Marine Geology 220 (2005) 131 – 151
www.elsevier.com/locate/margeo
Channelled erosion through a marine dump site of dredge spoils at
the mouth of the Puyallup River, Washington State, USA
Neil C. Mitchell *
School of Earth, Ocean and Planetary Sciences, Cardiff University, Main Building, Park Place Cardiff CF10 3YE, Wales, UK
Received 17 November 2004; received in revised form 4 June 2005; accepted 16 June 2005
Abstract
Channels are relatively common on river-mouth deltas, but the process by which they arise from river sediment discharge is
unclear because they can potentially be explained either by negatively buoyant (hyperpycnal) flows produced directly from the
river outflow or by flows generated by repeated failure and mobilisation of sediment rapidly deposited at the delta front.
Channels eroded through a dump site of dredge spoils are described here from multibeam and older sonar data collected in
Commencement Bay, at the mouth of the Puyallup River. Shallow channels on the seaward upper surface of the dump site, away
from any flows that could have been produced by delta front failures, suggest that at least some hyperpycnal flows were
produced directly from the positively buoyant river outflow up to 200 m from the edge of the river mouth platform. The form of
channel bed erosion is revealed by the longitudinal shape of the main eroded channel compared with the adjacent dump site
profile. It suggests that the channel evolved by its steep front retreating, rather than by simple vertical entrenchment or
diffusive-like evolution of the profile, a geometry interpreted as evidence that repeated failure of the bed occurred in response to
shear stress imposed by bottom-travelling flows. Model calculations based on shear strengths back-calculated from the
geometry of channel wall failures suggest that, if the main channel were eroded solely by hyperpycnal flows, their generation
was remarkably efficient in order to create flows vigorous enough to cause channel bed failure. Besides the sediment
concentration and discharge characteristics that have been considered to dictate the ability of rivers to produce hyperpycnal
flows, it is suggested that the timing of floods with respect to the tidal cycle should also be important because extreme low tides
may be needed to ensure that coarse sediment is transferred vigorously to the edge of river mouth platforms.
D 2005 Elsevier B.V. All rights reserved.
Keywords: hyperpycnal flows; underflows; multibeam sonar; slope stability; channel erosion
1. Introduction
* Tel.: +44 29 2087 5051; fax: +44 29 2087 4326.
E-mail address: [email protected].
0025-3227/$ - see front matter D 2005 Elsevier B.V. All rights reserved.
doi:10.1016/j.margeo.2005.06.032
Channels have often been described eroded into the
surfaces of small coastal submarine fans near river
mouths or in mine tailing deposits (Boe et al., 2004;
Bornhold and Prior, 1990; Ferentinos et al., 1988;
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N.C. Mitchell / Marine Geology 220 (2005) 131–151
Hay, 1987; Mosher and Hamilton, 1998; Normark and
Dickson, 1976; Prior and Bornhold, 1989; Prior and
Bornhold, 1990; Prior et al., 1981a,b; Syvitski et al.,
1987; Syvitski and Farrow, 1983). For example, the
apices of fjord-head and -side deltas commonly contain a down-slope-oriented shallow swale and ridge
topography in gravel and coarse sands which leads
down-slope into channels with well-defined walls in
the middle fan (Bornhold and Prior, 1990; Prior and
Bornhold, 1989, 1990; Prior et al., 1981a,b). Flutes
and truncation of fan stratigraphy observed from submersible (Bornhold and Prior, 1990; Prior and Bornhold, 1989, 1990) demonstrate that the channels were
created by erosion rather than by inhibited deposition
relative to interfluves. Levees and channel-floor bedforms suggest that they are commonly traversed by
turbidity currents at least in the lower fan (Syvitski et
al., 1987).
The precise origins of the erosive flows and their
relationships to their adjacent rivers have been
unclear, however. Some channels lead up-slope to
embayments in the upper delta (Prior et al., 1981a,b;
Syvitski and Farrow, 1983) where arcuate faults and
blocks suggest that erosive flows originated by slope
failure and disintegration of the failed material (Prior
et al., 1981a,b). Failure could also result from stresses
caused by surface waves or loading associated with
flushing of river mouth bar sands onto the upper delta
slope during floods, potentially involving pore-water
over-pressure effects if the deltas contain clay-rich
sediments deposited during quiescent periods (Prior
et al., 1981a). Alternatively, channels may be eroded
more directly from river outflow during floods if the
density of parts of the outflow can overcome seawater
density to produce hyperpycnal flows or if flow inertia
can carry a slurry of boulders, gravels and sands onto
the upper fan (Prior and Bornhold, 1990). Coarsegrained deposits representing river-mouth sand bars
produced by flood-generated hyper-concentrated
flows have been described by Mutti et al. (1996)
based on outcrop studies.
The confusion is also illustrated by studies of submarine canyons heading near river mouths. Some
potentially different origins for sedimentary flows in
the Mediterranean Var Canyon are discussed by
Mulder et al. (1998a) who suggested that rapid
deposition and failure of river delta front sands contributes significantly to sedimentary flows, as well as
hyperpycnal flows. Recordings of turbidity currents in
some channels have been found to be poorly correlated with river discharge, favouring storms or other
causes of flow initiation (Bornhold et al., 1994; Puig
et al., 2004, 2003). On the other hand, turbidity
currents recorded in Monterey Canyon were found
to contain low salinity water and did correlate with
discharge of the adjacent Salinas River (Johnson et al.,
2001). Cores recovered from Capbreton Canyon lying
near the Adour River, France apparently recorded a
storm-generated turbidite rather than a hyperpycnal
flow (Mulder et al., 2004). Sediment cores from
Sepik Canyon, which penetrates into the Sepik
River valley, Papua New Guinea, show many finegrained turbidites (Walsh and Nittrouer, 2003) but it is
unclear if they originated directly from river outflow
or from failure of deposits. A difficulty in addressing
this problem is that there is unfortunately a lack of
observations of processes occurring in the critical
region where flood waters pass from the river mouth
and cross the delta front. The different origins of the
erosive flows are usually difficult to distinguish from
morphology because sonar data from around river
mouths (e.g., Klaucke and Cochonat, 1999) can
show submarine channels heading near the river
mouth but allow for either origin. Puyallup River
mouth data (Gardner et al., 2001b) collected by the
National Oceanographic and Atmospheric Administration (NOAA) and the US Geological Survey
(USGS) provide an unusual geometry that is shown
in this paper partially to address this question as the
data show evidence for channelled erosion on the
outer slope of the dump site in areas where they
cannot have been generated by delta front failure.
