Studies and reports in hydrology 29 · 2014-12-28 · 2.2 Recharge of Groundwater 15 2.2.1 Recharge...

Studies and reports in hydrology 29

Recent titles in this series

20. Hydrological maps . Co-édition Unesco-WMO. 21 * World catalogue of very large floods/Répertoire mondial des très fortes crues. 22. Floodflow computation. Methods compiled from world experience. 23. Water quality surveys. 24. Effects of urbanization and industrialization on the hydrological regime and on water quality. Proceedings

of the Amsterdam Symposium, October 1977/Effets de l'urbanisation et de l'industrialisation sur le régime hydrologique et sur la qualité de l'eau. Actes du Colloque d'Amsterdam, octobre 1977. Co-edition IAHS-Unesco/Coédition AISH-Unesco.

25. World water balance and water resources of the earth. (English edition). 26. Impact of urbanization and industrialization on water resources planning and management. 27. Socio-economic aspects of urban hydrology. 28. Casebook of methods of computation of quantitative changes in the hydrological régime of river basins due to human

activities 29. Surface water and groundwater interaction.

* Quadrilingual publication : English — French — Spanish — Russian.

For details of the complete series please see the list printed at the end of this work.

Surface water and groundwater interaction

A contribution to the International Hydrological Programme

Report prepared by the International Commission on Groundwater

Edited by C . E . Wright

lunssco

The designations employed and the presentation of material throughout the publication do not imply the expression of any opinion whatsoever on the part of Unesco concerning the legal status of any country, territory, city or area or of its authorities, or concerning the delimitation of its frontiers or boundaries.

Published, in 1980 by the United Nations Educational, Scientific and Cultural Organization 7 place de Fontenoy, 75700 Paris

Printed by Imprimerie de la Manutention, Mayenne

I S B N 92-3-101862-0

O Unesco 1980

Primed in France

Preface

The "Studies and Reports in Hydrology" series, like the related collection of "Technical Papers in Hydrology", was started in 1965 when the International Hydrological Decade ( I H D ) was launched by the General Conference of Unesco at its thirteenth session. The aim of this undertaking was to promote hydrological science through the development of international co-operation and the training of specialists and technicians.

Population growth and industrial and agricultural development are leading to constantly increasing demands for water, hence all countries are endeavouring to improve the evaluation of their water resources and to m a k e more rational use of them. The I H D was instrumental in promoting this general effort. W h e n the Decade ended in 1974, I H D National Committees had been formed in 107 of Unesco's 135 M e m b e r States to carry out national activities and participate in regional and international activities within the I H D programme.

Unesco was conscious of the need to continue the efforts initiated during the International Hydrological Decade and, following the recommendations of M e m b e r States, the Organization decided at its seventeenth session to launch a new long-term intergovernmental programme, the International Hydrological Programme (IHP), to follow the decade. The basic objectives of the I H P were defined as follows: (a) to provide a scientific framevork for the general development of hydrological activities ; (b) to improve the study of the hydro-logical cycle and the scientific methodology for the assessment

of water resources throughout the world, thus contributing to their rational use; (c) to evaluate the influence of man's activities on the water cycle, considered in relation to environmental conditions as a whole; (d) to promote the exchange of information on hydrological research and on new developments in hydrology; (e) to promote education and training in hydrology; (f) to assist M e m b e r States in the organization and development of their national hydrological activities.

The International Hydrological Programme became operational on 1 January 1975 and is to be executed through successive phases of six years' duration. I H P activities are co-ordinated at the international level by an intergovernmental council composed of thirty M e m b e r States. The members are periodically elected by the General Conference and their representatives are chosen by national committees.

The purpose of the continuing series "Studies and Reports in Hydrology" is to present data collected and the main results of hydrological studies undertaken within the framework of the decade and the new International Hydrological Programme, as well as to provide information on the hydrological research techniques used. The proceedings of symposia will also be included. It is hoped that these volumes will furnish material of both practical and theoretical interest to hydrologists and governments and meet the needs of technicians and scientists concerned with water problems in all countries.

Contents

1. INTRODUCTION 11

1.1 Purpose and Scope of the Report 11

1.2 Formation and Activity of the Working Group 12

1.3 Relationship with other IHP Working Groups 13

2. DEFINITION OF THE INTERACTION I4

2.1 Part of Hydrological Cycle Considered 14

2.2 Recharge of Groundwater 15

2.2.1 Recharge by Precipitation 15

2.2.2 Recharge by Rivers and Canals 17

2.2.3 Recharge by Lakes 19

2.2.4 Artificial Recharge 20

2.3 Groundwater Component of River Flow 20

2.4 Influence of the Interaction on Water Quality 22

2.4.1 Surface Water to Groundwater 23

2.4.2 Groundwater to Surface Water 25

3. METHODS OF ASSESSING THE INTERACTION 27

3.1' Channel Water Balance 27

3.1.1 Compilation for River Reaches 28

3.1.2 Equations for River Reaches 28

3.1.3 Compilation for River Systems 29

3.1.4 Equations for River Systems 29

3.1.5 Computation of the Elements 30

3.1.5.1 Exchange Between Rivers and Aquifers 31

3.1.5.2 River Flow, Intermediate Inflow, Abstractions and Returned Water 32

3.1.5.3 Channel Regulation 32

3.1.5.4 Precipitation 32

3.1.5.5 Evaporation from Surface Water 32

3.1.5.6 Change of Stored Moisture 33

3.1.5.7 Ice Formation and Melting 33

3.1.5.8 Areal Definition of Flood Plain Reaches 33

3.2 Hydrograph Analysis 33

3.2.1 Flow Separation 33

3.2.2 Graphic Separation of River Hydrograph 36

3.2.3 Recession Curve for River Hydrograph Separation 39

3.3 Groundwater Table Fluctuations 41

3.3.1 Temperate Areas 42

3.3.2 Arid Areas 47

3.4 Use of Isotopes as Tracers 50

3.4.1 Introduction 50

3.4.2 Stable Isotopic Composition of Natural Waters 51

3.4.3 Environmental Tritium Concentration of Natural Waters 52

3.4.4 Recharge of Groundwater by Rivers 53

3.4.5 Recharge of Groundwater by Lakes 57

3.4.6 Groundwater to Surface Water 58

3.5 Use of Mathematical Models 59

3.5.1 Purpose of Modelling 59

3.5.2 Groundwater Recharge 60

3.5.3 Spring-Aquifer Interaction 61

3.5.4 Rivers and Canals 62

3.5.5 Lake-Aquifer Interaction 65

3.5.6 Example of Finite Element Analysis 66

3.5.7 Rainfall-Runoff Models 69

4. ACCURACY OF METHODS OF ASSESSMENT 74

4.1 Surface Water Flow 74

4.1.1 Temperate Areas 75

4.1.2 Arid and Semi-arid Areas 77

4.1.2.1 Measurement of Flood Flows 79

4.1.2.2 Measurement of Low Flows 80

4.2 Aquifer Characteristics 81

4.2.1 Hydraulic Conductivity 82

4.2.1.1 Laboratory Determination of Hydraulic Conductivity 82

4.2.1.2 Field Determination of Hydraulic Conductivity 82

4.2.2 Transmissivity 83

4.2.3 Specific Yield 84

4.2.4 Coefficient of Storage 84

4.2.5 Infiltration 86

4.3 Relative Accuracy of the Methods of Assessment 86

4.3.1 Channel Water Balance 87

4.3.2 Flow Separation 88

4.3.3 Mathematical Models 89

5. CASE STUDIES 93

5.1 Temperate Area: Great Ouse Pilot Scheme, UK 93


5.1.2 Description of the Pilot Scheme Area 94

5.1.3 Measurements 97

5.1.4 Natural River Flow and Groundwater Level Relationship 98

5.1.5 Analysis of Group Pumping Tests, 1971 101

5.2 Temperate Area: The Moscow Artesian Basin, USSR 102


5.2.2 The Moscow Basin 102

5.2.3 Analyses 103

5.2.4 Future Situation 104

5.3 Arid Area with Irrigation: Chu Valley, USSR 106


5.3.2 The River Chu 106

5.3.3 Description of the Study Reaches 106

5.3.4 The Channel Water Balance 107

5.3.5 Summary 107

5.4 Arid Area: Groundwater Replenishment by Surface Water, Tunisia 109


5.4.2 Aguifer Recharge in' the Kairouan Plain 109

5.4.3 Recharge by Surface Runoff from the Zeroud Wadi 109

6. CONCLUDING REMARKS AND RECOMMENDATIONS 113

6.1 Concluding Remarks 113

6.2 Recommendations 115

REFERENCES 116

SELECTED PAPERS FROM 1979 SYMPOSIA 123

Intvoäuation

i. Introduction 1.1 Purpose and Scope of the Report

There has been a tendency in past years for separate departments to develop specialising in

either surface water or groundwater systems. For this reason the understanding of the inter

action between surface water and groundwater and techniques for its analysis have tended to be

less well advanced than those for either discipline. In recent years the traditional division

between the disciplines has tended to be reduced with the result that some useful advances have

been made in understanding the interaction between surface water and groundwater. This report

does not attempt to review all the relevant research of recent years but rather to emphasise

and illustrate the importance of the subject.

Improvements in understanding the interaction can provide information useful in the

management of water resources. For example existing schemes may be operated more efficiently

and new techniques may be considered in planning the future development of resources. In all

areas where water is a relatively scarce commodity there is a positive requirement to define

the interaction accurately. One of the main purposes of this report is to assist developing

countries, especially those in arid areas, in the management of their water resources.

However, there are likely to be benefits arising as a result of accurately defining the inter

action in other regions such as those where the demand for water represents a high proportion

of the total resource and where changes in the interaction caused by man have a marked

beneficial or detrimental effect.

Detailed consideration has been confined to one part of the hydrological cycle, the

interaction between surface water and groundwater. In temperate regions the main aspect of

this process is the flow of groundwater to rivers, and in arid regions the flow is frequently

in the reverse direction with surface runoff recharging groundwater. Subject areas covered by

other Working Groups such as irrigation, groundwater models, water quality and low river flows

are referred to but not covered in detail. For example the effect of irrigation is included

in a case study, and sections of the report contain a discussion of groundwater models and

water quality. Changes in water quality due to the effect of man in polluting either surface

water or groundwater is a major topic that is referred to only briefly. Methods of assessing

the interaction between surface water and groundwater are described together with an

11

Introduction

assessment of their accuracy. Four case studies are included. Two describe investigations in

temperate regions and two describe aspects of the interaction process in arid regions.

Publications referred to in the text are listed in the references and an additional list of

selected relevant papers is given from symposia held during 1979 at Dortmund (FRG), Vilnius

(USSR) and Canberra (Australia).

1.2 Formation and Activity of the Working Group

The second session of the Intergovernmental Council of the International Hydrological

Programme (IHP), was held in June 1977, when the decision was taken to:

invite the Secretariat in co-operation with the International Commission on

Ground Water (ICGW) of the International Association for Hydrological

Sciences (IAHS). to prepare a technical report 'Improvement of methods of

assessment of the interaction between groundwater and river flow' and

report on the progress of this project to the third session of the Council.

Two sessions of the Working Group have been organised by the ICGW Secretary with the

assistance of the IHP Secretariat of Unesco. The first session was held at the Unesco

headquarters in Paris from 12 to 16 June 1978 and the second was held at Dortmund from 7 to

11 May 1979. The Working Group was composed as follows.

First Session Second Session

Mr V V Kuprianov (USSR) Mr 0 V Popov (USSR)

Mr J Soveri (Finland) Mr J Soveri (Finland)

Mr C E Wright (Chairman, UK) Mr C E Wright (Chairman, UK)

Mr H Zebidi (Tunisia) Mr M Ennabli (Tunisia)

In addition the following experts were invited to attend the sessions :

First Session Second Session

Mrs N Kapotova (USSR)

Mr C Pollett (Australia)

Mr J A Rodier

Mr G Castany

Mr M G Bos

(IAHS)

(IAH)

(ICID Committee on Irrigation Efficiencies)

The report has been prepared by the members of the Working Group together with the following

invited authors :

Mr C van den Akker (Netherlands)

Mr D A Kraijenhoff van de Leur (Netherlands)

Mr B R Payne (IAEA)

Mr J A Rodier (France)

Mr K R Rushton (UK)

12

Introduction

Mr H J Colenbrander, Mr C E Wright and Mr Y N Bogoyavlensky were responsible for the final

editing.

1.3 Relationship with other IHP Working Groups

Parts of the subject area of this report could overlap or are closely linked to the work of

other IHP Working Groups. The subjects associated with these Working Groups are listed to

enable further information to be obtained if required.

Project 5.1 Assessment of quantitative changes in the hydrological regime of river basins

due to human activities (1975-1980) - Preparation of a casebook on methods of

computation (1975-1979).

Project 5.4 Investigation of water regime of river basins affected by irrigation (1975-1980)

Preparation of a technical report (1978-1980).

Project 7.3 Investigations of processes of quality and quantity changes of groundwater

resources due to urban and industrial development.

Project 8.1 Physical and mathematical models for investigation and predicting the changes in

groundwater regimes due to human activities.

Project 8.2 Study of groundwater recharge, including water quality aspects.

Part of the terms of reference for project 7.3 include a review of the present knowledge

of the interaction between surface water and groundwater in the urban environment. Therefore

this report (project 3.6) does not include a section on the urban environment.

13

Definition of the Interaction

2. Definition of the interaction 2.1 Part of Hydrological Cycle Considered

The interaction between surface water and groundwater is a part of the hydrological cycle that

has been examined in some detail in recent years. There are two main aspects of this process,

firstly the flow of groundwater to support river flow and secondly the flow from rivers to

groundwater. The former is a common occurrence in temperate regions whereas the latter occurs

widely in arid regions. Figure 1 is a simplified conceptual model that illustrates the subject

area of this report. There is considerable scope for modifying the figure to allow for

local conditions. For example in highly permeable areas the surface storage component could

be negligible and therefore be omitted.

poration

Capillar

Rise

Precipitation

1 Surface

Storage

< Infiltratior

Storage in

Unsaturated Zone

i

i

1

Groundw

Rechar

Groundwater

Storage

ater

ge

Overland flow

Interflow

Base flow

<

Direct

Runoff

Total

Runoff

Figure 1 A conceptual model

River flow is derived essentially from precipitation less evaporation and the routes by

which precipitation becomes river flow are shown in Figure 1. In a natural river system with

14

definition of the Interaction

negligible abstractions and discharges there are two main components of river flow, namely

direct runoff and base flow. Direct runoff may be subdivided into channel precipitation,

overland flow and interflow, whereas base flow is that part of river flow that is derived from

groundwater. Groundwater flow is defined as flow within the saturated zone. In catchments

with more than one aquifer the base flow component may be subdivided according to the

contributing formation. The proportion of' direct runoff or base flow in total river flow may

vary substantially from one basin to another and from month to month because of the effect of

different soil types, geology, land use, topography, stream patterns and changes in

precipitation, evaporation and temperature.

In temperate regions groundwater recharge is derived mainly from precipitation less

evaporation, where evaporation is defined as including transpiration and interception losses

from vegetation. However, in arid regions, where annual potential evaporation exceeds

precipitation, groundwater recharge is frequently derived from temporary rivers that are in

flood. More generally both flood water and base flow from mountain rivers can recharge

aquifers in the foothills and adjacent relatively dry low lying areas. In addition groundwater

recharge may occur from lakes, canals and excess irrigation. If the groundwater table is near

to the surface of the ground, then the capillary rise may enable evaporation to deplete

directly the groundwater storage. The infiltration process and the movement of water in the

unsaturated zone are not discussed in detail in this report.

The storage, flow and quality characteristics of surface water and groundwater are

frequently dissimilar. For this reason the interaction is important in water-resource

development since advantage may be taken of the differing characteristics to increase yields or

improve the quality of water supplies. Changes in one part of the hydrological cycle may

induce beneficial or detrimental changes in another part of the cycle. A definition of the

water balance and its elements or component parts has been given by Brown et aZ.t (1972) .

2.2 Recharge of Groundwater

2.2.1 Recharge by Precipitation

The main source of groundwater recharge is generally directly from precipitation particularly

in those areas where annual average precipitation exceeds potential evaporation. Evaporation

may deplete water held in surface storage, in the soil or in the aquifer as shown in Figure 1.

Groundwater recharge occurs when the residual precipitation (precipitation less actual

evaporation) has infiltrated to the groundwater table. This may occur from several hours to

several months after the precipitation event. If the precipitation is in the form of snow then

infiltration is delayed indefinitely until there is a thaw.

To fully understand the characteristics of aquifer storage it is necessary to investigate

the characteristics of precipitation, evaporation, temperature and the unsaturated zone which

collectively determine the temporal distribution and rate of recharge. In some parts of

western Europe, consecutive monthly totals of precipitation may be regarded as independent

(random) events that are uncorrelated with past or future monthly totals, whereas evaporation

has a strong seasonal (cyclic) pattern that is repeated year after year. In such areas

15


significant random and cyclic components are observed in time-series recharge data. Where

precipitation and evaporation data display different properties the characteristics of recharge

data will vary accordingly and in colder regions will be significantly influenced by temp

erature as discussed in section 3.3.

In arid areas the direct recharge of groundwater from rainfall is likely to be

insignificant because of several factors.

1. for most of the year, rainfall is relatively small compared with potential

evaporation,

2. storm intensity frequently exceeds the infiltration capacity of the ground

surface resulting in overland flow,

3. the unsaturated zone tends to dry out and may therefore absorb a

significant volume of infiltrating water,

4. semi-permeable crusts may form in the unsaturated zone comprising fine

sediments that impede infiltration.

