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Dust blowing from the Sahara Desert over the Atlantic Ocean towards the Canary Islands

Sediment transport is the movement of solid particles (sediment), typically due to a combination of gravity acting on the sediment, and/or the movement of the fluid in which the sediment is entrained. Sediment transport occurs in natural systems where the particles are clastic rocks (sand, gravel, boulders, etc.), mud, or clay; the fluid is air, water, or ice; and the force of gravity acts to move the particles along the sloping surface on which they are resting. Sediment transport due to fluid motion occurs in rivers, oceans, lakes, seas, and other bodies of water due to currents and tides. Transport is also caused by glaciers as they flow, and on terrestrial surfaces under the influence of wind. Sediment transport due only to gravity can occur on sloping surfaces in general, including hillslopes, scarps, cliffs, and the continental shelf—continental slope boundary.

Sediment transport is important in the fields of sedimentary geology, geomorphology, civil engineering, hydraulic engineering and environmental engineering (see applications, below). Knowledge of sediment transport is most often used to determine whether erosion or deposition will occur, the magnitude of this erosion or deposition, and the time and distance over which it will occur.

Mechanisms

Sand blowing off a crest in the Kelso Dunes of the Mojave Desert, California.

Toklat River, East Fork, Polychrome overlook, Denali National Park, Alaska. This river, like other braided streams, rapidly changes the positions of its channels through processes of erosion, sediment transport, and deposition.

Congo river viewed from Kinshasa, Democratic Republic of Congo. Its brownish color is mainly the result of the transported sediments taken upstream.

Aeolian

Aeolian or eolian (depending on the parsing of æ) is the term for sediment transport by wind. This process results in the formation of ripples and sand dunes. Typically, the size of the transported sediment is fine sand (<1 mm) and smaller, because air is a fluid with low density and viscosity, and can therefore not exert very much shear on its bed.

Bedforms are generated by aeolian sediment transport in the terrestrial near-surface environment. Ripples^[1] and dunes^[2] form as a natural self-organizing response to sediment transport.

Aeolian sediment transport is common on beaches and in the arid regions of the world, because it is in these environments that vegetation does not prevent the presence and motion of fields of sand.

Wind-blown very fine-grained dust is capable of entering the upper atmosphere and moving across the globe. Dust from the Sahara deposits on the Canary Islands and islands in the Caribbean,^[3] and dust from the Gobi desert has deposited on the western United States.^[4] This sediment is important to the soil budget and ecology of several islands.

Deposits of fine-grained wind-blown glacial sediment are called loess.

Fluvial

In geology, physical geography, and sediment transport, fluvial processes relate to flowing water in natural systems. This encompasses rivers, streams, periglacial flows, flash floods and glacial lake outburst floods. Sediment moved by water can be larger than sediment moved by air because water has both a higher density and viscosity. In typical rivers the largest carried sediment is of sand and gravel size, but larger floods can carry cobbles and even boulders.

Fluvial sediment transport can result in the formation of ripples and dunes, in fractal-shaped patterns of erosion, in complex patterns of natural river systems, and in the development of floodplains.

Sand ripples, Laysan Beach, Hawaii. Coastal sediment transport results in these evenly spaced ripples along the shore. Monk seal for scale.

Coastal

Coastal sediment transport takes place in near-shore environments due to the motions of waves and currents. At the mouths of rivers, coastal sediment and fluvial sediment transport processes mesh to create river deltas.

Coastal sediment transport results in the formation of characteristic coastal landforms such as beaches, barrier islands, and capes.^[5]

A glacier joining the Gorner Glacier, Zermatt, Switzerland. These glaciers transport sediment and leave behind lateral moraines.

Glacial

As glaciers move over their beds, they entrain and move material of all sizes. Glaciers can carry the largest sediment, and areas of glacial deposition often contain a large number of glacial erratics, many of which are several metres in diameter. Glaciers also pulverize rock into "glacial flour", which is so fine that it is often carried away by winds to create loess deposits thousands of kilometres afield. Sediment entrained in glaciers often moves approximately along the glacial flowlines, causing it to appear at the surface in the ablation zone.

Hillslope

In hillslope sediment transport, a variety of processes move regolith downslope. These include:

• Soil creep
• Tree throw
• Movement of soil by burrowing animals
• Slumping and landsliding of the hillslope

These processes generally combine to give the hillslope a profile that looks like a solution to the diffusion equation, where the diffusivity is a parameter that relates to the ease of sediment transport on the particular hillslope. For this reason, the tops of hills generally have a parabolic concave-up profile, which grades into a convex-up profile around valleys.

