hydraulics of mountain rivers gary parker, university of illinois

1

National Center for Earth-surface DynamicsShort Course

Morphology, Morphodynamics and Ecology of Mountain Rivers

December 11-12, 2005HYDRAULICS OF MOUNTAIN RIVERS

Gary Parker, University of Illinois

River in Taiwan: Image courtesy C. Stark

2



December 11-12, 2005TOPICS COVERED

This lecture is not intended to provide a full treatment of open channel flow. Nearly all undergraduate texts in fluid mechanics for civil engineers have sections on open channel flow (e.g. Crowe et al., 2001). Three texts that specifically focus on open channel flow are those by Henderson (1966), Chaudhry (1993) and Jain (2000).

Topics treated here include:• Approximations for the channel• Shields number, Einstein number, generic bedload equation • Boundary resistance in mountain streams: Chezy and Manning-Strickler forms• Skin friction and form drag in mountain rivers• Backwater and the backwater length• Normal (steady, uniform) flow• Calculations of flow and sediment transport using the normal flow assumption

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December 11-12, 2005SIMPLIFICATION OF CHANNEL CROSS-SECTIONAL SHAPE

River channel cross sections have complicated shapes. In a 1D analysis, it is appropriate to approximate the shape as a rectangle, so that B denotes channel width and H denotes channel depth (reflecting the cross-sectionally averaged depth of the actual cross-section). As was seen the lecture on hydraulic geometry, natural channels are generally wide in the sense that Hbf/Bbf << 1, where the subscript “bf” denotes “bankfull”. As a result the hydraulic radius Rh is usually approximated reasonably accurately by the average depth. In terms of a rectangular channel,

H

B

channel floodplainfloodplain

H

BH

21

H

H2B

HBRh

1

10

100

0.0001 0.001 0.01 0.1

S

Bb

f/Hb

f AltaBritIdaColo

4



December 11-12, 2005THE SHIELDS NUMBER:

A KEY DIMENSIONLESS PARAMETER QUANTIFYING SEDIMENT MOBILITY

gDRb

3

c

2b

2g

34

~

DR

D

b = boundary shear stress at the bed (= bed drag force acting on the flow per unit bed area) [M/L/T2]

c = Coulomb coefficient of resistance of a granule on a granular bed [1]D = characteristic grain size (e.g. surface median size Ds50)

Recalling that R = (s/) – 1, the Shields Number * is defined as

It can be interpreted as scaling the ratio impelling force of flow drag acting on a bed particle to the Coulomb force resisting motion acting on the same particle, so that

The characterization of bed mobility thus requires a quantification of boundary shear stress at the bed.

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December 11-12, 2005THRESHOLD OF MOTION IN MOUNTAIN STREAMS

The threshold of motion is often expressed in terms of a Shields curve (Shields, 1936). Let denote the value of * at the threshold of motion, and Rep denote grain Reynolds number, defined in the lecture on hydraulic geometry as

Based Neill’s (1968) work on coarse sedimentg, Parker et al. (2003) amended the Brownlie (1981) fit of the original Shields curve to the form

The asymptotic value of for large Rep, i.e. coarse sediment, is 0.03. Consider the case of quartz (R = 1.65) in water at 20C ( = 1x10-6 m2/s). The smallest value of Ds50 for the data set introduced in the lecture on hydraulic geometry is 27 mm, in which case Rep = 17,850 and = 0.0289. So an approximate value of of 0.03 is appropriate for most coarse-bedded mountain streams.

In point of fact, there is no sharply-defined threshold of motion. The value of 0.03 should be interpreted to be a value below which the bedload transport rate is morphodynamically insignificant, not precisely 0.

c

DRgDpRe

]1006.022.0[5.0 )7.7(6.0pc

6.0p ReRe

cc

6



December 11-12, 2005SHIELDS NUMBER AT BANKFULL FLOW IN MOUNTAIN STREAMS

It will be shown in Slide 26 that for the case of mountain streams the Shields number at bankfull flow can be estimated as

where all parameters were defined in the lecture on hydraulic geometry. A plot of versus

is given to the right. The data are those from the lecture on hydraulic geometry. The average value of is 0.0486, i.e about 1.62 times the Shields number at the threshold of motion. Most alluvial gravel-bed streams move size Ds50 at bankfull flow.

