Experiments in Wedge-Shaped Deep Sea Sedimentary Deposits Part II - Evolution and Bedforms - JSR, 2009

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Journal of Sedimentary Research, 2009, v. 79, 608–628

Research Article

DOI: 10.2110/jsr.2009.065

EXPERIMENTS ON WEDGE-SHAPED DEEP SEA SEDIMENTARY DEPOSITS IN MINIBASINS AND/OR ONCHANNEL LEVEES EMPLACED BY TURBIDITY CURRENTS. PART II. MORPHODYNAMIC EVOLUTION OF

THE WEDGE AND OF THE ASSOCIATED BEDFORMS

BENOIT SPINEWINE,*1,2 OCTAVIO E. SEQUEIROS,¤1 MARCELO H. GARCIA,1 RICK T. BEAUBOUEF,{3 TAO SUN,3

BRUNO SAVOYE,14 AND GARY PARKER5

1Ven Te Chow Hydrosystems Laboratory, Department of Civil and Environmental Engineering, University of Illinois at Urbana-Champaign, 205 North Mathews Avenue,

Urbana, Illinois 61801, U.S.A. 2Fonds National de Recherche Scientifique, Rue d’Egmont 5, B-1000 Bruxelles, Belgium

3ExxonMobil Exploration Co., Houston, Texas 77252, U.S.A.4IFREMER, Labor atoire Environn ements Sedimentaires, BP 70, 29280 Pl ouzane, Fra nce

5Department of Civil and Environmental Engineering and Department of Geology, University of Illinois at Urbana-Champaign, Urbana, Illinois 61801, U.S.A.

e-mail: [email protected]

ABSTRACT: Rapidly decelerating sediment-laden flows typically emplace confined sedimentary deposits. In the fluvialenvironment, when sediment-laden rivers reach lakes the decelerating flow emplaces a subaerial delta with distinctive topset,foreset, and bottomset deposits. In the submarine environment, turbidity currents undergoing rapid deceleration commonlyemplace sedimentary wedges (i.e., deposits thinning in the downstream direction). Froude-supercritical turbidity currents havean intrinsic self-regulating mechanism for deceleration, in that the faster they flow, the more they incorporate ambient seawater through mixing at their interface. In addition, special topographic configurations, such as the entrance into a zone of much lower slope and/or lateral confinement, or the passage into a confined minibasin, may trigger sudden flow deceleration byforcing a transition to subcritical flow through an internal hydraulic jump. The present paper and its companion present

experiments on a generic configuration aimed at studying the emplacement of wedge-shaped sedimentary deposits bycontinuous supercritical density currents. The deceleration is achieved both by natural entrainment of ambient water and by thepresence of an obstructing barrier downstream. Lightweight plastic sediment was used as an analog for sand, and wastransported mostly as bedload, but with some suspension, by a saline underflow. The saline underflow served as a surrogate fora turbidity current driven by fine mud that does not easily settle out. The companion paper is focused on the flow patternsassociated with the decelerating current. The present paper focuses on the depositional sequences. The decelerating supercriticalflows produced a wedge with a distinct pattern of aggradation and progradation. In addition, a foreset-like structure isattributed to the presence of an internal hydraulic jump forced by the downstream barrier. Although they do not reproduce anyspecific field-scale setting, the experiments are deemed a good generic model for several wedge-shaped submarine deposits invarious settings, from slope aprons to deposits in minibasins or on the external flanks of channel levees. The paper alsodocuments the regimes of bedforms associated with the diverse flow regions. It provides the first evidence for the formation of trains of the upstream-migrating sediment waves known as cyclic steps, similar to those commonly observed on channel leveesand also along the thalwegs of some steep canyons. In addition, the experiments provide convincing evidence for the formationof downstream-migrating antidunes as well.

INTRODUCTION

When sediment-laden flows decelerate substantially over shortdistances, they typically create conditions prone to the gradualdeposition of part of their sediment load, and as a consequence

gradually emplace wedge-shaped sedimentary deposits. In the case of

fluvial systems, sharp declines in flow velocities can be forced by the

presence of a reservoir or a hydraulic jump. When a river reaches a

standing body of water such as a lake or reservoir, it emplaces a delta

with a coarse-grained topset and foreset and a fine-grained bottomset

(e.g., Vanoni 1975; Kostic and Parker 2003). When a supercritical open-

channel flow is forced to go through a hydraulic jump into a subcritical

regime, such as when a bedload-dominated mountain stream is

obstructed by a weir (Bellal et al. 2003) or an open check-dam (Busnelli

et al. 2001) it may locally emplace a similar prograding deltaic structure

with a distinctive topset and foreset.

* PresentAddress: Department of Civil and Environmental Engineering, Universite

catholique de Louvain, Place du Levant 1, 1348 Louvain-la-Neuve, Belgium

{ Present Address: Hess Corporation, Houston, Texas 77002, U.S.A.

1 Deceased.

¤ Present Address: Shell International Exploration and Production B.V., Kessler

Park 1, Rijswijk 2288 GS, The Netherlands

Copyright E 2009, SEPM (Society for Sedimentary Geology) 1527-1404/09/079-608/$03.00



The present paper and its companion (Sequeiros et al. 2009) documentlaboratory experiments aimed at reproducing a generic configuration forthe emplacement of sedimentary wedges in the submarine environment by

turbidity currents that are sustained but undergoing rapid spatial

deceleration. Turbidity currents are density underflows driven bysuspended sediment. As opposed to rivers, swift deep-sea turbiditycurrents have an intrinsic mechanism for deceleration in that a) they mayincorporate ambient water into the current through mixing at their upperinterface, and b) incipient deposition reduces the downslope pull of gravity acting on the current, in turn favoring further deceleration.

The experiments involved currents that transported and depositedsediment but were driven mainly by dissolved salt. Deceleration of thecurrent resulted naturally from mixing with the ambient water, and wasfurther enhanced by the presence of a barrier partly obstructing the flowfarther downstream. The companion paper focused on documenting theflow patterns associated with the decelerating current and the pondingdue to the downstream barrier. The present paper documents a) thedepositional sequences that led to the emplacement of the sedimentarywedge, and b) the regimes of associated bedforms that were observed.

The generic experiments are not intended to be a precise model of anyspecific field-scale setting. Nevertheless, they provide useful consider-ations that relate to several observed features of wedge-shaped submarine

deposits. Such deposits have been identified in diverse settings, from slopeaprons in regions of lower slope along stepped profiles (e.g., Prather andPirmez 2003; Prather 2003) or in confined submarine minibasins (e.g.,

Winker 1996; Badalini et al 2000; Beaubouef and Friedmann 2000), toconstructional levees of elongated submarine channels (Pirmez 1994;Nakajima and Satoh 2001; Migeon et al. 2004).

Sustained deep-sea turbidity currents carrying a mixture of mud andsand down steep canyons into a zone of lower slope and/or less lateralconfinement undergo at the slope break a sharp deceleration. Thedensimetric Froude number Frd introduced in the companion paper

(Sequeiros et al. 2009), may play an important role in the morphology of

the associated deposits. If the incoming currents are Froude-supercritical(Frd . 1) and the slope break is sufficiently significant, the decelerationmay be further enhanced by the formation of an internal hydraulic jumpthat rapidly transforms the flow into a slower Froude-subcritical regime(Frd , 1) (Garcia and Parker 1989). Whereas a transition from asubcritical flow to a supercritical flow happens gradually by a smoothpassage through a critical section for which Frd 5 1, the passage fromsupercritical to subcritical flow conditions, mediated by an internalhydraulic jump, usually occurs as a sharp transition.

A similar process occurs when swift turbidity currents enter submarineminibasins, i.e., confined topographic depressions of the seafloor. Uponentry in a minibasin, the head of the current may be reflected against its

outer rim and generate an internal hydraulic jump that travels backtowards the entrance of the basin. The deceleration of the flow results inwedge-shaped foreset deposits that thin in the downstream direction

(Toniolo et al. 2006). The ponding of the flow in the minibasin itself maybe so severe that it effectively suppresses the turbulence required tosustain sediment particles in suspension, resulting in the formation of a

settling interface associated with water detrainment (Lamb et al. 2006)and the emplacement of a thin bottomset constructed from the finestsediment sizes (Toniolo et al. 2006a; Toniolo et al. 2006b). Overall,turbidity currents undergoing internal hydraulic jumps might emplacewedge-shaped deposits that have conceptual similarities with subaerialdeltas emplaced by rivers connected to a lake or reservoir.

Constructional levees of elongated submarine channels typically consistof wedge-shaped deposits that thin away from the channel axis. Thelevees are constructed by deposition from thick currents overflowing themain channel. The flows are potentially Froude-subcritical in the channelitself, but the thin overflows become Froude-critical (Frd 5 1) near thelevee apex and supercritical farther down the levee (Pirmez and Imran

2003). They are typically net-depositional, and as such build up andsteepen the levees over time.

