groundwater hydraulics - territory stories: homeorigin of groundwater hydrologic cycle factors...

Technical Report WRD75027

Viewed at 00:02:44 on 18/02/2010 Page 1 of 258.

I I I I I I I I I I I I I I I I I I I I

GROUNDWATER HYDRAULICS

._, -.- -; --.. 'LJ...--.....J.. I ( I'.,., " f v ! -I 'T -- ~\' r_ • ...."

) .,

-:-.~:'-- ...... -.- ' .

. --.- .... _-:_" ".

- ""

LECTL'RES PRESENTED BY C. P. HAZEL OF THE IRRIGATION ANV WATER SUPPLY COMMISSION, Q.UEENSLANV TO THE AUSTRALIAN WATER RESOURCES COUNCIL'S GROUNDWATER SCHOOL, AVE LAWE. AUGUST i975

"



..


TABLE OF CONTENTS

INTRODUCTION

SECTION I : PROPERTIES OF WATER AND WATER BEARING MATERIALS

l.0 INTRODUCTION

1.1 FLUID MECHANICS

1. 1.1 HYDROSTATICS

1.1.2 HYDRODYNA.'!ICS

1.2 SOIL MECAHNICS 1. 2. r GRAIN-VOID RELATIONSEIP 1.2.2 FLUID FLOW PROPERTIES

1. 2.3 SOIL PRESSURES

1. 2. 4 PROPERTIES OF ROCK TYPES

SECTION I I: OCCURRENCE OF GROUNDWATER

2.0 INTRODUCTION

2. I

2.2

2.3

2.4

SECTION

3.0

3.1

3.2

3.3

ORIGIN OF GROUNDWATER

HYDROLOGIC CYCLE

FACTORS AFFECTING THE ABSORPTION OF '1IATER

VERTICAL DISTRIBUTION OF SU3-SURFACE WATER

I I I : AQUIFERS

DEFINITIONS

AQUIFER FUNCTIONS

TYPES OF AQUIFER FORMATIONS

HYDRAULIC PROPERTIES

SECTION IV: GROUNDWATER FLOW

4.0 INTRODUCTION

4.1 DARCY'S LAW

4.2 HYDRAULIC CONDUCTIVITY CK)

4.3 RELATION OF K TO PARTICLE VELOCITY

I

3

3

3

3

6

12

13

13

15

17

21

21

21

22

24

27

27

29

29

29

36

36

36

39

39



I I I I I I I I I I I I I 1 I I I I I I

4.4 RSL.A.TION OF K TO INTRINSIC P:t:RMEABILITY

4.5 REYNOLD'S NUMBER

4.6 RANGE OF VALIDITY OF DARCY'S LAH

4.7 GROUNDHATER FLO'w RATE

4.8 FLOI-I ANALOGIES

4.9 TYPES OF GROUNDWATER FLOH

SECTIO)l v: BORE DISCHARGE TESTS 5.0 INTRODUCTION

5. I DEFINITIONS

5.2 MEASUREMENTS AND MEASURING APPARATUS

TESTING NON FLOWING BORES CONSTANT DISCHARG:t: TLST CONSTANT D~A.WDOWN TEST STEP DP~WDOWN TEST

5.3

5.3.1

5.3.2 5.3.3 5.3.4 5.3.5

5.3.6

5.3.7

5.3.8

5.3.8.1

5.3.8.2

5.3.8.3

STEP DR.".WDO!,iN TEST (EXTEND:t:D rIRST STEP) VARIAB' ,S DISCHARGE !VARIABLE D?~;TNDO'·iN TEST

MULTIPLE AQUIFER TESTING USE OF OBSERVATION BOR:C:S APPLICABILITY OF TESTING PROCEDURES

STOCK AND DOHESTIC BORES IRRIGATION BORES INVESTIGATION BORES

5.3.8.4 TOI-IN WATER SUPPLY AND INDUSTRIAL BORES

5.3.9 PUMP STOPPAGES

5.4 TESTING FLOI-IING BORES

5.4. I ANTECEDENT CONDITIONS

5.4.2 5.4.3

5.4.4

5.4.5 5.4.6

5.4.7

5.4.8

RISK IN CLOSING LOI-I PRESSURE BORES FLOI-I RECESSION TEST (CONSTANT DRAi,DOI-IN)