The simplest hyperpycnal flows are produced
when the suspended sediment load of a river causes
its density to be greater than that of seawater (Bates,
1953), illustrated by an Icelandic jökulhaup shown in
Mulder et al. (2003) which plunged abruptly at the
coastline leaving only a minor sea surface plume. A
study of river suspended sediment–discharge relationships ("rating curves") suggested that hyperpycnal
behaviour is most likely for rivers draining mountainous areas with small catchment areas (Mulder and
Syvitski, 1995), such as the Puyallup River of
Washington State (Fig. 1). Apparently, hyperpycnal
flow can also occur when the river suspended load is
less than the 42 kg/m3 required for bulk negative
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N.C. Mitchell / Marine Geology 220 (2005) 131–151
48˚ 0N
Puget Sound
1965 Seattle-Tacoma
47˚ 40N
Commencement
Bay (Fig. 2)
47˚ 20N
Seismic magnitude scale
2001
Nisqually
4
Puyallup
gauge station
47˚ 0N
5
6
Mount Rainier
46˚ 40N
46˚ 20N
123˚ 0W
122˚ 0W
121˚ 0W
Fig. 1. Location of the Commencement Bay area of Washington State, USA. Solid circle marks the Puyallup River gauging station. Solid star
symbol locates Mt Rainier. The dashed outlined area around Commencement Bay locates the data in Fig. 2. The open box symbols represent
epicentres of earthquakes of magnitude 4.0 and greater for the period 1974 to present from the Advanced National Seismic System (http://
www.anss.org). The 1965 Seattle–Tacoma earthquake epicentre (Ichinose et al., 2004) is also shown. Scale and globe are placed over areas with
no data.
buoyancy in seawater because, for example, data from
bottom moored instruments in the Monterey Canyon
recorded hyperpycnal flows more frequently than
expected from the rating curve for the Salinas River
(Johnson et al., 2001). The specific processes
involved during this behaviour have rarely been
observed in the field. Suspended sediment concentration measurements in the outflow of the Sepik River,
Papua New Guinea (Kineke and Sternberg, 2000;
Kineke et al., 2000) and the outflow of the Yellow
River (Wright et al., 1988) showed them to have
separated into buoyant (hypopycnal) surface plumes
and hyperpycnal currents, but data on the region of
flow separation are generally lacking.
Laboratory experiments (Hoyal et al., 1999; Maxworthy, 1999; Parsons et al., 2001) suggest that the
separation into bottom and surface plumes can occur
by particles settling from the hypopycnal plume to
form descending denser fingers within the underlying
salt water which coalesce to create the hyperpycnal
current. This process has been suggested to reduce the
critical river sediment concentration needed to initiate
hyperpycnal flow to 5 kg/m3 (Mulder et al., 2003) or
even b 1 kg/m3 (Parsons et al., 2001, in press). As
freshwater signals have been recorded in hyperpycnal
currents (Johnson et al., 2001; Kineke et al., 2000),
the process presumably can also involve water from
the over-riding plume becoming entrained in the negatively buoyant fluid. Further experimental results of
McLeod et al. (1999) illustrate that stratification of the
flow entering seawater is likely to be important, as
hyperpycnal flows form from the particle-rich basal
layer. Kineke et al. (2000) also suggested that the
mixing interface between fresh and salt water can
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N.C. Mitchell / Marine Geology 220 (2005) 131–151
promote flocculation of clay-grade particles and hence
more rapid settling of the larger aggregated particles
produced.
Mulder et al. (1998b) and Skene et al. (1997)
modelled bed erosion by a hyperpycnal flow in the
Saguenay Fjord, Canada assuming that the eroded
material was cohesive with a linear shear strength
depth profile. Erosion depth was varied with u 2 / a,
where ’u’ is flow velocity and ’a’ is the rate of shear
strength variation with depth, as normally compacting
cohesive sediments typically have increasing shear
strength with burial depth (Skempton, 1970). The
model illustrates that erosion by hyperpycnal flows
is likely to be limited to the depth to which the shear
stress due to the flow and the sediment shear strength
are balanced if erosion occurs by bed shear failure, as
suggested later in this paper.
Studying the form of erosion in specific cases,
such as the Puyallup channels, may be helpful in
understanding how turbidity currents erode channels
more generally. Multibeam sonar data of deep submarine channels have revealed the presence of
knickpoints (locally steep reaches or inflections in
gradient). Many of these on steep continental slopes
originate from variations in lithological competence
or tectonics (Bourillet et al., 2003; Farre and Ryan,
1985; Huyghe et al., 2004; McHugh et al., 1993;
Mitchell, 2004, 2005; Ramsey et al., 2003), but the
results presented here may be relevant to canyons of
the lower slope where sediment dams produced by
canyon wall failures are evacuated from channels.
Knickpoints in high resolution multibeam data of the
Monterey Canyon, for example, do not appear to be
as prominent or as common as would be expected
from the volumes of observed canyon wall slope
failures if they were to fill the canyon floor (Greene
et al., 2002). The question then arises as to whether
this is because failed wall material disintegrates
rapidly to form mobile debris flows and turbidity
currents, which travel to the lower fan, or whether
dams of the material are eroded by subsequent flows
travelling down-canyon. If the latter, the form of
erosion is important to whether knickpoints produced by canyon-blocking material would be preserved or smoothed out by successive down-canyon
flows, analogous to the behaviour of moderately
steep knickpoints in bedrock channels which are
predicted to be advective or diffusive depending on
the bed erosion law operating (Whipple and Tucker,
2002).
This paper primarily concerns morphological
observations from high-resolution multibeam echosounder data collected in 2001 on the NOAA vessel
Rainier and survey launches (Gardner et al., 2001b).
Those data and evolution of the southern part of
Commencement Bay in legacy soundings are described first, before alternative origins for channels
imaged by the data are discussed. Model calculations
for the evolution of the channel floor and the discharge/sediment concentration are also presented.
They suggest that, if the channels were carved primarily by hyperpycnal flows, flow generation may
have been remarkably efficient as a significant shear
stress imposed by hyperpycnal or other flows on the
channel bed appears to be necessary to explain its
inferred retrogressive evolution. The influence of tidal
bay water level on river water outflow velocity is also
discussed because it may be important for mobilising
coarser grained sediment across the river mouth platform and hence for initiating strongly erosive hyperpycnal flows.
2. Background to the area and river sediment
Commencement Bay and its surrounding area are
shown in Figs. 1 and 2. The new multibeam data
cover the Puyallup River delta front (Dragovich et
al., 1994) where the river empties into the bay. The
Puyallup River and its tributaries drain the west slopes
of Mount Rainier (star symbol in Fig. 1), originating
from glaciers in the upper parts of the mountain. A
geological map (Fiske et al., 1963) shows that the
upper reaches of the river traverse basaltic andesite
flows and volcanic epiclastic and pyroclastic deposits.
Large debris flows initiated by glacial outbursts,
storms and subglacial eruptions have carried poorly
sorted material down to the lower flanks of the volcano, including the lower Puyallup River valley (Dragovich et al., 1994; Swanson et al., 1992). Their
deposits contain abundant clays of possible hydrothermal origin (Vallance and Scott, 1997). For 13 km
upstream from the river mouth, the Puyallup River
is constrained by artificial levees, ensuring rapid outflow during floods and rapid remobilization of seasonally accumulated bed sediments.