Turbidity currents transport sediment of various grain sizes, fromcoarse sand to fine mud. The mud serves to help drive the turbidity

current. The finest fractions of mud, however, do not readily settle out.This fine mud is not explicitly modeled in the experiments reported here,

but instead is replaced with a saline suspension. Thus the ‘‘fine mud’’ istransported under bypass conditions without interacting with the bed.The sand, however, is transported according to the local flow conditionsand may be deposited on the bed. Lightweight plastic particles were usedas an analog for the coarser sediments in the sand range. These plasticparticles were transported for the most part in a region close to the bed,

but they were also seen to be regularly entrained higher up intosuspension in regions of swift flow with high-amplitude bedforms. The

transport of plastic particles dropped to nearly zero as it entered theregion of slow ponded flow caused by the downstream barrier. That is,100% of the plastic particles were deposited before reaching the barrier,gradually forming the sedimentary wedge. By contrast, the trapping of the ‘‘fine mud’’ was perforce zero, inasmuch as the saline flows always

overtopped the barrier at the downstream end of the study reach.Besides evidence for wedge-shaped deposits, the experiments also

revealed distinctive patterns of bedforms. In the literature, fieldobservations of bedforms presumably created by turbidity currents have

been widely reported and discussed. The dramatic increase in resolutionof available methods for imaging the bed topography and subsurfacestratigraphy in the deep-sea environment has led to the observation thatbedforms are commonly present in the path of submarine turbiditycurrents. They are found at a variety of scales and shapes: large gravelwaves have been observed within the Var submarine channel (Piper andSavoye 1993), and more recently the migration of sand waves has beenrecorded along the axis of Monterey Canyon (Smith et al. 2007). Away

from confined channel thalwegs, large-scale upslope-migrating sedimentwaves are widely observed along the outer flanks of channel levees

(Normark et al. 2002) of many deep-sea depositional systems around theglobe, including, e.g., the Amazon (Pirmez 1994; Flood et al. 1995) andthe Zaire (Migeon et al. 2004) submarine fans and the Var sedimentaryridge (Migeon et al. 2006); similar sediment waves were also observed aswidespread fields in more open areas along the continental slope and rise,including the Magdalena turbidite system (Ercilla et al. 2002), although

they might be also attributed to bottom contour currents rather thanclassical turbidity currents (Wynn and Stow 2002). Large cyclic scours off the Shepard bend along the Monterey Channel (Fildani et al. 2006) or in adistributary channel off a bend of the Eel Canyon, California (Lamb et al.2008), have been proposed as resulting from turbidity currents throughprocesses similar to those creating sediment waves along the outer flanksof channel levees. In fact many such bed undulations previously reportedas resulting from submarine slides are being reconsidered as sedimentwaves associated with migrating bedforms created by density currents

(e.g., Lee et al. 2002, Schwehr et al. 2007, Berndt et al. 2006, Urgeles et al.2007).

Relatively few experimental studies have addressed the issue of

bedform regime associated with turbidity currents or have sought torelate the regimes of bedforms to the properties of the prevailing turbiditycurrents (but see, e.g., Kubo and Nakajima 2002). The key mechanismsgoverning their formation and migration are poorly understood, most of the knowledge being inherited from findings pertaining to subaerial flows

in the riverine and coastal environments. However, the internal verticalstructure of turbidity currents differs significantly from both river flowsand coastal wave-induced bidirectional flows, and the transfer of

available knowledge to turbidity currents is not necessarily straightfor-ward.

The present paper provides a preliminary description of the regimes of bedforms associated with supercritical turbidity currents. Three types of

WEDGE-SHAPED SEDIMENTARY DEPOSITS AND BEDFORMS EMPLACED BY TURBIDITY CURRENTS, PART 2 609J S R



bedforms were identified, here referred to as cyclic steps, antidunes, and

ripples. Cyclic steps are bedforms that are locked into an upstream-

migrating cohort by hydraulic jumps, as schematized for subaqueous

flows in Figure 1. They have been characterized in bedrock streams by

Parker and Izumi (2000), and in alluvial streams by Taki and Parker

(2005) and Sun and Parker (2005). Recently Kostic and Parker (2006),

Fildani et al. (2006), and Lamb et al. (2008) have provided strong

evidence that some ‘‘sediment wave’’ fields observed in the submarine

environment (e.g., Lee et al. 2002) may be cyclic steps produced by

turbidity currents.

The antidunes reported here are identified as such in terms of the

relative phasing of the undulation of the upper flow interface relative to

the bed. According to a standard definition (Kennedy 1963), antidunes

are features of Froude-supercritical flow for which the interface (the

water surface in the case of rivers) is in phase with the bed. Antidunes

may migrate either upstream or downstream (Carling and Shvidchenko

2002). The antidunes referred to herein migrate downstream. The term

‘‘ripples’’ used here refers to very small bedforms that developed before

installation of the barrier, and which could not obviously be classified in

terms of flow mechanism.

Summarizing, the objectives of the present paper are twofold: first, the

paper aims to present experimental evidence for the formation of wedge-shaped deposits associated with decelerating turbidity currents, as may

occur, e.g., at slope breaks, in minibasins, or on channel levees; second,

the paper aims at using the experimental data to provide a preliminary

description of the regimes of bedforms observed. As such it provides the

first evidence for the formation of cyclic steps associated with turbidity

currents, as well as evidence for the formation of trains of downstream-

migrating antidunes.

The paper is structured as follows. To begin with, the experimental

conditions of the tests, which were presented extensively in the

companion paper, are restated briefly. In addition to flume characteristics

and flow conditions, the properties of the sediment deposits are discussed.

Then, the depositional sequences are analyzed. The discussion identifies

various stages of bed development, including a profound remolding of the

upstream reach as large-scale bedforms emerge. The progressiveformation of a sedimentary wedge is documented in terms of bed profiles

averaged over those bedforms. Next, a dedicated section focuses on the

characteristics of the bedforms, the patterns of flow associated with them,

and their mechanisms of migration. Finally, a discussion on deep-sea

analogs and limitations follows before conclusions are drawn.

EXPERIMENTAL CONDITIONS

Initial and Boundary Conditions

The experiments were carried out at the University of Illinois Ven Te

Chow Hydrosystems Laboratory. They involve the release of continuousbottom density currents in a channel 15 m long, 45 cm wide, and 1.4 m

high, featuring a sloping sediment bed and a rigid tilted weir (barrier)

located 9 m downdip from the inlet where the currents were introduced

(see fig. 2 of the companion paper by Sequeiros et al. 2009). The weir was

placed at an angle of 45u relative to the bed slope. It obstructed the flowapproaching it, so providing an idealization of the downstream end of a

submarine minibasin. Two sets of experiments were performed. In the

first set, the vertical height of the weir was 32.5 cm, and in the second setit was 41.9 cm. Set 1 consisted of 24 consecutive tests, each with a

duration of about 20 minutes. Set 2 consisted of 33 consecutive tests, eachwith a duration of about 30 minutes.

An initial sediment bed was leveled at a constant slope of 6%. At the

upstream end, saline water and plastic sediment were fed at a constant rate.The density current so created was supercritical in terms of the densimetric

Froude number Frd as it emanated from the inlet, but the presence of the

obstructiondownstream forcedan internaltransitioninto subcritical flowviaan internal hydraulic jump upstream of the point where the current

overflowed the weir. As the flowgraduallyentrained water fromthe ambient

water above and then underwent the internal transition, flow velocities werereduced, causing deposition of sediment both upstream and downstream of

the jump. This deposition progressively built up the bed. Due to the limitedcapacityof theupstream tankwherethe saline suspensionwas prepared, each

testin a setwas sustained fora durationof 20–30 minutes.Multipletestswerethusrepeatedto simulatea muchlonger, sustained flowevent(, 8 hrforSet1 and , 16.5 hr for Set 2). At the beginning of each test the discharge was

increased very gradually from zero to its design value, in order to fill theidealized minibasin ata slowpaceand topreventbedalterations relatedto the

passage of the head of the density current.

The two sets of experiments differed in terms of inflow discharge, salinity,supply of sediment, and height of the downstream obstruction. The relevant

parameters are discussed in more detail in the companion paper, and are

recalled here in Table 1. The tests of Set 1 were found to be stronglydepositional even in the very proximal reach close to the inlet. With this in

mind, the conditions for Set 2 were adjusted to have a much lower rate of sediment supply, a higher inflow salinity to provide additional bed shear

stress, and a higher weir acting as the obstruction. In addition, the inflow

discharge for Set 2 was slightly reduced to accommodate a longer testduration of 30 minutes. The values for the excess density of the inflow given

in Table 1 reflect a minor effect due to a temperature difference between thesaline flow and the ambient clear water of a few degrees Celsius, but its

impact on the magnitude of the downstream pull of gravity was negligiblecompared to that caused by the dissolved salt (Sequeiros et al. 2009).

Sediment Properties

The plastic sediment had a specific density of 1.53 and no cohesion.The mean size was approximately 210 microns. Applying the formula by

Dietrich (1982), the estimated fall velocity in clear water is 9.2 mm/s.Constraining in terms of an equal fall velocity, the plastic sediment is thus

equivalent to a quartz-density particle with a diameter of 113 mm, or a

value of 3.15 in the W sediment size scale, i.e., a fine to very fine sand.

Although the grain-size distribution is rather narrow, measurablefractions of the plastic sediment size distribution consisted of materialfiner than 100 microns and coarser than 250 microns. The grain-size

distribution of the mixture is given as figure 4 in the companion paper.

The sediment was dry-fed through silo openings distributed over thechannel width upstream of the main channel inlet, and were well mixed

within the saline current as it entered the flume.

A reasonably accurate estimate of the porosity of a bed deposit isrequired in order to perform mass balances using sediment infeed rates

and deposit volumes. Such estimates were obtained under well-controlledconditions for dry and submerged deposits using a large cylindrical test

tube. They are reported in Table 2. An upper bound for the porosity

corresponds to the lowest packing (usually referred as the rlp, or random

loose packing) and is obtained for a bed formed by sediment falling outloosely onto a horizontal surface, either in emergent or submerged

FIG. 1.—Sketch of upstream-migrating cyclic steps bounded by hydraulic jumps.