As hillslopes steepen, however, they become more prone to episodic landslides and other mass wasting events. Therefore, hillslope processes are better described by a nonlinear diffusion equation in which classic diffusion dominates for shallow slopes and erosion rates go to infinity as the hillslope reaches a critical angle of repose.^[6]

Debris flow

Large masses of material are moved in debris flows, hyperconcentrated mixtures of mud, clasts that range up to boulder-size, and water. Debris flows move as granular flows down steep mountain valleys and washes. Because they transport sediment as a granular mixture, their transport mechanisms and capacities scale differently from those of fluvial systems.

Applications

Suspended sediment from a stream emptying into a fjord (Isfjorden, Svalbard, Norway).

Sediment transport is applied to solve many environmental, geotechnical, and geological problems. Measuring or quantifying sediment transport or erosion is therefore important for coastal engineering. Several sediment erosion devices have been designed in order to quantify sediment erosion (e.g., Particle Erosion Simulator (PES)). One such device, also referred to as the BEAST (Benthic Environmental Assessment Sediment Tool) has been calibrated in order to quantify rates of sediment erosion.^[7]

Movement of sediment is important in providing habitat for fish and other organisms in rivers. Therefore, managers of highly regulated rivers, which are often sediment-starved due to dams, are often advised to stage short floods to refresh the bed material and rebuild bars. This is also important, for example, in the Grand Canyon of the Colorado River, to rebuild shoreline habitats also used as campsites.

Sediment discharge into a reservoir formed by a dam forms a reservoir delta. This delta will fill the basin, and eventually, either the reservoir will need to be dredged or the dam will need to be removed. Knowledge of sediment transport can be used to properly plan to extend the life of a dam.

Geologists can use inverse solutions of transport relationships to understand flow depth, velocity, and direction, from sedimentary rocks and young deposits of alluvial materials.

Flow in culverts, over dams, and around bridge piers can cause erosion of the bed. This erosion can damage the environment and expose or unsettle the foundations of the structure. Therefore, good knowledge of the mechanics of sediment transport in a built environment are important for civil and hydraulic engineers.

When suspended sediment transport is increased due to human activities, causing environmental problems including the filling of channels, it is called siltation after the grain-size fraction dominating the process.

Initiation of motion

Stress balance

For a fluid to begin transporting sediment that is currently at rest on a surface, the boundary (or bed) shear stress ${\displaystyle \tau _{b}}$ exerted by the fluid must exceed the critical shear stress ${\displaystyle \tau _{c}}$ for the initiation of motion of grains at the bed. This basic criterion for the initiation of motion can be written as:

${\displaystyle \tau _{b}=\tau _{c}}$.

This is typically represented by a comparison between a dimensionless shear stress ${\displaystyle \tau _{b}*}$ and a dimensionless critical shear stress ${\displaystyle \tau _{c}*}$. The nondimensionalization is in order to compare the driving forces of particle motion (shear stress) to the resisting forces that would make it stationary (particle density and size). This dimensionless shear stress, ${\displaystyle \tau *}$, is called the Shields parameter and is defined as:^[8]

${\displaystyle \tau *={\frac {\tau }{(\rho _{s}-\rho _{f})(g)(D)}}}$.

And the new equation to solve becomes:

${\displaystyle \tau _{b}*=\tau _{c}*}$.

The equations included here describe sediment transport for clastic, or granular sediment. They do not work for clays and muds because these types of floccular sediments do not fit the geometric simplifications in these equations, and also interact thorough electrostatic forces. The equations were also designed for fluvial sediment transport of particles carried along in a liquid flow, such as that in a river, canal, or other open channel.

Only one size of particle is considered in this equation. However, river beds are often formed by a mixture of sediment of various sizes. In case of partial motion where only a part of the sediment mixture moves, the river bed becomes enriched in large gravel as the smaller sediments are washed away. The smaller sediments present under this layer of large gravel have a lower possibility of movement and total sediment transport decreases. This is called armouring effect.^[9] Other forms of armouring of sediment or decreasing rates of sediment erosion can be caused by carpets of microbial mats, under conditions of high organic loading.^[10]

Critical shear stress

Original Shields diagram, 1936

The Shields diagram empirically shows how the dimensionless critical shear stress (i.e. the dimensionless shear stress required for the initiation of motion) is a function of a particular form of the particle Reynolds number, ${\displaystyle \mathrm {Re} _{p}}$ or Reynolds number related to the particle. This allows the criterion for the initiation of motion to be rewritten in terms of a solution for a specific version of the particle Reynolds number, called ${\displaystyle \mathrm {Re} _{p}*}$.