STATIC TEST (RECOVERY) DYNAMIC (STEP DRAHDOHN) TESTS - GEN:t:RAL

OPENING DYNP~IC TEST CLOSING DYNAMIC TEST ORDER OF TESTS

40

4I

4I

42

42

43

44

44

46

46

48

48

49

50

50

51 ,. _I :"2

52 53

53

53 54

54

55

55

55

55

56

58 58

59

60




SECTION VI: EVALUATION OF A0UIFER PROPERTIES -TEST DATA FROM OBSERVATION BORES

6.0 INTRODUCTIOH

6. I SELECTING THE TYPE OF ANALYSIS

6.2 CONFINED AQUIFER TEST ANALYSIS

6.2. I CONSTANT DISCH.I\RGE TESTS

6.2.1. I STEADY STATE FLOW (THIEl1 EQUATION)

6.2. [.2 NON-STEADY STATE FLOH

6.2.1.3 MODIFIED NON-STEADY FLOW EQUATIONS

6.2.2 VARIP.BLE DISChARGE TESTS

6.2.3 OTHER METHODS

6.3 SEMI-CON:rINED AQUIFER TEST ANALYSIS

5.3.1 GENERAL

6.3.2 CONSTANT DISCH.A.RGE

6.3.2. I STEADY STATE FLOltl

6.3.2.2 NON-STEADY STATE FLOW

6.4 UNCONFINSD AQUIFERS vlTTHOUT DELAYED YISLD

6.4. I CONSTANT DISCHARGE

Ii. 4. i. I GENER"'L

6.4.1.2 STEADY STATE FLOTt! (THIEM DUPUIT HETHOD)

6.4.1.3 NON-STEADY STATE FLOW

6.4.2 JACOB'S CORRECTIONS FOR DRA\,]DOHNS IN THIN UNCONFINED AQUIFERS

6.4.3 LOHMP.N' S SUGGESTIONS

62

62

62

63

63

63

69

80

81

92

92

92

94

94

98

IOI

IOI

101

1- T U,

103

103

106

6.5 IINrONFINED AQUIFERS I




7.2.1.3

7.2.1.4

7.2.2 7.2.2.I

7.2.2.2 7.2.2.3

7.2.2.4

MODIFIED STERNBERG ANALYSIS

REMARKS

f~VALUATION OF NON-LINEAR HEAD LOSSES

DRAWDOWN METHOD

PRESSURE DI"ERENTIAL METHOD

RANGE 0, INTERCEPTS

STEP DRP."DOitJN Pli~!PING TEST

7.3 INTERMITTENT PUMPING TEST ANALYSIS

SECTION VIII: EVALUATION OF AQUIFER PROPERTIES WITHOUT PUMPING TESTS

8.0 INTRODUCTION

8.1 AREAL METHODS 8.1.1 NUMERICAL ANALYSIS

8. 1.2 FLOW-NET ANALYSIS

8.2 ESTIMATING TRANSMISSIVITY

8.2. I GENERP.L

8.2.2 SPECIFIC CAPACITY OF BORES

8.2.3 ROUGH METHOD

8.2.4 LOGS OF BORES

8.2.5 LABORATORY ANALYSIS

8.3 ESTIMATING STORAGE COEFFICIENT AND SPECIFIC YIELD

8.3. I CONFINED AQUIFERS

8.3.2 UNCONFINED AQUIFERS

8.3.3 WATER BALANCE

8.3.4 BAROMETRIC E,FICIENCY 8.3.5 TIDAL EFFICIENCY

SECTION IX: CORRECT! ONS AND EFFECTS TO BE ALLOWED FOR WHEN ANALYSING

9.0 INTRODUCTION

9.1 DELAYED YIELD ,ROM STORAGE

9.2 INCREASED DRAT,vDOWN CAUSED BY DEWATERING

9.3 ANOl'lALIES IN DR.A.WDOWN READINGS

9.4 PARTIAL PENETRATION

I30

130

132

I33

135

136

136

155

162

162

I62

162

165

168

168

168 170

171 172

172

172

172

I73

I73

174

176

176

176

I76

177

I78



I I. I I I I I I I I I I I I I I I I I I

9.5

9.6

9.7

ANTECEDENT CONDITIONS

POSSIBLE DEVELOPI1ENT DURING PUHPING

PROXIHITY OF BOUNDARIES

9.7.1 METHOD OF IHAGES

9.7. I . I RECHARGING BOUNDARY

9.7. 1.2 "IMPERMEABLE" BARRIER

9.7. 1.3 DR.,oMDOWN INTERFBENCE FROM DISCl-LA,.RGING BORES

9.7.1.4 APPLICATION OF IHAGE-BORE THEORY

9.S

9.9

WATER TEMPERATURE VARIATIONS IN HOT BORES

OTHER FACTORS TO BE CONSIDERED

SECTI ON X: APPLICATION OF AQUIFER PROPERTIES

10. I ESTIMATION OF LONG TZRM PU"'lPING RATES

10.1. I CONSTANT DISCHARGE PUMPING TEST

10. I. 2

10.2

10.3

10.4

TO.5

IO.6

10.7

VARIABLE DISCHARGE TESTS

DRAINAGE PROBLEHS

VOLUHE STORED IN AN AQUIFER

VOLUME REHOVZD FROM STORAGE

DOWN VALLEY GROUNDHATER FLOH

LEAKAGE

DRAHDOWN INTERFERENCE FROM DISCHARGING

SECTION XI: GROUNDWATER MANAGEMENT

I I. 0 ANNUAL YIELD

II.! CONTROL OF GROUNDWATER USE

BORES

!1.2 CONJUNCTIVE USE OF SURFACE AND GROUNDHATER

11.3 GROUNDWATER RECHARGE

11.3. I NECESSITY FOR REC}-,.A.RGE

11.3.2 NATURAL RECHARGE

11.3.3 ARTIFICIAL RECl-LA,.RGE

rT.4 SEA WATER INTRUSION IN COASTAL AQUIFZRS

I!. 4. I GENERAL

1[,4.2 GHYBEN-HERZBERG CONCEPT

179

ISO

ISO

lSI

lSI

lS4

184

187

lS9

189

190

I90

190

194

194

195

195

196

197

I9S

200

200

203

203

204

204

205

206

212

2I2

213



,. I I I I J I I I I I I I I I I I I I I

IT.4.2. I

IT.4.2.2

11.4.3

11.4.3.1

Il.4.4

11.4.4.1

Ir.4.4.2

11.4.5

1l.4.6

GENERAL DERIVATION

THE DYNAMIC CONCE?T DERIVATION

LOCATION OF THE INTERFACE

CONFINED AQUIFER UNCONFINED AQUIFER

STRUCTURE OF THE INTERFACE CONTROL OF INTRUSION

GLOSSARY OF SYMBOLS USED

REFERENCES

COMMO~ METRIC UNITS

~\ETRIC i"iULTIPLES AND SUB-MULTIPLES

RECOMMENDED METRIC UNITS

METRIC UNITS CONVERSION

INDEX

213

213

215 215

217

217

219

220

221

222

225

232

234

235

238

239

.'




I.

INTROVUCTrON

We have become aCC1.U>;tomed ;to ;the be-Uen ;tha;t ;the mO.l;t common .I oUltce 06 Menu£. 6Jtuh wa;tVt on auA I~ane;t -



2. 10 b.u N9 out U-6 en.:t..teMa;tional. IYLlJ.ti.:tute 60tt Land Redama;tioN and Implt(1V(?me'Lt, WageMNgen, The Nethv-veando, 1970.

In pl'U!.pM-




3.

SECfICN I

PFtf£RTIES OF WATER A'ID WATER BfAAING fi1ATERI~LS

1.0 INTRODUCTION

The ~cience 06 GlWundwa.tvt Hydr..a.Ll.li~ .if, bcw ed u.pon .:the nWl.da-men.:tal pltop erct



4.

The unit of density is kilogram/cubic metre (kg/m3). The density of ~Iater at 4°C is 1000 kg/ffi3.

Densities at other temperatures from O°C to IOOoe are given in figure 1-7.

[; eci 'ic Gravit (s. G.) 01" Relative is the its denslty to t e denslty 0

(R.D.) of a sUbstance

I I I I

Being a ratio, it is dimensionless. I Diupluc,,"ment (s) is a change in position in a specified direction.

It is then a vector quantity. II The unit of displacement is the metre (m).

Velocity (v) is the quantitative description of the motion of a body. Since it is related to direction as well as speed of motion, it is a vector quantity. It;s a rate of change of position in a specified direction. ds

Instantaneous velocity = dt.

The units of velocity are metres per second (m/sec), or metres per day (m/day).

jl"",eZeration (a) is the time rate of change of velocity, and is a vector yuantity.

- t 1· dv Ins antaneous acce eratl0n = dt

file units of iOcc21eration are metres per second per second (m/sec 2).

J1cceZerati01': due to cravitu (g) is the acceleration produced 011 a body by the earth's gravitational field and for most purposes is assumed to be constant at 9.80 m/sec2. It does, however, vary from place to place on the earth's surface.

Force (F) is that which produces or tends to produce a change in the state of w~tion of a body.

By Newton's second law of motion:-

F = Ma

The force exerted on a body by the earth's gravitational attraction is referred to as its weight (w):-

W E Mg ------1.2

The weight of a unit volume is referred to as the specific weight (y):-

y = pg

The adopted unit of force is the newton, defined as that force which when acting on a mass of I kg will produce an acceleration on it of I m/sec2 •

I I I I I I I I I I I I I I



I I I I I

• I I I I I I I I I I I I I I

5.

A dyne is defined as the force required to move a mass of I gm with an acceleration of I cm/sec2.

The weight of a body having a mass of I kilogram is then 9.80 newtons (N).

Pressure (p) is defined as the perpendicular force per unit area exerted on a surface with which a fluid is in contact.

The surface with which the fluid is in contact may, of course, be either a solid boundary or an imaginary plane passed through the fluid for purposes of analysis .

F P ; A ------1.3

The adopted unit for pressure is t~e pascal (Pa), which IS defined as I newton per square metre (I N/m ). The co~ unit will be the kilopascal (kPa) which is 1000 pascals.

Strictly speaking, the magnitude of pressure should be expressed in terms of pascals above absolute zero. However, it is generally more convenient to use atmospheric pressure as a reference, the relative intensity p then rejJresenting the difference between abSOlute intensity Pabs and atmospheric intensity Pat·

i.e. P ; Pabs - Pat ------1.4

Under norm~l conditions, Pat; lor kilopascals (kPa). ; IDIO millibars (rrb) in meteorology.

Pressure wZj also be expressed in terms of the number of metres of water (or mercury) that a certain pressure would support. Hydraulic head is measured in rretres.

The pressure at a point A in a fluid h metres below the surface is given by:-

where

p ; pgh ; a head of h ~~tres of water

p ; pressure in pascals p = density in kg/m 3

h = depth in metres

g ; acceleration due to gravity; 9.8 m/sec 2.

Pat is then equal to a head of 10.3 metres of water.

------ 1.5

Pressure is the sarre at all points in the same horizontal plane within the fluid at rest. Any increment of pressure applied at any point in a confined fluid is at once transmitted equally to all parts of the fl ui d.



6. 1. 1. 2 HYDRODYNAMICS

Hydrodynamics is the study of fluids in motion.

This section deals with the motion of fluids and the various tvpes' of flow which may be encountered.

Steady FZow

Consider a fluid in motion in Fig. 1.1. particle of fluid at a given point, a, will have VI' at b a velocity vZ' and at c at velocity v3 .

At a given time a a particular velocity If, at all other times,

the velocity of whatev~r particle of fluid is at the point a, remains constant at Vi' that at b remains Vz and that at cremains v3 , then the flow is said to be steady.

The path followed by a fluid particle in steady flow is called a streamline. Streamlines are characterised by the property that the tangen4 at any point on the streamline gives the direction of flow of the f'uid at that pOint. Particles of fluid may not flow from one streamline to another.

A bundle of similar streamlines is called a tube of flow e.g. steady flow in a pipe.

Non-Steady FZow

When the velocity of the particular fluid particle at a given point in a moving fluid varies with time, the flow is described as non-steady.

Continuity Prinaip~e

Consider steady flow in a pipe of varying cross-sectional area -as in Fig. 1.2. Let the velocity of flow normal to plane A be v i and at B be vZ" Let the cross-sectional area of lhe pipe at A be Al , and at B be Az• ~

STEADY FLC/r'I Fig. I. I

r-~

r- ~i- - ---_ ..... ---- ---r-+ ----- v Z __ -- _ --_ .. ----- -----

A B

CCl'lTINUI1Y PRINCIPLE Fig. 1.2

Since the fluid is assumed to be incompressible the mass of fluid passing A in a given time t must equal the mass passing B in the same time, or there would be an accumUlation of mass between A and B, i.e.

Mass passing A in time t = Mass passing B in time t Vi Al t = Vz Az t


j;; JIb ZlEJIhRlWt




or ----.-- I. 6

Thus in the case of steady flow in a pipe of varying cross-section, the highest velocity occurs at the smallest section.

Since the fluid undergoes an acceleration in moving from A to B, from the second law of motion (F = M a) the pressure at A must be higher than the pressure at B.

This principle is used in the venturi \·Ihich is used to measure flows in pipes. Small tubes tapped into a constricted pipe at positions such as A and B enable the difference in pressure to be recorded. This pressure difference is proportional to the velocity squared, so the meter can be calibrated to read flow directly.

Bern~~ZZi'S Theorem

By the application of the principle of conservation of energy to the flo'li of a fl ui din a tube of flow, Bernoul1 i 's Theorem can be derived, viz

...E... p 9

+ v2

2g + h = constant ------1.7

or is someti rres wri tten

where p = p = 9 = v = h = h" = I

=

pressure

Pz PS

2 Vz + ...--- + 29

density of fiuid acceleration due to gravity velocity

------ 1.8

potential or elevation head above datum level head lost in overcoming friction when the fluid moves from the point I to point 2.

It must be emphasised that Bernoulli's Theorem is strictly applicable only to streamline flow.

In equation I.7 -

..L = pg the pres sure head

v 2 2g = the vel acity head

h = the elevation head

The total head at any pOint is the summation of these three terms.

The summation of tne elevation head and pressure head gives the potentiometric head, and the concept is useful in analysing flows in pipes and flow in underground water.



8. A popular misconception is that flow takes place from areas

of high pressure to areas of low pressure. This is not so. Flow occurs from areas of high potentiometric head to areas of low potentiometric head. With a suitable arrangement of elevation heads if. is possible for a fluid to flow from an area of low pressure to onp. of high pressure. e.g. the pressure in a storage tank located 011 il hill is smaller than the pressure in the pipeline some distance down the hill. However, because of the greater elevation head on the hill, the potentiometric head is greater at the tank site and water i~ delivered along the pipe.

Viseosity

So far, only "ideal fluids" have been considered i.e. fluids which cannot transmit shearing stresses and on which no work is done in changing their shape.

In actual fact, no fluid is ideal and all possess to some degree the property of viscosity or internal friction.

