabular nature of thick sandstone
beds and the extensive draping of thin-bedded heterolithic intervals up the channel margin.
R.B. Wynn et al. / Sedimentary Geology 179 (2005) 153–162156
Fig. 4. The digiscoping set-up, with a digital camera connected to a tripod-mounted telescope by an aluminium sleeve. This outfit is relatively
cheap and portable, and is a rapid and effective tool for achieving high-quality images of distant outcrops. Features b10 cm in size can be
resolved at distances of several hundred metres.
R.B. Wynn et al. / Sedimentary Geology 179 (2005) 153–162 157
At the present time, the most suitable cameras for
digiscoping are in the Nikon Coolpix range as these
models have a rotating body and internal zoom lens.
The camera is attached to the telescope using a ma-
chined aluminium bracket and thumb screws (Fig. 4).
With the camera lens set at 4� magnification and a
fixed telescope lens of 20� or 30� magnification
overall magnifications of 50–100� are achievable,
far exceeding the capability of the most powerful
SLR zoom lenses. In practical terms this means that
features b10 cm across can be resolved when the
viewer is at distances of 100 m or more. Individual
images are saved at high-resolution and can then be
arranged into digital photo-mosaics using standard
computer graphics packages such as Adobe Photo-
shop and Illustrator.
4. Digiscoping results and bed geometries
The whole Fontanelice cliff section is shown in
Fig. 2, with the fill of Channel A clearly incised by
Channel B, which is offset-stacked to the southwest.
The base of Channel A is no longer visible, although
an older image published in Ricci Lucchi (1981)
indicates that the channel margin is a low-angle
dterracedT erosion surface. Accessible sections of the
Channel A fill show massive tabular sandstones up to
a few metres in thickness, separated by thin hetero-
lithic intervals. Overall, the sandstone : shale ratio
appears to be ~9 :1.
The base and fill of Channel B (Figs. 2 and 3) is
superbly exposed but completely inaccessible; the
detailed external geometry of features such as large-
scale erosional scours and marginal bed pinch-outs
therefore provide an ideal target for digiscoping.
Consequently, the analysis of external bed geometry
will focus on this channel. The fill of Channel B is
also dominantly composed of massive tabular sand-
stones, with most beds being 1–2 m thick but a
couple reaching 4 m in thickness. These thick sand-
stones are separated by thin heterolithic beds (b1 m
thick) that appear to be fine-grained siltstones and
mudstones. Overall, the sandstone/shale ratio is
~9 :1.
4.1. Bed geometry and marginal pinch-outs
The external geometry (in strike section) of thick
sandstone beds in Channel B is roughly tabular (Figs. 2
R.B. Wynn et al. / Sedimentary Geology 179 (2005) 153–162158
and 3). Marginal pinch-outs are commonly abrupt (Fig.
3), although in some cases they are complicated by the
pre-existing dterracedT erosion profile of the channel
base. For example, a digital photo-mosaic illustrates
the pinch out of a thick sandstone bed at the channel
margin (Fig. 5); this pinch out is irregular due to the bed
onlapping a step in the channel erosion profile.
Figs. 3 and 5 also illustrate how thin-bedded het-
erolithic intervals are draped up the channel margin,
and are only completely eroded out at the channel
base and at pronounced steps along the margin. These
thin-bedded intervals reach their greatest thickness on
flat sections of the terraced channel margin. They are
not continuous with adjacent massive sandstones, but
are interbedded between them.
In summary, the irregular pinch-outs and onlap of
beds comprising the channel fill suggest that the
channel has been subject to multiple phases of cut-
and-fill. The tabular nature of the thick sandstone
beds, and their abrupt marginal pinch-outs, contrasts
with the more extensive draping character of interven-
ing thin-bedded heterolithic intervals.
Fig. 5. Digital photo-mosaic (a) and line drawing interpretation (b) showi
Photo-mosaic is composed of 32 individual digital images. For location see
topography generated during initial channel incision. Thin-bedded heter
individual sandstone bodies in the two channel fills. The only places wher
erosion profile, producing a sand-on-sand contact.
4.2. Scours, amalgamation surfaces and soft sediment
deformation
Deposition of thick sandstone beds in Channel B
was accompanied by local scouring and soft sediment
deformation. Individual erosional scours are up to 1 m
deep and 5 m across (Fig. 6), and commonly cut
through underlying shales to generate amalgamation
surfaces. Many of these surfaces can be picked out by a
line of aligned nodules (Figs. 3 and 6), while in other
places the edge of the amalgamation surface is indicat-
ed by truncation of a thin shale unit (Fig. 6). It is likely
that other, less obvious, amalgamation surfaces are also
present in some of the apparently uniform thick sand-
stone beds. In many cases, thin shales were heavily
deformed and contorted during deposition of overlying
sandstone beds, and have been injected upwards over
distances of a metre or more (Fig. 6).