50s

bf50bf DR

SH

250s50s

bf

DgD

QQ̂

50bf

50bf

Bankfull Shields Number versus Dimensionless Discharge

0.001

0.01

0.1

1

1.0E+02 1.0E+03 1.0E+04 1.0E+05 1.0E+06 1.0E+07

Qhat

bf*

AltaBritIdaColoaveragethreshold

Q̂

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December 11-12, 2005THE EINSTEIN NUMBER:A DIMENSIONLESS QUANTIFICATION OF BEDLOAD TRANSPORT RATE

DgDR

qq b

qb = volume bedload transport rate per unit width [L2/T]D = characteristic grain size [L]R = (s/) – 1 1.65 for natural sediment [1]

One standard approach to the quantification of bedload transport is the specification of the functional form

An example appropriate for the bedload transport of gravel of uniform size is the modified form of the bedload transport equation of Meyer-Peter and Müller (1948) by Wong (2003):

In field gravel-bed rivers, however, a) the gravel is rarely uniform and b) gravelis commonly transported at Shields numbers below 0.0495 (see previous slide).

)(fq

0495.0,0

0495.0,)0495.0(97.3q

2/3

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December 11-12, 2005

LetQ = flow discharge [L/T]B = water surface width [L]H = cross-sectionally averaged depth [L]U = Q/(BH) = cross-sectionally averaged flow velocity [L]u* = (b/)1/2 = shear velocity [L/T]

Two dimensionless bed resistance coefficients are defined here: the Chezy resistance coefficient Cz given as

and the standard bed resistance coefficient Cf (= f/8, where f denotes the D’arcy-Weisbach resistance coefficient);

Note that as the bed shear stress increases, Cf increases and Cz decreases.

BOUNDARY RESISTANCE IN MOUNTAIN STREAMS

2/1fC

u

UCz

22

b

2

f )Cz(UU

uC

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December 11-12, 2005RESISTANCE RELATIONS FOR HYDRAULICALLY ROUGH FLOW

6/1

sr

2/1f k

HC

u

UCz

Mountain streams are almost invariably in the range of hydraulically rough flow, for which resistance becomes independent of kinematic viscosity . Keulegan (1938) offered the following relation for hydraulically rough flow.

90sks Dnk

s

2/1f k

H11n5.2C

u

UCz

where ks = a roughness height characterizing the bumpiness of the bed [L]. A close approximation is offered by the Manning-Strickler formulation:

Parker (1991) suggested a value of r of 8.1 for gravel-bed streams.

The roughness height over a flat bed of coarse grains (no bedforms) is given as

where Ds90 denotes the surface sediment size such that 90 percent of the surface material is finer, and nk is a dimensionless number between 1.5 and 3. For example, Kamphuis (1974) evaluated nk as equal to 2.

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December 11-12, 2005COMPARISION OF KEULEGAN AND MANNING-STRICKLER RELATIONS

r = 8.1

Note that Cz does not vary strongly with depth. It is often approximated as a constant in broad-brush calculations.

1

10

100

1 10 100 1000

H/ks

Cz

Keulegan

Parker Version of Manning-Strickler

6/1

sk

H1.8Cz

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1.00

10.00

100.00

1.00 10.00 100.00

Rb/ks

Cz

ETH 52

Gilbert 116

Parker Version of Manning-Strickler

6/1

s

b

k

R1.8Cz

TEST OF RESISTANCE RELATION AGAINST MOBILE-BED DATA WITHOUT BEDFORMS FROM LABORATORY FLUMES

Here Rb denotes the bed component of hydraulic radius Rh rather than depth: the flumes were too narrow to allow the approximation Rh H.