The coefficient of dynamic viscosity or absolute viscosity is defined as the ratio of the intensity of shear, T, to the rate of deformation

where

i.e. n: dv/dy

T)

T = dv/dy =

dynamic viscosity intensity of shear velocity gradient in the transverse direction

------I. 9

If we consider two plates, in Fig. 1.3, of surface area A separated by a thickness y of fluid and moving at a velocity v relative to each other under an imposed force F,

v F

"

Fluid

Fig. 1.3

then n = ~ • ~ ------1. TO

where total tangential force app~ied, i~ N.. 2 F =

A = Area over which the force 1S appl1ed, 1n m . (Nsrr-2) n = coefficient of dynamic viscosity in decapoises v = relative velocity. m/sec

distance between layers at which velocity is measured in y =

The coefficient of dynamic viscosity depends on the fluid and its tempe ra tu re.

m.

: lil AALiAtIM


mag.




In ten11S of uni ts -

= force newtons) m) distance velocity m/sec) area

" 2 n 11i 11 have the demensions of ne~/ton second/m and the unit is

knovm as the decapoi se.

Another viscosity coefficient, the kinematic coefficient of viscosity is defined by:-

Kinematic coefficient, v

La7rrinar ar..d T'vf.rbuZent Fw..;.)

= absolute viscosity (n) ------!. Ii mass densi ty (p)

In laminar fim", the fluid particles move along parallel paths in layers or iaminae. The magnitudes of the velocities of adjacent laminae are not necessarily the same. Laminar flow is sovemed by the equation given as the definition of dynamic viscosity above, i.e.

------1.12

The viscosity of the fluid is dominant and thus suppresses any ten cency tC'ila rds tu rbu 1 ence.

/\bove a certain cri~ical veloCity, the viSCOSIty of the fluid is insufficient to damp out ;;ul'oulence and the f,uid particles !r.()V2 :~ a haphazard fashion where it is impossible to trace the motion of an individual particle. Such motion is called turbulent flow.

The shear stress for turbulent flow can be expressed as -

T = (ll + z) dv dy ------ L 13

where z is a ractor depending on the density of the fluid and the fluid motion,

In pipeflow the transition from laminar to turbulEnt flo~>/ is characterised by ~/ell knol'm values of Reynold's Number (N R) which expresses the ratio of the inertial to viscous forces. Thus, there is a lower limit critical number around 2100 below which flo~1 in pipes is alwilYs laminar.

~y analogy, in flow through porous media a Reynold's Number nas bf'{'ll establ ished as

where

N = R

vD \)

------1.14

v = the SpeClT1C discharge i.e. discharge per unit area D = a characteristic length. In pipe flow 0 is the internal

diameter of the pipe. In flow through porous media 0 is related to grain size.



IO.

v = the kinematic viscosity of the fluid.

Because the grain sizes are so variable in flow through porous 1I,,'elia no one value for Reynold's Number can be set as the dividing line itf"'.Wf'fm laminar flow and turbulent flow. It has been established that this transition occurs normally for a Reynold's Number in the range I to 10.

where

For non-circular cross-sections - (in open channel flow):-

N = v(4R) R v

R = the hydraulic radius, and is equal to the ratio of the cross-sectional area to the wetted perimeter.

Surface Tension (cr)

Surface Tension is a property associated with the free surface of any liquid or the interface between any two non-miscible liquids.

It is well known that many insects are able to walk on the surfnce of liquids in apparent contradiction to Archiw~des' Principle. This ~roperty tends to suggest that there is a kind of membrane or skin that envelopes all liquids. In fact this very nearly describes what is actually the situation.

A molecule in the interior of a fluid is acted upon by atiractive forces in all directions and the vector sum of these is zero. f:lJ'.'cver, at the surface, a molecule is acted upon by a nett inward col:es i ve force perpendi cul aI' to the surface. Hence work is requi red to bring molecules to the surface.

The surface tension of a liquid is the work that must be done to bring enough molecules from inside the liquid to form one new unit area on the surface.

The intermolecular forces also come into play when a liquid is in contact with a solid object - in this case there is an attraction of the molecules of the solid for those of the liquid.

Hence. for example. if pure water is placed in a clean glass container it will be observed that the water surface in contact with the glass turns up and lies flat on the glass. Thus the attraction of glass for water is greater than that of water for itself.

On the other hand mercury placed in a glass container will be observed at the surface contact. to be drawn away from the glass indicating that the attraction of glass for mercury is less than that of mercury for itself.

Capillarity or the rise or fall of a liquid in a capillary tube is caused by surface tension and depends on the relative magnitudes of the cohesion of the liquid and the adhesion of the liquid to the walls of the containing vessel. Liquids rise in tubes they wet (adhesion> cohesion) and fall in tubes they do not wet (cohesion> adhesion).

I I I I I

• I I I I I I I I I I I I I I




where

whence

where

II.

\-2r~ / ~.s---

I

p - atmDspheric

T------ - --- p < atmospheric Density p, surface tension

From Fig. 1.4 -

CAP I L!..ARY RISE Fig. 1.4

Weight of the column of water = Force exerted by Surface Tension 2

1T r pghc

= 21T r rJ cos a ------1. IS

= 2 rJ cos a rpg

r = radius of tube (metres) p = density of fluid (kg/m 3) 2 g = gravitational acceleration (9.8m/sec )

surface tension (newtons/metre) capillary rise (m)

------ I. 16

a = angle of contact between solid liquid and gas.

For pure water in clean glass

a = 0 ar.d cos a = I

At 20°C,

cr = 0.073 N/m p " iOOO kg/m3

h = c

1.5 r

_5 x 10 rn

he = capillary rise in metres r = radius of tube in metres

------1. 17

Surface Tension is dependent on temperature; so then is capi1l arity.

At ggOC, for pure water in a clean glass,

cr = 0.0591 N/m



12.

p = 959 kg/m 3 h = 1.26 x IO- 5m c r

Capillary rise in granular material is comparable to a bundle of capillary tubes of various diameters,'· see Fig. 1.5.

. - - . ::==-~:---... -

I~( c=.- ;; '::-;-; ,. .

~_=--=~~_~~~_,~-:~ f3~----~==~--~ =' ~ _:::::::::. -. -::. --==:':--:::':7 - ~:...'. _ "

. - -~. - ----=- ~.~ :: -_.-._"- -:. ~; ~...,.,::- . --=..-~ - ~ .::,

CAPILLARY RISE IN TUBES Fig. l.5

Capillary rises in samples having essentially the same porosity (4I:r,j after 72 days (Atterberg, cited in Terzaghi, 1942) are given in Table I-I. (Note that h is nearly inversely proportional to grain size). c

TABLE I-I

CAPILLARY RISES IN GRANULAR MATERIAL

_. Material Grain Size h

( 1111l) c (em) --

Fine Gravel 5-2 2.5 Very coarse Sand 2-1 6.5 Coarse Sand 1-0.5 13.5 Medium Sand 0.5-0.2 24.6 Fine Sand 0.2-0.! 42.8 Silt 0.1-0.05 105.5 Silt 0.05-0.02 *200

* still rising after 72 days

1.2 SO I L i'EQ-IftN I CS

This study of Soil Mechanics will be limited to those soil properties relevant to groundwater hydraulics. In general then, the study will be limited to porous media. Porous media are comprised of two distinct parts, a granular matrix and interconnected voids. In a saturated porous medium, water (or some fluid) fills all of the voids or pore spaces.

-


------------------------




I3.

1. 2. I GRAIN-VOID FElATICNSHIP

Porosity (e) of the soil mass is defined as the ratio of volume of the voids to the total volume of the mass.

where

Vv e = VT

Vv = volume of voids

v = total volume T

------1.18

Primary porosity is related to granular material, while secondary porosity refers to the opening in joints and faults in hard rocks, and solution openings in iimestone, dolomite, gypsum or other soluble rocks.

Poros Hy is commonly expressed as a percentage and has no un; ts.

The porosity of a soii obviously depends on the properties of its constituent grains, some of which are enumerated below.

(il

( i i )

( iii)

Shape of the grains. vo j ume of voi ds, the grains the lower the increas2 porosity~

Since porosity is a function of the more intimate the contact between porosity. Angularity tends to

Size of arains. Provided grain size is uniform, the actual size will have no effect on the porosity.

Degree of assortment. A ~ii de result in a smaller volume of porosity. On the other hand, produce a higher porosity.

range of grain sizes will voids and hence a lOvier uniform sized grains will

(iv) Type of packing or arrangement of grains. Consider the idealized case of uniform sized spherical grains. For square packing a porosity of 4i.64% is achieved, while for rhombic packing the porosity will be only 25.95%.

In the same ~Iay, for random size and shape of grains, the porosity is controlled by the packing c.f. maximum and minimum density.

Voids Ratio (e) is defined as the ratio of the volume of the voids to the volume of solids.

'I v

V-------1.19 e =

s

1. 2.2 fLUID flDW PROPERITES

Coe[tioient of Permeabi Uty or Hydraul.ia Cor.duativit,! (K)

The law governing laminar water flow through soils is Darcy's La., "nd may be expressed as;- (See Section 4).



where

and

14.

Q = -KiA ------ r. 20

Q = i =

A = K =

rate of flow (in cubic metres per day) hydraulic gradient or head loss per unit distance travelled (non dimensional) the cross-sectional area through which the flow occurs

(in square metres) the coeffi ci ent of permeabi 1 ity or hydraul i c conduc;:i vity

(i n metres/day)

The coefficient of permeability should not be confused ~/ith intrinsic permeability (See Section 4).