During the relatively few days that rainfall exceeds evaporation in arid areas, the

storm intensity is frequently sufficient to induce surface runoff thus effectively removing

the potential recharge water to a location downstream. Any water that does infiltrate tends

initially to reduce the soil moisture deficiency then evaporate rather than recharge ground

water. Where rainfall is infrequent and irregular, direct recharge from precipitation is

likely to be even less frequent.

Aquifers may be divided into two types, fissured and arenaceous, depending upon whether

the storage of water is essentially within fissures or intergranular. However, some aquifers

may be a mixture of both types with, for example, storage contained substantially within the

granular interstices, but flow mainly through fissures. The delay between a precipitation

event and the consequential rise in the water table is dependent upon the aquifer properties

discussed in section 3.3. Where permeable soils overlie highly fissured deposits such as the

Karst Limestones, high intensity rainfall may infiltrate rapidly to depths from which

evaporation is negligible. An example of this phenomenon has been described by Downing and

Williams (1969) for the Lincolnshire Limestone of eastern England and Rushton (1976) estimated

that rapid recharge through 'swallow' holes and fissures contributes up to 40 per cent of the

total recharge to this aquifer. In these conditions groundwater recharge may be derived in

approximately equal proportions from precipitation (directly) and surface runoff.

Long period records of weather conditions, river flows and groundwater levels are

valuable aids in the analysis of water resources. Provided that the records are accurate the

longer the record the more accurately defined are the annual and monthly means and the

variation about the mean. In addition long terms trends and cycles may be detected. For

example in a study of the 1972 to 1973 drought in northern Nigeria some trends and cycles were

detected in the hydrological data (Sonuga, 1977) . Unfortunately long term records are not

16


available in many areas and the assessment of rainfall characteristics may be complicated by a

high variability of daily, monthly and annual rainfall within relatively small areas (Balek,

1978) .

Where long period weather records exist and a suitable model is available, it may be

possible to synthesise long sequences of aquifer recharge data. An abstract from such a

synthesised record is shown in Table 2.2.1 (Morel and Wright, 1978) which illustrates the

random and seasonal components of aquifer recharge for the Chalk of eastern England (West

Suffolk). This area has an annual average rainfall of 600 mm of which 450 mm is evaporated.

Recharge occurs mainly during the four months December to March, but may be negligible if

winter rainfall is insufficient to restore the soil to field capacity. The exceptionally dry

weather of 1972-73 and 1975-76 resulted in negligible groundwater recharge for periods of

18 months. From the long synthesised record it is apparent that such events may occur in this

area on average three or four times every 100 years.

Table 2.2.1 Typical values of monthly aquifer recharge in eastern England (1965-76)

(millimetres)

Year

1965

1966

1967

1968

1969

1970

1971

1972

1973

1974

1975

1976

Jan

0

20

23

39

57

49

70

51

0

0

54

0

Feb

0

54

36

22

49

50

10

30

0

46

13

0

Mar

12

0

0

1

37

23

23

24

0

2

64

0

Apr

10

9

12

0

0

29

0

0

0

0

22

0

May

0

0

21

0

20

0

0

0

0

0

0

0

Jun

0

0

0

0

0

0

0

0

0

0

0

0

Jul

0

0

0

0

0

0

0

0

0

0

0

0

Aug

0

0

0

13

0

0

0

0

0

0

0

0

Sep

0

0

0

56

0

0

0

0

0

0

0

0

Oct

0

0

0

24

0

0

0

0

0

0

0

0

Nov

24

24

0

28

0

0

2

0

0

88

0

0

Dec

91

67

40

42

0

0

15

0

0

21

0

17

Total

137

174

132

225

163

151

120

105

0

157

153

17

2.2.2 Recharge by Rivers and Canals

Recharge may occur whenever the stage in a river or canal is above that of the adjacent ground

water table, provided that the bed comprises permeable or semi-permeable material. This type

of groundwater recharge may be temporary, seasonal or continuous. Also it may be a natural

phenomenon or induced by man. For example intermittent recharge may occur in arid regions

when temporary rivers are flowing in valleys that are usually dry (Besbes et al., 1978),

seasonal flow can occur to and from bank storage (Popov, 1969), and there may be a

continuous flow to groundwater from rivers and canals (Smiles and Knight, 1979). In New

Zealand several groundwater bodies near the coast are recharged mainly by seepage from river

beds, and it is probable that similar processes occur in other places around the world

(Woudt et al., 1979). When there is seepage from a canal or ditch overlying a shallow water

17


table, water-logging of the soil at points some distance from the canal is a distinct

possibility (Bruch, 1979).

Man can induce groundwater recharge from rivers by lowering the water table adjacent to

rivers or by raising the river stage. The former is a relatively common occurrence which may

be caused by groundwater abstractions for supply or by mine drainage, and the latter may be

caused by reservoir releases (Kemp and Wright, 1977), weirs or other engineering works. A

serious deterioration in groundwater quality may result if the recharge water is saline or

significantly polluted. This is discussed in section 2.4.

Since the replenishment of groundwater by temporary rivers is frequently the main source

of aquifer recharge in arid regions, much of this section describes the process in such areas

and additional information is contained in two case studies. However, groundwater recharge

from rivers also occurs in other regions where the geological conditions are favourable,

especially where there are Karstic rocks.

In arid areas groundwater recharge from precipitation is generally limited because of

high rates of potential evaporation and other factors described in section 2.2.1. On the other

hand the replenishment of groundwater by rivers in flood is frequently the major source of

recharge. Temporary rivers are formed in the valleys, or wadis, following intense storms in

the hills which are sufficiently severe to generate surface runoff. These temporary rivers

may terminate either in spreading zones where the flood water infiltrates to the aquifer below,

or in chotts or sebkhats which are low lying areas where temporary lakes are formed. Water

that accumulates in these depressions evaporates leaving behind its salt content. In both

cases the aquifers are recharged mainly in the foothills, or piedmont zones, where the

surface runoff is concentrated and where topographical conditions and soil permeability tend

to be more favourable for infiltration to the saturated zone.

Several factors combine to enable recharge to take place in the piedmont zones:

1. in such areas there is a thickness of permeable detritus comprising sand,

gravel and talus (detritus fallen from a cliff face),

2. the beds of the wadis are higher than the groundwater table,

3. water may flow horizontally through the banks,

4. the surface water spreads out over the ground thus accelerating the process

of infiltration and subsoil saturation,

5. the finer sediments that could impede infiltration are carried to the

downstream periphery of the recharge zone.

Groundwater recharge from temporary rivers is very irregular in both time and space,

just like the storms that produce it. In contrast to direct recharge from precipitation it is

relatively localised and concentrated with a rapid and divergent groundwater flow at the

point where the valleys open out into the plain. After each flood event there is a period

18


during which the aquifer is recharged causing a rise in the water table in that area. The

observed changes in the piezometric surface are the result of the superimposition of the

recent and previous flood events, so that the effect of recharge from a specific flood is

superimposed upon the preceding recession of the water table.

The rise in groundwater levels is related to the size of the flood. However, there is a

delay in the response of the water table due to two factors. Firstly there is a delay due to

the thickness, permeability and porosity of the unsaturated zone, and secondly the horizontal

propagation of the flood wave in the saturated zone is related to the diffusivity of the

aquifer.

A recession of groundwater levels follows the rise caused by infiltrating flood water.

When flow through the unsaturated zone ceases, there is a recession in groundwater levels

until such time as the next major recharge episode. In areas close to the wadis the variabil

ity of inflows may be sufficient to prevent groundwater flow from reaching the steady state

condition, and the effect of major floods may be apparent even after several months. Further

away from the recharge zone the amplitude of groundwater level fluctuations decreases and the

flow approximates to or reaches a steady state condition.

In areas where aquifers are recharged by rivers or canals, the safe groundwater yield

(Q) may be expressed as (Bochever, 1979) :

Q = Q e + Q± 2.2(1)

where Q is that part of the yield derived from natural groundwater sources and Q. is the total

inflow from other sources such as rivers. The determination of Q. is dependent primarily upon

analyses of the interaction between groundwater and river water.

2.2.3 Recharge by Lakes

In the United States a large number of small reservoirs are being built and small lakes are

increasingly being used as a focal point in urban planning. This has given rise to pollution

and amenity problems that for their solution require some understanding of lake hydrology of

which the interaction between lakes and groundwater is an integral part (Cherkauer, 1977) .

In many studies of lake hydrology the precipitation, evaporation and inflow/outflow data

are available. However, evaporation assumptions in particular may lead to errors in the water

balance. If the residual is allocated to groundwater effects then serious misunderstandings

could arise concerning the interaction of lakes and groundwater. To investigate this inter

action, numerical model simulations were carried out (Winter, 1976).

Most natural lakes in the United States are caused by glaciation, and the studies by

Winter (1976) apply especially to those conditions. The lines separating the various types of

flow system, or divide, were obtained for various situations. In each case the groundwater

levels surrounding the lake were assumed to be at a higher level than the lake surface, and a

point was located on the divide where the head is a minimum. This minimum head may occur

19


beneath the shoreline on the downstream side of the lake and is called the stagnation point.

The relationship of the head at the stagnation point to the lake level is fundamental to

understanding the interaction of lakes and groundwater. If the head at the stagnation point

is greater than the lake level it is impossible for water to move from the lake to groundwater.

If a stagnation point is located then the divide is continuous, the lake cannot leak, and it

is the discharge point for the groundwater flow system. Alternatively if there is no stag

nation point then the lake can leak through part or all of its bed.

2.2.4 Artificial Recharge

To increase the natural replenishment of aquifers, man has used artificial recharge in

addition to those methods described in section 2.2.2 (rivers and canals). Natural infiltration

may be augmented in two ways. The first is through surface works, including recharge lagoons,

ditches, the building of low dams to cause flooding of riverside tracts, and excess irrigation.

These methods are, as with natural infiltration subject to evaporation losses and may occupy

large areas of land. The second means of augmentation is to inject the recharge water directly

into the aquifer through shafts and boreholes. While this method avoids evaporation losses and

reduces land use, there is the disadvantage that recharge water often requires extensive

treatment before injection to avoid serious clogging of the recharge wells. The normal source

of recharge water is surface runoff, but treated effluents and cooling water have been used.

Artificial recharge dates from early in the nineteenth century in Europe and near the

end of that century in the United States. More recently the experience in the United States

has been summarised by Todd (.1960) , in Israel by Harpaz (1970) and in the United Kingdom by

Rodda et dl., (1976). In arid and semi-arid regions, such as parts of the western United

States, salinity increases have been observed in both groundwater and surface water due to the

effects of irrigation practices. Much of the irrigation water is lost by evaporation, but some

recharges the aquifers and provides an increment to river flow. In these areas the

concentration of dissolved solids tends to increase and may reach a level intolerable to many

crops. In several places this effect is so pronounced that the quality of water rather than

the amount available restricts water use. This has led to the development of computer models

to predict changes in dissolved solid concentrations in response to varying hydrologie stresses

(Konikow and Bredehoeft, 1974).

The effect of excess irrigation upon aquifer recharge is such an important issue in arid

and semi-arid climates that a case study (section 5.3) concerning this subject is included in

this report. However, the subject area of irrigation and groundwater recharge is covered by

other IHP Working Groups (5.4 and 8.2) and it is therefore not covered in further detail in

this report.

2.3 Groundwater Component of River Flow

The groundwater component of river flow is derived from continuous and intermittent flows from

aquifers that drain to the river under varying degrees of hydraulic connection. It is the term

used to describe that part of river flow that has been formed by the complicated processes that

20


result in groundwater inflow. The main features of the interaction between surface water and

groundwater may be identified as a specific part of the hydrological cycle. On a regional

scale the characteristics of these main features, including groundwater inflow to a river, may

display a marked regularity in both space and time.

It is customary to subdivide the component of river flow derived from groundwater into

continuous base flow from the main aquifers and intermediate flow, or sub-surface runoff, from

temporary storage. However, it is frequently difficult to estimate quantitatively the varying

properties of base flow, short-term groundwater flow and surface flow, or direct runoff, that

are present in the measured river flow. These proportions tend to change due to the different

rates of recession that characterise each flow component. The recession of short-term ground

water flow is more rapid than that of base flow, but slower than that of surface runoff.

However, the accurate estimation of each component of flow can be completed for specific rivers

only on the basis of complex water balance investigations in representative and experimental

basins (Toebes and Ouryvaev, 1971; Brown et al., 1972). Because of this it is more usual to

group together short-term groundwater flow and surface flow as direct runoff.

The groundwater component of river flow may be subdivided according to its origin (i.e.

its genetic parts) with the detail depending upon the availability of hydrological and

hydrogeological information. Improvements in the methods of assessing the interaction between

surface water and groundwater should be based upon more subjective and detailed separation

of the groundwater component of river flow. Firstly, the base flow component should be

identified by considering the river flow and basin characteristics. For this purpose

conceptual models of groundwater flow to rivers are proposed based upon the classification

shown in Table 2.3.1.

In constructing conceptual models it is important to consider the extent to which

aquifers contribute to river flow. Also care should be taken to differentiate between the

single-aquifer and multi-aquifer system. Difficulties may be encountered when estimating the

base flow components of multi-aquifer systems, since it is then necessary to identify the

contribution from each aquifer on the basis of available hydrogeological information. In the

absence of sufficient information it is good practice to take account of the contribution from

the main aquifer, or at least to estimate the total groundwater inflow without attempting

further division.

Conceptual models that are used to estimate the components of groundwater runoff should

be based upon the available and essential observational data. If such data are not available

then an objective schematization of the complex natural conditions (i.e. that occur in the

interaction between surface water and groundwater) should be adopted together with the

application of simplified schemes (Toebes and Ouryvaev, 1971; Popov, 1969b; Dobroumov et al.,

1976) .

In some river valleys the position of the water table may vary in relation to the river

stage at different points along the valley. This may give rise to local groundwater inflow,

river water outflow and underflow through permeable deposits beneath the bed of the river. By

21


carefully siting river gauging stations it may be possible to minimise some of the

complications arising from these local conditions.

Table 2.3.1 Classification of groundwater discharge to rivers

Class Type Source of recharge

Hydraulic Connection

Present (P), Absent (A)

1. From unconfined aquifers

A. Intermittent Groundwater (inter-flow)

Temporary perched water ('Verkhovodka') in mountain rock

Water of raised bog

Water from intermittent springs and geysers

Intermittent flow from aquifer overlying permafrost

Melt water from groundwater frozen at surface ("aufies")

Return water or bank storage (flowing period)

Phreatic groundwater

Continuous flow from aquifer overlying permafrost

Groundwater flow between aquifers

Water flow below permafrost zone

B. Continuous

P, A

P, A

P

A

P, A

P, A

P, A

2. From confined aquifers (artesian)

A. Open flow

B. Close flow

Water of fen soil

Water of constant springs (spring flow)

Confined water, upper spring water discharging directly into the channel

Confined water moving into the overlying aquifer

P

A

P, A

2.4 Influence of the Interaction on Water Quality

Abstractions from surface water and groundwater for supply purposes are limited by both

quantity and quality considerations. Whenever there is a flow of water between the surface

and aquifers, in either direction, there is a relationship between the quality of water in the

two systems. Where pollution is tending to increase due to man ' s activities an understanding of

22

definition of the Interaction

the interaction is essential to reduce the effects of such pollution. Groundwater can pollute

surface water and surface water can pollute groundwater. Alternatively there may be

improvements in quality.

2.4.1 Surface Water to Groundwater

The various methods of natural and artificial recharge of groundwater are described in

section 2.2. Whenever groundwater is recharged by the infiltration of surface water, the

quality of the former depends to some extent upon the quality of the latter. In natural

conditions the recharge of groundwater from surface water tends to cause some reduction in the

quality of the groundwater with a consequential decreases in its usefulness. However, this

reduction in quality may be minimal because the infiltrating water receives some purification

caused by physical, biological and chemical processes, as it passes through the unsaturated

zone.

Infiltrating water is mechanically filtered and some substances are adsorbed. Biological

purification takes place either by oxic or anoxic dissimilation. The microbes on the soil

particles tend to exert a greater purifying effect the longer the water remains in the soil

stratum, and the slower the water flows. The most important chemical reactions are those

involving carbon, nitrogen, calcium, iron, manganese and sulphur. These reactions depend upon

the redox properties of the substances, but biological interference may change the approach to

equilibrium conditions as determined from thermodynamically known potentials. Thus the

reactions may have an importance which differs from that in the purely physico-chemical system.

The purifying activity in surface waters always depends upon the oxygen content. The

activity of microbes reduces the oxygen concentration with a resultant rise in the carbon

dioxide concentration. Firstly, the microbes use up the dissolved oxygen, then use organically

bound oxygen and the oxygen in nitrates and sulphates. Nitrate will be reduced only when the

oxygen content is less than 0.5 mg/1.

Organic substances in the infiltrating water disintegrate rather quickly as shown by

decreasing permanganate numbers. However the humic fraction does not disintegrate but forms

humâtes with metal compounds which become bound to the soil. In the Nordic countries a problem

exists because much of the surface water is derived from swamps and contains much soluble humic

material which is not retained by the soil but filters through to the groundwater. Table 2.4.1

contains the mean concentration of various substances in surface water and groundwater in

Finland.

Agricultural fertilizers may have some influence on the quality of groundwater. Organic

nitrogen is readily oxidized to nitrate after passing the ammonia stage, and groundwater may

contain all the oxidizing stages of nitrogen. Phpsphorus readily becomes closely bound to the

soil and thus groundwater tends to contain very little phosphorus.

Iron and manganese frequently detract from the usefulness of groundwater. These elements

have a solubility which depends on the redox potential and the pH value. Because biological

processes determine the redox state, certain organisms will influence the solubility of iron

23


and manganese.