The resolution limitations of digiscoping mean that
from the vantage position (several hundred metres
south-east of the outcrop) features b5 cm across are
not clearly resolvable. Consequently, subtle water
ng pinch-out of a thick sandstone bed at the margin of Channel B.
Fig. 2. Note the complex bed geometry due to onlap of pre-existing
olithic intervals drape the channel margin and effectively separate
e connectivity is achieved is just below a marked step in the channel
Fig. 6. Digital photo-mosaic (a) and line drawing interpretation (b) showing large erosional scour, bed amalgamation and shale deformation/
injection near the base of Channel B. Photo-mosaic is composed of seven individual digital images. For location see Fig. 2.
R.B. Wynn et al. / Sedimentary Geology 179 (2005) 153–162 159
escape features and small shale clasts, if present,
cannot be observed.
5. Sedimentology of the channel fills
Examination of the few accessible areas at the base
of the studied outcrop provides some indication of the
sediment fill of Channel A, and helps to verify the
digiscoping observations. The dominant sediment fa-
cies are thick massive sandstones separated by thin
fine-grained sandstone, siltstone and mudstone inter-
vals. Similar findings were presented by Mutti et al.
(2002), and beds with identical sediment facies are
also found in a small road outcrop on the opposite
side of the Santerno Valley, that apparently also
exposes a section of one of the channel fills.
5.1. Thick (N1 m) massive sandstones
Thick massive sandstone beds commonly appear
ungraded throughout, although both normal and inverse
grading occur locally at the base of some beds. The
sandstones are poorly sorted and grey-brown in colour.
They generally show a massive, structureless appear-
ance, although faint planar laminations and coarse sand-
stone dstringersT are observed in some places (Fig. 7).
Bed bases are usually erosional, sometimes down to a
few tens of centimetres, and are commonly loaded into
underlying fine-grained intervals; they can also form
irregular amalgamation surfaces with underlying thick
sandstone beds (Figs. 7 and 8). Bed bases contain
abundant randomly scattered coarser grains of 1–6
mm and rare larger pebbles up to 12 mm. Flute marks
are noted at the base of some beds and are infilled by
coarse lag deposits of 2–3 mm grain size. A key feature
is that bed tops are very sharp with a distinct grain size-
break, but may be disrupted by bioturbation.
The overall lack of grading is supported by both
visible and measured grain size analyses, with the
modal grain size of 100–200 Am remaining constant
throughout sampled intervals. The mud content
(here taken as b20 Am) is also remarkably constant,
ranging between 7% and 12% for all samples.
Fig. 7. Digital photo showing thick massive sandstone beds in road section opposite Fontanelice Channel outcrop (outcrop orientations indicate
that these beds are probably part of the channel fill). Note the presence of bed amalgamation and intervening thin-bedded siltstone and mudstone
intervals.
R.B. Wynn et al. / Sedimentary Geology 179 (2005) 153–162160
5.2. Thin (b50 cm) siltstones/fine sandstones and
homogeneous mudstones
The finer-grained intervals are normally graded
and bioturbated, and contain rippled greyish fine
sandstone/siltstone overlain by dark grey homoge-
neous mudstones (Figs. 7 and 8). These deposits
occur immediately on top of massive sandstone
beds, and display sharp flat bases. The siltstones and
fine sandstones always show cross- or contorted lami-
nations with unidirectional ripples. Modal grain size is
50–70 Am. The dark grey homogeneous muds are a
few centimetres thick and have gradational contacts
with the underlying rippled silts and fine sands. Bed
Fig. 8. Digital photo showing base of massive sandstone bed with possible sheared layer and clear evidence for loading and injection into
underlying fine-grained intervals.
R.B. Wynn et al. / Sedimentary Geology 179 (2005) 153–162 161
tops are sharp but often eroded, loaded and/or injected
by overlying thick sandstone beds (Figs. 6, 7 and 8).
6. Depositional processes of the channel fills
Massive sandstone beds in the Fontanelice Channel
fills show little or no vertical grain-size grading, have
sharp bed tops and display a tabular geometry in strike
section with abrupt marginal pinch-outs; these fea-
tures all point towards deposition from hyperconcen-
trated density flows (Mulder and Alexander, 2001).
However, some beds also show features more consis-
tent with concentrated density flows, such as faint
planar laminations and erosional bases with flute
marks (Mulder and Alexander, 2001). A more detailed
interpretation would require vertical grain size profiles
through a series of massive sandstone beds, but this is
not possible at this location due to the largely inac-
cessible nature of the outcrop. The depositional pro-
cess of the fine-grained intervals overlying most
massive sandstone beds is more straightforward; the
presence of normal grading, unidirectional current
ripples and homogeneous mud caps suggest that
they are deposits of relatively dilute turbidity currents.