Darcy's Law may also be written -

~ = v = -Ki ------1.2I

However careful distinction must be made between the superficial velocity v and the actual seepage velocity v ~/here

s

Q =

~/here

Q = A = v =

A = v

v = s

whence

where

Av = A v v s

total discharge rate cross-sectional a rea of porous average discharge velocity cross-sectional area of voids actual seepage velocity

v = ev s

e = porosity

------1.22

medium

------1.23

Hydraulic Conductivity may be determined in the laboratory by the use of permeameters, or in the field by pump testing. Because of the problems associated with obtaining undisturbed samples and of re~acking disturbed samples in the laboratory, coefficients of permeability obtained from laboratory tests must be considered unreliable.

in practice, most aquifers are non-homogeneous and anisotropic (i.e. the material does not have like properties on all orientations of planes, generally resulting from stratification), and the hydraulic conductivity will vary with location and direction of flow. The aquifer characteristics obtained from pumping tests represent the average of values around the discharging bore.

Hydraulic conductivity and porosity are essentially unrelated e.g. clay generally has a high porosity and low hydraulic conductivity while sand has a low porosity but high hydraulic conductivity.

Intrinsic Per-meabiZity (kJ

The Velocity of laminar flow through a porous medium may be described almost exactly by the equation -


.... _---------------------------




IS.

v = k. ~ 1'1 ------1. 24

where ip = the pressure gradient '1 = the dynami c vi sees i ty of the fi ui d k = the intrinsic permeability (square micrometre)

This may be written

where i = hydraulic gradient

v = Kpg; r,

------ 1. 25

This equation is of the sa~~ form as the Darcy equation -

v = Ki

Henc'2

The intrinsic peT'ffieability (k) is related solely of tne porous medium. The coefficient of pe~eobility or conductivi1:Y (K) is reliited not only to the properties of medium but also to the properties of the fluid.

to the properties r"""'r~~'l ;c .1 ........ ,...!".l • the porous

From extensive ~aboratory testing, Hazen found that the vlefficient of permeabiiity of saftds in a loose state depended on ~j;o quantities he called the effective 9rain size and the uniformity coe ffi ci ent.

The effective grain size 010 of a sample is a grain size

diameter such that 10 percent of the particles are finer and 90 percent coarser.

a sample

1. 2.3

o is defined as that diameter such that n% of the particles in n

are finer.

The uniformity coefficient U is defined as 0

60 U = D10

SOIL PRESSliRES

Pore Water Pressure

------1. 27

If the pores or interstices of a porous medium are fillec with water then this water is subject to the same principles as outlined in section I.i.I, hydrostatics.

The hydrostatic pressure in the water is given by

P = pgh w

------1.28



where 16.

Pw = hydrostatic or pore water pressure

p g h

= = =

density gravitational acceleration depth below the potentiometric level

IntergranuZar Pressure

If a load is applied to an unsaturated porous medium it is transmitted from grain to gr~in at ~:le p0ints of contact. The pressure at the points of contact, i.e. the intergranular pressure, is dependent on the force applied and the area of contact. The application of a force results in a larger area of intergranular contact with a resultant slight deformation of the matrix and a reduction in voids ratio.

TotaZ Pressure

If the porous medium ;s saturated and a confining layer placed over its surface and a load ;s applied, see Fig. 1.6, then the IOnd is taken partly by the pore water and partly by the grains.

PRESSURE DISTRIBUTION Fig. 1.6

,/ Confining Membrane

Saturated Porous Medium

If the total pressure applied to the porous medium, i.e. force/area, is Pt then at the interface of the porous medium and confining layer

p = t Pw + Pg ------1.29

where Pt = total pressure

Pw = that part of the pressure borne by water

Pg = that part of the pressure borne by gra ins

When the applied load is in fact the material overlying an aquifer, then changes in applied load such as atmospheriC pressure changes or tidal changes will result in changes in pore water pressure with resultant changes in potentiometric level.

Likewise, a lowering of the potentiometric lev~l by pumping results in a lowering of the pore water pressure and a resultant





I7.

increase in load to be carried by the grains and a slight compression of the aqui fer matrix. An understanding of this transfer of pressures is necessary if the concept of storage coefficient is to be understood.

1.2.4 PROPL'

Technical R

eport WR

D75027

View

ed at 00:02:44 on 18/02/2010P

age 25 of 258.

TABLE I-II

SUIfJAAY OF THE ARJTH~;:T1C HEAN OF PROPERTIES FOR ALL ROCK TYPES

Property

Strata Coefficient of Dry S.G. Poros ity Poros ity Specific Specific Penreabl11ty Unit of Un- Repack Retention Yield Weight Solids di s turbed (X) (X) (X) Re- Wert. Hariz. (g/ee) (X)

pack m/day) (m/day) (m! day)

SEDII£NTARY Water- Sandstone Flne 0.2 0.29 I. 76 2.65 33 13 21 IlOCKS Laid ~dium 3.1 1.68 2.66 37 10 27 Dep -osits SI1 tstone N I. 61 2.65 35 43 29 12

Claystone N I. 51 2.66 43 ~

CD

Shale 2.53 2.73 6 Cl ay N N I. 49 2.67 42 48 38 6 Silt .0025 .08 I. 38 2.66 46 46 28 20 Sand ~e 2.5 3.B 1.55 2.67 43 32 8 33 ium 12 14 1.69 2.66 39 35 4 32 Coarse 45 28 I. 73 2.65 39 34 5 30 Gravel Fine 450 1.76 2.68 34 7 29 ~edium 273 1.86 2.71 32 7 24 Coarse [50 I. 93 2.69 28 9 2[ -Wind Loess .08 1.45 2.67 49 46 27 18 Laid Aeo I ian Sand 20 1.58 2.66 45 38 3 38 Dep- Tuff O. [6 1.48 2.50 4[ 2[ 2[ osits

i . .L._ . _____ . ----

'. --------------------

Technical R

eport WR

D75027

View

ed at 00:02:44 on 18/02/2010P

age 26 of 258.

- - - - - - - - - - - - - - - - - - - -'.

TABLE 1-:. (Cont.,l I , ,

Property - ,-

Strata Coefficient of Dry S.G. Porosi ty POtos i ty Speci fic Specific Penreabil ity Unit of Un- Repack Retention Yield

---- Weight Soli ds disturbed (%) (q (%) Re- Vert. lIod z. (g/cc) (% ) pack (m/day) (m/day) (lnl day)

- ---r------ ---- - ---SEDIMENTARY Ice- Till Clay 2.65 ROCKS Laid S11 t I.lS 2.70 34 2S 6 (Contd.) Depos its Sand 0.5 I 1.88 2.69 31 14 16

,

Grave 1 30 l. 91 2.72 26 12 16 ~ cD

Washed 5i It .2 I. 38 2.72 49 9 40 -Drift Sand 38 14 l. 55 2.69 44 36 3 41

Gravel \>04 1.60 2.68 39 41 -- --- ,-~~- -- ------- -

Chemical Li me-stone I 1.8 I. 94 2.75 30 13 14

and Dolomite 2.02 2.69 26 Organic Peat 5.7 0.13 I. 54 92 49 44 Depos its

. -- .-~---~-. -.----- .. -- --IGNEOUS fleathered ~ranite 1.4 I. 50 2.74 45 ROCKS Weathered Gabbro O. [6 I. 73 3.02 43

Basa It 0.008 2.53 3.07 I7 -~-. ._,--

METAI'IlRPHIC Schist 0.16 I. 76 2.79 38 17 26 ROCKS STate N 2.94

--.- ----~-.--- -- .... _- .----- ------- --------NOTE: N indicates Negligible



20, Fig, 1.7

VnpOlH Pressure

(mrn of Hg) I;::OU 1000 800 600 400 20 [! 0

Surface fp.nSlon A 5~IO-2 80"0-2 75"0-2 70,,0-2 65,,0-2 60"0-2 55,,0.-2 (N/m)

Dynnmlc VI10-3 05:




2.0 INTRODUCTION

2I.

SECTION I I OCCURRENCE OF GROUNDWATER

Ma:tu,taX undeJtg!r.OWtd ,~u etYlty, :they h.ave ene,moUl.> eapacU;[u and do not beeome uogged INWt \ ift and weed!.> a~ do iak.u and ItU eJtvoJ.]u,. They ate Iteidi..veiy ,[nex.peJ1l.>-i V(' to tap; :they tose Ut:tie. oJt no wa.t:vr. by evaporca;t.(.on; ::they ean wp)Jiy wa..:teJt oveJt '.lefty ta"cge Meal w.U:hou:t .:the neeu;.,ay 06 br.UJ.a.tng eilaHnW, pJ..pet.tnel.> Oft otheJt oo;(:/ubtLaon 6y,oteJM; and,'[6 pJtOperJ'.y ma>1aged, :theDr.. pe/uod 06 Ul.>e6r..r..inu6 hal no noltueeabte UmU.

GJtOUYIdwateft oeeu,t6 J..n the POItU Olt J..,,;telt6UeC6 o~ Jtoek.6. In ~rmt-c.on6J..ned aqr..r..tnerc6 [6ecUon 3) taltge volumc; 06 wator.. may be 6toJted in tite. .6VM-peJtV.toUl.> ta.ye1t6 above and/otr. beiow .:the maJ..n aqr..r..t6e/L Wah a Itrdudton .



22.

Connate Water is w,'lter trapped in the interstices of a sedimentary rock. at the time it was deposited. It may, for example, have been derived from the ocean or from fresh water sources, dependi~q on the locality in which the sedimentary rock was formed. Although of little importance from

" the point of view of significant quantities of groundwater being obtained from this source, it is nevert,leless most important in its effect on water quality in various rocks.

Meteoric Water is water derived from the atmosphere, generally in the form of rain or sometimes snow and hail. It is the basic source from whi

Technical R

eport WR

D75027

View

ed at 00:02:44 on 18/02/2010P

age 30 of 258.

--------------------

(

ABSORPTION

-; .. ",- . :---.'- ' . . .