135
N.C. Mitchell / Marine Geology 220 (2005) 131–151
weight % per grade
80
60
o
122o 28' c
a
b
C Si Sa Gr
d
C Si Sa Gr
e
122o 27'
122o 26'
122o 25'
122o 24'
40
20
0
80
C Si Sa Gr
f
60
40
20
0
80
60
47o 18'
C Si Sa Gr
g
C Si Sa Gr
C Si Sa Gr
C
C Si Sa Gr
i
h
40
20
0
Si Sa Gr
C
Si Sa Gr
0
14
47o 17'
47o 17'
b
12
0
c
60
80
10
0
a
40
20
f
47o 16'
r
ive
pR
llu
ya
Pu
47o 16'
d
e
g
i
h
1 km
122o 28'
122o 27'
122o 26'
122o 25'
122o 24'
Fig. 2. Map of the 2001 multibeam sonar data collected in Commencement Bay, Washington State (Gardner et al., 2001b). White depth contours
are shown every 20 m. Grey background is an image of the city of Tacoma. The figure was produced by scientists of the US Geological Survey
(http://walrus.wr.usgs.gov/pacmaps/ps-puy.html). Also shown are sediment sample locations (a–i) which correspond to the grain size histograms
shown in the upper left inset (C, Si, Sa and Gr represent clay, silt, sand and gravel grades). The sediment texture data were derived from the
Washington Department of Ecology SedQual database and represent samples taken in (a–c) 1999, (d) 2004, (e) 1987, (f) 1991, 1995, 1996, and
(g–i) 1984.
The dump site was reportedly used for depositing
dredge spoils from the adjacent waterways until the
mid-1980s (Gardner et al., 2001b). Existence of a
mound in the following legacy soundings suggests
that dumping began before 1936. Although few sediment samples are available from the dump site itself,
erosion of the debris flow deposits as well as other
material derived from Mount Rainier suggest that
sediment carried by the river likely comprises a poorly
sorted mixture of both sands and finer grade material.
According to an unpublished report by the company
HartCrowser, the river bed materials presently comprise mostly silty sand but with all grades from clay to
sand represented. One sample taken by the author at
low tide from the outlet of the Puyallup River was
found to be a stiff, silty sand. Port of Tacoma staff
observed an event in the late 1980s in which wooden
pilings breached the surface (S.P. Palmer, pers comm.,
2003), which could have resulted from one or more of
the slope failures. Hence the dump site material is
136
N.C. Mitchell / Marine Geology 220 (2005) 131–151
UTM distance north (km)
most likely a heterogeneous mixture of grain-sizes but
also including some man-made objects and the river
sediment is also heterogeneous.
This view is confirmed by grain size analyses of
surface sediment samples taken between 1981 and
2004 which are summarized in the top-left of Fig. 2
(derived from the Washington Department of Ecology
SedQual database). As these are surface samples, it is
not entirely clear how representative they are of the
dump site material but nevertheless they show silty
5237
5236
5237
10
0
5236
5235
5235
1936
1936-1974
5237
5237
-24 -12 0 12 24 m
5236
5236
5235
5235
1974
1974-1996
5237
5237
5236
5236
5235
5235
1996
1996-2001
5237
5236
5237
100
5236
5235
5235
2001
541
1974-2001
542
543
544
541
542
543
544
UTM distance east (km)
Fig. 3. (left) Bathymetry maps of southern Commencement Bay, Washington State, derived from data collected by the US National Ocean
Service in 1936, 1974, 1996 and 2001. The modern coastline taken from Fig. 2 is shown in each figure. Contour interval is 20 m, with 100 m in
bold. Depths are relative to Mean Low Low Water. Coordinates are Universal Transverse Mercator zone 10 distances in km (WGS84 ellipsoid).
(right) Depth changes calculated from the bathymetry data. Contour interval is 6 m.
N.C. Mitchell / Marine Geology 220 (2005) 131–151
sand on the edge and within the channel of the dump
site (locations d and e), sandy silts in more distal areas
within the bay (locations a–c) and other sandy silts
around the margins of the dump site (locations f–i).
For discussion of possible triggering of submarine
failures described later, the open boxes in Fig. 1 show
the epicentres of all earthquakes in the region of
magnitude z 4.0 from 1974 to present derived from
the Advanced National Seismic System (http://
www.anss.org), representing the period since the
1974 data (described below) were collected by the
National Ocean Service (NOS). The location of a
further large (M w 6.7) earthquake, the 1965 Seattle–
Tacoma earthquake, is also shown. Modelling by
Ichinose et al. (2004) of strong motions caused by
the largest earthquake (M w 6.8 Nisqually of 2001)
suggest that peak horizontal ground accelerations in
Commencement Bay reached around 0.1 g. Their
modelling of the 1965 earthquake, however, suggests
that ground accelerations may have reached 0.3 g in
the bay. Although there is little evidence for bathy-
137
metric changes associated with the 2001 earthquake,
slope failure associated with the 1965 earthquake may
explained some features observed in the 1974 bathymetry dataset.
3. Bathymetry data origins and processing
All the bathymetry data in Fig. 3 were collected by
the NOS. Those for 1936–1996 were provided by the
hydrographic survey database of the National Geophysical Data Center and those for 2001 were provided by
the USGS (Gardner et al., 2001a). The original data
were collected by lead line (1936), single-beam echosounder (1974 and 1996) and multibeam sonar (2001)
with the coverage shown in Fig. 4. Depths were
reduced consistently to Mean Low Low Water
(MLLW) but horizontal datums differed—the 1936–
1996 data in North American Datum 1927 needed to
be translated northwards to coincide with the 2001
data, which were provided in the World Geodetic
Fig. 4. Sounding density maps for the bathymetry data shown in Fig. 3. Contours and modern coastline are from Fig. 3.
138
N.C. Mitchell / Marine Geology 220 (2005) 131–151
Fig. 5. Enlarged view of the 2001 multibeam data. The data are contoured every 10 m and coordinates are Universal Transverse Mercator
(UTM) distances in km (UTM zone 10 referenced to the WGS84 ellipsoid).
System 84 datum. Coordinates in Figs. 3–6 are Universal Transverse Mercator (UTM) zone 10 distances
in km. Although the Global Positioning System (GPS)
is known to have been used in 1996 and 2001, the
navigation technique and accuracy are unfortunately
unknown for the earlier surveys. Because of the likely
differing resolutions of navigation data, the 1936 and
1974–1996 soundings were first binned at 20 and 10
m, respectively, before interpolating with a minimum
surface curvature algorithm (Smith and Wessel, 1990)
and contouring (Wessel and Smith, 1991). The maps
on the right-hand side of Fig. 3 show elevation changes
for the periods marked.