The observed sequence of (hyper)concentrated
density flow deposits overlain by turbidity current
deposits is repetitive (Fig. 3), suggesting a genetic
link between the different deposit types. If this is
the case then dilute turbidity currents were presum-
ably linked to the (hyper)concentrated density flows
responsible for depositing the massive sandstones
(Mulder and Alexander, 2001). Dilute turbidity cur-
rents may have formed through fluid entrainment and
mixing at the top of the (hyper)concentrated density
flow during its passage downslope, or through distur-
bance of unconsolidated fine-grained seafloor sedi-
ments during initial slope failure and erosion/
pressure wave disruption at the head of the flow.
The resulting stratified flow was therefore composed
of two layers: 1) a sand-rich (hyper)concentrated den-
sity flow that was confined to the channel floor and
deposited thick massive sandstone beds that pinch-out
abruptly at the channel margin, and 2) an overlying
finer-grained, more dilute turbidity current that depos-
ited thin-bedded rippled siltstones and homogeneous
mudstones that taper gradually and drape several
metres up the channel margin.
7. Conclusions
This study shows that digiscoping is a relatively
cheap and efficient technique for imaging distant and/
R.B. Wynn et al. / Sedimentary Geology 179 (2005) 153–162162
or inaccessible outcrops, opening new opportunities
for data collection in previously well visited outcrop
areas. It is especially useful for analysing gross exter-
nal geometries at an individual bed scale, but is also
capable of resolving features b10 cm across at dis-
tances of several hundred metres. The application of
digiscoping to the Fontanelice Channels outcrop,
combined with more traditional sedimentological
techniques, has revealed that the channels were filled
by two-layer stratified flows, with a lower (hyper)-
concentrated density flow overlain by a more dilute
turbidity current. The resulting channel fill deposits
are composed of tabular massive sandstones with
abrupt marginal pinch-outs, overlain by thin-bedded
turbidites that taper gradually and drape up the chan-
nel margin.
Acknowledgements
This study forms part of the UK-TAPS Marnoso
Project. Financial support for this project has been
received through a NERC Ocean Margins LINK
grant (NER/T/S/2000/0106) and industry sponsors
(ConocoPhillips, BHP Billiton and Shell UK). We
are particularly grateful to Juli Ericsson and Geoff
Haddad of ConocoPhillips for their long-standing
support and scientific input to the project. The
reviewers, Finn Surlyk, Franco Ricci Lucchi and
Gareth Keevil, are thanked for providing critical
reviews that greatly improved the initial manuscript.
Maarten Felix provided editorial assistance. Harriet
Wimhurst-Brookes is thanked for undertaking grain
size analysis of selected massive sand beds.
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=Sedimentary Geology 17Bed geometry used to test recognition criteria of turbidites
and (sandy) debrites
L.A. Amy a,b,*, P.J. Talling a, J. Peakall b, R.B. Wynn c, R.G. Arzola Thynne a
aCentre for Environmental and Geophysical Flows, Department of Earth Sciences, University of Bristol, Bristol, BS8 1RJ, UKbSchool of Earth Sciences, University of Leeds, Leeds, LS2 9JT, UK
cChallenger Division, Southampton Oceanography Centre, European Way, Southampton, Hampshire, SO14 3ZH, UK
Received 20 April 2004; accepted 6 April 2005
Abstract
The origin of thick-bedded deep-water sandstones has generated much controversy in recent years. Two fundamentally
different models have been proposed for beds with the same internal sedimentary characteristics: (1) progressive particle
settling from the base of a turbulent flow—the bturbidity currentQ model and (2) en-masse freezing of a higher-concentration
flow—the bsandy debris flowQ model. These models predict beds with very different geometries; turbidites thin gradually
whereas debrites have abrupt terminations. Previous studies have relied upon sedimentary recognition criteria (i.e., sedimen-
tary features in small-scale outcrop or core) to interpret depositional mechanism. In this study, depositional mechanism is
deduced from bed geometry gained from extensive correlations of individual sandstones preserved in a classic turbidite
system (Marnoso-arenacea Formation, Italy). This approach allows recognition criteria for turbidites and submarine debrites
to be independently tested. We find that tabular and tapered sandstones (turbidites) have distinctly different internal
characteristics to beds with abrupt margins (debrites). Turbidites are relatively well sorted, often exhibit grading and traction
structures and have relatively low matrix mud contents. They may also contain massive division, floating clasts and inverse
grading. Debrites are moderate-to-poorly sorted, ungraded, structureless, contain floating clasts and have elevated matrix mud
contents. These findings have implications for the assessment of submarine gravity flows deposits and reservoir rock
characterization.
D 2005 Elsevier B.V. All rights reserved.