The data in question are from all the experiments without bedforms used by Meyer-Peter and Müller (1948) to develop their bedload relation.

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December 11-12, 2005FORM DRAG AND SKIN FRICTIONThe formulas of the previous slide hold only for flat, coarse granular beds. River beds are rarely flat. Lowland sand-bed streams typically contain dunes. Dunes are not common in mountain streams, but such bedforms as bars, pool-riffle sequences and step-pool sequences are common. All such features, as well as planform irregularity, contributed added resistance, so Cz is usually lower, and Cf is usually higher, than predicted by the equation of the previous slide.

Bars in the Rhine River, Switzerland. Cour. M.

Jaeggi

Step-pool pattern in the Hiyamizudani river, Japan. Cour. K. Hasegawa

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December 11-12, 2005FORM DRAG AND SKIN FRICTION contd.The component of the resistance coefficient Cfs due to the flow acting on the grains themselves is known as skin friction, and the extra component Cff is denoted as form drag (due to bedforms), to that the total resistance coefficient Cf is given as

Elbow River, Alberta, Canada at low flow and 100-year flood. Cour. Alberta

Research Council

fffsf CCC

Cfs can be computed from the relations of Slides 10 and 11.

In bar-dominated mountain streams, form-drag is prominent at lower flows, but is muted at flood flows (see images to the right). In steeper streams with pool-riffle patterns, and in particular step-pool patterns, form drag is likely significant even at flood flows.

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December 11-12, 2005BACKWATER

If (most) alluvial streams are disturbed at a point, the effect of that disturbance tends to propagate upstream. For example, the effect of a lake (slowing the flow down) or a waterfall (speeding the flow up) is felt upstream, as illustrated below.

Backwater from a lakeBackwater from a

waterfallBackwater effects are mediated by a dimensionless number known as the Froude number Fr, where

gH

UFr

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December 11-12, 2005BACKWATER contd.

For most alluvial mountain streams, the Froude number at bankfull flow Frbf satisfies the condition

As seen from the lecture on hydraulic geometry. For the case Fr < 1 backwater effects propagate upstream, so that the effect of a disturbance is felt upstream of it.

)flowlsubcritica(1bf Fr

0.1

1

10

0.0001 0.001 0.01 0.1

S

Fr b

f

AltaBritIdaColo

16



December 11-12, 2005BACKWATER contd.

Supercritical flow (Fr > 1) does occur in very steep mountain streams, and in particular streams with step-pool patterns and bedrock streams. In the case of a supercritical flow the effect of a disturbance propagates downstream rather than upstream.

Dry Meadow Creek, Calif., USA. Cour. M. Neumann

Stream in the interior of British Columbia, Canada. Cour. B.

Eaton.Sustained supercritical flow over an alluvial or bedrock bed is unstable, and usually devolves into a series of steps punctuated byhydraulic jumps at formative flow.

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December 11-12, 2005THE BACKWATER LENGTH

The characteristic distance Lb upstream (in the case of the more usual subcritical flow) or downstream (in the case of supercritical flow) to which the effect of a disturbance is felt is known as the backwater length. LetHd = the flow depth at the disturbance [L]S = down-channel slope of the river [1].The backwater length is then given as

Taking Hb to scale with Hbf, some estimates of the backwater length are given below.

Estimates of the backwater length obtained in this way average to 2.1 km, 1.1 km, 0.2 km, and 5.2 km for the data sets for Alberta, Britain, Idaho and Colorado introduced in the lecture on hydraulic geometry. That is, backwater lengthstend to be very short in mountain streams.