Table 2.4.1 Changes in composition of water from a sandy soil in Finland based upon

analyses for snow-melt, lysimeter water and groundwater

Chemical Determinands

pH

Electrical Conductivity

NO - N

NH. - N

NO - N

PO, - P 4

CI

Total

Hardness

S 0 4

Na

K

Ca

Mg

Mn

Cu

Pb

Unit

mS/m

M / l

ug/1

M / l

ug/1

mg/l

mmol/1

mg/l

mg/l

mg/l

mg/l

mg/l

M / l

ug/1

ug/i

Snow-melt

S w

4.4

2.3

410

230

5

8

0.6

0.02

2.0

0.3

0.2

0.4

0.1

25

3

8

Lysimeter Water

L w

7.4

22.6

73

3

1

7

1.0

0.53

28.0

1.6

1.5

9.3

0.9

72

7

4

Ground-Water

G w

7.4

23.0

73

3

1

7

1.0

0.8

2.6

1.5

0.8

5.4

0.8

110

190

30

Change in Value

L -S w w

3.0

20.3

-337

-227

-4

-1

0.4

0.51

26.0

1.3

1.3

8.9

0.8

47

4

-4

G - L w w

0

0.4

0

0

0

0

0

0.27

-25.4

-0.1

-0.7

-3.9

-0.1

38

183

26

G -S w w

3.0

20 .7

-3.37

-227

-4

-1

0.4

0 .78

0.6

1.2

0.6

5.0

0.7

85

187

22

The influence of the-quality of surface water on groundwater is primarily determined by

the time lag and distance of flow through the unsaturated zone. In general the quality of

groundwater is quite good if the delay is two or three months or more, depending upon the

composition and permeability of the soil and underlying aquifer. When groundwater is recharged

from watercourses, the quality of the groundwater tends to improve with increasing distance

from the recharge area. However, in arid areas salinity may increase where evaporation occurs

from groundwater, such as in 'sebkhats' as described in section 2 . 2 . 2 . Geological factors

including the structure of the aquifer and the mineral composition of the soil and bedrock also

influence water quality. More substances will dissolve from minerals formed at high temp

eratures than from minerals which are less easily attacked and have crystallised at low

temperatures. Minerals that are easily attacked include micas, dark thermal minerals and

limestones which readily dissolve in water containing carbon dioxide. If the water is very

hard, calcareous deposits may form. Iron and manganese dissolve in reducing environments, but

may be precipitated when the oxygen content of the water rises. Most light halic minerals,

such as quartz and feldspar, which are the main minerals in granite, gneiss and quartzite will

withstand chemical attack best, and little will dissolve from these minerals.

24


Two relatively common forms of groundwater pollution due to the activity of man arise

from waste disposal and saline intrusion along coastlines and estuaries. Pollution from waste

disposal occurs from a wide range of man's activities such as domestic sewage, industrial

effluent and waste disposal tips. In addition serious pollution can arise due to accidents

during transportation of chemicals if these occur adjacent to an aquifer. The storage of

radio-active waste poses special problems due to the long life of the pollutant, its potency

and the uncertain rate of groundwater flow at appreciable depths below the surface of the

ground.

Saline intrusion is likely to occur if the water table is lowered by groundwater

abstraction at sites adjacent to the ocean of other salt water environments. For example there

are long stretches of coastline in England along which aquifers are in contact with the sea,

and the pumping of groundwater has resulted in saline water moving inland at a number of sites

including the Humber, Mersey and Thames estuaries and along parts of the south coast. In

Israel seawater may penetrate at depth to the Jordan-Dead Sea Rift Valley (Kafri and Arad,

19 79). Aquifers can also be contaminated by the upward flow of fossil brines where these occur

at depth below freshwater.

2.4.2 Groundwater to Surface Water

During prolonged periods of dry weather a high proportion of river flow tends to be derived

from groundwater seepage. Thus the quality of groundwater frequently tends to dominate the

quality of dry weather river flows. Groundwater is generally of good quality but if it is

polluted then there is the risk of surface waters becoming polluted, especially during low

flow conditions when there is a minimum of dilution of base flow. A relatively common example

of river pollution by groundwater is that caused by the discharge of mine drainage to water

courses. This type of pollution may occur when minewater is pumped or when there is a natural

overflow from a disused mine.

Mine drainage can effect both the flow regime and the quality, tending to be relatively

constant throughout the year and during dry periods may contribute significant flows to rivers.

However, in England the major effect results from the quality of mine drainage (Rae, 1978) .

The River Pollution Survey of England and Wales (1970) shows that a large percentage of

polluted and poor quality watercourses are in the coalfields. This is in part due to mine

drainage. In a typical mine-drainage water the concentration of chlorides, sulphate, calcium,

total dissolved solids and occasionally iron will be several hundreds of milligrams per litre.

This tends to decrease rapidly downstream of the discharge point leaving a ferruginous deposit

on the bed. Although this deposit may not be totally destructive to the local biological

system it is unsightly and may inhibit the use of rivers for water supply.

The hydrograph of total river flow can be divided into its main components of base flow

and direct runoff as described in section 2.3. The characteristics of flood waters are

frequently different to those of low flows (Subramanian, 1979). If each component tends to

retain its own characteristics of quality and temperature then it is possible to construct

mathematical models of river flow based upon hydrograph separation techniques and water

25


quality considerations. Conservative or nearly conservative determinands such as alkalinity

(as CaCo.) and ortho-phosphate (as PO.) have been modelled with some measure of success.

In arid regions the available water, because of its scarcity, may be used several times

for various purposes. The re-use of water can cause quality problems which may be associated

with the cycling of water from the surface to groundwater and then back to the surface.

Excess water applied for irrigation purposes may infiltrate to the water table, reach the

surface water channels as base flow then be abstracted and used again for irrigation. This

has caused severe water quality problems and reduced crop yields because of the build up of

salt in the soil.

Quality and quantity changes may occur in surface water as a result of changes in land

use, such as changes to or from arable, forest or urban environments. In arid and semi-arid

areas a significant increase in salinity may occur in surface runoff after natural vegetation

has been removed for agricultural or other purposes. This process has been observed for

example in south-western Australia (Peck and Hurle, 1973). The removal of forest cover could

reduce evaporation, increase aquifer recharge and increase stream flow; but the associated

rise in the water table could cause some pollution of the aquifer by bringing the water table

above a zone containing saline deposits.

Another form of groundwater and consequential surface water pollution may occur from

inorganic fertilizers, sewage effluent and atmospheric sources. In England and Wales

atmospheric sources provide the greatest amount of nitrogen annually followed by animal and

human wastes and inorganic fertilizers (UK, CWPU, 1977). Although inorganic fertilizers

contribute the least to the total it is this source that has caused concern because of its

steady increase from two per cent of the total in 1933 to rather more than 25 per cent in 1972.

The slow build up of nitrogen, or other substances, in groundwater can create surface water

quality problems especially at times of low flow.

26

Methods of Assessing the Interaction

3. Methods of assessing the interaction 3.1 Channel Water Balance

The interaction between surface water and groundwater may be determined by analysing their

regime features throughout a drainage basin. International guides have been published that

enable such studies to be carried out based upon water balance investigations (Brown et al.,

1972:, Sokolov and Chapman, 1974; Toebes and Ouryvaev, 1971). Generalized features of the

interaction are reflected sufficiently in water balance calculations for channel networks, to

enable objective studies of the interaction to include the compilation and analysis of the

Channel Water Balance (CWB) for specific river reaches and river systems (Anon, 19 77a).

To estimate the CWB elements, observational data are required and the most rational

method of calculation must be used consistent with the characteristics of each channel reach

being studied-. The independent determination of the CWB elements provides the most

comprehensive information concerning the relationship between surface water and groundwater

and the characteristics of their interaction.

The elements of the CWB equations are determined from a consideration of the character

istics of the regime for each reach and river system. Accordingly, various observational data

for estimating the water balance and solving the CWB equation are obtained from valleys, flood

plains and channels. In the absence of observational data the corresponding CWB elements can

be determined by less rigorous methods. The values of elements that are within the limits of

the error of their definition are not included in the CWB computation. Channel Water Balance

computations may be based upon a month, a year and for typical periods of the hydrological

year. All the elements necessary for the CWB computation are defined in terms of the mean

discharge for a given period with an indication of the quadratic error (see section 4.3.1).

Examples of the CWB compilation have been described for examining various hydrological

problems including studies of the interaction between surface water and groundwater (Anon,

1977a; WMO, 1975) . In the majority of cases the most appropriate method for estimating

groundwater flow and river flow is that based upon the CW3 using hydrometric data (WMO, 1975).

27


3.1.1 Compilation for River Reaches

The elements of the channel water balance are calculated by using the appropriate equation for

each type of river reach. Thus four reach types may be defined taking into account natural and

artificial factors.

1. without flood plain, reservoir and water intake,

2. without flood plain or reservoirs but with water intake for irrigation

or other purpose,

3. with flood plain or reservoir and without water intake,

4 . with a considerable flood plain.

When using the CWB method to study the interaction between groundwater and river water,

the reaches should be chosen with homogeneous conditions of water exchange between rivers

and aquifers. This enables a simple interpretation to be given to the CWB estimate in a

design period. Therefore a further four types of reach should be identified:

a. with a continuous groundwater inflow to the river,

b . with a continuous outflow of river water to groundwater,

c. where groundwater inflow may alternate with river water outflow (eg with bank

storage phenomenon and groundwater table depressions adjacent to the river and

below the river stage during the low flow season),

d. with sub-channel stream flow.

3.1.2 Equations for River Reaches

The CWB for reaches of type 1, i.e. without flood plains, reservoirs or intakes, is computed

from equation 3.1(1).

2l +Qlr - 2 2 í Q g l í Q u í 2 0 i Q w = 0 3.1(1)

where Q . and Q ? are the discharges at the upstream and downstream cross sections respectively,

Q^ is the intermediate inflow. Er

0 is the channel regulation discharge, minus when water accumulates in the reach

and plus during the abstraction periods,

Q , is the allowance for ice formation (minus) or ice melting (plus),

Q is the exchange between the river and aquifers, plus for inflow of groundwater

to the river and minus for the reverse flow.

28


0 is the residual or remainder term that characterises the discrepancy in the o

water balance equation due to computation errors and incomplete account taken

of the CWB elements.

The sign of the residual term is defined on the basis of the relationship between the

CWB elements thus :

*o *2 *1 *Er - *w - *gl - *u

3.1(3)

The CWB for reaches of type 2 is computed using equation 3.1(2).

0, + Q„ - Q 0 +Q - Q„ + Q + Q + Q = 0 3.1(2) *1 Er "2 *<r *ß _ «gi _ *u - *o

where 0 is the total abstraction at a water intake in the reach, oc

Q is the total water returned to the river, p


Q, + Qv + Q - Q n - Q - Q„ + Q + Q + Q + Q + Q + Q = 0 *1 *£r *p *2 *EL Tit - *w - *gl - *u - *AM - AG - o

where Q is the river flow due to channel precipitation,

Q„T is the total evaporation from the water surface and transpiration from EL

vegetation along the reach that draws directly on channel storage,

Q is the evaporation from the flood plain and reservoir banks,

Q is the discharge corresponding to changes in the soil moisture storage

of the zone of aeration,

Q is the change in the groundwater storage in the flood plain and reservoir

banks.

The values of Q. and Q, are negative when the storage increases and positive when AM AG

storage decreases.


*1 Er p 2 « Qß * QEL - QEt Í Qw i Qgl Í Qu

- *AM - AG - o 3.1(4)

3.1.3 Compilation for River Systems

To study the interaction between surface water and groundwater for river systems, the CWB can

be compiled for the main watercourse to the downstream outflow point of the basin.

3.1.4 Equations for River Systems

The CWB for the main part of the river system is computed using equation 3.1(5)

29


EQEr + EQP - Q2 - Z^ + ZQ& - £ Q E L - 2QEt + E ^ + 2Qgl + EQU

± Q AM± 0 A G Î 0 o = ° 3 - 1 ( 5 )

where EQ is the sum of the inflows to the main part of the river system above the downstream Er

cross section or outlet,

EQ is the total water added to the main river system from precipitation on the surface P

of the river channel, reservoirs and flood plain along the reach being studied,

Q„ is the discharge at the downstream cross section or outlet,

EQ is the total abstraction at water intakes along the main river from its mouth up to

the outlet,

EQ is the total returned surface water to the main river up to the outlet, P

EQ is the total evaporation from surface water, including transpiration from EL

vegetation in reservoirs and along the flood plain of the main river from the mouth

to the outlet, that draws directly on surface storage,

EQ„ is the total evaporation from exposed or dry flood plains or from the banks of Et

reservoirs.

EQ is the total discharge due to channel regulation and runoff controlled by reservoirs

and the inundation of flood plains, located along the main river above the outlet,

EQ is the total water discharge due to ice formation and ice melting,

ZQ is the total water discharge involved in the water exchange between the main river

and aquifers along the reach up to the outlet,

£QAM an(3 £QAr

a r e the total water discharges corresponding to changes in the moisture

content of the soil and sub-surface zone of aeration, and groundwater in dry flood

plain reaches and reservoir banks along the main river above the outlet,

Q is the remainder term of the equation.

If there is no flood plain or reservoir in the main river up to the outlet, the CWB for

the main part of the river system is calculated by equation 3.1(6).

ZQEr - Q2 - ZQ« + Z S ± ZQw i ZQgl ± ZQu ± Qo = ° 3" 1 ( 6 )

3.1.5 Computation of the Elements

General information concerning the computation of the CWB elements is given in this section.

A more detailed description is given by Anon (1977a), Sokolov and Chapman (1974) and WMO (1975).

30


3.1.5.1 Exchange between Rivers and Aquifers

The Channel Water Balance method may be used for the assessment of that part of the interaction

between surface water and groundwater that relates to the exchange between rivers and aquifers.

Moreover the method enables the principal elements in the equation to be determined which are:

1. groundwater inflow to the river,

2. outflow from the river to groundwater,

3. water discharged to or from bank storage,

4. sub-channel stream flow.

The CWB elements concerning the various types of groundwater exchange included in the

equation are defined on the basis of the analysis of hydrological and hydrogeological

information including in particular observational data of river and groundwater stages. The

direction of the flow or exchange between rivers and aquifers depends upon the differences in

stage and slope of the groundwater table adjacent to the channel. Groundwater exchange is

estimated by hydrodynamic computations, water balance and other methods depending upon the

natural conditions and the availability of hydrological and hydrogeological data (Anon, 19 77b;

Bochever et al., 1969; Brown et al., 1972; Kudelin, 1969; Popov, 1969).

To provide detailed quantitative estimates of the interaction between river water and

groundwater is generally a complicated problem that requires for its solution special field

observations of the hydrogeology and of the regime (Toebes and Ouryvaev, 1971; Brown et al.,

19 72). Therefore it is expedient to estimate the exchange between rivers and aquifers using

hydrometric methods of differences by solving the CWB equation for the relevant elements. Thus

suitable channel reaches are chosen bounded by two cross sections within which the inflow to or

outflow from the river may be estimated (Kudelin, 1979; WMO, 1975). The remaining elements of

the CWB equation can be calculated when the values significantly exceed the respective

computational errors (see section 4.3) (Anon, 1974).

In the computation of the CWB a proportion of the runoff may be derived from bank

storage, Qh_f and some allowance may be necessary when the duration of a flood exceeds the

limits of a design period. If a change in river stage takes place during a design period

(such as one month) causing some increase in bank storage of a backwater type (Popov, 1969) ,

then quantitative estimates of Q, may be obtained from equation 3.1(7). In such cases bank

storage increases because of groundwater flow rather than outflow from the river to ground

water.

QbC = Qc - AQ 3.1(7)

where Q is the initial outflow from bank storage in a quasi-stationary regime (before flood or

spring flood),

AQ is the mean additional discharge of a design period as determined by water level

31


changes in the channel (Anon, 1977a; Shestakov, 1973).

Along the flood plain, or inundating reaches, the augmentation of groundwater storage is

defined at the expense of river water infiltration provided that the inundation occurs during

a design period and after the water levels have risen, i.e. a backwater and infiltration type

of bank storage is available (Popov, 1969). This value is introduced into the CWB equation

for reaches of types 3 and 4 and is computed by methods that have been described in some

detail (Anon, 1977a; Brown et al., 1972).

When computing the CWB for reaches with considerable sub-channel stream flow its values

at the upstream and downstream cross sections must be estimated. In addition its magnitude

relative to the error, or remainder, term in the CWB equation should be examined together with

probable errors in the discharge measurements. The systematic study and computation of sub

channel flow and its related problems has been described (Anon, 1977a).

3.1.5.2 River Flow, Intermediate Inflow, Abstractions and Returned Water

Techniques for the measurement of discharge and the procedure for computing runoff are given

by Toebes and Ouryvaev (1971), Sokolov and Chapman (1974) and WMO (1975). The natural

intermediate inflow between two cross sections is computed by summing the water discharge

adjacent to the flood plain in the reach, from rivers, streams or valleys using available

observational or theoretical data. The CWB computation allows for abstractions at water

intakes and returned water from various sources (Anon, 1977c; WMO, 1975).

3.1.5.3 Channel Regulation

Channel regulation water is that which accumulates in the channel, flood plain or reservoir

because of an increase in the stage. It is a particularly important component of the balance

during a flood if the design period is relatively short. However, with longer design periods

of up to a year the channel'regulation value may be approximately zero especially in areas

where the hydrological regime has a pronounced annual cycle. Various methods of computing the

channel storage suitable for the CWB calculation have been described (Anon, 1977a).