4. Observations of bathymetric evolution
Before describing the 2001 morphology, the evolution of the southern Commencement Bay from dumping, river sediment deposits, erosion and landslides is
described using Fig. 3 and the close-ups of the river
mouth data in Fig. 6. The 1936 lead line soundings
clearly delimit the early dump site centred around
UTM 542.9 km E, 5235.5 km N, separated from the
seaward extension of the river bed by a 200-m-wide
depression. The 1974 echo-soundings show this
depression largely infilled by the prograding river
mouth sediments, which have moved the break in
slope a little greater than 100 m northwards. The
dump site area grew generally in relief from dumping
and from this progradation, leaving the largely positive relief change shown in the 1936–1974 depth
difference map (upper-right of Fig. 3). Also revealed
in the 1974 data, an early trace of a submarine channel
can be seen running northwest from UTM 542.9 km
E, 5235.45 km N. The channel is 100 m wide at its
southeast end, narrowing to 50 m to the northwest,
although it is poorly resolved in these data.
The 1996 sounding data reveal a well-defined
amphitheatre at the head of the channel centred
around UTM 542.8 km E, 5235.5 km N. A second
amphitheatre is also clearly resolved on the north side
of the dump site around UTM 543 km E, 5235.6 km
N. The northern break in slope of the delta top has
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N.C. Mitchell / Marine Geology 220 (2005) 131–151
50
0
0
UTM distance north (km)
5235.5
1936
0
1996
50
50
5235.5
0
1974
542.5
2001
543.5 542.5
543.0
543.0
543.5
UTM distance east (km)
Fig. 6. Closeups of the Puyallup River mouth bathymetry. Contour interval is 2 m, with 10 m contours in bold. Also shown by the dotted lines
are the !1 and + 1 m depth contours. Depths are relative to MLLW. The modern coastline is reproduced from Fig. 3.
advanced more than 100 m farther northwards, despite
a submarine landslide that reportedly occurred in the
north slope in 1992 (Gardner et al., 2001b). The
shallow bank between the outermost artificial walls
of the Puyallup River channel and the delta front has a
shallow depression of 1 m depth elongated parallel to
the river trend with flanking shallows at the north and
west margins of the bank. The 2001 multibeam data
reveal a general deepening of the upper channel from
1996 (suggested by the depth changes in Fig. 3 and
from the longitudinal sections in Fig. 8). Apart from
loss of a small ridge in the main channel amphitheatre
at UTM 542.9 km E, 5235.47 km N, other differences
around the amphitheatres are largely attributable to
differences in data resolution.
Surveys of the Puyallup River bed between 1943
and 1951 by the U.S. Army Corps of Engineers
(USACE, 1951) showed that the delta front at the
river mouth was prograding by ~20 m/year. Given
the narrow ~200 m width of the Puyallup channel, that
rate is comparable volumetrically to the long-term
progradation rate of the 3-km-wide delta front filling
the Puyallup valley, which has averaged 2.3 m/year
based on the 13 km advance of the front since a dated
mudflow (Osceola) was emplaced at 5.7 ka (Dragovich et al., 1994). Compared with those rates, however, the 100 m extensions of the delta front 1936–
1974 and 1974–1996 seem somewhat modest, implying that some loss of the delta front slope may have
occurred during these periods by slope failures, such
as the 1992 slope failure reported by Gardner et al.
(2001b) and an April 1943 slide that removed part of a
river training wall (USACE, 1951).
5. Morphological observations from the 2001
multibeam data
The 2001 data (Fig. 5) show a well-defined 100-mwide channel (along section B–BV shown) floored by
large sedimentary dunes. In section B–BV, the dunes
can be seen with varied spacings of 20–80 m and
heights of 3–6 m. In (Figs. 2, 3 and 6), other channels
can be seen on the north side of the dump site and one
small channel on the south side. Embayments in the
upper east wall of the main channel and adjacent slope
140
N.C. Mitchell / Marine Geology 220 (2005) 131–151
farther east mark slope failures (e.g., immediately
northeast of UTM 542.8 km E, 5235.5 km N marked
by the dashed lines in Fig. 5). Importantly, four shallow
channels can be seen radiating to the northwest from
the top of the dump site (UTM 542.8 km E, 5235.2 km
N) at around 10 m depth. These can be seen as small 1–
3 m deep depressions east of the main channel in Fig. 7
(400–500 m in profiles 5–8 from the bottom).
The long-profile of the channel in Fig. 8 reveals a
general decline in gradient with distance from the
river mouth (gradients are mostly 3–58, with a maximum of 88) and with vertical relief declining from
nearly 20 m. This is also emphasized by the crosssections in Fig. 7, which show a moderately flat
channel floor with sharp margins. The wall failures
occur at 400–500 m in the three lowermost profiles in
the figure. Northwest of UTM 542.5 km E, 5235.7 km
N (Fig. 5), the channel disappears, possibly a result of
the carving sediment flows losing power because of
flow spreading and loss of sands from suspension
(Akiyama and Stefan, 1988), mixing with bay waters
or the decline in bed gradient (Fig. 8). The flows then
deposited their loads. From the two white outlined
areas in the lower-right map of Fig. 3, the volume
eroded from the channel, 2 " 105 m3, was roughly ten
times smaller than that deposited, 3 " 106 m3 (probably a minimum value as it was obtained only over the
area in Fig. 3 though it may also include deposits not
associated with the channelled flows).
6. Interpretations
The origin of the dump site channels is interpreted
here in terms of some of the processes outlined in the
introduction for erosion on fjord delta fronts. Largescale slope failure seems an unlikely origin for the
main channel as landslide chutes tend to be more
irregular and associated with arcuate headwall faults
in the fjord examples (Prior et al., 1981a). The steepest
Relative depth (m)
A
A'
0
50
100
0
100
200
300
Distance (m)
400
500
Fig. 7. Transverse cross-sections showing development of the channel away from the river mouth. (Sections located in Fig. 5).
141
N.C. Mitchell / Marine Geology 220 (2005) 131–151
NW
B'
Gradient
0
Depth (m)
20
40
0.10
Channel margins
N
S
Average
Channel axis
1974
1996
2001
0.15
0.05
Channel gradient
SE
0.00
B
60
80
0
200
400
600
800
Distance from B' (m)
Fig. 8. Longitudinal profiles of the channel and adjacent margins. The channel axis was sampled along B–BV shown in Fig. 5 from the gridded
bathymetry for 1974, 1996 and 2001. The channel margins were sampled from the 2001 bathymetry along the north and south lines parallel to
B–BV shown in Fig. 5 (dashed where reconstructed across landslide embayments). Gradient (dotted line) was calculated over a 100 m
lengthscale along the channel axis by regressing depth on distance.
elevation change occurs to the west of the dump site,
rather than parallel to the main channel. Repeated
cycles of deposition and smaller-scale slope failure
around the upper channel-head amphitheatres may,
however, have contributed some erosive flows for
the channels’ excavation. Deposition-failure cycles
would need to have led to relatively little net change
in the upper slope because the edge of the west channel-head amphitheatre, for example, has remained near
UTM 543 km E between 1974 and 2001.