S

HL d

b

River Bed Hbf (m) S Lb (m) South Fork Clearwater River, Idaho, USA Gravel 1.06 0.0055 0.2 km Minnesota River, Wilmarth, USA Sand 4.6 0.00019 24.2 km Fly River, Kuambit, PNG Sand 9.45 0.000051 185.3 km

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December 11-12, 2005NORMAL FLOW

Parameters:

x = downstream coordinate [L]H = flow depth [L]U = flow velocity [L/T]qw = water discharge per unit width [L2T-1]B = width [L]Qw = qwB = water discharge [L3/T]g = acceleration of gravity [L/T2] = bed angle [1]b = bed boundary shear stress [M/L/T2]S = tan = streamwise bed slope [1]

(cos 1; sin tan S) = water density [M/L3]

As can be seen from the lecture on hydraulic geometry, the bed slope S of most river, even most mountain rivers, is sufficiently small to allow the approximations

1cos,Stansin

xB

x

gHxBS

bBx

H

Normal flow is an equilibrium state defined by a perfect balance between the downstream gravitational impelling force and resistive bed force, in the absence of any perturbation due to backwater. The resulting flow is constant in time and in the downstream, or x direction. The approximation of normal flow is often a very good one in mountain streams.

19




UHBBqQUHq www Conservation of downstream momentum:Impelling force (downstream component of weight of water) = resistive force

xBxSgHBsinxgHB b

gHSb

Reduce to obtain depth-slope product rule for normal flow:

NORMAL FLOW contd.

Conservation of water mass (= conservation of water volume as water can be treated as incompressible):

xB

x

gHxBS

bBx

H

gHSu

20




1

10

100

1 10 100

Hbf/Ds50

Cz b

f

AltaBritIdaColo

that Cz can be estimated as

This is how Czbf was estimated in the lecture on hydraulic geometry, as shown to the right. If Cz can be estimaged, the flow velocity U is then given as

This relation is known as Chezy’s law.

CHEZY RESISTANCE COEFFICIENT AT BANKFULL FLOW

Using the normal flow approximation, it is found between the relations

gHSu

gHSBH

Q

gHS

UCz

u

UCz

gHSCzU

Regression of all four data sets:210.0

50s

bfbf D

H39.4Cz

21



December 11-12, 2005MANNING-STRICKLER RESISTANCE RELATION FOR PLANE BED

Using the normal flow approximation and the form for Cz for a plane bed in the absence of bedforms given in Slide 11, it is found that

or solving for U,

This is known as a Manning-Strickler resistance relation, where Manning’s “n” (a parameter that should be relegated to the dustbin due to its perverse dimensions) is given as

(But you must remember to use MKS units for n, whereas the equation for U works for any consistent set of units).

6/1

sk

H1.8

gHS

UCz

2/13/26/1

s

SHk

g1.8U

6/1sk

g1.8

n

1

22



December 11-12, 2005MANNING-STRICKLER RESISTANCE RELATION FOR MOUNTAIN RIVERS AT BANKFULL FLOW

The regression of the data of Slide 21 for mountain rivers at bankfull flow yields the relation

where Ubf = Qbf/(HbfBbf), or thus

This represents a generalized Manning-Strickler relation for mountain streams.

210.0

50s

bf

bf

bfbf D

H39.4

SgH

UCz

2/1710.0bf210.0

50sbf SH

D

g39.4U

23



December 11-12, 2005FORM DRAG VERSUS SKIN FRICTION AT BANKFULL FLOW

In order to compare form drag versus skin friction at bankfull flow, it is necessary to estimate the roughness height ks. Here the Kamphuis (1974) relation

Is used in conjunction with the reasonable estimate

Then defining Cf,bf, Cfs,bf and Cff,bf as the values of the total resistance coefficient, the resistance coefficient due to skin friction and the resistance coefficient due to form drag, respectively at bankfull flow, it is found that

The fraction of resistance that is form drag Fform is thus given as

420.0

50s

bf22bfbf,f D

H)39.4(CzC

90ss D2k

50s90s D3D

)3/1(

50s2

3/1

bf,s D

H

)1.8(

)6(Cf

bf,fsbf,fbf,f CCCf

bf,f

bf,fsbf,fform C

CCF

24



December 11-12, 2005FORM DRAG VERSUS SKIN FRICTION AT BANKFULL FLOW contd.