${\displaystyle \tau _{b}*=f\left(\mathrm {Re} _{p}*\right)}$

This can then be solved by using the empirically derived Shields curve to find ${\displaystyle \tau _{c}*}$ as a function of a specific form of the particle Reynolds number called the boundary Reynolds number. The mathematical solution of the equation was given by Dey.^[11]

Particle Reynolds number

In general, a particle Reynolds number has the form:

${\displaystyle \mathrm {Re} _{p}={\frac {U_{p}D}{\nu }}}$

Where ${\displaystyle U_{p}}$ is a characteristic particle velocity, ${\displaystyle D}$ is the grain diameter (a characteristic particle size), and ${\displaystyle \nu }$ is the kinematic viscosity, which is given by the dynamic viscosity, ${\displaystyle \mu }$, divided by the fluid density, ${\displaystyle {\rho _{f}}}$.

${\displaystyle \nu ={\frac {\mu }{\rho _{f}}}}$

The specific particle Reynolds number of interest is called the boundary Reynolds number, and it is formed by replacing the velocity term in the particle Reynolds number by the shear velocity, ${\displaystyle u_{*}}$, which is a way of rewriting shear stress in terms of velocity.

${\displaystyle u_{*}={\sqrt {\frac {\tau _{b}}{\rho _{f}}}}=\kappa z{\frac {\partial u}{\partial z}}}$

where ${\displaystyle \tau _{b}}$ is the bed shear stress (described below), and ${\displaystyle \kappa }$ is the von Kármán constant, where

${\displaystyle \kappa ={0.407}}$.

The particle Reynolds number is therefore given by:

${\displaystyle \mathrm {Re} _{p}*={\frac {u_{*}D}{\nu }}}$

Bed shear stress

The boundary Reynolds number can be used with the Shields diagram to empirically solve the equation

${\displaystyle \tau _{c}*=f\left(\mathrm {Re} _{p}*\right)}$,

which solves the right-hand side of the equation

${\displaystyle \tau _{b}*=\tau _{c}*}$.

In order to solve the left-hand side, expanded as

${\displaystyle \tau _{b}*={\frac {\tau _{b}}{(\rho _{s}-\rho _{f})(g)(D)}}}$,

the bed shear stress needs to be found, ${\displaystyle {\tau _{b}}}$. There are several ways to solve for the bed shear stress. The simplest approach is to assume the flow is steady and uniform, using the reach-averaged depth and slope. because it is difficult to measure shear stress in situ, this method is also one of the most-commonly used. The method is known as the depth-slope product.

Depth-slope product

For a river undergoing approximately steady, uniform equilibrium flow, of approximately constant depth h and slope angle θ over the reach of interest, and whose width is much greater than its depth, the bed shear stress is given by some momentum considerations stating that the gravity force component in the flow direction equals exactly the friction force.^[12] For a wide channel, it yields:

${\displaystyle \tau _{b}=\rho gh\sin(\theta )}$

For shallow slope angles, which are found in almost all natural lowland streams, the small-angle formula shows that ${\displaystyle \sin(\theta )}$ is approximately equal to ${\displaystyle \tan(\theta )}$, which is given by ${\displaystyle S}$, the slope. Rewritten with this:

${\displaystyle \tau _{b}=\rho ghS}$

Shear velocity, velocity, and friction factor

For the steady case, by extrapolating the depth-slope product and the equation for shear velocity:

${\displaystyle \tau _{b}=\rho ghS}$

${\displaystyle u_{*}={\sqrt {\left({\frac {\tau _{b}}{\rho }}\right)}}}$,

The depth-slope product can be rewritten as:

${\displaystyle \tau _{b}=\rho u_{*}^{2}}$.

${\displaystyle u*}$ is related to the mean flow velocity, ${\displaystyle {\bar {u}}}$, through the generalized Darcy-Weisbach friction factor, ${\displaystyle C_{f}}$, which is equal to the Darcy-Weisbach friction factor divided by 8 (for mathematical convenience).^[13] Inserting this friction factor,

${\displaystyle \tau _{b}=\rho C_{f}\left({\bar {u}}\right)^{2}}$.

Unsteady flow

For all flows that cannot be simplified as a single-slope infinite channel (as in the depth-slope product, above), the bed shear stress can be locally found by applying the Saint-Venant equations for continuity, which consider accelerations within the flow.

Example

Set-up

The criterion for the initiation of motion, established earlier, states that

${\displaystyle \tau _{b}*=\tau _{c}*}$.

In this equation,

${\displaystyle \tau *={\frac {\tau }{(\rho _{s}-\rho )(g)(D)}}}$, and therefore

${\displaystyle \frac{{\mathrm{\tau <!--\; \tau \; -->}}_b}{({\mathrm{\rho <!--\; \rho \; -->}}_s-<!--\; -\; -->\mathrm{\rho <!--\; \rho \; -->}}}$Zdroj:https://en.wikipedia.org?pojem=Sediment_transport
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