610 B. SPINEWINE ET AL. J S R



conditions. A lower bound for the porosity is obtained by compacting thedeposit through shaking and applying compaction shocks; this is the

highest packing (referred to as rcp or random close packing) attainable inthe absence of permanent external compaction.

Measurements

A range of measuring techniques was used to acquire flow data during

the tests. These techniques were detailed in the companion paper(Sequeiros et al. 2009). They include sets of siphons to estimate currentsalinity at various distances above the bed, an ADV probe to obtain

profiles of streamwise flow velocity along the vertical, and sidewallimaging with still and video cameras. The evolving bed profile along the

channel axis was tracked with an ultrasonic echosounder. This trackingwas done after selected tests of the first set and after each test of thesecond set. In addition, a laser-light-sheet imaging technique was used to

recover the full three-dimensional topography of the bed after the secondset of tests.

DEPOSITIONAL SEQUENCES AND WEDGE FORMATION

Selected image mosaics and measured bed profiles for the two sets of experiments are plotted in Figures 2 and 3. It may be recalled that theexperimental conditions for the two sets were different in terms of inflow

discharge, salinity, and rate of sediment supply. Also, the height of thedownstream obstruction was different, and consequently the position of the internal hydraulic jump within the flume differed. The reader is

referred to the companion paper (Sequeiros et al. 2009) for a detaileddescription of the experimental conditions. Despite the differingconditions for the two sets, the same general tendencies were observed:

a) substantial aggradation along the proximal region and in the region of the internal hydraulic jump upon entry in the ponded minibasin, resultingin a steepening of the bed slope; b) only residual deposition in the distalponded region; and c) development of flow-dependent regimes of

bedforms. In what follows, however, the description of the pattern of morphodynamic bed adaptation within the minibasin is performed solelyon the basis of the second set of tests. This is because for this set bed

FIG. 2.—Selected flow mosaics (for tests 04 and 11, respectively) and bed profiles for tests of Set 1. All dimensions are in meters. The image mosaics are distorted

vertically by a factor of 2, and the bed profiles by a factor of 4. The difference in shading of the flow between the two mosaics simply reflects different feeding of dye tovisualize the flow (see color version online).

TABLE 1.— Experimental conditions for the two sets of tests.

SetNumber of

testsHeight of

obstruction [cm]Inflow

discharge [l/s]Sediment supply

[g/min]Inflow

salinity [g/l]Excess inflow

density Dr/r [-]Duration of individual

tests [min]

1 24 32.5 3.0 640 38 0.027 , 202 33 41.9 2.0 161 49 0.034 , 30




profiles were measured after every single experiment, i.e., every half-hour

of continuous flow, thus allowing a more detailed analysis than was

possible with the data of Set 1.

For Set 2, the early position of the internal transition to subcritical flow

ranged approximately from x 5 2.5 m to x 5 3 m. Starting from the

initial sloping bed, several stages of evolution can be distinguished:

1. Initiation of the sediment wedge by aggradation and steepening of

the proximal supercritical reach. This stage was associated with the

formation of short-wave, downstream-migrating antidunes (e.g.,

Carling and Shvidchenko 2002);

2. Change of regime of bedforms in the supercritical reach, with theemergence of larger-scale cyclic steps, growing in amplitude and

slowly migrating upstream (e.g., Taki and Parker 2005; Fildani et al.

2006);

3. Formation and steepening of a foreset-like structure at the transition

into the ponded region, resulting from the deposition of most of the

remaining sediment at the internal hydraulic jump, continued

deposition causing the foreset of the sediment wedge to steepen

and prograde downstream and the hydraulic jump to migrate along

with it.

These three stages are discussed below in more detail. Each of them is also

associated with a specific regime of bedforms, which is the subject of a

subsequent section.

Step 1: Initiation of the Wedge: Aggradation and Steepening

The evolution of the bed profile over each of the first four hours of flow

is illustrated in Figure 4. The profiles themselves have been averaged over

a one-hour period (by simply averaging the profiles measured over two

consecutive experiments) to focus on the general trend of evolution rather

than on rapid changes that are due to small-scale migrating bedforms.

These bedforms are discussed more specifically in the next section. The

upstream supply of sediment was in excess relative to the transport

capacity, and consequently the bed featured a coherent pattern of

aggradation. The aggradation rate decreased progressively downstream

from circa 1.68 cm/hr at x 5

1 m to less than 0.25 cm/hr at x 5

4 m. Asa result, the longitudinal bed slope increased from circa 6% to 7.5%. In

fact, through deposition the bed slope was aiming towards new

equilibrium conditions prescribed by the upstream feed of sediment. This

initiated a wedge-shaped deposit, concurrently aggrading, steepening, and

extending downstream as larger bed slopes increased the capacity of the

current to transport the sediment farther downstream.

FIG. 3.—Selected flow mosaics (for tests 11 and 32, respectively) and bed profiles for tests of Set 2. All dimensions are in meters. The image mosaics are distortedvertically by a factor of 2, and the bed profiles are distorted by a factor of 4 (see color version online).

TABLE 2.— Measured porosity of sediment deposits.

Dr y deposi t Submerged deposi t

Porosity for rlp 0.572 0.598Porosity for rcp 0.496 0.525




In addition to this global trend, of interest is the peculiar behaviorobserved in the first meter at the upstream end of the channel. Significantdeposition at x 5 0.5 m was immediately followed downstream by azone of reduced deposition around x 5 0.75 m (See Fig. 4). After 8 tests,this local low in the bed profile was followed by a high point at x 5 1 mand a second milder low at x 5 1.25 m. As explained below, thisbehavior may initially be triggered by the changing configuration of theupstream boundary conditions as the bed aggrades. It was also thetriggering mechanism for the emergence of the cyclic steps that arediscussed in more detail below. Their emergence resulted in a completechange of flow regime all over the channel down to the internal transitioninto the ponded region.

Evolving Configuration of the Upstream Boundary

At the upstream end of the flume, an unusual pattern of aggradationwas observed. The saline inflow was supplied through pipe openingsdistributed along the flume width. The resulting jet-like flow passedbeneath a gate, and it was highly Froude-supercritical as it flowed into theflume. As sketched in the left panel of Figure 5, when the sedimentdeposit was still thin, the flow passed smoothly over the sediment bedcreated by previous flow deposits and remained supercritical throughoutthe first half of the channel. However, after several experiments theproximal bed deposits grew in thickness. The jets still prevented

deposition in the immediate vicinity of the gate, but the thick deposit

farther downstream created an adverse bed slope. As a result, the jetsbecame submerged and the flow turned subcritical in that region. Theflow then reaccelerated as it passed over the apex of the proximal depositand then flowed over the steep slope. In terms of boundary condition, thesupercritical flow that subsequently developed in the channel was thencontrolled by a critical section with Frd 5 1 at the apex of the farthestproximal deposit. As compared to the initial situation, this effectivelyreduced the Froude number of the supercritical reach to a value onlyslightly higher than 1, an effect that is believed to have triggered theemergence of large-scale bed undulations referred to as cyclic steps below.However, this forcing of a Froude number of 1 at an overflow point is infact by no means unrealistic, in that it can also be expected to occur in the

submarine environment. Such cases include a) the crest of a ridge where aturbidity current overflows from one minibasin to another, and b) the

point of lateral overflow of a turbidity current from a channel to its levee(Pirmez and Imran 2003; Fildani et al. 2006).

Step 2: Emergence and Migration of Cyclic Steps

As noted earlier, after a stage of consistent channel aggradation andsteepening the morphodynamics underwent a transition in the mostproximal region of the deposit. The transition was signaled by a tendencyfor incision into antecedent bed deposits at a given location, so creatingincreased deposition just downstream (Fig. 4). This process evolved into the

formation of a train of large-scale upstream-migrating cyclic steps, each

FIG. 4.—Initial stages of bed evolution illustrating coherent aggradation and steepening of the upstream reach. The initiation of cyclic steps at the upstream end is alsoapparent. The bed profiles pertain to tests of Set 2.

FIG. 5.— Evolution of the flow conditions at the upstream boundary. Left: initially, the jet-like underflow runs smoothly over the initial deposits and remainssupercritical everywhere. Right: as the bed aggrades, the deposit forces the formation of a jet submerged in a pool of subcritical flow. The flow then undergoes a transitiontowards supercritical as it flows over the crest of the proximal deposits.




bounded by internal hydraulic jumps (Kostic and Parker 2006; Fildani et al.

2006). As the first step matured, several new steps were formed in sequence

downstream, with an average wavelength of about 70 to 80 cm. Figure 6

illustrates their evolution over a number of experiments. The formation of

these steps came as a surprise during the experimental campaign. Given

their unexpectedly large impact on the evolution of the flow and the bed, a

full section is devoted to their characterization below.

Step 3: Foreset Formation and Progradation of the Wedge

At the beginning of Step 3, aggradation had continued all along the

upstream reach of the flume. Notwithstanding the emergence of cyclic steps,

substantial sediment transport was observed all the way down to theinternal hydraulic jump that separated the swift current upstream from the

ponded region downstream. From this moment onwards, the evolution of

the wedge-shaped deposit started to be significantly influenced by the

internal hydraulic jumpand the downstream barrier. At the transition to the

ponded region itself, the depth-averaged flow velocity decreased rather

suddenly. In addition, the configuration of the downstream barrier caused

the principal flowdirection to progressively detach fromthe near-bedregion

and move upward toward the point of overflow. All together, the transition

was associated with a strong downstream decrease in the sediment transport

rate, with only the very fine fractions of the grain sizes being transported

across the transition and deposited in the ponded region itself. Hardly any

sediment at all was seen to pass over the downstream obstruction.