3.1.5.4 Precipitation

Discharge derived from precipitation on the surface of channels, reservoirs and flood plains

must be considered in the calculation of the CWB. This includes snow melt within these areas.

If there is a considerable flood plain or reservoir and a relatively insignificant river flow

then this is likely to be a major component. Alternatively in reaches with a negligible flood

plain, precipitation and melt water have only a small influence on the CWB and are thus not

included in the calculations.

3.1.5.5 Evaporation from Surface Water

This section includes all evaporation from surface water regions, such as from vegetation

that draws on surface water (riparian areas) dried reaches of the flood plain and reservoir

banks. A considerable proportion of natural flows may be lost by evaporation from the water

32


surface and transpiration by vegetation in reaches with extensive inundations of the flood

plain. However, in reaches without flood plains evaporation may account for less than one per

cent of the river flow and is then not included in the CWB calculation.

During a period of flood plain inundation and reservoir filling a proportion of river

water in the design reach is accumulated in the soil and zone of aeration thus increasing its

moisture content. After the water has receded from the flood plain and when reservoir levels

are drawn down, some of the accumulated water is evaporated and this has to be included in the

CWB calculation. These elements of the CWB have to be estimated in the most objective way on

the basis of the experimental data after investigating the water balance of the appropriate

water body. Methods of computing the discharge for the elements have been described (Anon,

1977a).

3.1.5.6 Change of Stored Moisture

During inundations of the flood plain and reservoir filling some of the water accumulates in

the soil and sub-surface zone of aeration, and a further part increases groundwater storage

below the flood plain. In a dry period some of the water is evaporated, flows away or

increases groundwater storage. This causes a change in the moisture content in the zone of

aeration and in groundwater storage which has to be estimated from observational data for

soil, sub-surface moisture and groundwater levels (Anon, 1977a; Brown et at., 1972).

3.1.5.7 Ice Formation and Melting

In the autumn and winter periods a proportion of channel water may form into ice, which on

melting in the spring increases the flow in the design reach. The effect on river discharge

due to ice formation and melting is determined from changes in the volume of ice as indicated

by observational data (Anon, 1977a).

3.1.5.8 Areal Definition of Flood Plain Reaches

To determine particular CWB components in reaches with flood plains (Q„,» Q x , Ç) , Q„„, Q,„) gj. bu at ¿AM Ala

it is necessary to include the following elements in terms of water discharge: precipitation,

ice, evaporation from inundated and exposed or dry reaches of flood plains, change in moisture

content of the soil and sub-surface zone of aeration and groundwater storage. The areal

definition of flood plain inundations and their exposed reaches enables these components to be

calculated.

3.2 Hydrograph Analysis

3.2.1 Flow Separation

Improvements in the methods of assessing the interaction between surface water and groundwater

are closely connected with the development of computation techniques to determine the various

genetic components of river flow. Flow separation should be carried out as objectively as

possible and only in association with independent estimates of the genetic components from

detailed water balance studies for representative and experimental basins (Brown et al., 1972;

33


Toebes and Ouryvaev, 1971) . However, difficulties arise due to the absence of methods to

estimate the separate components directly. Various techniques may be used to analyse the

outflow from small drainage basins but these tend to lack a systematic basis. Therefore in

regional hydrological studies schematic flow separations are used (Kudelin, 1969; Toebes and

Ouryvaev, 1971).

The more complex hydrological and hydrogeological methods of flow separation may be

used either in their analytical or graphic versions to assess the methods of calculating the

various genetic classes of the groundwater component of river flow (Dobroumov et dl., 1976;

Kudelin, 1969; Popov, 1969). The main points of this procedure are now described.

River discharge is measured during a period when flow is essentially derived from ground

water, such as during seasons of low flow and when the flow is receding slowly and has

specific properties. From the results of this study quantitative estimates may be made of the

water exchange between the river and aquifers for the design periods. These typical river

discharges are then transformed into groundwater recharge, or flow to rivers, for selected

periods to enable the coefficients of the relevant equations defining the groundwater discharge

(interannual dynamics) to be determined for the basin under study. These coefficients may be

derived using different methods based upon hydrometeorological and hydrogeological data

(Dobroumov et al., 1976). When estimating the regional inflow of groundwater to rivers by

this procedure, the choice of cross section for estimating the typical water discharge and the

dynamic coefficients of groundwater inflow should be made with regard to the natural

discreteness of the flow formation in the basin under study (Popov, 19 75; Popov, 1978).

For aquifers that are drained by rivers the spatial discreteness of groundwater outflow

is closely linked to the different groundwater stages. The drained levels conform to

particular base levels of erosion in the basin, with the groundwater divide defining the

limits of typical homogeneous characteristics that occur in the interaction between river water

and groundwater. The characteristics of the inflow of groundwater to rivers may vary from one

groundwater stage to another depending upon the characteristics of the aquifer and the degree

of hydraulic connection between the aquifer and the river. Various combinations of the above

methods may be used to investigate the regularity of the interannual dynamics of groundwater

flow to a river and thus its component parts may be estimated.

Contemporaneous discreteness makes it possible to propose the following comprehensive

equations :

i = n W = E WTT. 3.2(1) ru Hi

i = l

where W is the sum of the groundwater flow to the river,

W . is the groundwater flow to the river for each drained groundwater stage above the

outlet.

The groundwater flow to the river may then be determined for successive increments of

reduction in the groundwater stage.

34


To estimate the interannual dynamics of groundwater flow to rivers is the most

important and complicated requirement of this method. Fully objective results may be achieved

only from analyses of hydrogeological and hydrometeorological information together with an

examination of the relationship between the interannual dynamics of groundwater flow to rivers

and a number of factors.

Analyses of the main features of the formation of the groundwater component of river flow

show that its interannual dynamics are primarily determined by changes over a period of time

in the main parameters of the surface water and groundwater regime. These parameters are:

Groundwater stage G, and stage H in the river basin, slope Y of the water table controlling

flow to the river, river stage H and groundwater levels H where there is an hydraulic

connection, and also the velocity of their change in a flood period, V„ and V„ respectively

(Dobroumov et al., \°ilb; Popov, 1969a; Popov, 1969b; Popov, 1975).

The computation of the interannual dynamics of groundwater flow to rivers for the main

types of regime is illustrated in Figure 2 and may be obtained from the solution of

equations 3.2(2) and 3.2(3) (Dobroumov, 1976).

Model for descending regime:

W = f ru G(t), Hu(t), Yru(t) 3.2(2)

Model for backwater regime :

Vrl W = f ru

Hr V, Hru

(t), Yru(t) , G(t) 3.2(3)

The accuracy of such calculations may be obtained by using appropriate probability and

statistical methods as described by Zektser (19 77).

Q

Figure 2

\ <̂ 2 ^3 \ <\z V

The main types of groundwater inflow for various hydrogeological conditions

35


Characteristic values of river discharge:

ql is the groundwater inflow to the river prior to the beginning of the river stage

rise; in most cases this inflow may be adopted to be equal to the river discharge

during the antecedent low-flow period.

q2 is the groundwater inflow corresponding to the peak of the total streamflow

hydrograph,

q3 is the groundwater inflow after the flood when river flow comprises only

groundwater.

Descending type of regime: a - qi<q2; <32<t^3

b - q i < q 2 ; q2<q3

Backwater type of regime: c - q-i^n' <32<<^)' < 1̂<< 3̂

d - q.x<12i q2<0; q ^ ^

where 1 is the groundwater inflow to the river; 2 is the outflow from the river to

groundwater.

In general this method requires field investigations with periodic hydrometric surveys

of river flow (Ratner, 1972). New approaches to improve the methods of estimating the

regional groundwater inflow to rivers are based upon the fact that previous methods of runoff

separation must be replaced by models designed with due regard to the actual characteristics

of the interaction process in the basin under study.

3.2.2 Graphic Separation of River Hydrograph

In practice this method is generally the most appropriate for determining the surface and

groundwater component of river flow. The method consists of plotting a line on the hydrograph

which separates the two components, and the groundwater inflow is then computed by determining

the area beneath the line by planimetry.

When plotting the line of separation it is assumed that the initial rise in the hydro-

graph corresponds to the beginning of the surface inflow to the river. Then sometime after

the flood peak, a decrease in the rapid recession rate indicates an end to the rapid surface

runoff and there follows a period when flow is derived essentially from soil storage and base

flow. However hydrograph separation can be very complicated because the hydrograph may

represent the surface runoff from successive storms and the dynamics of groundwater inflow may

be complex. Practically all of the numerous procedures that exist for hydrograph separation

differ from each other in their determination of the co-ordinates of the line of separation in

a design period (Brown et al., 1972, Popov, 1975; Popov, 1978; Sokolov and Chapman, 1974;

Toebes and Ouryvaev, 1972).

All methods of hydrograph separation provide approximate solutions. Their accuracy

depends upon the extent to which the conceptual models used for the separation reflect the

36


actual interaction process, the characteristics of the interannual groundwater inflow and the

regime type. A primitive scheme of hydrograph separation is a straight line drawn from the

point where the hydrograph begins to rise to the point where the rapid recession ends (Roche,

1966). This scheme corresponds with that for the descending type of groundwater inflow regime

that has a slow recession, possibly some increase in a flood period but otherwise an

insignificant fluctuation in time.

One of the simplest methods of separating the groundwater component is by means of a

straight line (horizontal) or smooth curve through the low flow points on the hydrograph for

winter and summer periods ('cut' method). This enables an estimate to be made of the ground

water inflow for descending and backwater regimes without bank storage which typically show

little change in groundwater inflow during a flood period.

Estimates of the groundwater component of river flow, Q , may be based upon minimum

discharges during a specified season. This is an analytical version of the 'cut' method and

is calculated as follows:

Qw + Qs Qn=—T-Kd 3.2(4)

where 0 and Q are the minimum discharges in the winter and summer low flow seasons of 30 days

duration, and

K, is an empirical coefficient.

The coefficient K, is determined for separate river basins from the relationship of

discharges in the low flow seasons as defined by equation 3.2(4) and also from detailed

analyses of groundwater inflow values derived from hydrometeorologlcal and hydrogeological

data for selected gauging stations in the basin under study (Popov, 1970) . In areas where

homogeneous conditions exist for the formation of groundwater inflow, including even some of

the larger river basins, K, has a constant value which enables the groundwater component of

river flow to be determined with satisfactory accuracy for 95% of the time.

It has been proposed for complicated hydrographs, such as those for mountain rivers,

that the separation line should be drawn through the low points as an 'envelope curve' as in

Figure 3 (Friedrich, 1954; Natermann, 1951). A similar separation scheme is recommended for

flood flow regimes that have a duration of a year or nearly a year, because these have no

clearly defined low flow season and the base flow could be approximately zero. The use of an

envelope curve should be based upon an objectively chosen conceptual model of groundwater

inflow to the river with a descending regime, together with a careful analysis of the

hydrometeorological conditions in the river basin under study (Amusia, 1974).

Schematic hydrograph separation is the method used for rivers with a descending regime

of groundwater inflow and substantial flow variability over a long period as shown in Figure 4

(Amusia, 1974). In such cases it is not possible to compute the co-ordinates of the

separation line during the flood period from independent hydrogeological data such as spring

flows or hydrometric analysis of the regime. Instead the separation has to be based upon the

37


Q M*/C I

Figure 3 Hydrograph separation based upon low envelope curve

low flow characteristics from a long flow record, and hydrographs are examined from years with

average flows and typical interannual flow distributions. The separation line, defining the

groundwater component, is plotted through points representing the beginning and end of the

flood period and an intermediate point when the maximum groundwater inflow occurs. The

abscissa representing the maximum groundwater inflow should coincide with the midpoint of the

flood period, and the ordinate should be selected to equal the mean monthly maximum flow in a

low flow season as determined from a long flow record and a month when flow is mainly formed

from groundwater inflow.

Among the various methods that may be used for hydrograph separation, there are few

that are based upon conceptual models of groundwater inflow. The majority of methods therefore

are subjective.

The characteristics of the mean long-term groundwater inflow may be determined from the

hydrograph separation. To achieve this it is expedient in practice to complete the hydrograph

38


(V/c

100 .

90 •

30 •

70 •

60 •

50 .

U0 .

30 •

20 •

10 -

0 ^ I ' M ' III ' IV ' V ' V I ' VII ' V I M ' IX. ' X ' X I ' XII '

Figure 4 Hydrograph separation with maximum groundwater inflow during a flood period

separation for each observational year, or water year, and analyse data from four selected

years. Two of these years should contain data with mean flow conditions, and the other two

should be dry and wet years equivalent approximately to 1 in 4 years frequency of

occurrence. The average percentage value of the groundwater component in the total river flow

of these four years is assumed to be the mean long-term value. Then the mean long-term

groundwater inflow is estimated from the above percentage value and the mean total river flow

derived from long-term data. If the selection of these four years is thorough and contains

data with typical interannual flow distributions, then the proportion of groundwater in the

total annual river flow should be estimated accurately (Anon, 1973).

3.2.3 Recession Curve for River Hydrograph Separation

•Recession curve' analysis is widely used in hydrograph separation (Amusia, 1974; Appolov,

1974; Brown et al., 1972; Toebes and Ouryvaev, 1971). The term 'recession curve' is defined

as the descending limb of a hydrograph, when no floods occur and when river flows decrease

uniformly at a rate that depends upon the rate of depletion of groundwater and surface water

storage in the basin. That part of the recession curve that reflects the uniform decrease of

flow as a result of depleted groundwater storage is termed the groundwater 'depletion curve'.

The beginning of the depletion curve coincides with that point of the recession curve

where river flow comprises groundwater inflow with negligible surface runoff. There are two

main ways for determining this point:

1. by estimating the time lag for surface flow to reach the outlet on the basis of

hydrometeorological information such as air temperature, precipitation and ice

39


phenomena in rivers (Appolov, 1974; Toebes and Ouryvaev, 1971; WMO, 1975).

2. with the aid of graphical analyses of river discharge for successive time

increments, Q, = f (Q.), based upon recession curves, which have 11 + 1) t

characteristic rates of recession, or slope coefficient for each flow component

as in Figure 5 (De Wiest, 1965; Sokolov, 1974).

Figure 5 Hydrograph separation based upon depletion curves

When characteristic recession curves have been derived for river water of various

genetic types such as surface, groundwater and base flow, the curves may be applied to hydro-

graphs to determine the main components of flow (Amusia, 1974; Toebes and Ouryvaev, 1971).

Depletion curves can be expressed with certain assumptions in the form of exponential,

hyperbolic or parabolic equations, the equations of Bussinesko and Maye being used frequently:

*t o -Œt

ö t = Qo

(1 + 3t>"

3.2(5)

3.2(6)

where Q is the initial discharge,

Q is the final discharge after a time interval, t, and

œ and ¡3 are depletion coefficients.

v£ When <* and ß are constant in time, the value of the logarithm of p. in equation 3.2(5)

and/r^ - 1 in equation 3.2(6) vary linearly with time. In practice, the method of hydrograph

separation using depletion curves is carried out graphically by plotting on the hydrograph

the descending limb that represents the groundwater inflow (Toebes and Ouryvaev, 1971).

Each annual depletion curve may be extrapolated graphically and combined to form the long-

term average depletion curve of groundwater inflow. In flow ranges where curves overlap the

average values can be determined by graphical or analytical averaging. The 'normal' recession

curves are plotted for each year and generally cover a considerable range of flows with due

allowance made for the lag-time of surface runoff (Toebes and Ouryvaev, 1971).

40


To facilitate the extrapolation of recession curves one can use a semi-logarithmic scale,

the logarithmic one for discharge and linear one for time, which changes the depletion curve

into a straight line or a series of straight lines (De Wiest, 1965; Nutbrown and Downing,

1976).

During flood periods the depletion curve may be used to define the separation line and

estimate the period of maximum groundwater inflow. For this purpose the depletion curve is

extrapolated forwards in time at the start of the flood and back in time at the end of the

flood. The period of maximum groundwater inflow may be checked from regime data or from data

on the velocity of groundwater flow compared with that for river .flow. The chosen maximum

point is then connected by a straight line or smooth curve with the extrapolated depletion

curve.

In some cases it is necessary to estimate the separation line in a flood period by

extrapolating two depletion curves. The first has a relatively slow recession rate and is used

during the period when the total river flow is increasing, and the second is used to define

the groundwater recession after the period of maximum groundwater inflow as shown in Figure 6

(Riggs, 1953; Snyder, 1939). Use of the depletion curve analysis together with a water

balance equation for the river basin enables the flow separation to be completed in some

detail by analytical methods such as those described by Kalinin (1957).

Qn

2 5 •

za /

15 /

io A /

Figure 6 Determination of the depletion curve (1) and recessxon curve (2)

Improvements in the techniques of hydrograph separation using recession curves and

groundwater depletion curves must be based upon additional substantiating information,

including verification of the genetic classes of the flow components, conformity with the

hydrogeological basin characteristics as well as hydrometeorological and hydrogeological data.

3.3 Groundwater Table Fluctuations

Groundwater storage tends to fluctuate continuously in response to external factors in both

confined and unconfined aquifers. Corresponding fluctuations occur in the groundwater table

41


where variations in level may be either relatively rapid, such as seasonal, or long-term with

a duration of several years. The short-term variations may be caused by changes in atmospheric

pressure, earth tremors, precipitation or seepage from surface water. In addition human

interference such as the regulation of watercourses, drainage, and earth-works may cause short-

term or permanent and rapid changes in level.

This section is concerned with the variations in groundwater level caused by seasonal

changes in the hydrological cycle and with long-term changes caused by climatological

variations, mainly through their causal connection with surface water storage.