Settling of sands and silts from the positively buoyant flooding river outflow into underlying seawater as
it spread beyond the edge of the shallow river-mouth
platform is suggested to have led to hyperpycnal
flows, such as by the types of mechanisms described
by Parsons et al. (2001). The 20 m descent into the
canyon-head amphitheatre floors provided potential
energy for descending particle-laden fluid to gain
turbulence and retain sand in suspension. Those captured sediment-laden flows then passed down the
channels below the amphitheatres. The alignment of
the main channel with the river channel implies that
the direction of flow was influenced by momentum of
the river outflow. The sediments providing negative
buoyancy were presumably relatively coarse grained
because the loss of erosive channel relief and deposition immediately below the dump site suggests rapid
settling and deposition.
The shallow channels on the upper northwest
smooth surface of the dump site are evidence that
negatively buoyant flows continued to be sourced
by the overriding positively buoyant flood waters
for 200 m from the platform edge. They are difficult
to explain by other mechanisms (e.g., wind-driven
downwelling would be expected to produce similar
channels elsewhere around the bay). The channels are
much shallower than the main channel, however, so
most erosive flows were generated within 200 m of
the platform edge. As seawater was carried away from
the delta front area by shear from the overlying flood
river outflow and underlying hyperpycnal flows, the
laterally open aspects of the two amphitheatres could
have allowed seawater to be recharged from the sides.
The longitudinal shapes of the channel margins
and axis shown in Fig. 8 are different—whereas the
channel margins are convex upwards in long-profile,
the channel axis is concave upwards. This geometry
is interpreted as suggesting that the channel floor
evolved by the steep front of the dump site retreating. The fine continuous lines in Fig. 9 (top profile)
show a simple kinematic model developed to illustrate such an evolution. The curves were constructed
using a simple finite difference calculation in which
the channel was represented by a 1-m array of
elevation values, starting with the channel margins
long-profile as the initial condition. During each
142
N.C. Mitchell / Marine Geology 220 (2005) 131–151
0
E~S
3
20
40
60
Depth (m)
0
E~S
20
2
40
60
0
E~S
20
1
40
60
0
100
200
300
400
500
Distance (m)
Fig. 9. Kinematic models for the development of the channel profile. In each model, the channel evolves from an initial state represented by the
channel margins profile (upper bold line) by successive removal of layers of sediment according to the rule E~S n (where S is slope and S N 0)
where n = 1 to 3 in the three graphs. A small 5-m averaging filter was applied during each iteration to stabilise the solutions. A boundary
condition of fixed topography at x = 490 m was applied, representing the loss of channel relief at that location.
iteration, the bed elevation was reduced by a small
thickness E (m) which was varied with the bed
gradient S (m/m) according to E~S 3 (S N 0, S calculated by differentiating the elevations) while the
right-hand edge of the model was constrained to
have zero erosion based on the observed loss of
channel relief. This scheme was not intended to
represent any particular theoretical erosion law, but
rather to illustrate that erosion was strongly accentuated by steep gradients, a clue that an effect of
gravity was involved.
An interesting possibility therefore is that erosion
was caused by bed failure under the combined stress
imposed by the flows and gravity acting on the bed
sediments. This idea is pursued in the following sections. A second possibility not pursued, but which
cannot be fully ruled out, is that bed erosion occurred
by abrasion by bedload particles. Abrasive erosion
should also be enhanced on steeper gradients because
the kinetic energy of saltating particles impacting the
bed is expected to be greater where flows are faster
(Hancock et al., 1998; Whipple et al., 2000). However, extreme bedloads can armour the bed, preventing abrasion (Sklar and Dietrich, 2001, 2004), which
seems likely in the Puyallup channels as large bedloads are suggested by the dunes revealed in the
multibeam data (Fig. 5).
7. Back-calculating dump site geotechnical
properties
As there are no geotechnical data available for
the dump site itself, sediment strengths were back-
N.C. Mitchell / Marine Geology 220 (2005) 131–151
calculated from the geometry of the slope failures
around the channel wall. Effects of seismic ground
shaking were ignored because no sediment was
observed at the water surface in the bay during
the 28 February 2001 Nisqually earthquake (S. P.
Palmer, pers. comm., 2004) and the amphitheatre
geometry appears to have changed little between
the 1996 and 2001 surveys carried out before and
after the earthquake (Fig. 6). The possibility of
excess pore pressures, such as those associated
with organic matter degradation, cannot be assessed
without suitable data, but such effects may also help
to explain how the retrogressive channel bed failure
occurred. Fig. 10 shows long-profiles taken down
the trends of two slope failures (dashed sections L1
and L2 shown in Fig. 5), along with gradients
calculated by regressing depth on distance over a
length scale of 100 m. These show that the steepest
parts of the failed surface have gradients reaching
0.235 (138).
The infinite slope approximation (Hampton et
al., 1996) was used to estimate bulk strength properties. First, the material was assumed to be perfectly incohesive so that the gradients of the side
wall failure (0.235) represent the angle of internal
friction l of the sediment, i.e. tanl = 0.235 or l =
138 (where l = internal friction angle). This value is
low but not improbable if the material contains
0
SW
NE
L1
0.2
20
0.1
40
Gradient
Depth (m)
(a)
Depth (m)
(b)
0
SW
L2
NE
0.2
20
40
Bathymetry
Gradient
0.1
Gradient
0.0
0.0
0
100
200
Distance (m)
Fig. 10. Longitudinal sections of two landslides (a) L1 and (b) L2
located in Fig. 5. The gradients shown by the dashed lines were
calculated by regressing depth on along-profile distance with a
lengthscale of 100 m.
143
abundant clays as suspected from the sediment sampling (Fig. 2). Second, the material was assumed to
be cohesive. Based on the gradient angles of the
wall slope failures (h = 138) and that the failures
could have been up to h = 10 m thick if they
occurred as single events (Fig. 5), the undrained
shear strength can be calculated from S u = qVghsinh.
Using a value of 893 kg/m3 for the buoyant density
q (from grain density q q = 2650 kg/m3 (quartz) and a
typical shallow sand porosity of 44% (Masselink and
Hughes, 2003)), the shear strength was estimated to
be 21 kPa.
8. Channel bed shear stress required to cause
erosion by bed failure
For bed shear failure to occur, the combined action
of gravity acting on the bed sediments and stresses
caused by the hyperpycnal flows must exceed the
strength of the bed sediments. A range of scenarios
was tested using the above geotechnical properties to
show that, whereas individual values may be uncertain, they all indicate that relatively large flow stresses
would be required to cause bed failure and hence, if
hyperpycnal flows were responsible, they were
remarkably vigorous.
For the case of noncohesive sand bed, the above
value of tanl = 0.235 was used and also an alternative
value tanl = 0.65 for typical noncohesive sand, to
give a range representing uncertainty. A thickness of
the bed failure h = 0.1–1.0 m was assumed, also a
broad range conservatively to illustrate uncertainty.