As a result, one can thus expect the formation of a steeper foreset

structure reflecting the rapidly decreasing sediment transport rates at the

transition, and a bottomset formed by the very finest sediment falling out

of suspension in the ponded region. However, the configuration of the

present experiments resulted in a behavior that differed somewhat from

the conceptual framework of an idealized topset–foreset–bottomset

sequence, for the following reasons:

N With salt being used as a substitute for fine mud transported by actual

turbidity currents, our saline flows were not expected to form a strong

bottomset. As opposed to fine sediments, dissolved salt does not

require any turbulence to remain in suspension, and does not deposit

on the bed. Only the finest fractions of our plastic sediment analog

were carried into the ponded zone, so resulting in a very thin

bottomset.N Bottomset deposits associated with the tail of dying turbidity currents

and with hemipelagic deposition in between successive turbidity

current events were not reproduced in the experiments.

N As opposed to the foreset of a subaerial delta, a foreset created by

turbidity currents undergoing an internal hydraulic jump imposed by

a downstream barrier can be expected to have a relatively low slope,

because intrinsic density stratification allows the turbidity current to

plunge more easily underneath the ponded region and maintain higher

shear stresses along the foreset. At the scale of our laboratory

experiments, the length scale associated with flow mixing at the

internal hydraulic jump was of the same order of magnitude as the

expected lengths of the topset and foreset reaches. As a result, the

differences in slopes between the topset, foreset, and bottomset

regions were not as strong as they might be in the field.

FIG. 6.—The profiles illustrate the intermediate stages of bed evolution. A) Formation of upstream-migrating cyclic steps in series along the far-proximal reach. B)

Downstream-migrating antidunes along the transition toward the subcritical ponded region. The plot of the upper panel is undistorted; in the lower panel the plot isdistorted by a factor of 4. Gray dotted arrows indicate the pattern of aggradation-migration. The bed profiles pertain to tests of Set 2.




N Whereas our experiments were purely two-dimensional, actual deep-

sea minibasins have a wider lateral extent. When debouching from a

confined channel into a much wider area of ponded flow created by

the minibasin, the turbidity currents undergo a lateral expansion. This

expansion further enhances the reduction of flow velocities and

encourages the proximal deposition of the coarse fractions of the

transported sediment to form the foreset.

N The emergence of the large cyclic steps previously mentioned resulted

in a more variable topset slope.

N Given that the foreset formation is a slow process presumably

associated with the passage of a large set of turbidity currents of long

duration, the number of experimental runs might not have been

sufficient for it to develop to a mature stage.

Despite these factors, the experimental results show reasonable

evidence for the formation of a foreset-like structure. In Figure 7 thetotal thickness of the bed deposit is plotted versus downstream distance

for selected representative times. If one averages the bed profile over the

cyclic steps, the upstream reach is seen to have quickly adjusted to a

pseudo-equilibrium, relatively constant slope, as indicated by the gray line

segments in Figure 7. The bed then continued to aggrade but kept

roughly the same average slope, with upstream-migrating cyclic steps in

superimposition. Globally, the associated deposits created a coherent

topset, with an average slope equal to 8.5%, as compared to the initial bed

slope of 6%. One can also notice how the cyclic steps were built in series

starting from upstream, with only one step present after 13 tests, two after

23 tests, and three after 33 tests, but with a fourth one likely in formation

at around x 5 3.75 m.

In the vicinity of the internal hydraulic jump, the local deposits induced

by the rapid decrease in sediment transport capacity gradually created asteeper foreset. As seen in Figure 7, the bed region extending from circax 5 3.75 m to 5 m gradually adopted a steeper slope. After 33 tests, the

slope averaged over this reach was equal to 12%, i.e., substantially higher

than the topset slope and twice the initial bed slope, but still much lower

than the angle of repose. A more careful examination indicates that this

foreset region did not adopt a linear profile, but rather showed an upward

concave profile with decreasing slope downstream. From x < 3.75 m to

4.5 m, the local slope is as high as 14%, whereas it is , 9.5% over the

region x < 4.5 m to 5 m. This trend is in agreement with a gradual

decrease in bed shear stresses as the inflow progressively decelerated and

mixed with the standing saline water in the ponded region.

Because the salt used as a substitute for fine particles in suspension

does not interact with the bed even in conditions of highly ponded flow,

the formation of a strong bottomset as part of the topset–foreset–

bottomset sequence was not anticipated. However, even though the grain-size distribution of the plastic sediment mixture was relatively narrow

(with a geometric standard deviation s g 5 1.15), the sediment was not

completely uniform. The preferential transport and deposition of

sediment particles according to their diameter likely caused a slight

pattern of downstream fining. Unfortunately, no samples of bed

sediments could be retrieved at the end of the tests, and hence

downstream fining could not be quantitatively assessed. It was clearly

observed, however, that the smallest fractions were transported in

sustained suspension along the swift supercritical reach and all the way

through the internal transition. They were then slowly deposited in the

ponded region. Though representing only a few percent of the total

deposit, the fine material was emplaced as a thin bottomset. As seen in

Figure 7, the bottomset had a remarkably uniform thickness: after a

transitional region from x < 5 m to 6 m where the bed slope was around

7.5%, the final bed slope remained equal to its initial value of 6%,

indicating that the deposits were created mostly by very fine sediments

raining uniformly out of suspension.

Once the foreset had reached an equilibrium slope profile, it could be

expected to prograde in the downstream direction as the topset continued

to aggrade. The formation of the foreset was, however, a slow process,

and the number of tests actually performed may not have been enough to

clearly document progradation. Nevertheless, it is clear from the analysis

of Figure 7 that the upstream reach over which significant sediment

deposition occurred was extending in the downstream direction as time

passed. After 13 tests, most of the sediment had been deposited in the first

3 to 4 m of the channel. After 33 tests, the main deposit extended

downdip to more than 5 m.

Another way to look at progradation is by estimating the spatial and

temporal variation of the associated sediment transport rate. Since nodirect measurement of sediment transport was done, an indirect measure

was obtained by using information from the measured bed profiles. The

calculation was aided by the fact that no sediment was observed to escape

the minibasin. Thus, by looking at differences between two distinct bed

profiles, one can obtain a time-averaged longitudinal profile of the bed

aggradation rate, gb/t, where gb denotes bed elevation and t denotes time.

This can be translated into a profile for the rate of change of the sediment

transport in the longitudinal direction, Qs/x, where Qs denotes the volume

sediment transport rate (in m3/s) and x denotes streamwise (downdip)

distance. The sediment transport rate is computed from the Exner

relation for the conservation of sediment:

LQs

Lx

~ {B 1 { lð ÞLgb

Lt

ð1Þ

FIG. 7.—Long profiles of the bed after selected tests of Set 2. In the plot, Dgb 5 gb 2 gb0, with gb0 referring to the initial bed profile. One can distinguish the gradualformation of a topset–foreset–bottomset succession, sketched as thick light-gray segments.




where B 5 0.45 m is the channel width and l denotes the bulk porosity

of the bed deposits. The actual porosity of the sediment deposits

produced by the experiments can be expected to be close to the random-

loose-packing value for submerged deposits, i.e., 0.598.

This porosity can be checked on the basis of the initial and final profile

(after 33 tests) of Set 2. Measured from the area between those two

profiles, the total bulk volume of sediment deposited in the channel bedwas computed to be 0.247 m3. Since there was virtually no outflow of

sediment over the downstream obstruction, the deposit must precisely

balance the total supply of sediment over the cumulated flow duration.

Knowing the constant sediment feed rate (160 g/min), the sediment

density (1530 kg/m3), and the flow duration (33 hr 3 0.5 5 16.5 hr),

one obtains an actual average porosity of 0.5725. This latter value is

indeed close to the estimated submerged random loose packing value of

0.598. The slightly denser packing of the deposits could be attributed to a

somewhat more effective rearrangement of the grains in the active layer

during transport and deposition.

Having a value for the porosity, one obtains a profile for Qs/x through

Equation 1 and two profiles at different times. The term Qs/x can then be

integrated to obtain the long profile of the actual volumetric sediment

transport rate Qs (averaged over the time interval in question). The values

for Qs are then converted to mass rates G s in g/min by using the sediment

density, rs 5 1530 kg/m3. A boundary condition for the integration is

obtained by specifying Qs 5 0 at the barrier. The result for G s so

obtained at the upstream end of the channel, when compared to the

known value of 160 g/min, gives an evaluation of the accuracy of this

indirect method.