. -

,"

..,~ --"-EYAf'('.fUnON QURlNG RAH"ALL

RAINFALL '---..../

{J ~ / _o'f"c'..\:

5O.l EVAPORATION ....... ,';

,

~ ((

E",PORATION lIT7 A DURING RAINFALL V

~-~-,

~ -. ,

~" "-- ...... ,.

RAINFALL ON SEA

~~ -"""--'--...

n~'~ RUNOF~

SEA

( ,

~-'

EVAPORATION

Ie cyrLE

N W

" ~. cCO

r-)

-,



24.

TyI'!' _nf.. SUr'[aee Material

Sandy soils will absorb rainfall but clays prevent easy access for the water.

'{'!fPC of Vegetation

Plants can greatly reduce water entering the aquifers by trans-plrlng and evaporating it, through their leaves. On the other hand, vegetation can increase infiltration by reducing surface runoff, improving soil structure, protecting the soil from compaction and by providing organic matter - holes.

CZmate

A cool climate with frequent rain is more favourable for under-ground water replenishment than a hot climate and infrequent rain.

2.4 VERTICAL DISTRIBUTION OF SUB SURFACE WATER

"ub SUr'faee Water

Water occurring beneath the land surface is called sub surface water. It is in two parts. The lower part within the "Zones of Saturat-ion" is usually referred to as groundwater. Above this, in the "Zone of Aeratior.", water is held between the particles of soil etc. This is kno~m as "Suspended Water".

7q~e of Aeration

As 11ater enters the soil it moves downward due to gravity. However, some of the water is held on the grains and between them by surface tension. This is suspended water, and is found in the "Zone of Aeration" - refer to Fig. 2.2. This zone extends from the surface to the actual water-table. It includes an upper band containing soil water, an intermediate zone and finally, just above the aquifer, a capillary zone. In the soil moisture belt water is evaporated from the soil as well as being used by the plants.

In the intermediate belt the water is held by molecular attraction and little or no movement occurs except when recharge occurs.

The thickness of the capillary zone depends on the nature of the material overlying the water bed. It can vary from a few centimetres to a few metres. The finer the material the greater the thickness. This zone is recognised by drillers as a sign of the proximity of water.

The study of the flow of water in this unsaturated zone is becoming more widely known as Soil Water Hydrology and will not be touched in any detail in these notes. Soil Water Hydrology can be very important when considering groundwater recharge.

~g~e of Saturation

In this zone all the openings in the rocks are completely filled with water. It is from this zone that supplies of groundwater are obtained. Refer to Fig. 2.2.





25.

In non-pressure called the water table. and groundwater flow.

or unconfined aquifers, the top of the zone, is The water table will fluctuate with recharge, use

In confined aquifers the water is prevented from rising by a confining layer. If a number of bores were drilled through this confining layer the heights to which the water would rise would represent the potentiometric surface of the aquifer.

For the most part the presence of groundwater is continuous in the zone of saturation. However, its availability depends upon the nature of the rock formation in which it occurs. For example, clay may be saturated but wi 11 not release the water to a bore or wel1. On the other hand a coarse saturated gravel would yield large quantities.

In some cases the rate at which the saturated material will yield water is so slow that it is not immediately obvious in a well or bore and drillers are inclined to say that the hole is dry.

This can explain the situation where a bore is drilled a few feet from a so called dry hole and yields a substantial quantity of water. In actual fact in both cases the material may be saturated but only in the latter case will it yield I,ater at a rate sufficient to give a successful bore.

Zone of Aeration

r'

,

Zone of Saturation <

t t , ","'N , t ,

SOIL " - . , ,

,

, . . WATER

,

~ , •. ij'k-Y// f'

,

, , ,

,

'1 Soil Water ~ j Zone

\ }

Intermediate

I Zone

~' -'------------~.----~ }

caPillary

, . CAPILLARY WATER Zone _ . - . - -. - - ,. . , ,

/ . . , GROU~DWA!ER. - , . -. - , . . . / . ~

Bedrock

VERTICAL DISTRIBLrrION OF SUB SURFACE WATER Fig. 2.2

Suspended Hater

The schematic representation in Table 2-1 gives the compiete range of divisions. These will all be present in areas of relatively deep water tables after prolonged dry spells. In other areas with higher water tables the uppermost divisions may not be present. Beneath lakes, swamps and streams only the zone of saturation will be present.



Pressure

Gas phase, equals atmospheric

Liquid phase, less than atmospheric

Less than atmospheric

Atmospheric Greater than atmospheric

26.

TABLE 2-1

DISTRIBUTION OF SUBSURFACE WATER

Zone Division(l)

Unsaturated Discontinuous Zone capillary saturation

Semi-continuous capillary saturation

Saturated(2) Continuous Zone capillary saturation

water table Unconfined groundwater

NOTES: (1) Cap.Uealty "ZOYLe6" 06 TeJtzaglU (1942 J (ZJ A6 lte.de6-i.n.e.d bq Hubbe;r;t (1940). See. a.t60

Lohman (1965, p.92).

Bore

S.W.L._





3.0 DEFINITIONS

Aauifer

27.

SECTION I I I

AQUIFERS

Lohman and others 1972, have defined an aquifer as a formation, group of formations or part of a formation that contains sufficier.t saturated perrr~able material to yield significant quantities of water to bores and springs. Under this definition, an aquifer includes both the saturated and unsaturated part of the permeable formation.

Con fined (or Pressure or Artesicrn) Aauifer

A confined aquifer is a comp1etly saturated permeable formation of Vlhich the upper and lower boundaries are impervious layers. Completely impervious layers rarely exist in nature and hence confined ilquifers are less corrmon than is often realised.

In a confined aquifer the water is under sufficient pressure to cause it to rise above the aquifer if given the oppc'-tunity f.g. if penetrated by a bore. The level to wh~ch the water rises is called the potention~tric level, the standing water level (S.W.L.) or static head. Pressure water or confined water relates to water contained in aquifers of this nature.

If the pressure is sufficient to raise the Vlater to the surface when a bore penetrates such an aquifer then the \-Iater may f1c~1 from thE bore without mechanical assistance. This type of bore is called a "flowing artesian bore".

Unconfined (Non Pressure 01' Water TciJZe) Aou.ifer

An unconfined aquifer is a permeable formation only partly filled with water and overlying a relatively impervious layer.

It contains ~Jater which is not subjected to any pressure other than its own weight. If a bore penetrates such an aquifer the water wi 11 rise up the bore no higher than the depth at ~Ihich it was first encountered.

The level at which water stands in a bore penetrating an unconfined aquifer, i.e. the standing water level, is known as the water table, and is the depth at which the water in the unconfined aquifer is at at~~spheric pressure. Water does occur in the aquifer above the water table but at pressures less than atmospheric.

A perched aquifer is an unconfined aquifer separated from an underlying body of groundwater by an unsaturated zone. its vlater table is a perched water table. It is supported by a perching bed whose permeability is low. Perched groundwater may be either permanent or temporary.

Hater in an unconfined aquifer is called unconfined or phreatic water.



28.

Semi-Confined (or leaky) Aquifer

The confining layers of many pressure aquifers are not completely impervious. The hydraulic conductivity of the confining layer may be very small when compared with that of the aquifer material, but as the radius of influence of a discharging facility increases, the area through which the confining layer is contributing water becomes very large and the volume of water contributed can be a very significant part of the total water discharged.

Such an aquifer is called a leaky or semi-confined aquifer. The flow of water from the confining layer to the aquifer is assumed to be vertical. The horizontal movement in this layer is negligible.

Semi-Unconfined Aauifer

If the hydraulic conductivity of the fine grained layer in a semi-confined aquifer is so great that the horizontal flow component in the covering layer cannot be ignored, then such an aquifer is intermediate between the semi-confined aquifer and the unconfined aquifer and may be called a semi-unconfined aquifer.

In general, such aquifers do not release their water instantaneously from storage and exhibit what is called delayed drainage. Such aquifers may be called unconfined aquifers exhibiting delayed drainage or delayed yield effects.

Schematic representation of the aquifer types is given in Fig. 3.I.

.