Keywords: Turbidity current; Debris flow; Bed geometry; Sedimentary facies
0037-0738/$ - see front matter D 2005 Elsevier B.V. All rights reserved.
doi:10.1016/j.sedgeo.2005.04.007
* Corresponding author. Tel.: +44 117 954 5235; fax: +44 117 925
3385.
E-mail address: [email protected] (L.A. Amy).
1. Introduction
Submarine sediment-laden density flows are mix-
tures of sediment and water that flow along the sea
or lake floor due to their excess density. These flow
events represent the major sediment transport pro-
9 (2005) 163–174
L.A. Amy et al. / Sedimentary Geology 179 (2005) 163–174164
cess from continental shelf to deep ocean, and their
deposits form thick sedimentary successions that
now host many of the world’s largest petroleum
reservoirs (Weimer and Link, 1991). The ability to
predict lateral variations in bed geometry and sedi-
mentary character (e.g., porosity and permeability)
using sedimentary models is critical to the economic
recovery of hydrocarbons.
The origin of the basal interval of thick-bedded
deep-water sandstones has generated much controver-
sy in recent years (Shanmugam and Moiola, 1995;
1997; D’Agostino and Jordan, 1997; Bouma et al.,
1997; Coleman, 1997; Lowe, 1997; Slatt et al., 1997).
Two contrasting depositional models have been pro-
posed. The bhigh-density turbidity currentQ model
proposes that particles gradually settle out of turbulent
suspension and progressively aggrade a bed (Kuenen
and Migliorini, 1950; Lowe, 1982; Kneller and Bran-
ney, 1995; Stow and Johansson, 2000). Turbidity
currents must have low enough concentrations for
individual particles to settle out from the flow.
Given high suspended load fallout rates, a relatively
high-concentration flow boundary may develop at the
base of the current with reduced turbulence and en-
hanced grain interaction (Lowe, 1982; Kneller and
Branney, 1995). These bhigh-concentration turbidity
currentsQ are considered by some authors to be a flow
type transitional between turbidity currents and debris
flows (e.g., concentrated density flows of Mulder and
Alexander, 2001). Importantly however, deposition
from such currents is still considered to occur pro-
gressively (Lowe, 1982; Kneller and Branney, 1995).
In contrast, a dsandy debris flowT model infers that
deposition occurs by rapid and en-masse freezing of a
high-concentration current (Shanmugam and Moiola,
1995, 1997; Shanmugam, 1996; Stow and Johansson,
2000). Debris flows undergo flow arrest and freezing
when the forces of shear resistance, viscosity and
friction become equal to the flow’s driving force
(Lowe, 1982; Postma, 1986; Mulder and Alexander,
2001). This type of behaviour occurs both in cohe-
sionless flows that have frictional strength due to
interlocking of grains, and cohesive flows that possess
a yield strength due to the presence of cohesive
sediment such as clays.
Distinguishing between a dturbidity currentT or a
ddebris flowT depositional model for sandstone beds is
important since the two models predict deposits with
markedly different geometries. Turbidity currents
should produce spatially extensive deposits with ta-
pered margins as shown by physical experiments
(e.g., Alexander and Morris, 1994; Hallworth and
Huppert, 1998). In comparison, debris flows should
form elongate deposits with abrupt lateral pinch-outs
as indicated by the deposits of subaerial (Major, 1997)
and subaqueous experimental flows (e.g., Hallworth
and Huppert, 1998; Mohrig et al., 1999; Marr et al.,
2001) and the deposits of natural subaerial (e.g., Law-
son, 1982) and very thick (10–50 m) submarine debris
flows (e.g., Aksu and Hiscott, 1989; Laberg and
Vorren, 1995). Some debris flows may deposit by
the progressive accretion of successive flow surges,
but their deposits still display a debrite-like morpho-
logy with abrupt margins (Major, 1997).