Two such sets of profiles are presented in Figure 8 for two stages of evolution. The first one is based on the bed aggradation between tests 1

and 5. The profiles for bed elevation (Fig. 8A) reflect the results at the

end of tests at hours 0.5 and 2.5. The profile for sediment transport rate

(Fig. 8B) was computed on the basis of the deposition between hours 0.5

and 2.5. The sediment transport rate vanished at a distance of about 3.5

to 4 meters down the channel. The profile is smooth and upward-

concave, with a stronger reduction in transport rates, and thus larger

aggradation rates, in the proximal region, indicating that the reach was

evolving toward a new pseudo-equilibrium configuration with a steeper

slope. This steeper slope eventually evolved to the topset slope. Values for

the first 0.5 meters of the channel are not available, because no bed-level

measurements were made in that region. But if one extrapolates the

concave-upward profile to x 5 0, one obtains an estimated value for G swhich is in the range of 160 to 170 g/min. This is very close to the actual

FIG. 8.—Bed profiles and estimated profiles of sediment transport rates along the channel, at two stages of bed evolution. A, B) Panels illustrating an early stage of development. The two bed profiles in A pertain to hour 0.5 and hour 2.5, and the corresponding rate of transport shown in B exhibits a smoothly decaying profile. C, D)Panels illustrating a later stage of development after maturation of the cyclic steps. Panel C shows two stepped bed profiles (one averaged for hours 10 to 11, and thesecond averaged for hours 15.5 to 16.5). Panel D shows the corresponding estimate for the rate of transport (averaged over the period from 10.5 hours to 16 hours).During this later time period, the sediment transport rate is significant much farther downstream as a result of progradation. Distances are in meters, and transport ratesG s in g/min.




rate of sediment supply ( 5 160 g/min), so giving confidence in the

calculated profile.

Panels C and D in Figure 8 show two bed-elevation profiles and one

profile for the sediment transport rate at a later stage of evolution. The

first elevation profile is an average based on the results for hours 10 to 11,and the other is an averaged based on results for hours 15.5 to 16.5. This

averaging acted to reduce small-scale bed variations due to rapidly

migrating ripples superimposed on the global trend of bed variation. The

profile of sediment transport rate (Fig. 8D) is thus based on the

deposition rate between 10.5 hours and 16 hours.

In Figure 8, very different behaviors are observed in the later profiles

(Fig. 8C, D) for bed elevation and transport rate, as compared to the

earlier profiles (Fig. 8A, B). At the later stage, the cyclic steps over the

supercritical reach have attained a mature state. Their pattern of

aggradation and upstream migration is very clear, with intenseaggradation on the upstream (stoss) side and very little aggradation on

the downstream (lee) side. As a result, the sediment transport rate

adopted a wavy profile, with a succession of substantial decreases over

the aggrading stoss sides of the steps, and ‘‘plateaus’’ over the lee sides.

However, when averaged over the cyclic steps from x 5 0.5 m to 3.5 m,the profile was essentially linear, indicating once again that the topset, in

a pseudo-equilibrium configuration, was aggrading uniformly over its

length. By contrast, along the foreset (x 5 3.5 m to 5 m), the profile is

concave-upward, indicating that the foreset was still steepening. Finally,

the integrated value of sediment transport rate at the upstream end again

compared favorably with the actual supply of 160 g/min. When the

elevation profiles of Figure 8C are compared to the profiles of Figure 8A,

the downstream progradation of the deposit is clear. Also, whereas thesediment transport rate initially vanished at around x < 3.5 m for the

earlier time (Fig. 8B), at the later time the sediment transport rate

remained substantial until x < 4.5 m to 5 m (Fig. 8D). This documents a

downdip extension of the topset–foreset sequence.

REGIMES OF BEDFORMS

The objective of this section is to take a closer look at the regimes of

bedforms that were observed during the experimental tests, by analyzing

their main features, their pattern of migration, and the characteristics of the flows that sculpted them. There have been a number of experimental

studies of turbidity currents and/or saline bottom flows to date that

discuss bedforms (e.g., Garcia and Parker 1991; Garcia and Parker 1993;

Kubo and Nakajima 2002; Fedele and Garcia 2001). To our knowledge,

however, the present study is the first one to a) identify upstream-

migrating sediment waves as cyclic steps, and b) identify a subset of

downstream-migrating bedforms as antidunes. In addition to longitudinal

bed profiles along the channel centerline and video sequences through the

sidewalls, the data available for the present sets of experiments also

include a detailed 3D mapping of the final bed topography at the end of

Set 2, obtained by means of digital imagery and a laser light sheet.

Downstream-Migrating Antidunes

In open-channel hydraulics, subcritical flows are by far more widely

encountered than their supercritical counterparts. The vast majority of

river flows are subcritical. Subaerial environments where supercritical

flows are encountered more often include steep mountains torrents over a

gravel bed or bedrock, and the coastal swash zone. The most common

bedforms associated with such supercritical currents are antidunes, which

are defined in terms of a water–surface profile (i.e., water–air interface)

that is in phase with the bed profile, in contrast with the out-of-phase

pattern that defines dunes in the subcritical regime (e.g., Kennedy 1963).

Antidunes generally migrate upstream. Downstream-migrating anti-

dunes have, however, been predicted theoretically (Kennedy 1963;

Engelund 1970) and observed (see Carling and Shvidchenko 2002 for asummary). Whereas upstream-migrating antidunes tend to be symmetri-cal, downstream-migrating antidunes tend to be asymmetrical, with asteep lee face similar to dunes. In the simplified framework of Kennedy

(1963) the direction of migration may be related to the evolution of flowthickness over the bedforms. If the flow thickness at the crest is higher

than at the troughs, the flow is decelerating over the rising edge of thebedform and accelerating over its trailing edge, thus causing depositionon the rising edge and erosion on the trailing edge. As a result thebedforms migrate upstream. Downstream-migrating antidunes arepossible whenever the flow thickness at the crest of the bedforms islower than at the troughs. Spatial lags between flow acceleration and theresponse in terms of sediment transport, especially for suspended load,may alter the universality of that principle. This notwithstanding, the

analyses of Engelund (1970) and Fredsøe (1974) predict both down-stream- and upstream-migrating antidunes for the case of dominantsuspension.

As mentioned above, downstream-migrating antidunes were producedin the present experiments. The flow was supercritical, and the dominant

mode of transport was bedload rather than suspension. Figure 9 showsclear evidence for their appearance. The image pertains to Test 6 of thefirst set of tests. The interface between the saline current, dyed in red, andthe ambient water above is easily ascertained, though the dye is lessclearly visualized in the upper part of the current which has mixed withthe ambient clear water. The fact that the profile of the interface is in

phase with that of the bed is evident. The antidunes observed in theexperiments had a coherent wavelength of about 18 to 20 cm, and awaveheight decreasing in the downstream direction from about 3 cm to2 cm. Their aspect ratio (wavelength divided by waveheight) is thus in therange from 6 to 9. The mean flow thickness is slightly larger than 5 cm.Estimating the discharge at x 5 2 m to be around 4 l/s (i.e. the suppliedvalue at the upstream end, 3 l/s, plus some entrainment), the depth-averaged flow velocity is of the order of 0.15 m/s, and the excess density

due to the salt is around 0.02 (the upstream value of 0.027 is discussed inthe companion paper, Sequeiros et al. 2009). The estimated densimetricFroude number is thus around 1.4, defining the flow as well into thesupercritical range.

During the initial stage of bed evolution, the inflow remained highlysupercritical over a substantial portion of the channel before the internalhydraulic jump defining the entrance into the ponded region. Antiduneswere observed throughout that reach (i.e., more or less the first 4 m of thechannel).

At a later stage, the aggrading bed caused a submersion of the inflow(Fig. 5), resulting in upstream Froude numbers dropping to values only

slightly larger than unity. This change was accompanied by a change inregime from antidunes to cyclic steps, which are discussed in more detailbelow. However, in that later stage, bedforms similar to the initialantidunes were still found to occur downstream of the cyclic steps, in the

foreset region just upstream and within the internal hydraulic jumpcaused by the downstream barrier. The bed slope was highest in thevicinity of the upstream region of this transition, as described in theprevious section. The incoming current remained attached to the bedthere, and plunged under the ponded saline region. Within this localized

region of plunging current over a steep slope, the apparent Froudenumber of the flow (i.e., the one based only on that plunging part of theflow, without considering the top cap of standing saline water) mightagain have reached values significantly higher than 1. This might explainwhy strong downstream-migrating antidunes were found in that regionalso.

Whereas Figure 9 shows only a view from the flume sidewall, the actualpattern of the antidunes along the channel width is depicted in Figure 10.

This figure is an oblique photograph taken after the same Test 6 of Set 1as Figure 9 but taken slightly downstream. It shows an irregular 3D




arrangement of the bedforms, with an alternation of local scours andhumps, suggesting that the flow pattern over the bedforms is itself strongly three-dimensional.

Returning to Figure 9, it is reasonably clear from the image that theflow thickness over the stoss side and crest of the antidunes is lower than

over the lee side and trough, so indicating downstream migrationassociated with erosion on the stoss side and deposition on the lee side.Indeed, this pattern of downstream migration is verified in Figures 11 and12. The images and bed profiles pertain to Test 10 of Set 1 aroundx 5 3.75 m, i.e. over the evolving foreset. The bed profiles indicateconcurrent aggradation as the bedforms progress downstream, with apartial preservation of the lee sides of the bedforms in the depositionalrecord. From the profiles, the migration rate was found to be about20 cm per hour and the aggradation rate was found to be 2 cm per hour.These bedforms can thus be called ‘‘climbing antidunes,’’ in analogy tothe climbing ripples described by, e.g., Allen (1973), Ashley et al. (1982),Jopling et al. (1968), or Jerolmack and Mohrig (2005).

The downstream-migrating antidunes associated with dominant bed-load transport may have a field analog in terms of the gravel wavesobserved in the Var Canyon (Piper and Savoye 1993). These bedformsshow the same asymmetry as those observed here, and it is most likelythat the gravel in question were moved as bedload. In addition, therelatively steep slopes on which these bedforms were observed (, 7% to3%) suggest supercritical rather than subcritical flow.