Potentiomet

~~~~1-__ 1Level __ -E2:~~~~~~~ « /['0_ Saturated

CONFINED

: . . : FINE GRAINED .. '1 . ~ '" • · . '.J




29.

An aquifuge is an impermeable formation ~Ihich neither contains nor tran$niits water. An example of an aquifuge is solid granite.

Cunt~ninq Bed.

U.S.G.S. Water Supply Paper 1988 suggests that the terlJl "ccnfining bed" should supplant the terms "aquiclude","aquitard" and "aquifuge". It is a body of impermeable material stratigraphically adjacent to one or mOl'e aqui fers.

3,! AQUIFER FUNCTIONS

An aqui fer has three important funct ions namely:-

Storace

It stores water - as a reservoir. The aquifer characteristic describing the ability of an aquifer to store water is the Porosity. The characteristic describing its abil ity to release water under gravity drainage is its Specific Yield. The property describing its elastic storage is the Storage Coefficient or Storativity, and this is related to the elastic properties of both the water and aquifer materiai.

,.. .. IT'OYi.S/111..SS'Z-on

It transmits water like a pipeline. The relevant aquifer characteristic is Transmissivity (called Transmissibility in some references l.

Mi.rInI?

It mixes water of different qualities. A poor quality water can be injected into an aquifer at one point, mixed with the local groundwat~r and the mixture withdrawn as useable water at another location.

3.2 TypES OF AQUIfER FORMATIONS

There are two general classes of formation which store and transmit I-Iater these are:-

POi'UUD Hocks

Spaces between grains of sand and gravel - unconsolidated sands and gravels and consolidated sandstones;

Fractured Rocks

These include crevices, joints and fractures in hard rock; and solution channels in limestone, and shrinkage cracks and gas bubbles in basJlt type volcanic rocks.

3.3 HYDRAULIC PROPERTIES

There are ten main hydraulic properties of the aquifer which should be understood. These are:-

(al HydrauliC Conductivity (b) Transmissivity lcl Storage Coefficient



(d) (e) ( f)

~fil W

Specific Mass Storativity Specific Yield Specific Retention Hydraulic Resistance Leakage Coefficient Leakage Factor Drainage Factor

30.

HydrauZio Conductivity (K)(see Sections I and IV)

The hydraulic conductivity of an aquifer is the rate at which water at the prevailing viscosity can be transmitted through a unit area of an aquifer, normal to the "direction of flow, under a unit gradient. It is referred to in engineering literature as the coefficient of permeability.

where

K=-.;!. 1 ------ 3.1

q = discharge rate per unit area normal to the direction of flow i = hydraulic gradient

Hydraulic conductivity is dependent primarily on the nature of the pore space, the type of liquid occupying it, and the gravitation attraction.

The dimensions of hydraulic conductivity are:-

Volume/time/area (Length/Time)

The units to be used are:-

m/day or m3/day/m2

I I I I I I I I I I

Transmissivity (T) I The transmissivity of an aquifer is the rate at which water at the I

prevailing viscosity can be transmitted through a unit strip of aquifer under a unit gradient.

where T = Kb

K = the hydraulic conductivity b = aquifer thickness

The dimensions of transmissivity are:-

Volume/unit time/unit width (Length2/Time)

The units to be used are:-

m3/day/m or m2/day

Storage Coefficient (5) (Confined Aquifers)

------ 3.2

The storace coefficient or storativity of an aquifer is defined as the volume of water which a saturated column of aquifer releases from or takes into storage per unit surface area per unit change in head.

Storage coefficient is related to the elastic properties of the

I I I I I I

water and of the soil matrix. It is not an indication of the total I volume stored in an aquifer.

I



•• I I I I I I I I I I I I I I I I I I I

3L For a confined aquifer the value of Storage Coefficient ranges

from about lQ-S to lQ-J or about 5 x 10-6 per metre of aquifer thickness.

C.E. Jacob (l940,1950), probably the first to consider flow in an elastic aquifer, derived the equation for Storage Coefficient as:-

\!he re

s = bpge (S + GeL ) 8 (~ + C = b 9 eE )

w S b = thickness of the aquifer p = density of the fluid (for water = I 000 kg/m3) g = acceleration due to gravity (9.8 m/sec 2) e = porosity -7 -1 S = compressibility of water, 4.8 x 10 (kPa) Ew= ~ bulk modulus of water ;:.08 x 106 kPa c = compressibility of soil matrix E = 1 = bulk modulus of soil matrix

s c

------ 3.3

C = a dimensionless ratio, which may be considered unity in an uncemented granular material. In a solid aquifer suer: as a limestone having tubular solution cavities, C is apparently equal to the porosity. The value for a sandstone douttl ess ranges between these limits, depending upon the degree of cementation.

Bear (I972, P.207), suggests that the above expression is in error and the correction expression for porous media should be:-

S = bpg

= bpg (se + c(I - el) (~ + (I E e))

w s ------ 3. ~

Bear's approach considers a constant control volume with both the fluid and soil matrix being free to move across the boundaries when the porous media is subjected to varying pressures. Jacob considered that the control volume itself undergoes deformation.

In either case, in the absence of other information a minimum value for Storage Coefficient can be determined by using the first part only of equation 3.3, i.e. the compression of water alone

= bpgs S. E m~n

3.5 w

It will be observed that S is directly proportional to b.

Storage Coefficient has no dimensions. From the following it can be seen that Storage Coefficient is best

thought of in terms of strain.

The Coefficient of Bulk Compressibility, ab' may be defined for a saturated porous medium as the fractional change in the bulk volume of the porous medium with a unit change in the external stress (0) exerted by the formation. d U

. ,-I b 3 6 1.e. Ob = U- do ------ . b

if the hydrostatic pressure (p) is held constant. \.here Ub = the volume of a fixed mass of porous medium.



32. Because groundwater work is ~enera11y associated with a

relatively constant external stress (a) and a variable hydrostatic pressure (p) another coefficient of bulk compressibility, ab' is often defined with respect to a unit change in p.

-1 d Ub a b = U

b dp

where a is now held constant.

------ 3.7

Two more kinds of compressibi 1 ity proposed as

in addition to ab' have been

and

and

where

(a) Rock (or solid) matrix compressibility, as' which is fractional change in volume of the solid matrix (Us) unit change in p,

the with

. -1 1.e.a = U-

s s

d U s dp

(b) Pore compressibility, Cu, defined as in pore volume (Up) with unit change

-1 d Up i.e. a p = Up dp

It follows then, since

U + U s P d U d U

s = dp + p dp

and since

Us = (1 - e) Ub' i.e.

i. e.

Then, by definition

-1 d Ub ~ = U

b dP

=_(1-6) Us

= (1 - e) as + e a p

1 _ (1 - e) Ub - Us

1 _ e --U-Ub P

d U p

dp

From the definition of S ( by Bear ),

5 = bpg {se + a (1 - e)} Sis now wri tten for a and a is now written ?or as

------ 3.8

the fractional change in pressure p.

------ 3.9

------ 3.IO





33.

The major bracketed term is the compressibil ity of the porous medium, and bP9 is in the form of a stress. Since the compressibility is the inverse of the bulk modulus, then the Storage Coefficient S is in the form of a strain of the porous mediu'ffi.

It can be thought of as the amount of deformation of "later and matrix per unit increase in hydrostatic pressure, or can be referred to simply as the elastic storage.

The Specific Mass 5torativity (5s ) or Specific Storage is definec aj the VOlume of water releas2d from or taken into storage per unit volume of aquifer per unit change in head.

i. e.

where

s (~ (1 - e) ) = pg + " t E w s



34.

HydrauZio Resistance (cl

The hydraulic resistance is a property of the confining layer of a semi-confined aquifer. It is the resistance against vertical flow and is defined as -

b' c = ~ ------ 3.13

where

b' = the saturated thickness of the semi-pervious layer K' = the hydraulic conductivity of the semi-pervious layer

for vertical flow.

If DarcY's law is applied to the confining layer then the hydraulic resistance may be thought of as the drawdown in the aquifer required to give a unit discharge per unit area from the confining layer.

If c = 00 the aquifer is confined.

Hydraulic Resistance has dimensions of Time.

The units used are days.

Leakaae Coefficient

The leakage coefficient is a the semi-confined aquifer. It is the is defined as -

property of the confining 1 ayer of inverse of hydraulic resistance and

where

leakage Coefficient = ~:

K' = Hydraulic Conductivity of the semi-pervious layer for vertical flow.

b' = saturated thickness of the semi-pervious layer.

leakage coefficient may be defined as the rate at which water will leak from a unit area of the confining layer per unit drawdown in the aquifer proper.

It has the dimensions of ti~e' The units of leakage coefficient are day -1

Leakage Factor (ll

3.14

The leakage factor is a property of the semi-confined aquifer.

It is defined as -

l = IREC

= IT'C ------ 3.15

"





where

35.

c = hydraulic resistance of the semi-pervious layer K = hydraulic conductivity of the aquifer material b = thickness of the aquifer T = transmissivity of the aquifer

. The leakage factor describes the distribution of leakage into a semi-confined aquifer. High values of L indicate that the influence of leakage will be smal1, i.e. a high resistance of the semi-pervious layer to fl O~I, as compared wi th the res i stance of the aquifer i tse If.

The dimensions of L are in length.

The units of L are metres.

Drainaae Pactor (B)

The drainage factor is associated with the delayed yield in unconfined aquifers and is similar to the leakage factor for semi-confined aquifers.

where

It is defined as -

K = b = T =

1 - -

or

B =/Kb as y

B =/h-y

hydraulic conductivity of the aquifer aquifer thickness transmi.ssivity of the aquifer

the Boulton delay index (an empirical constant)

S = the specific yield after a long pumping time y

------ 3.I6

Large values of B indicate a fast drainage. If B = 00 the yield is instantaneous with the lowering of the water table, so the aquifer would be unconfined without delayed yield.

The dimensions of B are in length, and the units are metres.



36. SECTION IV

GROUNDWATER FLOW

4.0 INTROVUCTION

Be60Jte a pW1lphtg .teot. oJt gJtounlioa..teJt ba.6ht can be anai.y6ed wm any degJtee 06 conMdenc.e; .u:.}A neceMa.Jty .to have a ba.6-i.c. undeJt6.tandhtg 06 glWunlioa..teJt 6£.ow, aqcU.oeJt .typu and hydJr.a.aUc pl1.0peJttiu 06 .the a.qcU. 6 eJt.

GJtoundwa..teJt ht W na.tuJta.t 6.ta..te ~ htvcvUably movhtg and ;t~ movemen.t ~ r.cl.a..ted .to u.tabfuhed hydJr.a.aUc p.iUndplu wh-Lch .(.11. :twm Me goveJr.ned by .the geolog-i.c.a£. 6Vtuc.:twc.e 06 .the ma..twai. .




This statement. that the rate of flow of water through porous media is proportional to the head loss and inversely proportional to the length of the flow path, is knol1n un1ve:-sally as Darcy's Law. Darcy's Law forms the basis for the present day kno\'liedge of groundwater flow.