The fraction of resistance that is form drag at bankfull flow in mountain streams is less than 0.5. The fraction is 0.2 ~ 0.3 in relatively deep mountain streams (Hbf/Ds50) > 20, but can be above 0.3 in relatively shallow mountain streams. Deeper streams tend to have lower slopes, and shallower streams tend to have higher slopes, as shown in the next slide.

0.1

1

1 10 100

H/Ds50

Ffo

rm

25



December 11-12, 2005SHIELDS NUMBER AT BANKFULL FLOW USING THE NORMAL FLOW ASSUMPTION

Using the definition of the Shields number * and the estimate for bed shear stress b from the normal flow approximation,

gDRb

gHSb

the following estimate is obtained for the Shields number at bankfull flow:

This is the origin of the estimate of Shields number used in the chapter on hydraulic geometry and in Slide 7 of this lecture.

A crude approximation of the plot to the right yields

so that S decreases as Hbf/Ds50 increases.

50s

bf50bf DR

SH

049.050bf

Bankfull Shields Number versus Dimensionless Discharge

0.001

0.01

0.1

1

1.0E+02 1.0E+03 1.0E+04 1.0E+05 1.0E+06 1.0E+07

Qhat

bf*

AltaBritIdaColoaveragethreshold

26




,UHqw

CALCULATING THE FLOW AT NORMAL EQUILIBRIUM: CHEZY FORMULATION

DR

HS

gDRb

DR

S

g

q

)Cz(

1 3/23/12w

3/2

3/2

w

gSCz

qH

,U)Cz(

,SHgCSHC

gU 2

2b2/12/1

z2/12/1

f

Between the relations

Thus if the water discharge per unit width qw, down-channel bed slope S, characteristic bed grain size D and submerged specific gravity R are known, and if the Chezy resistance coefficient Cz can be estimated, the flow depth H, flow velocity U, bed shear stress b and Shields number * can be computed as indicated above.

3/1w

3/2 gSq)Cz(U

3/2w

3/2b gSq)Cz(

it can be shown that

27



December 11-12, 2005CHEZY MANNING-STRICKLER

6/1

sk

H1.8Cz

For the case of plane-bed rough flow, the following formulation for resistance was given in Slide 10:

Both relations can be cast in terms of a generalized Manning-Strickler formulation, such that

The corresponding relation based on data for mountain rivers at bankfull flow is (Slide 20)

Assuming that ks = 2 Ds90 and Ds90 = 3 Ds50, the above relation can be cast into the form

210.0

50sD

H39.4Cz

210.0

sk

H92.5Cz

gn

sg k

HCz

28




Consider a generalized Manning-Strickler resistance relation of the form

where for example g can be estimated as 5.92, ng can be estimated as 0.210 and ks can be estimated as 2Ds90 for mountain gravel-bed streams at flood flows (Slide 27). The relations for H, U, b and * now become

where

gn

sg k

HCz

CALCULATING THE FLOW AT NORMAL EQUILIBRIUM: MANNING-STRICKLER FORMULATION

msg

n

g

wns

gS

qkH

ms

g

ms

n

ns

g)n1(

wk

gSqU

]n)2/1(1[

n

g

wns]n)2/1(1[

bms

msg

ms Sqk

g

RD

Sqkg

]n)2/1(1[n

g

w)nn(s

n)2/1(ms

ms

msgms

]n)3/2(1[(

)3/2(n

gms

29




For no particularly good reason, most formulations of bedload transport in gravel-bed streams have ignored form drag. For the sake of illustration, we do so here.

Consider a “flume-like” river with no form drag and containing “uniform” gravel of size D, roughness height ks (~ 2D) and submerged specific gravity R. The reach has bed slope S, and is conveying water discharge per unit width qw. For this case it is reasonable to assume ng = 1/6 and g = 8.1, i.e. the relation of Slide 10.

