Sediment transport

Sediment transport is movement of grains by a flowing fluid, through motion close to a bed or within a particle suspension. A flow's shear stress and fluctuating forces compete with grain weight, contacts and friction. The Shields parameter compares a bed-stress scale with the submerged grain-weight scale and helps organize the threshold for movement.

Bedload transport carries grains along or near the bed, while suspended sediment balances settling against upward mixing and entrainment. The transport budget changes the bed itself: mass conservation in the Exner equation relates downstream flux divergence to erosion or deposition. A flow entering an erodible reach can take a finite distance to establish its transported grain population, described by sediment transport saturation. These couplings make a bed an evolving part of the fluid problem rather than a fixed rough wall. Measuring both transported material and bed-height change tests whether the sediment budget closes.

In a reduced sediment-bed model with normalized discharge one, water momentum gives $s=1-u^{-1}+F^2(1-u^2)/2$, while concentration adaptation and bed exchange give $c_x=E(u)-c=-s_t$. Linearizing about $u=c=1,s=0$ yields $\tilde{s}=(1-F^2)\tilde{u}$ and $(1+ik)\tilde{c}=E^{\prime }(1)\tilde{u}$. Elimination proves the displayed dispersion relation.

For $E^{\prime }(1)>0$, the growth rate is $E^{\prime }(1)k^2/[(F^2-1)(1+k^2)]$, positive exactly on the supercritical side $F>1$ for nonzero $k$. The phase speed is $-E^{\prime }(1)/[(F^2-1)(1+k^2)]$, so unstable bedforms migrate upstream and are antidunes. At high wavenumber growth approaches a finite plateau rather than becoming unbounded. This closure selects no finite fastest wavelength; short-scale selection needs additional physics. The critical limit $F=1$ and a vanishing entrainment derivative are degenerate cases outside the displayed conclusion.

Bedload transport is motion of sediment grains along or close to the bed, represented in a depth-averaged model by a solid-volume flux per unit width. The Exner equation balances this flux divergence and changes in porous-bed solid storage against exchange with suspended sediment. Consistent mass units are necessary when adding this equation to a suspended mass-concentration balance.

Keeping sediment acceleration while neglecting water storage and a small gravity-friction residual gives a reduced model for the bed elevation $s$, water speed $u$ and sediment speed $v$. A prescribed water flux $q$ gives $h=q/u$; integrated water momentum gives $s=C-F^2{u}^2/2-q/u$. Around the normalized uniform state $h=u=q=1$, let $(1-v_0{)}^2=\kappa v_0$, $0<v_0<1$. A Fourier mode of wavenumber $k$ has growth-rate real partThus nonuniform modes decay for subcritical Froude number and grow for supercritical Froude number. The spatially constant mode is neutral, and the critical value needs terms neglected in this reduction. In a general uniform state the sign change is at the actual mean-flow Froude number equal to one, rather than at a velocity-scale parameter unrelated to that state.

Suspended sediment is particulate material carried within a fluid rather than solely along the bed. Its concentration balance includes advection, erosion into suspension and deposition from suspension. For water depth $h$ and suspended mass concentration $c$, the conservative balance $(hc{)}_t+(huc{)}_x={\rho }_s(v_E-v_D)$ reduces using water continuity to $h(c_t+u{c}_x)={\rho }_s(v_E-v_D)$.

A dune is a sediment bedform. In the elementary subcritical flow description of river dunes, bed undulations and free-surface undulations are out of phase. This phase relation identifies a hydraulic response; instability additionally requires a positive real part of the relevant dispersion relation. Contrast the in-phase surface of an antidune.

An antidune is a bedform associated with supercritical free-surface flow, in which bed and surface undulations are approximately in phase. In the elementary erodible-bed model with suspended sediment, an unstable branch can migrate upstream. Its growth and migration depend on the hydraulic and sediment-response parameters; the high-wavenumber antidune dispersion calculation makes the required supercritical regime explicit.

In the fast-water limit of an erodible-bed shallow water equations model, let $e=E^{\prime }(1)>0$, $q_1=q^{\prime }(1)>0$ and $\gamma >0$ be the suspended-load relaxation coefficient. Linearization and elimination give the dispersion relation $\lambda =-k^2[q_1+\beta e/(\gamma +ik)]/[3\delta +ik(F^2-1)]$. Its real part tends to the finite constant in the title, and phase velocity tends to $-q_1/(F^2-1)$. Thus the unstable upstream branch requires supercritical flow and enough sediment feedback to overcome friction; subcritical flow has the opposite propagation sign and decay at leading order. The near-critical limit needs a separate balance.

Sediment resuspension transfers deposited bed particles into a particle suspension. It can supply particle buoyancy flux to a particle-laden gravity current if erosion or upward mixing exceeds settling losses. Ambient fluid without particles dilutes concentration without supplying particle mass.

For downward settling velocity magnitude $W_s>0$ and eddy diffusivity ${\kappa }_T=Kz$, the total upward particle flux is $J=-W_s\varphi -Kz{\varphi }_z$. A stationary profile tending to zero at infinity has $J=0$, giving the displayed power law. The exponent compares settling with turbulent mixing. It describes a region above the bed; entrainment from the bed and viscous near-wall effects determine the reference concentration separately.