As water percolates thiOugh a permeable material, the indhidual water particles move along paths vlilich deviate erratically but only slightly from smooth cur'lGS known as "flow lines". If adjacent fl'lw lines are parallel the flow is saici to be "iinear" or "laminar".

The hydraulic principl:=s i~ljoived in laminar flow are illustrated by Fig. 4.1.

Pressure head at A

.,

J --> I I

--L __ _ , I

Elevation head

at A LAMINAR FLOW IN A POROUS MEDIUM

F-G il,. -.' .l •

Datum Level

Elevat:'on

I

head a' B

t

In Fig. 4.!, the points A and B represent the extremities of the flow lines. At each extremity a stand pipe, known as a potentiometric tube, has been instaHed to indicate the level to which the water rises at these points. The water level in the tube at B is designated as the potentiometric level at B and the vertical distance from this level to point B is the pressure head at B. The vertical distance between A and B represents the "position head" or difference in elevation heads above a set datum level.

If the water in the hydraulic system stands at the same elevation in the potentiometric tubes at ,!J. and S, the system is in a state of rest, regardless of the magnitude of the "position head". Flow can occur only if the potentiometric levels at A and B differ. The difference in potentiometric levels, "h" is known as the "hydraulic head" at A with respect to B, or merely referred to as the difference in potentiometric level between A and B. It can be observed that the difference in potentiometric level is equal to the difference in pressure heads at A and



38.

8 only if the position head is zero.

In Fig. 4.1, Al and 81 represent any two points at the same elevation in the potentiometric tubes rising from A and 8 respectively. Since the unit weight of the water is p g, the hydrostatic pressure at

w A1exceeds that at 81 by an amount Pw9h. The difference Pw9h between the hydrostatic pressure at the two points located at the same elevation is referred to as "excess hydrostati c pressure". It is this pressure that drives the water through the soil between A and B. The ratio -

. h u lp = PJl T = T ------ 4 4 .

in which "u" is the excess hydrostatic pressure, represents the pressure gradient from A to B. The ratio,

. i I h 1 = ~ = Pwg • ~ = T

i s knOl~n as the hydrau1i c gradi ent. It is a pure number.

------ 4.5

If water percolates through fine saturated sand or other fine grained completely saturated soils without affecting the structure of the soil, the discharge velocity or unit discharge is almost exactly determined by the equation -

-k v =-11 i p ------ 4 6 . in whi ch

v = average discharge velocity or unit discharge (m/s) 11 = the dynamic viscosity of water (Nsm-2) k = an empirical constant referred to as the intrinsic

permeability (m2 ) ip = pressure gradient (Palm)

The viscosity of water decreases with increasing temperature. The value k is a constant for any permeable material with given porosity characteristics, and is independant of the physical properties of the percolating liquid.

From equation 4.5 and 4.6 we obtain for the discharge velocity the expres s i on,

-k v = - P gi 11 w

------ 4.7

I I I I I I I I I I I I I I I

Most seepage problems encountered in Civil Engineering and Ground-water Hydrology deal almost exclusively with the flow of groundwater at I moderate depths below the surface and with leakage out of reservoirs. The temperature of the percolating water varies so little that the unit weight gpw is practically constant, and, in addition, the viscosity I varies within fairly narrow limits. Therefore, it is customary to substitute in equation 4.7 -

K = kp g

w 11

------ 4.8 I I I



I'

I I I I I I I I I I I I I I I I I I I

whence

where

39.

v = -Ki

v = average discharge velocity or unit discharge K = coefficient of permeability, or hydraulic

conducti vi ty i = hydraulic gradient

------4.9

In Civil Engineering, the value K is commonly called the "coefficient of pel1lleability". In Groundwater Hydrology K is called the hydraulic conductivity. Equation 4.9 is commonly known as Darcy's Law.

where

Darcy's Law may also be written in the form -

Q =-KiA ------4. IO

Q = discharge rate through an area A i = the hydraulic gradient A = the cross-sectional area normal to the direction of flow K = the hydraulic conductivity of the material

4.2 HYDRAULI C CCNDUCTIVITY (i0 From equation 4.3 the hydraulic conductivity is expressed by -

K = - Q A(dh/dl)

which shows that K has dimensions of distance divided by time, i.e. velocity, but is in fact a discharge per unit area. It is defi1ed as the rate at which water is transmitted through a unit area under a unit gradient.

Laboratory tests by the United States Geological Survey have yielded coefficients of permeability or hydraulic conductivity, varying from 9.1 x 10-7 to 4.5 x I02 metres per day, but for most natural aquifers values range between 0.04 and 25 metres per day.

The hydraulic conductivity of a porous medium refers to the ease with which a fluid will pass through it, and varies with the diameter and "degree of assortment" of the individual particles • .A well sorted gravel has a higher hydraulic conductivity than a well sorted coarse sand. However, gravel with a moderate percentage of medium- and fine-grained material may be considerably less permeable than a uniformly sized coarse sand. The reason is that the smaller particles fill the larger pore space, of the gravel and so reduce the area available for flow of water.

The hydraulic conductivity need not be the same in all directions. hydraulic conductivity in the vertical direction is normally less than the hydraulic conductivity in the horizontal directions. This is caused by deposition of materials in horizontal layers. An occasional relatively impermeable layer reduces the vertical hydraulic conductivity without appreciably affecting horizontal hydraulic conductivity.

4.3 RELATICN OF K TO PARTICLE VELOCITY

Because K has the dimensions of velocity (length/time), some might



I 40.

mistake this for the actual, or particle velocity, of the water. However, I from equation 4.10 we see that before simplification -

3 K=-zm 1 I

m day m m

which is the volumetric rate of flow through a given cross-sectional area I under unit gradient. For the average actual or particle velocity, we must know also the porosity of the medium. Thus,

where

4.4

_ dh Q = vA - -KA aT

v = K dh - e . dl v = average water particle velocity e = porosity, as a decimal fraction

~~ = hydraulic gradient = i

RELATI(XIl OF K TO INTRINSIC PEWEABILITY

------4.11

------4.12

Intrinsic permeability is a measure of the relative ease with which a porous w~dium can transmit a liquid under a potential gradient. It is a property of the medium alone dependent on the size and shape of the pores, and is independent of the nature of the fluid and of the force field causing movement.

intrinsic permeability, k, has the dimensions length 2, whereas in equation 4.I2, hydraulic conductivity, K, has the dimenSions, length/time.

where

The two concepts are related by the following expression -

k = f!l. ------4. I3 P9

or = Kv g

k = intrinsic permeability of the medium (length2) K = the hydraulic conductivity (length/time) n = dynamic viscosity of the fluid (mass/length time) p = the density of the fluid (mass/length3) g = acceleration of gravity (length/time2)

v = kinematic viscosity (length2/time)

------4.14

It should be emphasised that the intrinsic permeability characteristics of a porous material are expressed by k (length2) and not by K (length per unit time). The value of k is independent of the properties of the liquid, whereas K depends not only on the properties of the porous material, but also on the properties of the liquid.

The channels through which the water particles travel in a mass of soil have a variable and irregular cross-section. As a consequence, the real velocity of flow is extremely variable. However the average flow through such channels is governed by the same laws that determine the rate of flow through straight capillary tubes having a uniform cross-section.

I I I I I I I I I I I I I I I I




41.

If the cross-section of the tube is circular, the velocity of flow increases with the square of the diameter of the tube. Since the average diameter of the voids in soil at a given porosity increases practically in proportion to the grain size 0 it is possible to express K as

K = constant x 02 ------4. IS

From his experiments with loose filter sands of high uniformity (uniformity coefficient not greater than about 2), Allen Hazen obtained the empirical equation,

K(cm/sec) = C1 Dio ------4.16 in .,hich 0 10 is the effective size in centimetres of the sand and C,

(i/cm.sec) varies from about 100 to ISO. It should be noted that equation 4.16 is qpplicable only to fairly uniform sands in a loose state.

The coefficient of permeability or the hydraulic conductivity of uniform unconsolidated materials may then be thought of as being proportional to the square of the effective grain size.

4. 5 REYNOLD I S NUM3EF<

In pipe flow the transition from laminar to turbulent flow is character; sed by we 11 known values of Reynol d's Number (N p) whi ch expresses the ratio of the inertia forces to the viscous forces. There is a lower limit critical number around 2,100 below which flow in pipes is always laminar.

By analogy, in flow through porous media a Reynold's Number has been established as -

where

vO N =-R v

v = specific discharge i.e. discharge per unit area

------4.I7

o = characteristic length. In pipe flow 0 is the internal diameter of the pipe. In flow through porous w~dia 0 is related to grain size

v = kinematic viscosity of the fluid

Because the grain sizes are so variable in flow through porous rredia no one value for Reynold's number can be set as the dividing line between laminar flow and turbulent flow. It has been established that this transition occurs normally for a Reynold's Number in a range I to 10.

4.6 RANGE OF VALIDITY OF DARCY' S LA~I

It must be emphasised that Oarcy's law is only valid for laminar flow in porous media.

For very low velocities, laminar flow occurs. There is a condition known as prelaminar flow for extremely slow percolation. However, this will not be considered in these notes. The upper limit of the applicability of Darcy's Law is when the Reynold's Number is in the range of I to 10. A range rather than a specific number is given because the distribution of grain sizes of natural media for a specified average grain diameter is 1 imitless.



42.

4.7 GROUNDWATER FLOW RATE

The previous sections have shown that when the water table is sloping, or has a gradient, then groundwater movement must exist and the magnitude of the resultant flow is dependant on the hydraulic conductivity of the aquifer material and the magnitude of the gradient.

To obtain some idea of the magnitude of this flow, assume a gradient of 2 metres per kilometre and a hydraulic conductivity of 0.5 m3/day/m2.

assuming a hydraulic

v = K ~ 2 = 0.5 . IOOO

v = 0.001 metres per day conductivity of 250 m3/day/m2

2 rooo v = 250

= 0.5 metres/day

As can be seen, the rate of flow of groundwater under natural conditions is relatively small and varies markedly from place to place as the gradient or the hydraulic conductivity varies.