Upstream-Migrating Cyclic Steps

As noted previously, the downstream-migrating antidunes firstobserved in the upstream supercritical reach were later replaced by much

larger-scale, upstream-migrating cyclic bed undulations. These undula-tions were much too long and high-amplitude in relation to the typicalcurrent thickness to be categorized as either antidunes or dunes. Instead,

they appear to be in the same class as similar bedforms observed in open-

channel flows, which Parker (1996) has called cyclic steps.

In common with antidunes, cyclic steps require supercritical flow

conditions to form. Their defining aspect, though, is that once they are

formed, the flow over them is transcritical, with an alternation of Froude-

subcritical and Froude-supercritical regions bounded by Froude-criticalsections appearing close to the crests. The flow downstream of the crest is

supercritical over a short distance along the steep slope preceding the

trough and then undergoes a transition through a hydraulic jump

immediately downstream. The flow then remains subcritical until the crest

of the next step. In the subaerial setting, bedforms that can be identified

as cyclic steps have been found to occur in alluvium (Winterwerp et al.

1992; Taki and Parker 2005; Sun and Parker 2005), in cohesive sediments

(Reid 1989’ Parker and Izumi 2000) and in bedrock (Koyama and Ikeda

1998; Wohl 2000). Cyclic steps differ from antidunes in terms of the

following three aspects.

N Antidunes are short-wave phenomena, in that the ratio of wavelength

to flow thickness is of the order of 5 or smaller. Cyclic steps are long-

wave phenomena, in that the same ratio is of the order of 10 or larger.

N Antidunes tend to be ephemeral features that form, grow inamplitude, break, and then repeat this process in a different place.

Cyclic steps are well-organized, quasi-permanent features that march

upstream in an orderly train.

N Cyclic steps are continuously bounded by hydraulic jumps that

maintain their form, whereas antidunes do not have sustained

hydraulic jumps.

The most common bedform observed on the seafloor is known as the

‘‘sediment wave’’ (e.g., Lee et al. 2002; Peakall et al. 2000). They are

commonly observed at slope breaks (Prather 2003; Prather and Pirmez

2003) and also on the outer slope of levees of submarine channels (e.g.,

FIG. 9.— Downstream-migrating antidunescreated by a supercritical saline underflow,highlighting the in-phase pattern of the bed andflow profiles. The saline flow has been coloredwith red dye. The image pertains to Test 6 of Set1, and the ruler indicates the location of theimage at approximately 2 m down the channel(see color version online).

FIG. 10.—Oblique view of the antidunesacross the channel width after Test 6, Set 1,around x 5 3 m, highlighting the irregular 3Darrangement of the bedforms (see colorversion online).




Normark et al. 2002; Nakajima et al. 1998; Lewis and Pantin 2002).Seismic images indicate that these sediment waves migrate upslope enechelon. Recently Kostic and Parker (2006) , Fildani et al. (2006), andLamb et al. (2008) have provided strong evidence that in many cases such

sediment waves represent another manifestation of cyclic steps.

As noted above, the experiments reported here showed long-wave,upstream-migrating bedforms that can unambiguously be identified ascyclic steps. This is because both the orderly upstream migration and thesustained hydraulic jumps bounding the steps could be observed visually.

The cyclic steps we observed exhibit striking similarities with submarine

FIG. 11.—Closer views of the downstream-migrating antidunes (Set 1, Test 10). The imagesare tilted parallel to the underlying rigid bed, andthe ruler indicates the downstream coordinate inmeters. The top and bottom views are over atime interval of 700 seconds. The vertical grid-lines are placed every 5 cm and the horizontalgridlines every 2 cm. Black lines refer to ante-cedent bed profiles tracked on the sidewall with apermanent marker.

FIG. 12.—A sequence of profiles illustrating the evolution of the bed in the presence of downstream-migrating antidunes. The profiles were obtained from the images of Figure 11, plus some intermediate states obtained from video records. The profiles have been transferred into the global system of coordinates attached to the channel. Aprofile for the initial bed obtained from echo-sounder measurements along the channel centerline has been added for comparison. The vertical scale has been distorted bya factor of 2 for clarity.




sediment waves discussed above. They strengthen the case that such

sediment waves are indeed cyclic steps rather than antidunes. Inspired by

the model originally proposed by Parker et al. (1986), Kostic and Parker

(2006) presented a numerical model of turbidity currents that produces

cyclic steps, and Fildani et al. (2006) applied this model to explain, amongother things, sediment waves on the outside levee of the Shepard bend of

the Monterey Channel off California, USA. The present study, however,

is to our knowledge the first to reproduce cyclic steps created by density

flows in the laboratory and to document their formation and migration as

well as the properties of the associated flow. Indeed, the flow

configuration used in the experiments, designed to emplace wedge-shaped

deposits, turns out to provide at the same time an ideal setting to study

submarine cyclic steps. More specifically, the upstream half of the flume

provides a reasonable model describing flow onto a levee of a submarine

channel, a location where sediment waves are commonly observed in the

field (Wynn and Stow 2002, Migeon et al. 2000, Migeon et al. 2001). The

evolution in time of the upstream boundary discussed previously (Fig. 5),

with the submergence of the inflow jets and the forcing of a critical

section with Frd 5 1 at the apex of the proximal deposit, results indeed in

a configuration very similar to a turbidity current overflowing its levee. In

the context of a cascade of submarine diapiric minibasins, the presence of

a critical section is also a necessary transition as a ponded subcritical

turbidity current exits a minibasin, spilling over a steeper slope and

becoming supercritical before eventually heading towards the next

minibasin.

The first step observed in our experiments was formed right at the

upstream end of the channel (Fig. 4), and successive steps then

appeared in sequence, with the progressive evolution of one step

seeming to trigger the formation of an additional step downstream. The

upstream faces of the steps were shorter than the downstream faces, but

the slopes on both faces were of a similar order of magnitude. The steps

migrated upstream in coherent trains (Fig. 6). Once initiated, a cyclic

step grew in amplitude rather rapidly, i.e., in less than one hour of

continuous flow. By contrast, its subsequent upstream migration,

associated with continued aggradation, was a much slower process that

required many hours of flow.

Some quantitative geometrical characteristics of the steps can be

inferred from the bed profiles. Their wavelength was in the range from 70

to 80 cm. Their waveheight after reaching some degree of maturity was inthe range from 6.5 to 8 cm; hence their aspect ratio is roughly equal to 10.

Sediment waves encountered in the field tend to be substantially more

elongated, but values for the aspect ratio extend over a wide range from

25 on the Reynidsjup Channel to more than 200 for some of the waves on

the Var sedimentary ridge (Normark et al. 2002; Migeon et al. 2000;

Migeon et al. 2001).

The bed slopes of our cyclic steps were positive on the downstream side

and negative on the upstream side, with magnitudes as high as 0.2. The

rate of upstream migration, based on the position of the crests, was about

4 cm per hour. However, the rate of migration was not constant in time.

The steps were seen to migrate faster in their initial stage and slower once

they reached some degree of maturity. The upstream condition of an

overflow passing through a critical Froude number, as illustrated in

Figure 5, may play a role in preventing the steps from migrating farther

upstream.

The bed aggradation at selected locations as a function of time is

plotted in Figure 13. As expected, the rates of bed aggradation are lower

at more downstream locations. They range from as high as 16 mm/hr for

the first 5 hr of flow at x 5 1 m to as low as 1.5 mm/hr at x 5 5 m.

Figure 13 also reveals, however, sudden changes in aggradation rates in

the upstream profiles. The visible events of reincision of previous bed

deposits are associated with the formation of the cyclic steps.

The transcritical nature of the flow over the steps is illustrated in

Figure 14. The first panel of the figure features a photo mosaic of the flow

during Test 32 of Set 2. An approximate estimate for the profile of

densimetric Froude number (Fig. 14E) was obtained according to the

method outlined below.

As a first step, estimates of flow discharge at selected locations were

obtained by integrating all the available velocity profiles measured during

FIG. 13.—Time histories of bed aggradation at selected locations as functions of time. Bed aggradation is measured in terms of the deposit thickness relative to theinitial bed.




the last few tests of Set 2, when the cyclic steps were mature (tests 18–33).These estimates are plotted as black squares in Figure 14. In point of fact,instead of integrating the velocity profile within the density current itself,

more accurate discharge estimates could be obtained by integrating

within the region of the backflow in the ambient clear water and applyingconservation of water mass. Indeed, the flow in that region was much lessturbulent and more coherent and uniform along the channel width,because it was not locally affected by lateral fluctuations induced by thebedforms. Since the backflow recirculates—it represents the water that isactually entrained back into the main current through mixing in theupstream reaches—values obtained for the backflow were then added to

the known constant supply of saline water from upstream (see thecompanion paper Sequeiros et al. 2009) to obtain the total discharge of the saline current. From its supply value of 2 l/s upstream, the dischargeis seen in Figure 14B to increase gradually to a value slightly larger than4 l/s over the first 3.5 m and then to remain constant in the pondedregion, where the flow is too slow to entrain clear water from above. Asmoothed continuous profile of flow discharge Q was then fitted to theavailable data. It is represented as a black line on Figure 14B.

As a second step, depth-integrated values for the salinity were obtainedfrom discharge estimates, the inflow rate of dissolved salt, and thecondition of conservation of salt discharge throughout the channel. The

values for the salinity C (in units of g/l) so obtained and the smoothedprofile are shown in Figure 14C. The values have been normalized usingthe inflow value C 0.