Groundwater storage changes occur because of differences between inflow rates to and

outflow rates from groundwater. This difference will vary in space and time, particularly from

one climatological zone to another due to different precipitation and evaporation patterns. In

some areas groundwater recharge may be derived predominantly from precipitation directly, and

in other areas from the infiltration of surface water. Groundwater level fluctuations in a

given locality tend to occur in a regular manner so that they may be used as an index of the

regime characteristics. Thus the groundwater regime may be classified by methods such as those

of Konoplyantsev and Kovalevsky (1963) which is based upon a water balance analysis and takes

-into account the interaction between surface water and groundwater. The main types of natural

groundwater regime of the USSR have been determined using this method. These main types of

regime are :

1. short-term recharge, mainly during the summer

2. seasonal recharge, mainly during the spring and autumn

3. annual recharge, mainly during the winter.

Divisions may be made within each of these three types to distinguish between the

various levels of recharge, whether of the abundant type of the temperate regions, scarce

type of the arid regions or hydrological types related to recharge from river flow. The

general characteristics of groundwater formation and its regime type depends upon various

natural factors. In the following sections the main features of groundwater level fluctuations

are described for two regions. Firstly for a temperate region (Nordic area) and secondly for

an arid region (northern Africa).

3.3.1 Temperate Areas

In temperate and humid environments where annual precipitation generally exceeds the annual

potential evaporation, groundwater is recharged mainly by precipitation. Groundwater flow in

these areas will tend to be towards rivers, where it becomes the base flow component of river

flow. Groundwater thereby combines with direct runoff to form the total river flow.

Where groundwater is recharged mainly from precipitation the changes in groundwater level

depend upon the characteristics and state of the unsaturated aquifer and overlying soil.

Important factors are the degree of saturation, effective porosity, permeability and the

42


distance that infiltrating water has to travel to reach the groundwater table. These factors

are responsible for a delay in the reaction of the groundwater table to rain. In the Nordic

countries the delay is usually small as the soil layer is generally rather thin and

predominantly Quaternary. This causes the groundwater table to follow closely the annual

rhythm of seasonal changes.

In the temperate areas of the Nordic countries the annual rhythm consists of three or

four distinct phases. The snow melt causes the groundwater table to rise with a maximum level

reached in March to May depending upon the locality. This is often the maximum for the year.

After this the groundwater table falls at a regular rate because of the effect of summer

evaporation. During the summer months evaporation is relatively high and there is a

negligible recharge of groundwater from precipitation. Groundwater recharge occurs again in

the autumn when rainfall exceeds evaporation and soils reach field capacity. Then in the

winter the surface of the soil becomes frostbound, infiltration more or less prevented and

the groundwater table again falls.

The effect of the soil frost in Nordic countries is to interrupt the hydrological cycle

for periods of two to seven months. This effect can be observed in groundwater hydrographs,

such as Figure 7 which shows the mean seasonal groundwater variation patterns for a number of

representative stations (Nordberg and Soveri, 19 78). The line showing mean groundwater levels

is based generally upon 10 to 26 years of record, and monthly mean precipitation and air

temperature is taken from adjacent official weather stations for the same period of observation.

There are four main groundwater regime patterns in the Nordic area. Denmark and southern

Sweden are in zone four, where the characteristic curves show a net recharge during the autumn

and winter months and declining groundwater levels for the remainder of the year. Proceeding

north from this area the increasing negative influence of the winter on recharge is apparent.

Over much of Sweden and Finland winter precipitation is in the form of snow, and since this

generally falls on frozen ground no groundwater recharge occurs in those months. This is also

true for much of Norway, but data from that country is at present insufficient to show the

different zones. Moreover, the steep topography and the influence of the Atlantic ocean pose

special problems that will have to be studied when more data becomes available. For these

reasons the lines on Figure 7 showing the boundaries of groundwater zones are not continued

into Norway.

Groundwater patterns in zone three indicate a secondary decrease in levels in the winter

which is observed mainly in a belt through south central Sweden. The characteristic patterns of

zone two show a more dominant influence of the winter, with the major recharge occurring in

connection with snow melt, and a minor recharge in the autumn before precipitation is in the

form of snow. Zone two occurs across north central Sweden, around the Gulf of Bothnia, in

southern Finland and in the adjacent parts of thé USSR. In zone one the characteristic pattern

indicates long winters with net groundwater recharge only during and immediately after snow

melt. This zone occupies the inland parts of northern Sweden, Finland and the USSR, with a

boundary some distance from the Baltic Sea (Nordberg and Soveri, 1978).

43


Figure 7 Mean seasonal groundwater variations in the Nordic countries and adjacent parts

of the USSR related to precipitation and air temperature

44


The main features of the groundwater recharge patterns are controlled by precipitation,

by the air temperature governing snow melt and by evaporation. The long-term natural

behaviour of aquifers is controlled by the long-term ratio of recharge to discharge. Whereas

the discharge from an aquifer is highly dependent upon hydraulic properties that do not change

with time, the recharge is controlled by such changeable variables as precipitation,

temperature and evaporation.

The succession of annual mean groundwater levels illustrate the relationship between

discharge and recharge that can be related to climatological conditions. A method developed

by Konoplyantsev (1970) may be used, that minimises the local effects of various hydraulic

properties within and around different aquifers and relates the annual mean level to the

observed maximum amplitude for the observation period. This relationship is given in equation

3.3(1) .

Hi ~ Hmin 3 < 3 ( 1 )

H - H . max m m

where X is the coefficient of relative groundwater levels,

H. is the annual mean level, and

H . and H are the observed extremes, mxn max

At the present time the Nordic groundwater observations generally cover an insufficient

number of years to satisfy the requirements for statistical computations of long-term mean

values. Figure 8 shows groundwater hydrographs for selected representative stations.

The annual mean groundwater level is determined by the climatological influence and by

geological factors. Figure 7 shows that the seasonal variations differ considerably from the

southern to the northern part of the area examined. The main qualifier is the snow melt that

generally occurs four to six months later in the north than in the south. Therefore the

winter accumulation of snow and the mode of melting exert a major influence upon the mean

groundwater levels. When proceeding northwards there is generally an increasing time lag

between precipitation and groundwater recharge.

The hydrogeological properties of the aquifer and the unsaturated zone control the

response of the groundwater levels to precipitation and melting snow. A deep unconfined

aquifer may react with a lag of several months or several years. The information presented in

Figure 8 is derived from aquifers with a relatively small and uniform lag of up to two months.

When making comparisons between aquifers the different lag effects must be considered that

influence the annual mean groundwater levels.

From Figure 8 it is apparent that there is a trend of decreasing groundwater levels in

south and south east Sweden, and in the USSR south of the Gulf of Finland, during the first

part of the 1970s. The same trend may also be observed in south western Finland, Denmark and

southern Norway, although the trend is rather weak in Finland and Denmark. This period of

declining groundwater levels coincides with a period of deficiency in precipitation.

45


Figure 8 Multiannual groundwater variations in the Nordic countries and adjacent parts of

the USSR

46


The effects have been recorded in aquifers of many types, small and large, deep and shallow,

and confined and unconfined. In 1977 the decreasing trend was halted and some recovery took

place. Within areas of generally decreasing groundwater levels opposite trends have been

observed due to local conditions. In the remainder of the area shown in Figure 8 no clear

regional trends are detected. However several records from northern Sweden show increasing

levels during the 1970s (Nordberg and Soveri, 1978) .

The absolute change in groundwater level is qualified by the effective porosity and

permeability, and may vary from a few centimetres to several metres. This aspect is not

discussed in this report.

Periodicities in the multiannual variations have not been detected at a meaningful level

of significance possibly due in part to the rather short period of observation. Zaltsberg

(1977) reported periodicities of 4 to 5, 8 to 11, 16 to 17 and 25 to 27 years for groundwater

in western USSR and a periodicity of five years within an observation period of 11 years for

groundwater levels in Finland (Soveri, 1973). Rather longer periodicities were detected in

Lake Saima in Finland, and lake Vanern in Sweden (Nordberg and Soveri, 1978).

3.3.2 Arid Areas

In arid zones annual rainfall is always less than the potential evaporation. In most

instances there is not even a monthly water surplus. Because of the marked iregularity of i

rainfall a water surplus may be evident for just one day or a week during the wet period which

for example in Tunisia occurs from September to April. The response of the water table to this

discontinuous replenishment in the rainy season is a seasonal fluctuation in the piezometric

levels characterized by an annual minimum and maximum.

In parts of north Africa groundwater recharge occurs after the autumn rains, the water

table reaches the highest level in the spring and there is a recession throughout the summer.

When the water table is some depth below the soil surface infiltrating water may take some time

to pass through the unsaturated zone. Lowest levels often occur during September or October

and the highest levels in March or April. The volume of water involved in this seasonal

fluctuation of the water table defines the annual regulating capacity of an aquifer. Thus the

relationship between surface water and groundwater may be quantified by analysing the steady

state groundwater recession of the summer and the subsequent, water table recovery of the rainy

season.

In arid zones the direct infiltration of rainfall to groundwater does occur, but is

negligible compared with the inflow from rivers. Recharge from floods occurs mainly in the

piedmont zones. Consequently, in the case of an homogeneous aquifer, water table fluctuations

will generally be most apparent along the upstream part of an aquifer. This upstream zone

tends to be relatively permeable and it is in this region that the aquifer is periodically

recharged from surface flows during successive rainy seasons.

The rise of groundwater levels in the recharge zone after each flood is followed by a

gradual rise in the surrounding water table. The speed of this sideways movement of groundwater

47


depends upon the aquifer diffusivity and always occurs some time after the flood event. When

successive recharge events occur close together then recharge water from one flood will be

superimposed upon water from previous events, thus delaying the recession of groundwater levels

and adding to the groundwater storage. With increasing distance down gradient from the

recharge zone the fluctuations progressively decrease, until in the lower plains the effect of

floods is not easily measured. If there is no direct recharge by rainfall in the lower plains

then the effect of evaporation is to create almost steady state conditions with water table

fluctuations of less than one metre. In the recharge zone fluctuations may be several metres.

Where there are a sufficient number of piezometers and in cases where the aquifer is

approximately homogeneous it is possible to estimate the annual volume of recharge from surface

inflows. The areal extent of groundwater fluctuations of a given amplitude are plotted and

measured by planimetry to solve equation 3.3(2).

R = A. S. h 3.3(2)

where R is the volume of recharge

A is the area of groundwater fluctuation

S is the storativity or coefficient of storage, and

h is the mean amplitude of the fluctuations.

It is apparent that a lack of precision in determining the storativity (S) may

significantly reduce the accuracy of the estimated recharge. Indeed, this method demands a

considerable amount of field data. Also in the case of over-exploited aquifers it is necessary

to take into account the effect of pumping on the piezometer levels, because the demands of

water for agriculture are not related to the aquifer stage.

When the aquifer replenishment occurs mostly along one main river, the variations in

storage from a major replenishment event may be defined in a similar manner, but fewer

carefully sited piezometers are then necessary. In this case the volume of recharge water

during a rainy period may be estimated as follows:

1. identify for each piezometer and between two recession periods the aquifer

recovery after a flood (or successive floods),

2. draw similar recovery curves for the aquifer to determine the sections

saturated by flood water. This is achieved by plotting groundwater levels

along sections perpendicular to the river for successive time increments,

3. planimeter the recovery zones and estimate the specific yield.

This procedure ensures that the record of groundwater levels at each piezometer is

carefully analysed so that the flow increment corresponding to the rainy period is separated

analytically from the initial total flow of the aquifer when it was in a steady state

condition. To assist in the separation of the groundwater hydrograph it is necessary to know

48


the natural recession characteristics of the aquifer. With this information an estimate may be

made of the groundwater levels if aquifer recharge had not occurred. The general form of the

groundwater recession is similar to equation 3.2(6) and is:

qt = qo e"Œ t 3.3(3)

where q is the groundwater flow at time t,

a is the groundwater flow when t = o, and

--f~»-£ -oct

In a laminar flow regime we also find that, h. = h e , where h and h represent the

hydraulic head related to the flow base level H. This base level is the asymptote towards

which h converges as t increases. To calculate the curve h (t) the first step is to estimate

H at any time during the steady state period, then the recession constant °= from the

experimental curve of groundwater level recession. The values of H and = depend upon the

position of the piezometer and on the diffusivity T/S at that point, and their determination

enables the characteristic recession curve to be drawn for the piezometer site. The recession

curves for all the other sites may be derived from a simple relationship based upon the delay

in groundwater recession. However, this method requires rather a precise knowledge of the

distribution of the specific yield for the whole area affected by the recovery of groundwater

levels. In addition, in zones far from the river relatively small variations in the water

table over large areas may be imprecisely estimated with a consequential decrease in the

accuracy of the inflow estimates. The main advantage of this method is its ability to

distinguish the contribution of each major flood to the aquifer replenishment.

Observations of groundwater level fluctuations over many years have shown that long

period variations exist which reveal the multiannual irregular characteristics of inflows.

There is a tendency for the random variation in rainfall to be stronger than the seasonal or

periodical variation with the result that seasonal variations in groundwater levels are not as

large as multiannular variations. The aquifers are able to store the inflows of successive

rainy years which slowly seep towards the lower plains and enter the next phase of the

hydrological cycle. Similarly, successive dry years may occur and result in a serious

depletion of groundwater flow. In parts of north Africa a ten year period has been observed

in the sequence of successive dry and wet years, and it is frequently observed that the

amplitudes of multiannual fluctuations are five times greater than the seasonal fluctuations.

The relationship between surface water and groundwater is difficult to evaluate in areas

with pronounced multiannual cycles because the variation of long cycles is more difficult to

characterise than short high frequency cycles. Long cycle variations may not be noticeable if

the period of observation is relatively short. Various integration techniques, such as the

'moving average' curve which smooth the amplitude of the series by filtering the high

fluctuations, may be used to analyse the variations of long inflow series as shown in Figure 9.

The relationship between inflows and the aquifer response may be examined by using moving

average techniques when these two variables are plotted with the same time scale.

49


800H

600 -

10

c ¡5

OC

(O 3 C C <

E 400 (0 0)

5

3 year moving average

rainfall

Groundwater Level

(in 0.1 m )

V. / Ma

i \ V \ /

h H \

I/ il

\ /

*y u ,•• "i

10 year moving average rainfall

V.

1920 1940 1960 Year

Figure 9 Groundwater levels and 3 and 10 year moving average rainfall in the Haut-Mornag,

Tunisia

Simple relationships may be derived between inflows (based upon rainfall or flood

discharge) and aquifer fluctuations using linear regression models with delayed effects. Long

series of data are frequently available for deriving such relationships which allow for the

effect of the aquifer inertia. These analyses include deriving a relationship (convolution

product) between the replenishment and a delay or transfer function which has assumed

parameters of length and shape. The transfer function is obtained from a knowledge of the input

series and output series (deconvolution). In some analyses an adjustment is made to the input,

flood inflows, and in estimates of the base level of the groundwater flow. However, these

techniques are essentially accurate only when applied to data from one piezometer and are

less useful for studying the phenomenon over larger areas. More complex models are suitable

for examining the relationship between surface water and groundwater in aquifer systems. These

include those based upon finite difference or finite element techniques.

3.4 Use of Isotopes as Tracers

3.4.1 Introduction

There are a number of problems associated with the interaction of surface water and groundwater

that may be solved conveniently by the use of a tracer, or built-in tracer, that reflects the

50


origin of the water. For example the question may arise as to whether pressure variations in

an aquifer indicate actual movement of water or a large range in values of aquifer permeability.

Tracers may be used to identify the mass transfer of water either from surface sources to

groundwater or in the reverse direction, and are useful in studies of water pollution.

The intentional injection of a tracer in groundwater is unlikely to give more than

relatively localised information on flow patterns except in cases of preferred flow paths in

fractured rock systems. However, the water molecules have built-in isotopic tracers that

reflect the origin of the water, and thus may be used to identify the occurrence and extent of

the movement of water between surface and groundwater systems. These natural tracers are the

stable isotopes deuterium and oxygen-18 and also the radioactive isotope of hydrogen known

as tritium, which although part of the water molecule, is not truly conservative because of its

radioactive decay. The term 'environmental isotope technique' in this context is defined as

the use of the variations in the environmental isotopic composition of natural waters as a tool

in hydrology. Other commonly used environmental isotopes in addition to the three already

mentioned include carbon-14 and carbon-13. The carbon isotopes are not actually part of the

water molecule but occur in water as dissolved inorganic carbon species.

This section summarises the causes which give rise to the variations in the environmental

isotopic composition of natural waters, the principles of their application to surface water-

groundwater interactions and examples of their use.

3.4.2 Stable Isotopic Composition of Natural Waters

The three main isotopic species of water, with their abundance in mean ocean water, are 16 16 18

H„ 0 (99.73%), HD 0 (0.031%) and H 0 (0.199%). The variations of the isotope ratios D/H •i O i /- ^

and 0/ 0 are measured in an isotope ratio mass spectrometer. The ratios are not measured

absolutely but are expressed in a S notation as a relative deviation from a standard.

6= ' ^ ' R s t d ' - 1 Q 3 3.4(1, Rstd

where R and R are the isotope ratios (D/H or 0 / 0) of the sample and standard

respectively.

The standard adopted is Vienna-SMOW (Standard Mean Ocean Water) which approximates to

the mean isotopic composition of the oceans. It will be noted that the delta values are

expressed in per mil. The usual analytical errors are 0.1%» and 1.0%« for oxygen-18 and

deuterium respectively.

When water changes phase, such as from vapour to liquid or vice versa, there is an

isotopic fractionation which results in a difference in stable isotopic composition of the

newly formed phase as compared to the other phase. The fractionation is greater the lower the

temperature. So in the condensation process for example there is a continual preferential

removal of the heavy isotopes deuterium and oxygen-18.