It is possible to measure hydraulic conductivity both in the laboratory and in the field. Laboratory measurements are outside the scope of this course but types of field measurements such as pump ~ests, will be discussed later.

4.S FLOW AN~LOGIES

The equivalent of Darcy's Law, as expressed in equation 4.3, to the basic laws of heat and electricity flow has permitted the ready adaptation of solutions of complex heat and electrical flow problems to the flow of groundwater. The classic text "Conduction of Heat in Solids" by H.S. Carslaw and J.e. Jaeger, has become a source of information for the solution of problems in groundwater hydraulics.

The equivalents of flow of groundwater, heat and electricity is very useful in the modelling of groundwater problems ..

where

Heat Equivalent

Fouri er I sLaw

f = -K dv dl

f = a flux, or heat flow, per unit area K = the thermal conductivity v = temperature 1 = length of flow path

------4. IS





where

4.9

43.

Electrical Equivalent

Ohm's Law

I = C = R = E = 1 :

I = -c dE dl ------4.19

current (amperes) per unit area electrical conductivity (l/R), (mho em) per unit area electrical resistance (ohms) per unit area electrical potential (volts) length of flow path (centimetres)

TYPES OF GROUNIWATER FLa,./

Laminar Flow (Darcy Flow)

Darcy Flow is laminar and any head loss in overcoming frictional resistance for laminar flow is proportional to the discharge. This flow conforms with Darcy's Law.

Turbulent Flow (Non-Darcy Flow)

If the velocity become large enough, and Reynold's Number exceeds some value between I and ro, the fluid flow becomes turbulent and Darcy's Law no longer applies. Head loss associated with overcoming frictional resistance for turbulent flow is normally proportional to the discharge squared.

4.10 STATES OF GROUNDWATER FLOW

There are two general states of groundwater flow. These are Steady State flow and Non-Steady State flow.

Steady State Flow

Steady State or Equilibrium flow occurs when the discharge from the bore is in equilibrium with the recharge to the aquifer system. Examples of when steady state flow might occur are -

(al when the radius of influence from a bore located in the centre of a circular island intersects the sea;

(bl after extensive pumping in a leaky aquifer.

Steady State flow does not occur very frequently. In practice it is assumed to occur when the change of drawdown with times becomes negligible, after corrections have been made for outside influences, such as tides, which might affect the water levels.

Non-Steady State Flow

All flow which is not in equilibrium with the recharge to the aquifer is called a Non-Steady State flow. As long as there is a measurable change in drawdown \~ith time the flow is classed as non-steady. In practice, this is the most common type of flow associated with discharging bores.



5.0

44.

SECTION V BORE DISCHARGE TESTS

INTRODUCTION

Fo~ too tong, 6towing and non 6tow~g bo~e.o have been ~eg~ded by .6Ome au:thotU.:tLe.o .u b~g eompte;tety cU66~ent o~om one ano-th~ and .u a ~e.ouU -the -tec.hn).qu.e.o u.6ed -Ln ;te.o:u'ng -them have been cU66~ent. TIUt:, -




Fig. 5.1 GROUI«) SUFffiOC;£ Pump

A

"

AQUIFER

SHALE

NON FLOWING

POTENTiOMC:TRIC SUR&C£ - --H----- -- ---- '- FI""mQ 8""

CONE OF " OEPRESSION-------"

o !' .....

SHALE

AQUIFER

SHALE

FLOWING

FLOWING AND NON FLOWING BORES



46.

5.1 DEFINITIONS

Before proceeding with the various types of tests a number of terms will be defined:-

Standing Watep Level (S.W.L.): The depth from ground level to the water level in a non pumping bore outside the area of influence of any pumping bore.

Static Head: This is the flowing bore equivalent of S.W.L. It is the height above ground level that water at a particular temperature stand if the casing were extended upwards. It is expressed as meters or kilopascals.

DPawdown: The distance that the water level in a bore or well has been lowered from the S.H.l. during pumping. It is measured at specified times after pumping commenced.

Residual Drawdown: The distance that the water level in a well or bore remains lowered from the S.W.L. after pumping ceases. It is measured at specified times after pumping stopped.

Recovery: The amount by which the water level in a bore has risen at a given time after pumping ceased. It is the difference between the Residual Drawdown after the given time and the hypothetical drawdown if pumping had not ceased. When the water level returns to S.W.L. recovery is said to be complete.

AvaiZable DrawdOwn: For a particular pump installation this is the distance between the S.W.L. and the pump suction or the depth of water over the pump suction.

5.2 MEASUREMENTS AND MEASURING APPARATUS

The adequate evaluation of a pumping test relies very much upon the recording of a number of sets of measurements throughout the test. Each set comprises measurements of:

(il (i i ) (i i i )

Time Discharge Head

The values obtained for the aquifer characteristics cannot be of greater accuracy than that of the basic data. Care should be taken then in the measuring and recording of time discharge and head.

The head measurement may be expressed in terms of water level, drawdown, or for a flo~ling bore, the back pressure.

The time measurements are normally started from the beginning of the test and may be recorded as time of day or time since the test started. It is desirable to record the time of day as this not only gives the location of the set of measurements in anyone test, but also indicates the relationship between that test and any other test which may have been carried out on the same day. This allows correction for any antecedent pumping conditions.

Instantaneous discharge measurements are desirable, so the use of equipment such as the orifice meter or orifice bucket is advocated.





47.

Time

Time is recorded as time of day and date. In addition the time in minutes after a particular test starts can be recorded, but if an accurate record of time of day is kept, this is not essential.

TABLE 5-1

SUGGESTED DURATIONS FOR DISC~ARGE TESTS

Bore Use or Pumping Duration Recovery Duration Reason for Test (hours) (hours)

Stock and Domestic 4 - 6 2 - 4 Irrigation 24 6 Town Water Supply 100 , 24 Bore Remeasurement 2 (dynami c tes t) 2

-

rirrtel' Level OT PressW'e

Water levels in non flowing bores are measured by a variety of apparatus the most common being the type whereby an electrical circuit is campi "ted when the probe touches the water surface.

I n the case of fi owi ng bores it is not feas i b 1 e to measure water levels. The pressures measured are measurements of back pressure on the inlet ~ide of the control valve.

These pressures can be measured with a pressure gauge, which is not accurate or sensitive enough and needs to be calibrated quite frequently, or by means of a mercury manometer.

If the back pressure is very small a water tube could be used i.e. the back pressure is the height that water rises up the tube. Normally back pressures of bores being tested are not low enough for the use of a water tube.

Most bores have a reasonably high back pressure (which reduces as the flow increases) and the most accurate method of measuring it is by use of a mercury manometer. If the pressure is too high for one manometer two manometers can be used in tandem.

The use of mercury manometers, single or in tandem, will not be dealt with in these papers. If more information is required it is suggested that you write to the Secretary, Irrigation and Water Supply COllll1ission in Brisbane and request a copy of "Artesian Bore Testing -July 197I".



48.

Discha:t'ge

Discharge measurements are made using -

(a) Time to fill a container of known volume (b) Weir Boards (c) Orifice Bucket (d) Orifice Meter

These are written in ascending order of accuracy. The use of these various techniques is described in detail in "Artesian Bore Testing-July 1971" - (Queensland Irrigation and Water Supply Commission internal paper) .

Temperature

This is important in flowing bores only. When a hot bore stops flowing the water cools down and as it cools it becomes denser than the hot water so that a decrease in measured pressure results. This is often the cause of the falling off in pressure during the later stages of a static test. This effect has not been fully appreciated in the past but it is now obvious that a systematic programme of water temperature reading must be incorporated in any test programme.

The temperature of the water should be taken as near to the discharge point as possible and preferably in the mouth of the orifice meter at the following times -

5.3

5.3.1

(i) At the beginning of the flow recession test, just after opening;

(ii) At the end of the flow recession test, immediately before clvsing for the static test;

(iii) During the first minute of the first stage of the opening dynamic test;

(iv) At each set of measurew€nts in the first stage of the opening dynamic test;

(v) At the end of each other stage of the opening or closing dynamic test.

TESTING NON FLOWiNG BORES

CONSTANT DTSCEARGE TEST (See also Section 9.5)

As the naw€ implies this test involves pumping the bore at a constant discharge rate and measuring the varying drawdown throughout the test.

With the discharge held constant, drawdown measurements are taken during the test at the following times after pumping commenced -I, 2, 3, 4, 6, 8, 10, 20, 25, 3D, 45, 60, 75, 90, 100, 120 minutes then each half-hour to six hours and then hourly until the end of the test.

Rate measurements are taken at least at the start of the test, after IS minutes, 30 minutes and every half-hour thereafter, but at more frequent intervals if possible. However, care should be taken to maintain the discharge rate constant throughout the test by regular inspections of the tube on the orifice meter or orifice bucket if these are being used.





49.

At the end of the constant discharge test residual drawdown measurements are taken, if possible, at the following times after pumping ceases - i, 2, 3, 4, 6, 8, 10, 15, 20, 25, 30, 45, 60, 75, 90, IDO, I20 minutes, then hourly until the water level has come back to within 15 centimetres of the standing water level. If for any reason a reading· at any of the above times is missed, then a reading should be taken as soon as possible thereafter and the 5ctual time of this reading should be· recorded. It is more important to k~ow the time when the drawdown was measured than to have the drawdowns measured at the exact times laid out in these notes.

5.3.2 CONSTANT DRA'dDO'dN TEST

In this type of variable discharge test the drawdown is held at d constant value and variations in discharge are measured. This type of test may be required when it is not possible to measure drawdo~ms. The drawdown could be maintained at a constant level with the pump.

The drawdown is held constant by making sure that the pum~· breaks suction soon after the test begins .. The water level is ther. maintained at the pum? suctior. throughout the test.

Becaus

groundwater hydraulics - territory stories: homeorigin of groundwater hydrologic cycle factors...

Documents