Bed geometry is a good indicator of depositional
mechanisms. However, bed geometry is often difficult
to obtain due to the limitations of outcrop or seismic
resolution. Instead, interpretation of sedimentary beds
is usually based on relatively small, centimetre- to
metre-scale sedimentary features preserved in outcrop
or core. Mud-poor, relatively well-sorted sandstones
that contain traction structures and normal grading are
usually interpreted as the deposits of turbidity currents
(e.g., Bouma, 1962; Middleton and Hampton, 1976;
Lowe, 1982; Mutti, 1992). Sedimentary features com-
mon to debrites include (1) mud-rich matrix (if cohe-
sive), (2) poor sorting, (3) shear fabric at the base and
margins, (4) clasts protruding above the top of the bed,
(5) clasts floating in a matrix, (6) a lack of internal
structures, (7) fluid escape structures, (8) sharp upper
contact and (9) lack of grading or inverse grading
throughout or at the base of the deposit (Fisher, 1971;
Rodine and Johnson, 1976; Enos, 1977; Naylor, 1980;
Lowe, 1982; Major, 1997; Sohn, 2000; Marr et al.,
2001; Mulder and Alexander, 2001). However, many
of these features (especially characteristics 5–9) also
occur in metre-thick sandstone beds interpreted as
dturbiditesT. Consequently, it has been suggested that
many well-known dturbiditeT successions such as the
Jackfork Group in Arkansas and Oklahoma (Shanmu-
gam and Moiola, 1995, 1997), the Marnoso-arenacea
Formation in northern Italy (Shanmugam, 1997) and
the Annot sandstones Formation in southeast France
(Shanmugam, 2002) contain large proportions of
sandy debris flow deposits. This reinterpretation is
controversial since it includes structureless, moderate-
60
6970
67
45
66
6559
64
63
1, 80-82
3
2
77
85
84
83
75
10
76
7826
71
72&73
74
42
40
68
6
7
89
11
12
14
79
51&49
15
48
1316
17
5530
50, 52-5428
29 37
56
39
3634
33323120
21
22 23
24
62
44
43
61
57
58
25
Bagno diRomagna
Sansepolcro
Firenzuola
Modigliana
Faenza
Forli
Cesena
10km0
Main flowdirection
N
Measured section
Transect shown in Fig. 2
Main thrust fault
Extent of MA Fm.
MarnosoArenacea Fm.
KEY
BOLOGNA
FIRENZE
Emilia-Romagna
Marche
AdriaticSea
Toscana
Umbria
ITALY
MarnosoArenacea Fm.
Fig. 1. Location map of the study area in central–northern Italy showing the Marnoso-arenacea Formation (shaded area), the positions of
measured sections and line of the stratigraphic panel shown in Fig. 2.
L.A. Amy et al. / Sedimentary Geology 179 (2005) 163–174 165
L.A. Amy et al. / Sedimentary Geology 179 (2005) 163–174166
ly to well-sorted sandstones with low mud contents
previously interpreted as the Ta division in the Bouma
(1962) depositional model.
In this study, we use bed geometry as an indepen-
dent criterion for verifying recognition criteria of
turbidites and submarine debrites. The aims of this
study are (1) to distinguish turbidites and sandy deb-
rites in outcrop using bed geometry and (2) to verify
which sediment characteristics present in outcrop at
single locations are diagnostic of depositional flow
type.
2. Study area
The Miocene Marnoso-arenacea Formation crops
out extensively over central–northern Italy (Fig. 1).
The evaluation of bed shape is possible in this forma-
tion since distinctive limestone marker beds and out-
crop continuity allow correlations of individual event
beds over an area of 123 by 27 km (Ricci Lucchi,
1978; Ricci Lucchi and Valmori, 1980). These are the
most extensive correlations of individual beds within
any outcropping turbidite system known to the
authors. The formation mainly comprises basin plain
deposits, composed of interbedded deep-water sand-
stones and mudstones separated by hemipelagic
marls. They were deposited in the Apennine foredeep
basin and subsequently accreted into the Apennine
thrust belt (Ricci Lucchi, 1978; Ricci Lucchi and
Valmori, 1980). The Apennine foredeep basin con-
sisted of a northwest–southeast oriented elongate
trough over 150 km long and 50 km wide (Argnani
and Ricci Lucchi, 2001). Sediment was transported
into the basin by southward-directed flows carrying
detritus from a non-carbonate Alpine source posi-
tioned some 150–200 km northwest of the present
outcrops (Argnani and Ricci Lucchi, 2001). North-
ward directed flows carrying calcareous-rich detritus
were responsible for forming limestone marker beds
(Gandolfi et al., 1983).
A single 25–30 m thick stratigraphic interval of
Serravallian age, located between the most prominent
dContessaT marker bed and a stratigraphically higher
dColombineT marker bed (Ricci Lucchi and Valmori,
1980) was selected to study bed geometry. Recent
publication of 1:10,000 scale geological maps (pub-
lished by the geological surveys of Emilia-Romagna,
Toscana, Marche and Umbria), indicating the position
of the Contessa marker bed, has allowed over 70 new
sections to be measured (Fig. 1). These new sections
have allowed bed shape to be defined with much
greater precision than in previous studies.
During the time of deposition of the studied inter-
val, the basin plain area is believed to have had low
sea-floor gradients a character also exhibited by mod-
ern, deep-sea, abyssal plains (Ricci Lucchi, 1978;
Ricci Lucchi and Valmori, 1980). Low sea-floor gra-
dients are inferred from the high continuity of indi-
vidual beds, absence of channelisation and ability of
currents from different sources to flow in opposite
directions. Hence, in the basin plain area, bathymetry
is not believed to have significantly influenced the
behaviour of sediment gravity flows nor the geome-
tries of the beds they deposited. Hence, this area is an
excellent natural laboratory to study the deposits of
sediment gravity currents that flowed, unobstructed,
across a basin plain.