The Shields number can be computed from the previous slide as

and the volume bedload transport rate per unit width q can be estimated from Slide 7 as

CALCULATING BEDLOAD TRANSPORT AT NORMAL EQUILIBRIUM:

RD

Sqkg

]n)2/1(1[n

g

w)nn(s

n)2/1(ms

ms

msgms

2/3b )0495.0(DgDR97.3q

]n)3/2(1[(

)3/2(n

gms

30




In the case of an exponent ng of 1/6 (the standard Manning-Strickler exponent of Slide 10), the relevant relations reduce to:

MANNING-STRICKLER: STANDARD CASE OF ng = 1/6:

5

3nms

5/3

g

w6/1

s

gS

qkH

5/3

6/1s

g5/2

w k

gSqU

10/7

5/3

g

w6/1

s10/7b S

qkg

RD

Sqkg

10/75/3

g

w10/1s

10/3

31



December 11-12, 2005REFERENCESBrownlie, W. R., 1981, Prediction of flow depth and sediment discharge in open channels, Report

No. KH-R-43A, W. M. Keck Laboratory of Hydraulics and Water Resources, California Institute of Technology, Pasadena, California, USA, 232 p.

Chaudhry, M. H., 1993, Open-Channel Flow, Prentice-Hall, Englewood Cliffs, 483 p.Crowe, C. T., Elger, D. F. and Robertson, J. A., 2001, Engineering Fluid Mechanics, John Wiley

and sons, New York, 7th Edition, 714 p.Gilbert, G.K., 1914, Transportation of Debris by Running Water, Professional Paper 86, U.S.

Geological Survey.Jain, S. C., 2000, Open-Channel Flow, John Wiley and Sons, New York, 344 p.Kamphuis, J. W., 1974, Determination of sand roughness for fixed beds, Journal of Hydraulic

Research, 12(2): 193-202.Keulegan, G. H., 1938, Laws of turbulent flow in open channels, National Bureau of Standards

Research Paper RP 1151, USA.Henderson, F. M., 1966, Open Channel Flow, Macmillan, New York, 522 p.Meyer-Peter, E., Favre, H. and Einstein, H.A., 1934, Neuere Versuchsresultate über den

Geschiebetrieb, Schweizerische Bauzeitung, E.T.H., 103(13), Zurich, Switzerland.Meyer-Peter, E. and Müller, R., 1948, Formulas for Bed-Load Transport, Proceedings, 2nd

Congress, International Association of Hydraulic Research, Stockholm: 39-64.Neill, C. R., 1968, A reexamination of the beginning of movement for coarse granular bed

materials, Report INT 68, Hydraulics Research Station, Wallingford, England.Parker, G., 1991, Selective sorting and abrasion of river gravel. II: Applications, Journal of

Hydraulic Engineering, 117(2): 150-171.

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December 11-12, 2005REFERENCESParker, G., Toro-Escobar, C. M., Ramey, M. and S. Beck, 2003, The effect of floodwater

extraction on the morphology of mountain streams, Journal of Hydraulic Engineering, 129(11), 885-895.

Shields, I. A., 1936, Anwendung der ahnlichkeitmechanik und der turbulenzforschung auf die gescheibebewegung, Mitt. Preuss Ver.-Anst., 26, Berlin, Germany.

Vanoni, V.A., 1975, Sedimentation Engineering, ASCE Manuals and Reports on Engineering Practice No. 54, American Society of Civil Engineers (ASCE), New York.

Wong, M., 2003, Does the bedload equation of Meyer-Peter and Müller fit its own data?, Proceedings, 30th Congress, International Association of Hydraulic Research, Thessaloniki, J.F.K. Competition Volume: 73-80.

For more information see Gary Parker’s e-book:1D Morphodynamics of Rivers and Turbidity Currents http://cee.uiuc.edu/people/parkerg/morphodynamics_e-book.htm

hydraulics of mountain rivers gary parker, university of illinois

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