Since this calculation addresses the retreat of the
steep front of the channel bed, a gradient h = 78 was
used in the calculation. The critical bed flow stress
needed to cause failure s f was then calculated from
s f = qVgh(coshtanl ! sinh), which was obtained by
adding flow-imposed stress to the stability equations
for an infinite slope (Hampton et al., 1996). From the
range of values given, the range of expected flow
stresses is s f = 98–4600 Pa. These values are plotted
in Fig. 11 by the light grey band.
For the case of a cohesive bed, the critical
flow stress can similarly be calculated from S u =
qVghsinh +s f. Again choosing a range of failure
depths h = 0.1–10 m to illustrate uncertainty, the
critical flow stress was estimated to be s f = 10–21
144
N.C. Mitchell / Marine Geology 220 (2005) 131–151
Cohesive bed
10
τf (kPa)
1
Noncohesive bed
0.1
0.01
0.001
0
5
10
Flow thickness H (m)
Fig. 11. The curve in the figure shows the estimated bed shear stress
assuming that a flow equivalent to a quarter of the Puyallup River
peak discharge travelled down the dump site channel in contact with
the bed (Fig. 5). H is the hyperpycnal flow thickness. Grey bands
show the estimated shear stresses required for failure, which were
estimated by back-calculating dump site geotechnical properties
from the geometry of wall failures. As the bed shear stress only
begins to approach the bed shear strength for H b 1 m, the flow is
inferred to have been thin and vigorous if a hyperpycnal flow.
kPa. These values are plotted in Fig. 11 by the dark
grey band.
9. River discharge and sediment load
characteristics
As with many rivers (Mulder et al., 2003), sediment and water discharge data for the Puyallup
River suggest that its outflow mean density probably
does not exceed sea water density to become negatively buoyant. Daily mean discharge and suspended
sediment concentration data were provided by the
USGS for a gauging station at Puyallup located in
Fig. 1 (http://waterdata.usgs.gov). The USGS water
quality data show that the river water has only 54
mg/l dissolved solids on average so their effect on
water density was ignored. The river discharge
exceeded 1000 m3/s in November 1986 and February 1996 (Fig. 12). Suspended sediment concentration measurements were made only prior to 1994
and do not include the 1987 flood, so extreme
concentrations need to be estimated from historical
data. Suspended sediment sample data are plotted
against the daily mean discharge for the day of
sampling in Fig. 13.
From the outlier near 1000 m3/s in Fig. 13, the
sediment load may have reached 3 kg/m3 during the
floods, somewhat low compared with the minimum
4–5 kg/m3 density excess required for hyperpycnal
behaviour (Mulder et al., 2003), though larger than the
1 kg/m3 suggested by Parsons et al. (in press).
According to Wren et al. (2000), the bottle sampling
techniques typically used by the USGS can underrepresent the heaviest suspended loads which occur
within 10 cm of the bed, although the heavier loads
are unlikely to be carried with the buoyant plume so
the concentration data should nevertheless suit the
purpose here.
To predict sediment concentration (C), the two
lines in Fig. 13 show alternative extrapolations.
The solid line shows a regression of log10(C) on
log10( Q) for Q N 100 m3/s, which suggests that C
could have reached 9 kg/m3 for a discharge of 1000
m3/s. If a graph slope of unity is used, which is
consistent with rating curves of similar-sized rivers
(Johnson et al., 2001; Syvitski and Morehead, 1999),
and constrain the relationship with the data for
Q N 300 m3/s (dashed line in Fig. 13), C = 6 kg/m3
is predicted at 1000 m3/s. As the floods occurred
during winter months, a small temperature effect
could be included to allow for the river waters
being colder than Puget Sound water, e.g. 2 kg/m3
for a 10 8C temperature difference and 2.1 " 10! 4 1/
8C water expansion coefficient (Fischbeck and Fischbeck, 1987). There may also have been a further
effect of dilution of Commencement Bay water by
the flood waters themselves. The bay has a volume
of 1.15 " 109 m3 and the cumulative volume of the
flood waters over the 5 days of highest discharge in
1996 was 3 " 108 m3. Given that the NOAA database of salinity values for the bay shows a relatively
normal seawater mean value of 29.7x, dilution during floods could have reduced the density difference
between river and bay waters by 30% if there was
little exchange between the bay and the broader
Puget Sound (an extreme possibility as the bay is
quite open to Puget Sound). The combination of all
the above effects suggests that, at maximum, the
density anomaly of the river waters could have
been reduced to within 10 kg/m3 of that required
for bulk negative buoyancy.
145
N.C. Mitchell / Marine Geology 220 (2005) 131–151
(a)
Discharge Q (m3/s)
1000
800
600
400
200
Q (m3/s)
(b)
1980
November 1986
1000
Year
1990
February 1996
January 1990 Tide
4
2
0
500
0
25 30
5
2000
10 15 20 25 30
5
Tide h (m)
0
1970
10 15 20 Discharge
Day of the month
10
1
0.1
1
Fr
U (m/s)
(c)
0
25 30
5
10 15 20 25 30
5
10 15 20
Day of the month
Fig. 12. (a) Daily mean discharge of the Puyallup River at USGS water gauge station 17110014 (at Puyallup). (b) Enlargements of discharge
during the two largest floods in 1986 and 1996, and a smaller flood in 1990. Also shown are the bay water levels due to tides predicted using
NOAA software (tidal heights are relative to MLLW). (c) Outflow velocity U estimated from the Puyallup River discharge and its channel width
and depth assumed to be constrained by bay water level. Fr is the water outflow surface Froude number also computed from the discharge.
Extreme U and Fr are exaggerated because dynamic effects on the river outflow surface are ignored (see main text), but the graphs illustrate that
tidal level may be as important in dictating outflow velocity and hence for mobilising bedload at river mouths as river water discharge. Also
shown by the horizontal grey bars are maxima in U and Fr estimated from Q using the d’Arcy–Wiesbach formula for channelled flow which
provides more conservative estimates.
The maximum density difference is thus smaller
than that required for bulk negative buoyancy but the
outflow density excess would have been much larger
than the 1 kg/m3 minimum suggested by Parsons et al.
(2001). The outflow would have been initially hypopycnal (positively buoyant) before denser parts
became unstable and plunged. If transformation
occurred by progressive development of instabilities
due to settling of particles (Hoyal et al., 1999; Maxworthy, 1999; Parsons et al., 2001), this probably
occurred within a distance of 200 m from the edge
of the river-mouth platform after the river outflow
traversed the platform in contact with the bed as
seems likely from the shallow 1996 bathymetry
(Fig. 6).