Thirdly, values for the thickness of the current were obtained bytracking the bed and current profiles on the sidewall photo mosaic of Test32 (Fig. 14A). This introduces some bias in the method, because thedischarge, salinity, and current-thickness estimates required for thecomputation of the densimetric Froude number do not pertain preciselyto the same test, so that the cyclic steps have migrated over the period.One thus must make the approximate but reasonable assumption that thedischarge and salinity estimates obtained over different tests remain

reasonable estimates for Test 32. The evolution of current thicknessdepicted in Figure 14D illustrates the considerable variations between thetroughs and crests of the steps.

These three steps were used to generate the data for Figure 14E. Thefigure indeed shows the expected trends: subcritical conditions, withdensimetric Froude numbers down to about 0.5, are obtained in thetroughs of the steps; supercritical conditions with values up to 1.5 (above2 for the second step) are found on the lee sides of the steps, just after thecrest where a critical section is expected. Farther downstream after thesteps, the flow enters the truly ponded region and Froude numbers dropto very small values. Figure 14E is very similar to Figure 12B of Fildani

et al. (2006), which was generated from a numerical simulation of submarine cyclic steps.

This transcritical behavior of the flow over the steps directly relates to

the mechanism for their upstream migration. Figure 15 illustrates

schematically the observed flow pattern. The figure features an actualbed profile, measured after Test 28 of Set 2; the dashed line is a sketch of the associated flow profile, and the arrows qualitatively refer to the maintrends for the flow. Over the downstream edge of a step, the supercriticalflow is thin, rapid, and nearly parallel to the bed. Over the trough and the

upstream face of the next step, the subcritical flow is much thicker, withan interface with the ambient water that is almost horizontal. Theobserved cyclic steps consistently had a relatively flat lee face and a moreirregular stoss face. The bed almost invariably displayed a sharp changein curvature just upstream of the deepest point of each step. That may actas a ‘‘ski jump’’ where the flow undergoes a detachment from the bed.The region of highest flow velocities, indicated qualitatively by solidarrows in Figure 15, was indeed observed to deviate from the bed over the

stoss face of the steps, reattaching to the bed only near its crest. Theregions above and below this main streamline were associated with eddies

and recirculation zones, indicated by the dotted arrows in Figure 15.

Those regions did not contribute significantly to the discharge. Sediments

were effectively trapped and deposited irregularly along the stoss of the

steps, causing their upstream migration.

These various aspects can be seen clearly in Figure 16. The photographpertains to Test 30 of Set 2, at a time when the steps have reached a

mature stage. The photograph illustrates particularly well the flow

detachment associated with the ‘‘ski jump’’ at the trough, and the

associated bursting of sediment into suspension, as well as the

recirculation zone, where sediment was being trapped and deposited.

This local feature of flow detachment and sediment trapping is believed to

be one of the driving mechanisms for the upstream migration of the steps

in our laboratory experiments.

Downstream-Migrating Ripples in Subcritical Flow

Downstream-migrating bedforms were observed farther downstream of

the sediment waves and antidunes, i.e., in the truly ponded region of

subcritical flow extending from x < 5 to 9 m. Their amplitudes,

wavelengths, and rates of migration were much lower than for theantidunes and cyclic steps discussed earlier. For want of a better word,

they are here classified generically as ‘‘ripples.’’ They were formed at a

very early stage in the execution of each set, when an initial sediment bed

was emplaced before the installation of the barrier. They were thus

initially formed by a continuous supercritical current that extended all

along the channel, and they may represent small precursors of

downstream-migrating antidunes. The deposit in this region that was

recorded after the obstruction was in place represented mainly draping of

fine sediments raining out of suspension onto the ripples. Hence, their

geometrical properties may not exactly relate to the flow conditions in the

ponded region, and no firm conclusion could be drawn concerning their

exact nature.

3D Topography of the BedformsWhile the above discussion was based solely on 1D bed profiles taken

with an echosounder along the channel centerline, a laser-scanning

imaging method was also used to characterize the three-dimensional

topography of the bedforms at the end of the second set of tests.

The method is similar to the one developed by Spinewine et al. (2004) in

a different context, and relies on a purely optical method to track the bed

topography. A red diode laser (LasirisH 40 mW 660 nm) mounted above

the channel produced a light sheet that illuminated a thin (, 1 mm) cross

section of the bed. The trace of the cross section was captured by a digital

camera (NikonH D200) placed farther downstream at an angle of

approximately 40 degrees relative to the light sheet. The laser and camera

were mounted as a rigid assembly on a carriage which was translated

along guide rails over the length of the flume. The channel did not need to

be drained, because an underwater visual access of fixed geometry for thelaser and the camera was provided through a transparent acrylic plastic

viewing box inserted just below the water surface. The position of the

laser trace on the image is projected into a 3D coordinate system attached

to the channel, with the help of a calibration procedure (Spinewine et al.

2003) that allows recovery of the internal geometry of the camera and

corrects for the refraction across the air–acrylic plastic–water interface.

Overall, the resolution of the method is estimated to be less than 1 mm in

the vertical and transverse directions. Images taken by translating the

carriage at 1 cm intervals were processed and assembled to form a full

digital terrain elevation model of the bed.

Figures 17, 18, and 19 present the 3D topography of the bed from

upstream to downstream, along the three characteristic regions of

bedforms identified above. For each image, the vertical scale of the plots

has been stretched by a factor of 2 for better visualization. The color







scheme illustrates deviations from the mean bed profile averaged over the

bedforms, so that crests appear reddish and troughs appear bluish.

Figure 17 focuses on the cyclic steps. They are seen to be mainly two-

dimensional, with crests and troughs aligned essentially perpendicular to

the channel axis. The trailing edges of the steps, immediately following

the crests, are very smooth. The troughs and rising edges feature

disorganized 3D undulations of smaller magnitude. At the end of the

trailing edge, but just before the troughs, one can consistently identify the

sharp ‘‘ski-jump’’ feature that was discussed previously.

Figure 18 presents the 3D topography of the bed in the foreset region,

where downstream-migrating antidunes with a strongly 3D pattern were

observed. The figure highlights the irregular arrangement of the crests

and troughs in this steep foreset region. By contrast, the lower-amplitude

bedforms (ripples) observed farther downstream were associated with

milder slopes. As depicted in Figure 19, they show predominantly a 2D

pattern with crestlines more or less aligned at regular intervals across the

whole channel width.

DISCUSSION

The experiments reported here were aimed at reproducing a generic

configuration for the emplacement of wedge-shaped deposits by deceler-

ating turbidity currents, and were as such not intended to reproduce as a

whole any specific field-scale setting. As stated previously, however, the

structure of the wedge itself and the observed bedforms have features that

relate them to actual deep-sea deposits. Three tentative field analogs are

identified. They relate to a) deposits in diapiric submarine minibasins, b)

sediment waves on levees emplaced by laterally overflowing turbidity

currents in elongated submarine channels, and c) fields of downstream-

migrating antidunes along the thalwegs of steep submarine canyons.

FIG. 16.—Flow over a cyclic step. The imageshows clearly a smooth lee side, the ‘‘ski jump’’and the flow detachment associated with it, and

the adverse slope between the ‘‘ski jump’’ and thenext crest. The black lines document antecedentbeds before cyclic steps emerged. The photo isfrom Test 30 of Set 2 (see color version online).

FIG. 15.—Sketch of the flow pattern over thecyclic steps, and its impact on the mechanism of upstream migration. The solid line represents anactual measured bed profile (after Test 28 of Set2). The dashed line is an approximate sketch of the current profile. The solid arrows indicate thepath of the main flow, and the dashed arrows

indicate zones of flow recirculation that do notcontribute significantly to the flow discharge.

r

FIG. 14.—Estimated profiles of densimetric Froude number based on measured velocity profiles and sidewall imaging. From top to bottom: A) Photomosaic for Test32 of Set 2, with the saline current dyed in red/green, from which a depth profile was obtained. (the image aspect ratio is distorted by a factor of 2). B) Measured totaldischarge (c) and interpolated profile. C) Measured salinity profile (c) and interpolated profile. D) current thickness as tracked on the photomosaic in part A. E) Derivedprofile of the densimetric Froude number. The gray dashed lines indicate locations where available vertical profiles (tests 18 onwards) were used to estimate discharge andsalinity (see color version online).




FIG. 18.—3D view of the final bed from x 5 3.5 m to 5 m, illustrating the topography over the antidunes at the end of the experiments of Set 2. The shading illustratesdeviations from the mean bed profile averaged over the bedforms, with crests appearing reddish (bright in print version) and troughs bluish (dark in print version). Alldimensions are in meters. The vertical scale has been exaggerated by a factor of two. Note the irregular pattern of the crests, indicating 3D bedforms (see color version online).

FIG. 17.—3D view of the final bed topography from x 5 0.5 m to 3.5 m illustrating the topography over the cyclic steps at the end of the experiments of Set 2. Thevariation in shading documents elevation deviation from the mean bed profile averaged over the steps, with the crests appearing reddish (bright in print version) and thetroughs appearing bluish (dark in print version). All dimensions are in meters. The vertical scale is exaggerated by a factor of two for visualization purposes (see colorversion online).