51


In practical terms this means that precipitation at higher latitudes is more depleted

than that at lower latitudes and winter precipitation is more depleted than slimmer

precipitation. Also, a very important property in applied hydrological studies, there is a

negative correlation of the heavy isotope content of precipitation with increase in altitude.

At any given season continental precipitation is more depleted in deuterium and oxygen-18 than

coastal precipitation. An overall global picture of the environmental isotopic composition of

precipitation is available from data obtained from the IAEA/WMO isotopes in precipitation

survey which commenced in 1961 (IAEA 1969; 1970; 1971; 1973; 1975; 1979). In the formation of

precipitation, deuterium and oxygen-18 behave in a similar way and the average global

correlation is:

ÔD = 8ô180 + 10 . 3.4(2)

Whereas the condensation process takes place in thermodynamic equilibrium this is not the

case for the evaporation process where the relationship given in equation 3.4(2) no longer

holds and the slope has a value in the range of about 4 to 6. Thus one finds that waters which

have been partially evaporated have a more enriched stable isotopic composition than their

feed water and fall below the meteoric water line defined by equation 3.4(2). In this way such

waters have a characteristic isotopic composition which may be used to study interactions

between lake water and groundwater.

3.4.3 Environmental Tritium Concentration of Natural Waters

Environmental tritium concentrations are normally expressed in TU (Tritium Units) where T —18

1 TU = — . 10 . Tritium (half life, 12.26 y) is produced by cosmic radiation in the upper H

atmosphere and also by the detonation of thermonuclear devices in the atmosphere.

Concentrations of tritium in precipitation resulting from cosmic ray production have been

estimated to be- about 10 to 20 TU depending upon geographic location. However, as a result of

the injection of tritium into'the atmosphere from thermonuclear devices, concentrations in

precipitation in the northern hemisphere were of the order of thousands of TU by 1963.

Concentrations in the southern hemisphere were very much less owing to the smaller number of

tests in that hemisphere and also the larger proportion of oceans to land mass. Since 1963

the concentrations have markedly decreased but in most parts of the world they are higher than

the estimated level of tritium due to the effect of cosmic rays.

Once precipitation with a given concentration of tritium has infiltrated into the ground

the tritium gradually decays and for all practical purposes no detectable tritium is present

after about 40 or 50 years. Surface waters forming from runoff from precipitation always

contain tritium. However, during prolonged dry periods when the proportion of groundwater in

river flow is high, quite low tritium concentrations may be observed. Because of the facts

outlined above it will be appreciated that the use of environmental tritium by itself is of

limited use for investigating the relationship between surface water and groundwater. However,

it may be useful when interpreting stable isotope data as a qualitative indication of the time

since recharge.

52

Methods of Assessing the Interaetion

3.4.4 Recharge of Groundwater by Rivers

The solution of a problem by isotope analysis in cases where groundwater is recharged by a

river is based upon the fact that a river contains water which has been derived from

precipitation falling on ground at higher altitudes than the area where river-groundwater

interactions are being studied. As a result of the effect of altitude the stable isotopic

composition of the river water will be more depleted than that of the groundwater if this has

been derived from the infiltration of precipitation falling directly on the area under

investigation. Therefore the stable isotopic composition of groundwater is compared with the

stable isotope indices estimated for the two potential sources of recharge.

The magnitude of the altitude effect for oxygen-18 is about -0.2 to -0.3%° per 100 m

change in elevation, that is two to three times the measurement accuracy. Normally a river

will contain water which has been derived from precipitation that has fallen at different

altitudes in the basin up to the maximum of the watershed. By the time the river has reached

the study area it will have integrated all the inputs from different elevations and the final

stable isotope index of the river water will reflect the oxygen-18 contents at different

elevations weighted by their respective volumetric inflows. Therefore it may be difficult to

estimate the stable isotope index purely from the topography of the basin.

A study in Ecuador illustrates the use of the isotopes as tracers (Payne and Schroeter,

1979) . The project area is about 20 km east of Guayaquil extending from Bucay, where the

River Chimbo leaves the Andes mountains, to Yaguachi in the western extremity. Shallow

groundwater was sampled at the end of the dry season from 51 open and closed wells fitted with

handpumps in a deltaic fan of sedimentary deposits in which the groundwater is about 3 m below

the ground surface (Figure 10). The variation in the stable isotopic composition of the

shallow groundwater extended over a range of 4%o and 30%o for oxygen-18 and deuterium

respectively (Figure 11). This indicated that the shallow groundwater is a mixture derived in

varying proportions from two sources each having its own characteristic stable isotopic

composition.

Figure 12 shows a histogram of the frequency distribution of the SD values of the shallow

groundwater excluding two samples which exhibit a marked effect due to evaporation. The

distribution is highly skewed with the maximum frequency close to the stable isotopic index

estimated for recharge by the infiltration of local precipitation. The lower frequency occurs

at delta (i.e. 6) values close to that of the river water. The latter being estimated from

actual river samples and groundwater which was clearly recharged by the river. Analyses of

water from individual wells indicated the proportions and areal extent of the river water

recharge to the shallow groundwater.

Morgante et al (1966) studied the infiltration of water from the Isonzo river to ground-18

water in the Gorizia plain in north eastern Italy. The 6 0 values of groundwater from 32 wells

were measured at various times between December 1963 and April 1965, and the Isonzo river water 18

was sampled between March 1964 and April 1965 at different locations. The mean 6 0 value was 18

close to -10%». It was apparent that the 6 0 values for groundwater became more positive with

53

Methods of Assessing the Intevaotion

•-25.7 • -24.1

O - 2 7 7

• -24.5 •-22.1

-29.10 9 # . 2 9 1 # . 2 5 1

-26.3 # . 3 A 8

\ * - 2 9 . 2

o OPEN WELLS

# . 2 7 5 • HAND-PUMPED WELLS

X Y A G U A C H I _^2 # -23 .6

Ly_ -46.8 • -29 9 T V MILAGRO0 -2°4 • "g3

#VVs/^~^2/ l7J -25.3 -258 ^

" 4 A 8 ^ ^ é - 2 6 2 . 0 0 - 2 7 ° O - ^ -23.5»^N^.557 • 0 . 2 7 1 SAN -2A.8. \ 3 7 ¿ ^ 6 o 2 4 6 C A R L O S

-23°2 . -261 ̂ i ¿ ß ^ ^ -23.6 Ä ^

•-29.5

N

\

• -28.1

• -29.6

• -30 0-24.9

• -29.3

^ r^T-^O CHIMBO BUCAy

_ - 4 0 . 9 • ••36.4 g —

RIO CHANCHAN ^ T • -276 V

• -43.8 N.

0 5 10Km -•vi ^ ^ 1

•

Figure 10 Deuterium content of shallow groundwater bordering the Rio Chimbo in Ecuador

-10T

-6 -5 -4

OXYGEN-18 %o

Figure 11 Stable isotopic composition of shallow groundwater

54


10--

>-O -f w R

a5

o LU er ü.

H—H- + 15 -55 -50 -45 -40 -35

DEUTERIUM %o -30 -25 -20

Figure 12 Frequency distribution of deuterium content of shallow groundwater

increasing distance from the river and that infiltration of river water occurred to the south

and west. The frequency distribution of the groundwater data was highly skewed with a maximum

at -7.25%o, the index value for recharge derived from local precipitation. The low frequency In

occurred at the 5 0 value similar to that for river water.

Within the context of a detailed study of the environmental isotopic composition of

groundwater in the Winnipeg area, Fritz et al. (19 74) report on the influence of river

infiltration to particular wells. Two wells located close to the confluence of the

Assiniborne and Red rivers were believed to receive infiltrated river water on the basis of

water level and chemical data. Oxygen-18 data showed that the Red river undergoes seasonal

variations which exceed 10%», while the values of the two wells remained essentially constant.

This indicated that little or no river water entered these wells. On the other hand another 18 well located in the inlet structure of a major flood basin has a variable 6 0 composition.

18 During high stage of the Red river the 6 0 value of the well is identical to that of the

18 river, while in summer the 6 0 value is typical of groundwater in the area.

Similar studies have been reported in many other areas and environments. Thus studies

by Vogel et al. (1975) indicate that in many Andean drainage basins the infiltration of river

water into gravels and sands of outwash plains is an important process for groundwater recharge.

Brown and Taylor (1974) report on a study of the Kaikoura Plain in New Zealand and demonstrated

55


the importance of recharge by infiltrations from the Kowhai river. The technique is also

proving valuable in arid zones for assessing recharge by flash floods in wadis.

An increasing number of municipal water supplies are indirectly or directly linked to

rivers. In many instances an improvement in water quality is possible if river water

infiltration is induced and the water pumped once it has passed through river gravels and sand.

However, if the wells are too far removed from the river they draw on both river water and

local groundwater. In a study of the water supply system of the city of Bern, Siegenthaler

and Schotterer (1977) demonstrated the increase of the groundwater component with increasing

distance from the Aare river.

To accurately estimate the proportion of river water in groundwater depends upon the

accuracy of the estimates of the stable isotopic indices of the two potential sources of

recharge and the difference between these indices. An estimate of the river index may be made

on the basis of samples taken from the river. This should be done at different times and river

stage to ascertain whether there are any significant variations in the stable isotopic

composition. If variations are evident then the mean value weighted for discharge should be

used. A quicker and preferable approach is to sample groundwater close to the river where

groundwater levels indicate that river water is the source of recharge. The estimation of the

index for recharge by infiltration of local precipitation may be based on measurements of

groundwater away from the influence of the river or, if sufficient data are available, on the

peak value of the skewed frequency distribution. If the errors in the estimates of the indices

of the two potential sources of recharge are not greater than the analytical error, then the

accuracy in the estimate of the proportion is better than 10%.

An example of the use of tritium and carbon-14 is provided by a study to assess the

environmental impact of the construction of a nuclear plant in the upper reaches of the Reno

river (Carlin et al., 1975).. The investigation assessed the relationship between the Reno

river water and groundwater in the area of Bologna. At two pumping stations the tritium and

carbon-14 content of the groundwater decreased with increasing distance from the river. This

demonstrated the contribution of river water to the aquifer. At a distance of about 700 m

from the river the effect of this contribution by the river was found to be negligible. At

another pumping station which was commissioned more recently and is located further from the

river, no influence of river water was detected. However, it is conceivable that as a result

of prolonged pumping water derived from the river may ultimately also reach this station. By

monitoring the tritium and carbon-14 content this possibility would be checked.

Conclusions concerning the origin of recharge water are based on the analyses of ground

water samples. This implies the availability of sampling points tapping specific horizons

that are sufficient in number and spatial distribution to provide a representative coverage of

the area under investigation. In practice the limitations of the method are not caused by the

method itself, but by the availability of meaningful samples.

56


3.4.5 Recharge of Groundwater by Lakes

The interaction between lake water and groundwater may be examined by isotope analysis

provided that the stable isotopic composition of the lake is significantly different from that

of recharge derived from local precipitation. Lake water undergoing partial evaporation is

enriched in deuterium and oxygen-18 along evaporation lines with a slope of four to six and

may therefore be differentiated from samples falling on the meteoric water line of slope 8.

In some cases the evaporation effect on the lake water may be minimal, but if the inflow to

the lake is derived from higher elevations then the problem may be treated in the same manner

as for river-groundwater interactions.

18 As early as 1962 Gonfiantini et al. published 6 0 data on groundwater in the region of

In Lake Bracciano in Italy. The average ô 0 of groundwater was -6.0%» while that of the lake

was 0.69%o which demonstrated the independence of the two systems. The samples covered the

period June 1960 to March 1961 and little or no variability outside analytical error was

evident.

Lake Chala is a volcanic crater lake located at 840 m on the SE slope of Mount

Kilimanjaro on the border between Kenya and Tanzania. The lake has neither surface inflow nor

outflow and in connection with the possible expansion of a nearby irrigation scheme the problem

was to know the relationship between the lake water and springs in the area (Payne, 1970).

Samples of groundwater were analysed for their stable isotopic composition from ten locations,

at different depths and on two occasions nine months apart. No significant variation greater

than the analytical error was found. Samples of spring water were analysed also and in most

cases over a period of three years.

Figure 13 shows that the stable isotopic composition of the lake is markedly different

from that of the groundwater samples which fall on the meteoric water line. It was therefore

concluded that none of the springs studied received any significant contribution of water from

Lake Chala. Taking into account measurement errors the maximum possible cpntribution to

individual springs would not be more than a few per cent of their respective discharges.

Fontes et al. (1970) studied the influence of water from Lake Chad on the groundwater

system in the neighbourhood of the lake. The groundwater close to the north east side of the

lake has a stable isotopic composition which is indicative of an appreciable contribution of

lake water but only over a relatively short distance from the lake. Further away from the lake

the groundwater becomes more depleted in stable isotopic composition which reflects the

increasing proportion of recharge from rainfall that infiltrates into the dunes during the

summer rainy season.

A quantitative assessment of the relationship between lake water and a groundwater system

is not possible with a small number of samples if there is some variability in the composition

of the lake and groundwater. In fact it is the seasonal variation of the isotopic composition

of lake water which is a limiting factor in the sensitivity of the method.

57


20.0T Ó D % O

LAKE CHALA

D Ô180%o

30.0-1-

Figure 13 Stable isotopic composition of Lake Chala and groundwater in the area

3.4.6 Groundwater to Surface Water

The tritium concentration in the river Thet, United Kingdom, was measured and compared with

the concentrations in rainfall and groundwater (section 5.1, Case Study 1). Analyses of

samples showed a seasonal trend in the tritium concentration for both rainfall and river flow.

Equation 3.4(3) is typical of those derived based upon 31 samples taken in 1968.

T = 41 + 0.7 T F d 3.4(3)

where T is the tritium concentration in river water, in TU

T is the tritium concentration in rainfall, in TU r '

F is the proportion of direct runoff in river flow.

Although this equation accounts for less than 70 per cent of the total variance in T, it

quantifies an important factor, that the tritium concentration in river water tends to increase

as the proportion of direct runoff increases. This is because the concentration in rainfall

tended to be relatively high, between 70 and 250 TUs in the example given, and the

concentration in groundwater is typically just above zero. In addition summer rainfall tended

to have a higher tritium concentration than winter rainfall with lower values frequently

associated with the more intense rainfall events. Thus there is a tendency for concentrations

in major flood flows to decrease.

In the summer months the tritium concentration in river flow is typically low in spite of

the fact that the value for rainfall is high. This can be explained by hydrograph separation

exercises which indicate that summer flows generally have a high groundwater component with a

58


low tritium concentration, and only a small proportion of summer rainfall reaches watercourses

due to the effect of evaporation.

3.5 Use of Mathematical Models

3.5.1 Purpose of Modelling

Quantitative assessments of the interaction between surface water and groundwater should be

based on mathematical modelling. In recent years computers have become increasingly available

to solve complex or tedious mathematical problems, and this development has increased the

understanding of the interaction process. Models of the interaction should be based on

conceptual models of the hydrological cycle, for specific conditions of the water body under

study and may include for example a coupled surface and groundwater model (Cunningham and

Sinclair, 1979). In particular an adequate account should be taken of the mechanisms that

lead to the formation of surface water and groundwater.

One of the main purposes of modelling the interaction process is to investigate whether

assumed parameters and mechanisms are realistic. Thus, for example, if the assumed interaction

between a river and an aquifer is incorrect, then this will result in an inadequate model

response in terms of groundwater heads or flows when comparisons are made with field data. In

such cases it may be possible to use the model to investigate the type of mechanism that should

be introduced. A frequent problem encountered in practice is insufficient reliable field data

and a simplified mathematical model may then be used.

Another purpose in modelling is to investigate the effects of human interference in the

surface water or groundwater regime such as the effect of groundwater abstraction, river

regulation and changes in water quality. Water quality models are not discussed in detail in

this report since they are covered by other IHP Working Groups such as 6.3 and 8.1.

When considering the effect of the groundwater level, it should be noted that it has

little influence upon the recharge mechanism except for shallow aquifers. In contrast the

groundwater level is of critical importance in the interaction between aquifers and springs,

rivers, canals and lakes. When water flows to or from an aquifer the quantity flowing is very

sensitive to the difference in head between the surface source and the aquifer, differences

of one metre having a very significant effect. A head difference of one metre is small

compared with a head difference of say 100 metres which may occur in some areas. This must be

seen in the context of modelling the groundwater head when the accuracy is usually no better

than three per cent or approximately three metres.

Another important feature is that springs, rivers or canals are small in dimension

compared with the overall aquifer. Thus a spring, river or canal is in effect a mathematical

singularity. The flow patterns are complex in the regions of the interaction between surface

water and groundwater because of the small size of the region of the interaction compared with

the overall size of the drainage basin or aquifer. Therefore it is advisable to use a

detailed model of the aquifer. Proven techniques are available for modelling regional

groundwater flow using finite difference or finite element methods (Prickett, 1975). There is

59


no basic difference in the methods and therefore, for simplicity of presentation, reference

will be made initially to the finite difference technique followed by an example based upon

finite elements.

3.5.2 Groundwater Recharge

There are many alternative methods of evaluating recharge to an aquifer. Some are based on a

soil moisture balance (Penman, 1949) whilst others consider the flow in the unsaturated zone

(Van Keulen, 1975). However the interaction between surface water and groundwater in the

outcrop area is often more complex than the above models suggest and they tend to underestimate

the quantity of water entering the aquifer due to precipitation. For example, the soil

moisture balance method of Penman assumes that recharge occurs only when the soil is at field

capacity, i.e. when there is no soil moisture deficit. Consequently it is estimated that in

temperate climates there is negligible aquifer recharge during the summer months. However,

observation well data indicates that recharge does occur sometimes during the summer as shown

by a partial recovery in the groundwater head. Furthermore, the quantity of recharge as

calculated by the above method often leads to excessive declines in the groundwater head when

incorporated in a mathematical model. Alternatively this may be due in part to errors in

evaporation estimates (Anon, 19 78).