10. Effects of tidal height
Mobilisation of sediment across the river mouth
platform, and hence formation of hyperpycnal
flows, also likely depends on the bay’s tide because
a high water level implies that the discharge will
spread vertically and slow, and may even override
seawater on the platform causing some sediments to
deposit, whereas strong outflow velocity associated
146
Sediment concentration C (kg/m3)
N.C. Mitchell / Marine Geology 220 (2005) 131–151
1
0.1
0.01
10
100
Discharge Q (m3/s)
1000
Fig. 13. Sediment rating graph constructed from US Geological
Survey suspended sediment concentration (C) and associated discharge ( Q) measurements at Puyallup station 17110014. Discharge
values are the daily mean Q for the dates that samples were
collected (i.e. not instantaneous discharge). Sediment data were
collected 1955–1994. The two lines show alternative scenarios for
extrapolating to predict the river suspended load concentration
during the largest flood in 1996. The continuous line shows
log10(C) regressed on log10( Q) for Q N 100 m3/s, while the dashed
line shows the regression for Q N 300 m3/s but with slope constrained to be unity as for the Eel River (Syvitski and Morehead,
1999).
with extreme low tide will more likely cause coarser grained river sediment to be transported vigorously to the platform edge. The potential for
hyperpycnal flows to develop depends on the density of the base of the outflow (McLeod et al.,
1999), which in turn depends on availability of
river bed sediment, particle settling velocities and
flow stress (Bridge, 2003).
Fig. 12b shows the largest discharges associated
with the 1986 and 1996 floods. A lesser flood in 1990,
however, coincided with a more extreme tide and may
have mobilised sand across the platform more efficiently. Tidal heights shown in Fig. 12b were predicted using NOAA tidal modelling software and
represent water level with respect to MLLW at
NOAA tide gauge site 9446484, located at an adjacent
warf. The prediction error for July–December 1997
when observed tides were available had a standard
deviation of only 13 cm and mean offset of only 12
cm (observed higher on average than predicted water
level height).
Outflow velocity U can be roughly assessed from
the ratio Q / (W(h–h 0)) shown in Fig. 12c (where Q is
discharge, W = 200 m is the channel width, h is bay
water level with respect to MLLW and h = ! 0.5 m is
a typical depth with respect to MLLW of the channel
across the river mouth platform). This value assumes
that the outflow had the same water surface as the
bay water. It exaggerates the extremes of outflow
velocity because in practice the water’s viscosity
will have caused its surface to decline more gradually
into the bay. Hence, the extreme Froude numbers
shown for low tide conditions are also exaggerated
(calculated from Fr = U / M( g(h–h 0)) with g = 10 m/
s2) and the outflow is unlikely to have become as
supercritical as shown. Nevertheless, the graphs illustrate that floods coinciding with extreme low tides
could produce fast turbulent flows carrying sediment
across the river mouth platform. This effect could be
as important as the magnitude of river discharge in
ensuring that larger sediment grades are carried
across to the platform to the channel heads in the
delta front because U and Fr are as large for small
floods coinciding with spring tides as large floods
coinciding with neap tides.
To provide alternative constraints on U and Fr, the
d’Arcy–Weisbach formula for steady channelled flow
was used (Bridge, 2003). The formula represents the
balance between weight of the water driving flow and
friction around the margins. It states that the depthaveraged velocity U = (8gdS / f)1 / 2, where d is the
flow depth (m), S is bed gradient (m/m) and f is a
friction factor. Approximating d with Q / WU, the outflow velocity is predicted to be U = (8gQS / fW)1 / 3.
The velocities shown by the horizontal grey bars in
Fig. 12c were calculated with this relation using peak
discharges Q = 1100 m3/s for 1986 and 1996, and 750
m3/s for 1990, W = 200 m and S = 0.002 (1 m in 500 m
from the 1996 bathymetry in Fig. 6). The two bars in
each graph correspond to f = 0.02 and 0.1 for a sandy
plane bed and dune-covered bed, respectively. These
calculated values are also not particularly accurate but
they reinforce the view that outflows coinciding with
extreme low tides could be effective in creating vigorous sand-laden hyperpycnal flows, as even the
minim velocities of 2 m/s shown are able to mobilise
10 mm grains (Miller et al., 1977).
N.C. Mitchell / Marine Geology 220 (2005) 131–151
Unfortunately, a lack of observations during
floods prevents confirmation that the outflow was
actually supercritical (Fr N 1). If it were supercritical,
however, the loss of velocity associated with flow
expansion and the potential for flow thickening at
the platform edge suggests that the outflow could
have undergone a hydraulic jump there. The associated enhanced turbulence may then have promoted
mixing with underlying seawater and hence hyperpycnal flow generation. On the other hand, the densimetric Froude number of the river outflow-seawater
interface (FrV = U / (Dqgd / q)1/2, where Dq is the outflow density contrast with seawater of density q) was
almost certainly greater than unity for U N 2 m/s,
Dq / q ~0.02 and d ~1–3 m, implying a strongly
unstable lower interface.
11. Comparing flow stress with critical shear stress
inferred from the landslides
The peak recorded discharge of the river (N 1000
m3/s) can be combined with a simple argument of
flow geometry to infer the range of possible flow
stresses that caused channel bed failure. Assuming
that erosion was solely by hyperpycnal flows generated by particle settling, their total discharge was
assumed equal to the river discharge given their
apparent correspondence in laboratory experiments
(Parsons et al., 2001). The different channels provided a number of routes for the flows (Fig. 2),
with probably one fourth of the discharge having
travelled down the channel under study. The maximum specific discharge is therefore estimated to
be Q f ~2.5 m2/s given the channel width of 100
m. The speed of the flow can be estimated from
Q f divided by the flow thickness H. Fig. 11 shows
the variation in bed shear stress as a function
of flow thickness predicted using the relation
s f = q wC du 1002 with C d = 0.006 for rippled sand
(Soulsby, 1997) and where q w is the density of
water (1000 kg/m3) and u 100 is the flow speed (m/
s) at 100 cm from the bed. Approximating u 100 by
Q f / H thus allows us to calculate s f as a function
of H. As the bed shear stress only begins to
approach the estimated bed shear strength for
H b 1 m, the flow was likely thin and vigorous if
hyperpycnal.
147
12. Discussion of channel profile evolution
Although the kinematic models shown in Fig. 9
were not intended to represent stream-power erosion
models, it is nevertheless interesting that the geometry
of the channel is best explained with a strong exponent on gradient (E~S n with n = 3). The model run
with n = 1 (lower graph in Fig. 9) merely advects the
channel margins profile as expected because E~S (i.e.,
Bz / Bt~Bz / Bx, where z is topography, t is time and x
is down-channel distance) has a simple progressive
wave solution (e.g., Paola, 2000). It does not reproduce the observed curvature of the channel bed,
whereas the model with n = 2 provides an intermediate, almost acceptable solution. Weissel and Seidl
(1998) show how knickpoints with n N 1 have solutions in which the propagation speed is an increasing
function of gradient, and hence the solutions with
n = 2 and n = 3 reproduce the bed curvature because
the steep upper part of the channel propagates fastest.