3. Bed geometry
In this contribution, we concentrate on the geo-
metry of relatively thick sandstone beds. The corre-
lated stratigraphic interval contains 12, metre-thick
sandstone beds of which 10 occur in the stratigraphic
panel shown in Fig. 2. All of these sandstone beds
thin in a downstream direction between the most
northern and southern sections. Two cross-sectional
sandstone bed geometries are common: those that thin
gradually (tapered) and those that thin abruptly down-
stream (Fig. 2). Detailed examples of these beds are
shown in Fig. 3.
3.1. Tapered sandstones
Most of the metre-thick sandstone beds preserved
in relatively proximal sections display a gradual thin-
ning pattern (Figs. 2 and 3A). These beds thin from
N0.8 m to b0.5 m over ~20–40 km. Small changes of
b10–20% of their thickness are recorded between
neighbouring measured sections, spaced 1–10 km
apart. An isopach map for the sandstone thickness
of bed 7 is shown in Fig. 4A and illustrates the
localised thickening and thinning often seen in these
beds. Localised thickness changes may be related to
Thr
ustf
ault
Thr
ustf
ault
Contessa mudstone Contessa mudstone
Flow directionfor most beds
KEY
1
Datum
2
2.5
3
4
5
5.1
6
7
8
C1
m si vf f m c
Sect. No.
Bed No.
17.0 km 16.9 km21.1 km
25m
7.2 km 6.0 km
8 11 12 14 48 50 3755 60 70 455762 44 61 58
?
?
SENW
Marl
Mudstone
Poorly sorted silty-muddy sandstonecontaining mud clasts
Cover
Mudstone or marl (undistinguished)
Sandstone (well-to-moderately sorted)
Fig. 2. Stratigraphic panel showing correlation of thick beds between relatively proximal (left) and relatively distal (right) sections over a distance of some 68 km. The datum from
which sections are hung is the top of the Contessa marker bed’s mudstone. Measured sections are simplified to show only lithofacies and not internal bed structure. Flow direction of
depositional currents was from left to right except for carbonate-rich beds (e.g., bed C1) whose parental flow travelled in the opposite direction. Most thick sandstone beds display a
gradual reduction in thickness moving downstream. Beds 2.5 and 5.1, however, display an abrupt decrease in thickness downstream.
L.A.Amyet
al./Sedimentary
Geology179(2005)163–174
167
(B) BED 2.5: Sandstone showing abrupt thinning
Thin sand or mud ~10 cm thick
8 11 12 14 48 50 3755 60 455762 44 61 58
(A) BED 4: Sandstone showing gradual thinning
Lithologic units
Marl
Mudstone
Mudstone or marl (undistinguished)
Sandstone (well-to-moderately sorted)
Poorly sorted silty-muddy sandstone
Cover
Liquefaction lamination
Chaotic swirly texture
Folded / sheared?mud-poor sandstone
Sedimentary structures
Cross-lamination
Dunecross-bedding
Starved dune / dune top Planar stratified
Planar laminated
Overturned convolute laminated
Convolute laminated
Wavy:convolute or dune laminated
Clast horizon
Diagenetic nodules orsiltstone clasts
Scattered clastsmud / marl
?
FlatUndularLoadedSmall flutes (<2cm deep)Grooved / Tool marks
Base of bedClasts
Organics
Wavy surfaces / indistinct surfacesdune-scale wavy, dewatering or,corrugated related structures
Flow directionm si vf f m c
17.0 km
(Note transect broken by thrust fault)(Note transectbroken bythrust fault)
16.9 km21.1 km
1m
1m
7.2 km 6.0 km
KEY
Fig. 3. Sedimentary logs of two correlated sandstone beds. Bed 4 is an example of a bed that displays gradual downstream thinning (A). Bed 2.5 is an example of a bed that shows
abrupt thinning downstream (B). These two bed types are interpreted as the deposits of turbidity currents and cogenetic turbidity current–debris flows, respectively.
L.A.Amyet
al./Sedimentary
Geology179(2005)163–174
168
Fig. 4. Isopach maps of sandstone thickness of a turbidite bed (bed 7) displaying gradual thickness changes (A) and a cogenetic debrite-turbidite
bed (bed 2.5) displaying abrupt thickness changes (B). Sandstone thickness is defined as the portion of the bed with a grain size of siltstone and
coarser. Isopach maps were created using ODMk computer software. Grids use an inverse distance interpolation method. Interpolation does not
use data from different sides of major thrust faults.