A careful field investigation of outcrop areas indicates that a proportion of the

precipitation can enter an aquifer by-passing the soil zone. Sometimes this occurs when the

runoff from steeply sloping ground enters the aquifer via natural 'swallow' holes (Fox and

Rushton, 1976). In other situations natural drains lead directly into an aquifer. Though

field responses indicate that some water does enter certain aquifers by-passing the soil

moisture, it is essential to check such an assumed mechanism by incorporating it in a

mathematical model of the aquifer. Only then is it possible to check whether the correct

response occurs in the observation wells. Also the general trends in aquifer heads and flows

indicate whether the total recharge is of the correct order of magnitude (Spink and Rushton,

1979).

The depth of the water table below the soil zone can vary from a fraction of a metre to

hundreds of metres. Unless this distance is small, the passage through the unsaturated zone

can be important, particularly due to the effect of increasing nitrates in groundwater. Model

studies have shown that in some aquifers the flow mechanisms in the unsaturated zone are such

that the downward movement of water through the fissure system has a typical vertical flow rate

of 0.7 m/day whereas the vertical flow rate of the solutes is only 1 m/year (Oakes, 1977). In

this study the flow rate was estimated by an equation relating the response of the water table

to estimated infiltration at the surface. Approximately 15 per cent of the infiltration

reached the water table during the first month and the remainder in succeeding months. The

time taken for infiltration to flow through the unsaturated zone depends upon the distance

from the ground surface to the water table and the properties of the aquifer. In his

discussion on the rate of solute transport, Oakes considered the movement of both tritium and

nitrate profiles. If it is assumed that solutes diffuse from the infiltrating water into the

relatively static water in the rock matrix then this can account for the slow downward

60


movement of the solutes. This concept has been tested by experiments and model investigations.

3.5.3 Spring-Aquifer Interaction

The standard method of describing an overflowing spring is to assume that if the groundwater

head at the spring (h ) is above the elevation of the spring (Z ) then overflow occurs, and if a s

it is below no overflow occurs, and provided that there is a spring flow, then:

h = Z a s

However, if setting the aquifer head at the elevation of the spring results in water

entering the aquifer at the spring node, then the constraint is removed and Q = 0 , where Qs

3 is the overflow in m /day. These conditions can be applied during a numerical solution, but the procedure warrants careful examination.

It has been shown that a single node towards which radial flow occurs is equivalent to

a lake of radius 0.208 Ax, where Ax is the finite difference mesh interval (Rushton and

Herbert, 1966). Therefore when the groundwater head at the spring node is set equal to the

spring elevation and the grid spacing is 1 km, then the model represents a lake of radius

208 m, rather than a small spring. Clearly this will lead to an excessive spring discharge

unless the correction described by Rushton and Herbert is incorporated.

An improved method may be used to describe the characteristics of a spring if data is

available relating the flow from a spring to the groundwater head at a distance of

approximately 0.2 Ax from the spring. Typical field relationships are listed in Table 3.5.1.

From these values an expression can be deduced relating the flow from the spring to the excess

head. Under certain conditions a good approximation to the spring discharge for data in

Table 3.5.1 is given by:

Q = 130 (h - Z ) if h > Z when Z = 61 m 3.5(1) cL S 3. S S

Q = 0.0 if h « S 3.5(la) a s

In a mathematical model this can be included either as a flow dependent on the groundwater

head or as an equivalent hydraulic resistance.

Often a non-linear relationship exists between the flows and the groundwater head. This

can be represented by an expression of the form shown in equation 3.5(2).

Q = A (h - Z ) + B (h - Z ) for h > Z 3.5(2) a & d o cl a

where A and B are constants. This expression indicates that proportionally higher flows occur

as the groundwater head increases. The increase in flow occurring because of increased

transmissivity as the saturated depth increases. Under natural groundwater conditions

expressions of the form of equations 3.5(1) and 3.5(2) can be included in a mathematical model

without difficulty.

61


Table 3.5.1 Spring discharge related to groundwater head at 180 m from the spring

Groundwater head in observation

(m)

60.0

62.0

63.7

65.1

66.0

67.0

69.5

well Spring

discharge 3

(m /day)

0

130

350

530

640

790

1100

3.5.4 Rivers and Canals

The interaction between rivers or canals and aquifers introduces certain additional features.

Influent or effluent conditions can apply. A typical relationship is described by Prickett and

Lonnquist (1971) and is illustrated in Figure 14. Provided that the groundwater head (h ) is

greater than the stream bed level (ZL ) , equation 3.5(3) may be used.

Q = (k'/m') A (h - Z ) for h » Z^ 3.5(3) s a s a b

where k' and m' are the hydraulic conductivity and thickness of the stream bed deposits

respectively, and A is the plan area of the stream assigned to that node. If the groundwater

head falls below the base of the stream, the flow follows the vertical portion of Figure 14 (b)

thus:

Q = (k'/m1) As (Zb - Zg) for h & < Z^

groundwater head

water table

stream level

base level

sediment

/ / / / / / / / / / / / / / / / losing stream gaining stream

Figure 14 Influent and effluent stream conditions

62

Methods of Assessing the Intevaat-ion

There are a number of difficulties associated with equation 3.5(4). The first is that it

is difficult to ascertain the values of the parameters k' and m'. Equally important is the

difficulty of defining what is meant by the groundwater head at the stream h - An examination a

of the form of the flow pattern in the vicinity of a partially penetrating stream demonstrates

that significant changes in head may occur. In particular the groundwater head may vary

significantly on the vertical below the stream. Certain attempts have been made to allow for

the complex flow patterns near to a stream. Aravin and Numerov (1953) make allowances for this

effect by means of equivalent seepage resistances which are determined from analytical

solutions which represent the curvature of the streamline. Another approximation by Herbert

(1970) leads to similar results. These expressions for the additional resistance to flow can

be written in a form similar to equation 3.5(3) and can be included without difficulty in a

numerical model. In practice a non-linear relationship is often more appropriate with the flow

described by an expression of the form of equation 3.5(5) (Rushton and Tomlinson, 1979).

1 - exp { - D (h - Z ) } a. Ö

3.5(5)

where C and D are constants.

Although it is difficult to ascertain suitable values for the properties of the stream

bed, they do have a significant effect on the interaction between an aquifer and river or 3,

canal. For example when the inflow per unit length of a river is 0.225 m /day, a range of

head differences may be assumed to cause this leakage as shown in Table 3.5.2. The increase

in base flow in a specific reach is equivalent to the total leakage in that reach which may

be defined in terms of a leakage coefficient.

Table 3.5.2 Head differences related to leakage coefficient

Leakage Coefficient Head difference

(m) (m/day)

( k ' / m 1 ) A per unit length

0 .01 2 2 . 5

0.2 1.125

2 0.1125

5 0.045

The conditions with a head difference of 0.01 m approximate to those of a fully

penetrating river. Each of these values was used in turn in a finite difference model to

represent the flow to a river over a period of several years, and the results are shown in

Figure 15.

Using the four leakage coefficients relatively large and significant variations in

aquifer head are shown to occur compared with the small variations in the inflow along an 8 km

length of river. Indeed, when the recession part of the curve is examined in detail (Figure

16), the differences in flow are negligible for a relatively large variation in the leakage

coefficient. This result is particularly important because it indicates that an adequate

63

Methods of Assessing the Intevaction

85

80

•o 75 ra u

X

70

65

Leakage Coefficients

K ( m / d )

0 . 0 4 5

0 . 1 1 2 5

1.125

22.5

Year

Flow proportional to

head difference

v o> (D JÉ at

Year

Figure 15 Four year cycle of groundwater inflow to a river with variable head differences

and leakage coefficients

agreement between modelled and measured base flow during a recession period does not

necessarily mean that a satisfactory model of the river has been obtained. Also it indicates

why a model may appear to be satisfactory when a partially penetrating river has been assumed

to be fully penetrating if comparisons are made only between modelled and measured base flow.

64


In these circumstances the model could provide a poor representation of groundwater heads at

some distance from the river during recharge periods.

Leakage Coefficients

K( m/d)

*

o

Year 2

Figure 16 Annual cycle of groundwater inflow to a river with variable leakage coefficient

3.5.5 Lake-Aquifer Interaction

The interaction between a lake and an aquifer is different to that of a spring, river or canal

because the area of contact between the lake and the aquifer is much larger. Consequently

the flow patterns in the vicinity of the lake tend to be more complex, so that water may flow

from the aquifer into the lake in one region whereas in another region the same lake water may

be transmitted from the lake back into the aquifer. Two other features that often have a

dominant effect on the interaction are the hydraulic conductivity of the sediments at the

bottom of the lake and the ratio between the horizontal and vertical hydraulic conductivity

of the underlying aquifer.

Typical examples of the complexity of the flow patterns are given by Winter (1976). He

examines a number of single and multiple lake systems by the use of a mathematical model

65


representing a vertical section. A steady state finite difference approximation was used with

the lake represented as a region of high hydraulic conductivity. It is not possible to

summarise all the results since each example showed a different response. However, from the

large number of results included it is clear that sediments in the bed of a lake can cause a

significant change in the distribution of the groundwater head and thus in the flow pattern.

3.5.6. Example of Finite Element Analysis

The finite element method may be used to provide solutions to problems associted with the

interaction between surface water and groundwater in those cases where the groundwater flow

pattern is strongly schematised. It may be important to investigate the flow patterns, for

example, to determine the rate and direction of flow or the residence time of groundwater for

pollution control. Figures 17 and 18 show examples where the streamlines may be estimated

using the finite element method in the case of two dimensional steady state groundwater flow.

V////^/}//)(/////^VA

9^y^TF7^^7^7^7^^l^5^5Si5i5l5'??^7 yJVWWWWJs^^

Figure 17 Example of an interaction

between groundwater and surface

water

Figure 18 Example of an interaction

between groundwater and surface

water with entrance resistance

Darcy's law may be used with the principle of continuity to estimate the groundwater

heads in the two dimensional case provided that the transmissivity and the coefficient of

storage are known together with the boundary conditions. To examine the cases described

computer programs are available which rapidly and cheaply solve the rather complicated flow

problem by estimating the streamlines and groundwater heads. The following assumptions have

been made :

1. the water is homogeneous with a constant density and viscosity,

2. the porous medium is homogeneous, therefore the permeability k = k(x,y),

where x and y are the axes of the cartesian co-ordinates,

3. the porous medium is isotropic,

4 . no water is abstracted or recharged into the area except along the boundary.

66


The area being examined is divided into elements and with the aid of computer programs

the potentials are calculated (case A). Using these potentials and the estimated inflow and

outflow along the boundary another groundwater flow problem is created (case B), such that the

calculated potentials for case B are the streamlines for case A.

The porous medium may be non-homogeneous with a discontinuity in the transmissivity as

shown in Figure 19. Values of the transmissivity used in case B now have to be determined.

Perpendicular to the plane of division the flow towards and from the plane is equal, so with

the notation in Figure 19:

Vnl = V n2

hence

dh dh

k l . — .cos91 = k2 . — 1

cos 6,

also

b = dl. dl„

sin 6,

or dl sin 6 = dl sin 9

Between two equipotential lines

dhL = dh2

hence kl t g 91 k2 * 62

3.5(6)

Figure 19 Porous medium with a discontinuity

To calculate the streamlines for case A there is the necessity of first determining the

equipotential lines of case B. If the equipotential lines in case A are perpendicular to the

equipotential lines of case B then:

67


tg 91 . tg V1 = -1

tg 92 . tgï2= -1

where f. and ¥ are the angles with the normal in case B.

3.5(8) it follows that:

k 2 ^ \

which means that to solve the flow problem, case B, the reciprocal values have to be used to

those for flow problem A.

As an illustration of the technique a schematical cross section has been drawn of two

aquifers and a drainage ditch with various assumed hydrogeological parameters (Figure 20).

The section has an impermeable base at a depth of 25 metres below the phreatic surface and the

vertical line below the centre of the ditch can be considered as representing an impermeable

boundary in the flow problem assuming that conditions are symmetrical about this line. At an

horizontal distance of 50 metres from the centre of the ditch, fixed head conditions are

assumed to exist in both aquifers with the head at that point being constant at 1 m above that

of the ditch. The input to the upper aquifer from effective rainfall is assumed to be 1 mm/day.

Separating the two aquifers is a semi-permeable layer with a resistance of 10 days and a thick

ness of 2 metres. On either side of this semi-permeable layer the aquifers have a permeability

of 5.1 m/day over a thickness of 2 metres, and the remainder of the aquifers have a

permeability of 10 m/day.

The ditch illustrated in Figure 20 is assumed to have a bed and sides with relatively

low permeability resulting in an entrance resistance of 1 day. Surrounding this area of low

permeability is an additional, layer with a thickness of 1 metre and a permeability of 5 m/day.

From this information the groundwater head may be estimated in each part of the cross section

using the finite element technique. For the solution of this problem the program 'STRATRECT2'

was applied which is available at the National Institute of Watersupply in the Hague. This

program uses rectangular elements of various sizes to schematise the reality. To solve the

problem outlined above the flow region was divided into 272 nodes and 240 elements. The output

gave the groundwater head at every nodal point and outflow or inflow of water along the

boundary.

The distribution of the heads is shown in Figure 21. From a knowledge of the estimated

inflows and outflows along the boundary the stream function along the boundary can be

calculated. Thus as described previously this stream function may now be used to define a new

groundwater flow problem (case B). Again the program 'STRATRECT2' is used to calculate the

heads with the results as illustrated in Figure 22. These equipotential lines are now the

streamlines for the first groundwater flow problem.

The advantage of this method is that the value of the stream function is calculated at

all the nodal points in the flow region. This means that the direction and volume of water

3.5(7)

3.5(8)

From equations 3.5(6), 3.5(7) and

68


Figure 20 Cross section with given hydrogeological parameters

flowing in any part of the section may be determined readily in addition to the residence

times which are important for example in problems associated with water quality. Thus the

characteristics of the interaction between groundwater and surface water can be determined by

varying the soil parameters such as the entrance resistances of ditches, the resistance of

clay layers and the permeability of the porous media.

3.5.7 Rainfall-Runoff Models

Many types of mathematical model have been developed for estimating river flows from weather

information (Clarke, 1975). Some contain conceptual representations of the hydrological cycle

including the inflow of groundwater to a river or the flow of river water to an aquifer. Other

models contain both conceptual and regression components.

The relationship between geology and streamflow characteristics has been studied for many

years with particular attention given to examining the areal variability of low flows

(Schneider, 1965} Vladimirov, 1966). Recent studies have included the ranking of rook types

according to their ability to maintain low flows and determining the relationship between

various low flow measures and basin characteristics. Thus an estimation of the flow duration

curve, flow frequency curve and storage-yield relationships for ungauged basins have been found

69


Figure 21 Equipotential lines for case A

Figure 22 Equipotential lines for case B which are identical to the flow lines for case A

70


to depend mainly upon an assessment of the properties of the underlying geology (UK, Inst, of

Hydrology, 1980). In addition certain geological formations and soil types have been shown to

have a major influence upon the magnitude of the mean annual flood (UK, NERC, 1975).

Various types of geological index have been used to quantify the influence of geology

on river flows. One form of this index is based upon the variability of daily, monthly or

seasonal low river flows. Another is the Base Flow Index that has been derived from the

recession characteristics of rivers (UK, Inst, of Hydrology, 1980). In this report the

results are presented for a monthly model that has been used to relate river flows to rainfall

less evaporation in over 70 representative basins in England and Wales (Wright, 1978). The

equations that were derived for each river have been examined to quantify certain aspects of

the influence of geology on river flows. Two examples of this influence are now discussed,

the maximum significant lag and flow variability.

The equations relating river flow to weather conditions indicated the number of months

that flows continue to be significantly influenced by the weather conditions in a given month.

This may be termed the memory of the model and is the number of months, termed lags, that

positive values of rainfall less evaporation are significant in the multiple regression

analysis. Physically this is associated with the groundwater storage and outflow characteris

tics of the river basin, together with rainfall characteristics and the length of the model

calibration period as shown in equation 3.5(9).

LAG = 6.6 + 4G - 0.007 R + 0.007 C 3.5(9) a

where LAG is the maximum significant lag in the regression equation, in months,

G is the description of basin geology. G = 1 for relatively impermeable basins.

G = 2 if 20% or more of basin comprises aquifers other than chalk. In this context

an aquifer is defined as a rock associated with a flow variability equivalent to a

coefficient of variation of 0.3 or less as shown in Table 3.5.7. G = 3 if 20% or

more of basin comprises Chalk,

R is the annual average rainfall, in mm,

C is the duration of the model calibration period, in months.

In the data set used to derive the equation, values of LAG varied between 1 to 18 months;

G was 1, 2 or 3; R varied between 580 to 2000 mm; and C varied between 60 to 1000 months. a

Approximately 85 per cent of the total varaiance in LAG was accounted for by the equation.

Although the dominant independent variable in equation 3.5(9) is G, basin geology, a

relatively coarse grouping of values was used 1, 2 or 3. Basins with a substantial groundwater

storage (G = 2 or 3) were those which sustain river flows during prolonged dry periods.