Sediment transport saturation is relaxation of a transported grain population toward the equilibrium sediment flux supported by local flow conditions. A change in bed shear stress does not immediately give a fully adjusted grain speed and population. Close to equilibrium, a steady model isFor constant $q_{\mathrm{sat}}$ and saturation length $L_{\mathrm{sat}}>0$, this gives $q=q_{\mathrm{sat}}+[q(0)-q_{\mathrm{sat}}]e^{-x/L_{\mathrm{sat}}}$. The spatial lag matters for bedform growth.^[ref]

A published wind tunnel test used a $1\mathrm{m}$ wide, $0.5\mathrm{m}$ high and $4.5\mathrm{m}$ long tunnel. Repeated measurements of bed elevation after stopping the flow supplied an erosion rate; integrating that rate inferred downstream flux by mass conservation. The fitted near-equilibrium relaxation length was distinguished from the earlier grain-ejection stage, so the total development distance was not treated as the same parameter.^[ref]

The equilibrium sediment flux is the transport supported locally by a prescribed bed shear stress. A common idealization is $q_{\mathrm{sat}}={\varphi }_b\chi (\tau -{\tau }_{\mathrm{th}}{)}_{+}^{\gamma }$. Linearization above the sediment entrainment threshold requires perturbations small compared with the positive excess stress.

The saturation length measures downstream relaxation of a transported grain population toward its equilibrium sediment transport flux. The elementary steady relaxation law is $L_{\mathrm{sat}}{q}_x=q_{\mathrm{sat}}-q$; it introduces a phase lag between forcing and actual transport.

The Exner equation is local sediment mass conservation expressed in terms of solid volume: accumulation raises the bed and flux divergence lowers it. Here ${\varphi }_b$ is the solid volume fraction of the packed bed and $q$ is solid volume flux per unit width.

A leading fixed-discharge, level-surface approximation to a shallow erodible-bed flow gives perturbation equations $C_x=4U-C$ and $U_t+3\beta U_x=-4U+C$. The displayed Fourier growth rate has negative real part for every nonzero wave number, independently of bedload parameter $\beta$. Bedload changes the propagation speed but not this damping. The zero-wave-number mode is neutral, so absence of growing modes should not be confused with strict asymptotic decay of every spatially uniform perturbation.

The Shields parameter is $\Theta =\tau /[({\rho }_p-{\rho }_f)gd]$, comparing bed shear stress with a grain-scale submerged weight per area. Its critical value depends on the assumed drag coefficient, static friction, lift, and contact geometry.

The sediment entrainment threshold is the bed shear stress at which resting grains begin moving. An elementary force balance using spherical grains, tangential quadratic drag, and static friction gives ${\Theta }_{\mathrm{th},0}=4{\mu }_s/(3C_D)$ when the drag coefficient is defined relative to the bed shear velocity.

For locally bed-tangent drag force and an upslope angle $a$, a spherical-grain force balance gives ${\tau }_{\mathrm{th}}(a)={\tau }_{\mathrm{th},0}(\cos a+\sin a/{\mu }_s)$. The small-slope term is $\delta {\tau }_{\mathrm{th}}=({\tau }_{\mathrm{th},0}/{\mu }_s){\eta }_x$. A horizontal drag force instead changes the normal contact force and gives a different slope correction.

A bedform is organized relief on a mobile sediment bed, such as a dune or an antidune. Its shape changes the overlying flow, which changes the shear stress and the sediment transport; erosion and deposition then feed back on the shape. The pattern is therefore a coupled fluid-and-grain response rather than simply a frozen pile of material.

For a small sinusoidal bed disturbance, an upstream lead of the forcing relative to bed height can favor growth, while sediment transport saturation introduces a downstream lag. The Exner equation converts the resulting flux divergence into bed-height change. Their combination determines both growth and migration; the direction of migration cannot be inferred from a bed's appearance alone. Recording a bed profile at successive times separates changing amplitude from movement of a crest, while flow measurements test the assumed bed shear response.

Combining the Exner equation, sediment transport saturation, bed shear response, and inclined-bed sediment threshold gives a competition between upstream forcing and downstream relaxation. With $a={\tau }_{0}A$, $b={\tau }_{0}B-{\tau }_{\mathrm{th},0}/{\mu }_s$, $C=\chi \gamma ({\tau }_0-{\tau }_{\mathrm{th},0}{)}^{\gamma -1}$, and positive wavenumber $k$, the growth rate is $\sigma =C{k}^2(b-ak{L}_{\mathrm{sat}})/(1+k^2{L}_{\mathrm{sat}}^2)$. Thus $b>0$ permits long-wave growth when $a>0$, while a finite saturation length stabilizes shorter waves.

For a small sinusoidal bedform, a fluid-dynamical closure can be written $\hat{\tau }={\tau }_{0}k(A+iB)\hat{\eta }$ for $k>0$. The real coefficient $A$ is in phase with the bed height and $B$ describes an upstream phase lead. Neither coefficient follows from the Exner equation or sediment relaxation alone; they require a flow calculation or measurement.