Hancock et al. (1998) developed a scaling argument for quarrying of blocks on river beds in which
shear imposed by the flow is opposed by sliding
friction which depends on block weight (normal
stress). The argument suggested that the vertical thicknesses of joint-bounded blocks that can be quarried by
streams should vary with the square of the local flow
velocity. As the weight of a block depends on its
thickness, a stronger flow can quarry a thicker
block. Allowing for effects of gradient on flow velocity and other assumptions for how bedrock joints
develop over time, it implies an erosion law of form
E~S 2 / 3 (Whipple et al., 2000). As shear strength in
cohesive sediments also typically increases linearly
with depth (Skempton, 1970), a similar scaling argument may also apply. The apparently larger exponent
on gradient inferred here (n = 3) then potentially arises
because sediment can have non-zero shear strength at
the surface (in contrast to block friction in the Hancock et al. (1998) model, which is zero at zero depth)
and because increasing sediment shear strength with
burial depth likely inhibits deeper erosion, hence
tending to prevent the channel eroding back at constant elevation as implied by the n = 1 solution in Fig.
9. Whereas in the simple friction argument of Hancock et al. (1998) cycles of erosion and preparation of
the bed jointing can continue and lead to deep erosion,
unloading of compacted sediments does not usually
148
N.C. Mitchell / Marine Geology 220 (2005) 131–151
lead to their shear strength adjusting to a normal
profile with small cohesion at the surface, hence
deep erosion by shear failure in sediment is likely to
be inhibited.
A further interesting aspect of the long-profile
topography of the dump site channel is that it has
not evolved in a diffusive-like manner. Diffusion can
be predicted in a similar way to alluvial channels
(Paola, 2000), if materials in the bed are readily
detached and erosion is controlled by variations in
transport flux. The diffusion equation can be derived
by combining the original (Bagnold, 1963) result that
specific sand bedload transport flux Q b (kg/m/s)
should roughly follow Q b~u 3 with the expectation
that flow speed u should vary with the third power
of gradient from the Chezy formula for channelized
flows (e.g., Komar, 1969) coupled with a conservation
of flow discharge relation. Using a continuity relation
for the sediment (Bz / Bt = ! 1 / q sBQ b / Bx, where q s is
the sediment density in the channel bed (kg/m3)), this
implies a linear diffusion equation of form Bz / Bt = K /
q sB2z / Bx 2 (K constant). Transport-limited bed evolution may be more complex and at least non-linear in
practice because of the threshold of motion term and
suspended load transport (Soulsby, 1997), the effect of
gravity on bed and suspended loads (Damgaard et al.,
2003) and assimilation of ambient water and bed
material. Nevertheless, at least in simulations the latter
effects are gradual over 100 m lengthscales (Mulder et
al., 1998b; Skene et al., 1997) and flow density should
vary little because the river load dominates the
volume budgets. Aside from the topography of bedforms, some smoothing of relief over shorter lengthscales might therefore be expected even if bed
evolution were not strictly speaking following a diffusion equation. Erosion of these channels was therefore
probably not transport limited but detachment limited
as would be the case with bed failure.
13. Conclusions
From solely the morphological data presented, it is
strictly speaking not possible to infer how much of the
channel erosion occurred by hyperpycnal flows and
how much from repeated deposition and failure of the
upper delta front. However, the four shallow channels
on the upper northwest surface of the dump site, away
from the delta front, suggest that at least some of the
flows were hyperpycnal and that they continued
developing from the positively buoyant outflowing
flood waters 200 m from the edge of the river
mouth platform. Installation of equipment capable of
recording the pattern of sediment and water fluxes
during floods is needed to address this question in
more detail.
The outflow waters were predicted to have had
densities of 4–10 kg/m3 greater than clear freshwater
at bay water temperatures, which are still smaller than
needed to achieve bulk negative buoyancy in salt
water. The development of a plunging hyperpycnal
flow may therefore have occurred by mechanisms
such as described by Parsons et al. (2001), in which
particles settled from the positively buoyant river outflow to create a denser particle-laden flow in the
underlying salt water within the channel head
amphitheatres overridden by flood waters at the
edge of the river mouth platform.
A potentially important further factor for hyperpycnal flow initiation is tidal height because extreme
low tides may be needed to allow rapid flow across
the river mouth platform and mobilisation of larger
sediment grades. Deceleration of outflow during
higher water levels may instead lead to sand bars
being deposited. Considerations of tidal cyclicity
and river mouth bathymetry with schemes based on
river rating curves (Mulder et al., 2003) may be
necessary for a more accurate prediction of hyperpycnal flow frequency and intensity.
Erosion of the main channel led to retrogressive
retreat of the steep channel bed, rather than a diffusive-like evolution of the channel bed profile which
might be expected from bedload transport flux
modulated by how flow velocities vary with bed
gradient (Paola, 2000). Bed evolution appears to
have occurred with erosion rate E modulated by
gradient S according to E~S 3, i.e. with a much larger
gradient exponent than has been expected for quarrying in bedrock river channels (Whipple et al.,
2000). Bed failure is probably limited by increasing
sediment shear strength with burial depth. Scenarios
were assessed for the bed stress required for failure
based on sediment geotechnical properties back-calculated from slope failure geometry and using the
Puyallup River’s peak discharge to predict possible
stresses that could be created by hyperpycnal flows.
N.C. Mitchell / Marine Geology 220 (2005) 131–151
Although the calculations are not accurate, they suggest that, if hyperpycnal flows were the sole cause of
erosion, they would need to have been vigorous in
order to cause bed failure and hence the creation of
hyperpycnal flows from floodwaters may have been
quite efficient.
Acknowledgements
The multibeam sonar mapping was carried out
from the NOAA vessel Rainier under the direction
of Captain Daniel Herlihy and Lt. E.J. van den
Ameele of NOAA and of Jim Gardner of the USGS.
The USGS and NOAA generously provided access to
these multibeam, other sounding and associated data.
Jim Gardner is thanked for clarifying some technical
aspects of the data. Peter Dartnell very kindly confirmed permission to reproduce Fig. 2 and provided
coordinate data. Phil Allen suggested the idea of
calculating outflow Froude number. I thank also
Andrew Barclay for helping to locate sediment samples from Commencement Bay and Martin Payne of
the Washington Department of Ecology for providing
the sediment sample summaries shown on Fig. 2.
Gene Ichinose helpfully clarified aspects of seismic
modelling. I am in particular very grateful to Steve
Palmer of the Washington State Department of Natural Resources, Geology and Earth Resources for a
very enjoyable day in the field examining the bay and
Puyallup channel and sampling Washington cherries,
as well as providing some historical data and reports.
Thierry Mulder, an anonymous reviewer and editor
David Piper gave some very helpful feedback which
led to significant improvements of this paper. This
research was supported in part by a research fellowship of the Royal Society and a Cardiff University
travel award while the author was on research leave at
Scripps Institution of Oceanography.
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