It is speculated that the presence of a barrier obstructing the flow at the

downstream end of the study reach provides an analogue to the rim of aconfined submarine minibasin, in that the structure of the deposits is

strongly influenced by the sharp deceleration upon entry in the minibasin,

possiblyenhanced by the presenceof an internalhydraulic jump. As a result,

sustained deep-sea turbidity currents carrying a mixture of mud and sand

into a confined basin might be able to emplace a wedge-shaped deposit that

has a conceptual similarity to a subaerial delta emplaced by a river

connected to a ponded reservoir or the sea. The ponding would be due to

entrapment of the current within the minibasin, creating a calm body of

water from which the finer fraction of the mud would slowly settle out to

form a bottomset. The sand, and the coarser fraction of the mud, would be

deposited more proximally to form a steeper foreset, whose position would

be tightly linked to the position of the internal hydraulic jump.

The concept of wedge-shaped ‘‘deltaic’’ deposits in submarine

minibasins is not without some field support. Beaubouef et al. (2003)

have tentatively identified such a deposit in the terminal portion (Basin 4)

of a chain of four Pleistocene intraslope minibasins often referred to asthe Brazos–Trinity Intra-Slope System, on the north slope of the Gulf of

Mexico continental margin (Fig. 20). The process by which successive

minibasins are filled in sequence by sediments deposited from turbidity

currents is referred to as ‘‘fill-and-spill’’ (Winker 1996; Badalini et al.

2000; Lamb et al. 2004). Beaubouef et al. (2003) tentatively compare the

internal stratigraphic structure of ‘‘Basin 4’’ with a stratigraphic model

derived for parasequences of river deltas. Although conceptually similar,

the deposits produced in the present experiments are missing two

important elements: a lateral expansion and a well-developed muddy

bottomset. Our prismatic flume does not account for the two-dimensional

spreading of the current upon entry in a minibasin. The lateral expansion

could enhance the effect of the flow deceleration caused by the external

rim of the basin. On the other hand, a well-developed muddy bottomset

could probably be produced in the laboratory by a) using a deeper, longer

FIG. 19.—3D view of the final bed from x 5 5.5 m to 7 m, illustrating the topography over the ripples at the end of the experiments of Set 2. The shading illustratesdeviations from the mean bed profile averaged over the bedforms, with crests appearing reddish (bright in print version) and troughs bluish (dark in print version). Alldimensions are in meters. The vertical scale has been exaggerated by a factor of two. Note the regular pattern of the crest lines, indicating approximately 2D bedforms (seecolor version online).

FIG. 20.—Tentative field analog for a wedge-shaped deposit emplaced in the terminal mini-basin portion (Basin 4) of a chain of fourPleistocene intra-slope minibasins often referredto as the Brazos–Trinity Intra-Slope System.After Beaubouef and Friedmann (2000). Theturbidity currents enter the system at the upper-left corner of the image. The dark red channelsindicate the main pathways for turbidity currentsoverspilling basins 2 and 3 before spreading andemplacing the deposits in Basin 4 (see colorversion online).




model minibasin and b) either driving the bottom underflow with fine-grained sediment or adding this sediment to the saline inflow.

The cyclic steps observed in the experiments have qualitative analogs in

terms of a) the sediment waves that widely occur on the external flanks of

submarine channel levees (Fig. 21), and b) the cyclic scours that areobserved along the thalwegs of some steep canyons and distributary

channels created at partial channel avulsions. These two field analogues

share an important conceptual similarity in that they feature long-wave

bedforms, i.e., with wavelengths that are at least an order of magnitudelarger than the estimated typical current thickness. This is in contrast toshort-wave bedforms such as ripples, dunes, or antidunes. The two

analogues have also significant differences. Sediment waves on channel

levees are net-depositional, composed predominantly of mud, and show a

pattern of coherent upstream migration in orderly trains (Normark et al.

2002). Cyclic scours in distributarychannels off the Eel Canyon (Lamb et al.

2008) or the Monterrey canyon (Fildani et al. 2006) are net-erosional, and

their composition is sand-rich. Yet we speculate that both analogues aredistinct manifestations of the same universal bedform feature known as

cyclic step (Parker 1996). Our experiments unambiguously document, for

the first time in the laboratory, the formation of cyclic steps by turbidity

currents: a train of long bedforms bounded by internal hydraulic jumps,

showing a coherent pattern of slow upstream migration. Our cyclic steps are

net-depositional by design (the sediment supply was chosen to ensure thatthe flows were net-depositional), yet they show some events of reincision of

previous deposits on the lee sides of the steps. The plastic sediment that

composes them scale up to sand, and the flows do not contain mud, which is

instead modeled by salt, which does not interact with the bed. Overall, theobservations suggest that cyclic steps might be net-erosional or net-

depositional, with net-erosional steps favored for swift turbidity currents

over steep slopes, for which the mud would be transported in bypass

conditions and would not interact with the bed.

The downstream-migrating antidunes would appear to be tentativeanalogs for the gravel waves observed in the Var Submarine Canyon

(Piper and Savoye 1993). Figure 22 shows a photograph of the field of

gravel waves obtained from a SAR deep-tow side-scan sonar image across

the Var canyon. The waves have an average height of around 2 to 3 mand a typical wavelength of around 30 to 40 m. They are present in a

reach where the canyon slope is around 3%, clearly supporting that they

were emplaced by supercritical turbidity currents, and excluding the

possibility that they can be classified as dunes. Close-up observations with

a submersible ROV revealed that they are composed of a variety of coarse

grains up to cobble size. The size of the cobbles suggests that their

dominant mode of transport was bedload.

Although none of the field analogue discussed above is a perfect match

to our experiments, the analogy is deemed sufficiently convincing to

justify further research. Linking the geometrical characteristics of turbidity-current deposits in general, and bedforms in particular, to the

properties of the flows that sculpted them may provide a valuable tool for

the interpretation of deep-sea sedimentary systems, particularly in the

context of hydrocarbon-related problems. Such a tool should ultimately

aid in back-calculating the flows and extrapolating over wide areas the

stratigraphic and sedimentologic data available from a few seismic

profiles and well cores at selected drilling sites.

CONCLUSIONS

The results reported here build on the work of the companion paper

(Sequeiros et al. 2009). They document the following features of interest:

N Two sets of experiments document the formation of a wedge-shaped

sedimentary deposit associated with supercritical turbidity currents

rapidly decelerating in the downstream direction. The structure of thedeposit suggests a topset–foreset sequence that has a conceptual

similarity with deltaic deposits. Presumably the foreset is associated

with sharp deceleration enforced by an internal hydraulic jump. The

deposits have a tentative field analog, i.e., in the terminal basin portion

(Basin 4) of a chain of four Pleistocene intraslope basins oftenreferred to

as the Brazos–Trinity Intra-Slope System (Beaubouef et al. 2003).

N The experiments also produced upstream-migrating sediment waves

analogous to those commonly observed in the deep sea (e.g., Lee et al.

2002). Kostic and Parker (2006) and Fildani et al. (2006) used numerical

analyses to identify these sediment waves as falling within the rubric of

cyclic steps (e.g., Parker and Izumi 2000; Taki and Parker 2005; Sun and

Parker 2005), i.e., upstream-migrating bedforms bounded by hydraulic

FIG. 21.—Tentative field analog for the observed upstream-migrating cyclic steps: trains of sediment waves on the levees of the Toyama deep-sea channel. Also, buriedin the deposits are indications of shorter, steeper, and downslope-migrating reflectors (see black arrow) emplaced at an early stage of levee formation. Those might beindicative of early sand antidunes emplaced before the bedforms evolved into larger sediment waves, similar to our experimental observations. After Kubo and Nakajima(2002) and Nakajima and Satoh (2001).




jumps. The experimental results unambiguously allow identification of

the sediment waves observed therein as cyclic steps.

N The experiments also allowed, for the first time, identification of

downstream-migrating bedforms with slip faces as downstream-

migrating antidunes. The bedforms were formed predominantly by

bedload transport. They likely have field analogs in terms of thegravel waves in the Var Submarine Canyon off the Mediterranean

coast of France (Piper and Savoye 1993).

ACKNOWLEDGMENTS

Funding for this work from ExxonMobil Exploration Co. as part of theStratigraphy Tripod Project is gratefully acknowledged. The first author alsoacknowledges the support of the Fonds Special de Recherche, Universitecatholique de Louvain. Constructive feedback and suggestions from reviewersG. Postma and H.M. Pantin, as well as from Associate Editor Bill McCaffreyand Corresponding Editor John B. Southard, were highly valuable andappreciated. The authors also thank Enrica Viparelli, Eric Anders, MarianoCantero, Andy Waratuke, Rocio Luz Fernandez, and Martino Salvaro fortheir assistance during the experiments and helpful discussions.

This manuscript is dedicated to the memory of co-author Bruno Savoye,who passed away accidentally during the review process of this manuscript.Bruno has been an inextinguishable source of inspiration for linking processesobserved at laboratory scale with field observations. The community of sedimentologists and marine geologists has lost not only a friend, but also agreat contributor to research on deep-sea turbiditic systems in general, andbedforms and sediment waves in particular.

NOTATIONS

B channel width

C 0 concentration of salt at the inflow point

C layer-averaged salinity concentration

D Particle grain size [m]

Frd densimetric Froude number

Q saline current dischargeG s mass rate of sediment transport [g/min]Qs volumetric rate of sediment transport [m3/s]T time elapsed since start of each set (equivalent continuous flow)X downstream coordinate

W Particle size in the Krumbein phi logarithmic scale, W 5 2log2D, with D in mm.

r fluid densityrs density of sediment materialgb geometric standard deviationl bulk porosity of bed deposits g geometric standard deviation

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Received 27 March 2008; accepted 4 February 2009.


Documents

Experiments in Wedge-Shaped Deep Sea Sedimentary Deposits Part II - Evolution and Bedforms - JSR, 2009