L.A. Amy et al. / Sedimentary Geology 179 (2005) 163–174 169
L.A. Amy et al. / Sedimentary Geology 179 (2005) 163–174170
the compensation of subtle basin floor irregularities
caused by the depositional relief of previous deposits
or local basin subsidence. More significant changes in
bed thickness, ~40–50% of their thickness, are occa-
sionally observed. For example, bed 7 decreases by
~50% of its thickness between sections 57 and 58
over a distance of ~3 km, and so does bed 1 between
70 and 45 over a similar distance (Fig. 2). However,
such large changes are usually associated with the
gain or loss of the very fine sandstone or siltstone
found at the top of the bed and not the coarser grain
size fraction. The large spatial extent and gradual
downstream thinning of these beds suggests deposi-
tion by turbidity currents.
3.2. Sandstones with abrupt margins
Four out of the twelve correlated metre-thick sand-
stone beds display an abrupt change in their thickness
moving downstream. The abrupt margin of two of
these units, beds 2.5 and 5.1, occur in the stratigraphic
panel shown in Fig. 2. Bed 2.5 is shown in detail in
Fig. 3B. These sandstones thin by ~80% from ~1 m
to b0.2 m thickness, over 2–10 km. In some cases, the
sandstone bed pinches-out completely and only a
correlative mud cap may be identified. In plan view,
the thick portion of the bed is tongue-shaped (e.g.,
Fig. 4B), thinning in a direction both parallel and
perpendicular to the flow direction. The elongate
plan-form shape, limited spatial extent and abrupt
terminations of the thick portion of these sandstone
beds suggest deposition from a debris flow.
4. Internal character
4.1. Turbidite sandstones (beds with tapered margins)
The internal sedimentary character of the turbidite
sandstones varies between individual beds and local-
ities but a number of similar features are shared (Fig.
5A). They are relatively well sorted, contain traction
structures (cross, parallel and convolute lamination)
and commonly show normal grading. Beds commonly
display 1–10-cm-thick planar or gently undulating
grain size stratification in their lower part. In some
beds, the basal stratified unit is inversely graded and is
observed to occur in a step-wise fashion, as opposed
to a gradual grain size increase. However, traction
structures and normal grading are not always perva-
sive throughout the bed; beds commonly exhibit mas-
sive intervals lacking structure or grading (Fig. 3A).
Clasts are sometimes present and may occur individ-
ually or as a group scattered along a discrete horizon.
The total mud content of collected samples from
several turbidite beds was measured using SEM anal-
ysis (Talling et al., 2004). In this analysis, detrital mud
was not distinguished from diagenetic mud and hence
probably indicate somewhat greater values than that
derived from the depositional flow. Results show that
turbidites have a relatively low mud matrix content of
b12% and typically between 5–8% by volume.
4.2. Debrite-turbidite sandstones (beds with abrupt
margins)
Beds with abrupt margins display distinctive ver-
tical and lateral changes in character (Fig. 3B). The
thick portion of the beds in most locations is charac-
terized by a distinctive tripartite structure comprising
(1) a basal, massive or laminated unit, (2) a middle,
unstructured and relatively thick, clast-rich unit and
(3) a laminated upper unit (Fig. 5B). This tripartite
bed structure occurs upstream of where the sandstone
begins to pinch and is continuous for 20–30 km (Figs.
3B and 4B). In downstream sections, the sandstone
bed is much thinner (b20 cm) and it is characterized
by fine-grained laminated sand. In relatively proximal
northern sections, the tripartite bed structure is some-
times absent and instead the bed comprises structure-
less, normally graded or ungraded sandstone.
The middle interval in the tripartite bed structure,
previously referred to as a dslurried divisionT (RicciLucchi and Valmori, 1980), is distinctive in character.
It is moderate to relatively poorly sorted (containing
mud, silt sand and clasts) and ungraded with a rela-
tively sharp grain size discontinuity at its upper
surface. Outsize clasts of mud and hemipelagite,
0.2–30 cm long, are randomly distributed throughout
the unit and larger sand grains may be dispersed
within the finer grained matrix. The matrix has a
relatively high mud content, N15% and typically
20–22% by volume as assessed from SEM analysis
of samples (Talling et al., 2004). The high mud con-
tent ensures that outcrops are often friable and have a
distinct grey hue.
Fig. 5. Outcrop photographs and sedimentary logs of a turbidite bed (A) and a cogenetic turbidite-debrite bed (B). The turbidite bed has a
characteristic blocky weathering pattern. The debrite has a distinctive swirly weathering pattern and occurs sandwiched between thinner
cogenetic turbidite sandstones. See Fig. 3 for graphic log key.
L.A. Amy et al. / Sedimentary Geology 179 (2005) 163–174 171
The abrupt termination of the middle clast-rich
interval of tripartite beds suggests it is a debrite. The
basal and upper divisions of tripartite beds, however,
have nearly identical characteristics to those found in
tapered turbidite sandstones, such as traction struc-
tures and a low mud content (b10% by volume).
Similarly, the characteristics of relatively proximal
parts of these beds (i.e., low mud contents and normal
grading) suggest they may be turbidites and not deb-
rites although this interpretation cannot be justified
L.A. Amy et al. / Sedimentary Geology 179 (2005) 163–174172
using geometry. The abrupt lateral margins of these
composite ddebrite-turbiditeT beds reflect the pinch-
out of the clast-rich debrite interval; hence, a turbidite
interpretation for proximal and basal clean sandstones
within these beds does not compromise the geometric
data.
These tripartite beds indicate that individual sub-
marine events traversing distal basin plains may have
both turbidity current and debris flow components. A
detailed discussion of the origins of cogenetic turbi-
dite-debrite beds from the Marnoso-arenacea Forma-
tion is provided by Talling et al. (2004). Beds with a
similar tripartite structure are noted from other deep-
water systems and have also been interpreted as
cogenetic turbidite-debrite beds (e.g., Hickson,
1999; Kneller and McCaffrey, 1999; McCaffrey
and Kneller, 2001; Haughton et al., 2003; Talling
et al., 2004).
5. Discussion
5.1. Implications for recognition criteria
Certain sedimentary characteristics preserved in
outcrops are indicative of depositional process, as
determined independently using bed geometry. Turbi-
dite sandstones may be recognised on the basis of
being relatively well sorted and possessing relatively
low mud contents, normal grading and traction struc-
tures. Massive ungraded sandstones that are relatively
well sorted with low mud contents (often described as
the Bouma Ta division in other studies) also common-
ly occur within turbidite sandstones. In addition, tur-
bidite sandstones may contain floating clasts and
inverse grading locally at their base usually associated
with centimetre-thick stratification. The results of this
study suggest that all of these sedimentary features
can be formed by particles progressively falling out of
suspension from a relatively dilute and turbulent flow,
albeit in some instances through a relatively high-
concentration basal flow boundary. These findings
are consistent with the recognition criteria developed
over the last four decades for the deposits of sand-
bearing turbidity currents (e.g., Bouma, 1962; Mid-
dleton and Hampton, 1976; Lowe, 1982; Mutti,
1992). The results do not support the model that
relatively well-sorted and mud-poor massive sand-
stones and those that exhibit inverse grading or float-
ing clasts within this system are deposited exclusively
by (sandy) debris flows as suggested by Shanmugam
and Moiola (1995, 1997) and Shanmugam (1996,
1997, 2000, 2002). Field data indicate that debrites
are characterized by moderate-to-poorly sorted, un-
graded and structureless sandstones with floating
clasts and a relatively high matrix mud content, con-
curring with many previous models (e.g., Middleton
and Hampton, 1976; Lowe, 1982; Mutti, 1992).
Experimental studies have suggested that very
low mud contents are sufficient to suspend sand-
sized particles and cause debris flow behaviour.
Hampton (1975) calculated that sediment mixtures
of fine sand with clay contents as low as 2 wt.%, or
less, can move as debris flows. Marr et al. (2001)
produced sandy debris flows with mud contents of
0.7 wt.% using bentonite and 7 wt.% using kaolin-
ite. However, much higher mud contents were
found in turbidite beds of the Marnoso-arenacea
Formation; mud contents of up to 12% by volume
are recorded in turbidites, although measured values
include both detrital and diagenetic mud. Hence, the
interpretation of ancient sandstones as debrites
based on low absolute values of mud content, i.e.,
in the range of a few percent as suggested by
experiments, appears problematic if not simply be-
cause it may be difficult to distinguish detrital from
diagenetic mud. However, it was found that debrites
of the Marnoso-arenacea Formation did have rela-
tively high mud percentages (N20% by volume)
compared to cleaner turbidite sandstones. Hence,
relative values of mud content, as opposed to ab-
solute values are possibly a better indicator of
depositional mechanism.
The bed correlations show that submarine flows
are capable of depositing both debrite and turbidite in
the same event so that different deposit types may
pass laterally into one another over short distances
(Fig. 3B). Hence, recognition criteria may be used to
identify depositional mechanism locally but may not
always be valid away from points of control. For
example, a thin laminated turbidite sandstone may
represent the lateral equivalent of either a relatively
thick turbidite sandstone or debris flow unit. Future
work should be aimed at distinguishing those sedi-
mentary characteristics that can be used to predict the
character of beds laterally.
L.A. Amy et al. / Sedimentary Geology 179 (2005) 163–174 173
Acknowledgment
The research was funded by the United Kingdom
Natural Environmental Research Council and Conoco
(now ConocoPhillips) through the OCEAN MAR-
GINS-LINK scheme (grant number NER/T/S/2000/
01403). We thank Luca Martelli and Giani Zuffa for
advice concerning the field area and George Postma,
Guido Ghibaudo, Stan Stanbrook and Jaco Baas for
constructive reviews.
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