However it may not be easy to identify this characteristic if aquifers occur in wet areas with

relatively persistent rainfall (R > 1000 mm per annum). Aquifers in such areas may always be

full or nearly full, with the result that river flows in these basins comprise a very high

proportion of surface runoff. In contrast an identical aquifer in a drier area could dominate

71


the river flow characteristics with the major flow component being base flow. In the wetter

parts of England and Wales where there is a negligible surface, soil or groundwater storage,

the maximum number of significant lags may be one or two months. At the other extreme monthly

rainfall can significantly influence river flows for 15 months or more thereafter (LAG = 15+).

This characteristic is not unusual in the drier eastern part of England where groundwater is

derived from the Chalk.

Another feature to note in equation 3.5(9) is the effect of the length of the model

calibration period (C ). It is difficult to obtain a satisfactory equation relating river a

flows to weather conditions if a river has a high proportion of base flow and there is less

than five years of flow and weather data. Such a short calibration period is likely to be

unsatisfactory for several reasons including:

1. Probable lack of extreme weather conditions in a short calibration period, for

example lack of drought conditions and low flow data,

2. Limited range of groundwater levels and thus groundwater outflow in the data

set. This could apply in particular to aquifers such as sandstones that have a

high coefficient of storage,

3. An inadequate equation with too few lags, thus failing to define the long-

term aquifer storage and outflow characteristics»

A further analysis of the results from representative basins in England and Wales

indicated that the variability of monthly river flows was significantly related to various

factors, including the basin characteristics of annual average rainfall, soil type and solid

geology. These three factors were found to be of equal importance and collectively accounted

for over 80 per cent of the total variance in flow variability. For example flows have a

low coefficient of variation in those basins where the soils are permeable and there are

substantial aquifers (G = 2 or 3). Flow variability has been related to rock types for those

areas where annual average rainfall is between 580 and 1000 mm and the equivalent value for

evaporation approximately 450 mm. The results are shown in Table 3.5.3. The effect of flow

variability on model accuracy is discussed in section 4.3.3. It is necessary to emphasise

chat the coefficients in equation 3.5(9) and Table 3.5.3 are-valid for a certain part of

western Europe and different coefficients and variables are likely to be necessary elsewhere.

72


Table 3.5.3 Rock types related to flow variability

Rock type Example in England and Wales

Flow variability (Coefficient of variation)

minimum

0.3

0.4

0.2

0.2

0.4

0.1

0.15

0.2

0.15

maximum

0.4

0.5

0.4

0.3

0.5

0.2

0.3

0.4

0.3

Alluvium

Hill peat

Sand and gravel

Boulder clay - relatively permeable

Stiff boulder clay

Porous fissured limestone

Porous fissured limestone

Compact fissured limestone

Poorly cemented permeable sandstone

Poorly cemented permeable sandstone

Cemented fissured sandstone

Sandstones and mudstones

Clay - relatively impermeable

Hard compact sedimentary rocks

Igneous

Chalky Boulder Clay (East Anglia)

Upper and Middle Chalk

Great and Inferior Oolite Series

Carboniferous Limestone

Upper and Lower Greensand

Bunter Sandstone

Old Red Sandstone (Wales)

Coal Measures and Culm Measures

London, Weald and Oxford Clays

Lower Palaezoic and Pre-Cambrian

0.15

0.3

0.2

0.4

0.3

0.4

0.4

0.4

0.5

0.6

0.6

0.6

73

Accuracy of Methods of Assessment

4. Accuracy of methods of assessment 4.1 Surface Water Flow

There are various methods used to measure river flows such as those based upon floats,

chemicals, river sections calibrated by current meter, calibrated structures and ultrasonic

or electro-magnetic techniques (Herschy, 1978; ISO, 1973). Current meters are widely used for

the measurement of low flows, but there are difficulties in measuring very low velocities by

this method. The measurement of high flows may also require special techniques (Barnes, 1974;

Benson, 1968). A suitable site for flow measurement is not always easy to find and the method

of measurement employed may vary depending upon local requirements and the construction and

maintenance costs involved.

In recent years increasing use has been made of remote sensing from aircraft or

satellites to investigate a wide range of problems in hydrology. The subjact areas have been

listed by Moore (1979) and those of particular relevance to this report include:

inventory of springs and seepage areas

estimates of land surface permeability

delineation of aquifer boundaries

estimates of water table depth and saturated thickness

delineation of probable recharge and discharge areas.

For example a method of estimating spring discharge associated with an artesian basin in

an arid area has been described by Williams and Holmes (1978). The gauged spring discharge was

correlated with the area of associated swamp, and further deductions of spring discharge were

then made with the aid of areal photography. This section has been divided into two parts

because of the differing requirements and problems encountered in temperate and arid regions.

74

Aaaiœaay of Methods of Assessment

4.1.1 Temperate Areas

As character is reflected in handwriting, so the physical characteristics of a hydrological

system are expressed in its outflow hydrograph. For example the degree of smoothing and

persistence in a hydrograph is determined by the magnitude of active storage reservoirs in the

hydrological system. The water balance equation relates the changes in storage, including

gains from rainfall and snow melt on the'one hand and losses through outflow, evaporation and

other withdrawals on the other. After inflow ceases storage depletion occurs. Thus if the

only component of flow is groundwater runoff that appears as surface flow at the outlet, then

the hydrograph is a groundwater recession curve which indicates the extent of the active

groundwater storage and its accessibility to the drainage system.

In a temperate region, however, the regular supply of soil moisture usually supports an

active vegetation with roots reaching down to the deeper soil layers where periods of short

supply may be bridged by an upward capillary flow of groundwater to the root zone. Where this

type of groundwater depletion occurs there is no longer a straightforward relationship

between the volume of active groundwater storage and the outflow recession rate (Ineson and

Downing, 1964). This complication can be disentangled only by a separate calculation of

actual evaporation. Although considerable advances have been made recently with regard to the

role of interception, much research is still necessary before estimates of evaporation can be

considered sufficiently accurate, especially during flow recession periods. The rate of change

of groundwater storage can be estimated from the water balance equation if the other components

of the equation are known. These components include evaporation and surface runoff. Thus the

accuracy of channel flow measurements during periods of dry weather flow recession is an

important limitation on the accuracy of estimates of groundwater storage.

In the temperate regions the traditional hydrological interest has been centred mainly

on the measurement, prediction and regulation of high flows during the winter and the main

tenance of water levels in the summer. Unfortunately these management objectives have

resulted generally in the construction of weirs with a wide horizontal crest either fixed or

adjustable which have a limited range of accurate flow measurement. These weirs were

designed for the measurement of high flows with a certain percentage of error and like many

calibrated river sections are relatively less accurate in measuring low flows.

More recently, flow measuring structures such as those used for irrigation control that

maintain a proportional accuracy over a wide range of flows have become more widely used. In

irrigation schemes a structure is designed to measure a predetermined range of supply rates,

the objective being to measure quantities of water over certain intervals of time rather than

instantaneous rates. This objective is similar to that required of measuring flow as an

independent variable in the water balance equation. In this case also the flow rates are

integrated over certain time intervals and it follows that a proportional accuracy for the full

range of flows is required. However there remains a disadvantage with irrigation stuctures

that they have not been designed for the measurement of lowest flows. The lower limit of the

measuring range is set always by a certain minimum head over the crest or over the invert of

the control section. This is to avoid systematic errors such as those caused by floating

75


debris as shown in Plate 1, algal growth (Plate 2), minor sediment deposits, viscosity, and

surface tension.

However real this lower limit is, it is not very satisfactory in groundwater studies

because it puts severe restrictions on the applicability of existing structures for the

accurate measurement of low flows. Some managers tend to ignore the implication of this lower

limit and consider that an approximate measurement is better than none at all. However, such

data are occasionally used indiscriminately in water management studies which thus contain

misleading information concerning low flows and any conclusions concerning groundwater aspects

may be seriously in error.

To illustrate this point 24 monthly values of outflow from the Hupsel Creek, a

research basin in the Netherlands, have been plotted on a logarithmic scale as shown in

Figure 23. The discharge from this 650 ha area may range between 3000 and a few litres per

second, and is measured in an H-flume as shown in Plate 3. The flume was calibrated using

two methods. For higher flows a model to a scale of 1:5 was constructed including the

approach conditions, and for low flows a true scale model was used in combination with a

volumetric calibration in situ.

For the purpose of this exercise the true flow rates were converted into heads over an

imaginary weir with a horizontal crest width of 2 m. The performance of this imaginary weir

may now be assessed as shown in Figure 23. If the lower limit of the measuring range is set

at a head of 5 cm (H = 5 cm) as shown by the horizontal line, then the flow would have been

measured with insufficient accuracy for more than 50 per cent of the time.

Algal growth on the sill and sediment deposits could readily create a positive error

A = 2.5 mm or up to 10 mm in the measurement of head. The consequences of A = 5 mm are shown

in the hydrograph in Figure 23, the measuring error of total flow over the first five months

then being 75 per cent. Figure 24 shows how the flow duration curve for the same period would

have been changed.

Finally the logarithm of the recession rates as plotted in Figure 25 show an over-

estimation of nearly 40 per cent in the characteristic time for groundwater storage when

A = 5 mm. It follows that such low flow hydrographs when subject to the inaccuracies described

are quite misleading when used for the study of groundwater storage.

The problems associated with measuring the full range of flow out of a drainage basin

can indeed be considerable. Local circumstances such as a very wide range of flows in flashy

streams, relatively uniform flows from forested areas, sediment transporation sometimes

including boulders, floating debris including weeds and logs, air entrainment, water pollution,

low temperatures, lack of sufficient head, accessibility, supervision and maintenance, require

a particular optimized solution for each individual case. Sometimes relatively straightforward

solutions are possible such as separate gauges for measuring high and low flows either in

parallel or in series. Plate 4 illustrates a solution for measuring a wide range of flows with

a restricted available head, the possibility of floating weeds and some sediment transportation.

76


10'-9 8 7 6

5

L

o c 3

3-

2-

10'-9-8 7 6

5

i,-

3 -

2 - - -

10u-

5-

H = 5cm . _ ! . .

A=5mm i 1

i

J

ru

r-t

• i

i

I

t 72 V . 6 V . 19 V . 11V. error

i 1 1 1 1 1 1 1 1 1 1 1 1 1 1 1 1 1 1 1 1 1 1 r month 4 5 6 7 8 9 10 11 12 1 2 3 4 5 6 7 8 9 10 11 12 1 2 3 197« 1975 year

Figure 23 Measured outflows affected by an error in the crest level of 5 mm

In this case there is a good accessibility, supervision and maintenance.

Good solutions for the accurate measurement of the full range of flows may be expensive,

but poor quality data will result in costly errors in research, planning and management.

Groundwater studies in particular require an accurate measurement of low flows.

4.1.2 Arid and Semi-arid Areas

There are problems associated with measuring river flows in arid and semi-arid regions that are

not generally important in temperate regions. These problems arise due to the characteristics

of runoff which occurs in arid regions under certain conditions. In semi-arid regions, or in

arid regions during exceptionally wet years, flood flows may occur over a short season which

in some tropical zones may be two months in the summer or mainly in the autumn or spring as

in the north of the Sahara (Herschy, 1978, Chapter 13; Callede, 1977).

77

Aaauraay of Methods of Assessment

summer 1974

001 0.05 i 1 r

0 2 05 1 i 1 1 r

20 30 40 50 60 70 80 T 90

i 1 r 95 98 99

Figure 24

7o time equalled or exceeded

Flow duration curve affected by errors in crest level

An example of the extreme irregularity of runoff from one year to the next is that for

the Enneri Bardague at Bardai in the Sahara on the north side of the Tibesti massif where the

mean annual rainfall is 15 to 20 mm. During a period of nine years there were four years

without any runoff one with a flood of 425 m /sec and four years with one to three smaller

floods.

In addition hydrographie degeneration may occur, a phenomenon which has a considerable

importance in the measurement and evaluation of river flows. This phenomenon occurs when the

river channel is not well defined and there is a flow of surface water to groundwater. The

continuity of runoff is no longer obvious and minor floods in the headwaters may produce no

runoff further downstream. Where the slope of the channel reduces the river overtops its banks

and seeps into the flood plains. Effluent arms may leave the main channel and water is lost

78

Aoavœaoy of Methods of Assessment

= loge, cot got

A = 5mm

time units of 3hr

Figure 25 Recession curve affected by error in crest level

in marshy depressions. Also the main channel may reach an inner delta zone from whence only

a negligible runoff occurs in the downstream area.

Where hydrographie degeneration occurs some of the generally accepted terms used in

hydrology have little meaning. For example the concepts of a drainage basin and discharge in

terms of flow per square kilometre are hardly meaningful. The measurement of total runoff at

a given point should take account of these factors. In studies of groundwater recharge it may

be useful to estimate the downstream point of a river system where surface runoff occurs at

frequencies of once in 10 or 50 years.

4.1.2.1 Measurement of Flood Flows

The measurement of flood flows is important in arid regions where floods are frequently the

main source-of aquifer recharge. However there are few river gauging stations in such areas

because of the difficulties of access and operation. The very sparse network is supplemented

by hydrometric surveys carried out every few years. These surveys include the direct

measurement of flood flows or indirect measurements with investigations of flood level and

evaluation by the Manning or other formula. Investigations may include local enquiries,

observations of debris and the level of bank erosion.

Floods are often violent and difficult to measure accurately. For example in Tunisia 3 2

peak flows of 1 000 m /sec should be anticipated for a basin of 1 000 km in the mountainous areas, with the possibility of up to 10 000 m /sec in extreme cases (Cruette and Rodier, 1971).

In such conditions the high speed of the flow and the general instability of the bed make flow

measurement difficult at the permanent gauging stations. Cableways are required for current

79

Aaaurasy of Methods of Assessment

meter measurements, supplemented with floats when the flows are too high for current meters.

It should be possible to evaluate the flood flow within an accuracy of 10 to 15 per cent,

provided that there is a stable river control section, and a good organisation for flood

measurement including a regular channel shape and an adequate measuring reach for floats.

However, if the bed is unstable and it is impossible to follow the variation of the bed at the

calibrated river section, then the error can be substantially greater. In general the main

gauging stations are carefully sited where there are natural rock sills or alternatively

artificial controls are installed to stablize the bed. If it is not possible to stabilize

the bed then special allowances are made for variations in the depth of the river such as the

determination of the profile of the bed at the time of maximum water velocity. In most cases

the computation of flood flows should be accurate to within 20 to 40 per cent if they are based

upon systematic surveys of the traces of the flood event and local enquiries.

In flood plains the direct measurement of high discharges is not too difficult if a

strip along a cross section is carefully prepared and vegetation removed. The error in

measurement may then be within 15 to 30 per cent of the true value. However if flows are

estimated after the event from surveys of debris and erosion marks relating to the peak stage

then errors of 100 per cent are possible.

4.1.2.2 Measurement of Low Flows

If the recession is of very short duration, then the measurement of low flows is difficult

unless there are personnel located near the station to carry out this work. For events and

locations where the recession may continue for several weeks it may be possible to visit the

site several times. In most cases the measurement itself has to be carried out with great

care.

At first it is advisable to determine the reading of the water level recorder or staff

gauge at zero flow. This will probably be one measurement in arid regions which may be

determined with the greatest precision. However, it should be noted that runoff may cease at

the measuring section and reappear further downstream. This type of resurgence may occur if

the bed comprises alluvium and has a variable depth and permeability. If the surface flow is

continuous along the river then underflow at the measuring section may be checked by taking

several gaugings along that reach of the river.

The measurement of very low flows often demands special improvisation. For example the 3 / usual gauging station is seldom sufficiently sensitive to measure flows less than 0.2 m /sec

if it has been designed to measure floods in excess of 1 000 m /sec, even if an artificial

control has been installed. To adequately gauge low flows a small permanent canal may be

constructed and fitted with a portable weir. Alternatively a small temporary canal may be dug

and a miniature current meter used to measure velocities. Sometimes a second water level

recorder is required with a suitable scale to record the lowest levels, or the complete

recession may be established by interpolating between regular spot measurements from river

gauging. For small rivers it is advisable to check if the flow is subject to daily variations

80


if the valley is covered with vegetation. The amplitude of this variation can reach 5 to 15

per cent of the mean low flow discharge.

The error in discharge evaluation is typically between 10 and 20 per cent if there is a

stable calibrated river section with regular and precise measurements of discharge and with due

allowance for diurnal flow variations. This represents the best conditions. At other river

gauging stations where the bed is reasonably stable and gauging accuracy is average, the error

in measurement is frequently greater than 20 per cent. The same result is obtained if the

station has a poor sensitivity at low discharges, such as those stations where the control is

a very broad rocky sill. If the bed is unstable and there are no measurements of discharge,

then large errors are possible, and even with careful estimation errors may be two or three

times greater or smaller than the true discharge.

Large errors may also occur in those cases where there is a very stable bed but the

rating curve has not been established for low flows by measurements. The extrapolation of a

rating curve can lead to errors in discharge estimates of 50 to 100 per cent. It should be

noted that even with a rock control there may be a lack of stability for discharges less than

0.05 m /sec. At these flows a relatively small rock or the branch of a tree can significantly

alter the rating curve. For this reason the relative precision for evaluating low flows is

generally worse than for flood flows. To accurately assess the interaction between surface

water and groundwater every effort should be made to improve the accuracy of low flow

measurements.

4.2 Aquifer Characteristics

The relationship between surface water and groundwater is strongly influenced by the

infiltration characteristics of surface water and four aquifer properties. These properties

are:

Hydraulic conductivity (K) and transmissivity (T) that govern the rate of

groundwater flow; and specific yield (Sy) and storativity (S) that relate

to the volume of water held in storage.

Each of these properties are briefly discussed in turn. A more detailed description has

been given elsewhere (Brown et at., 1972). When water is abstracted from an aquifer the

drawdown in groundwater levels increases with time and decreases with distance from the

abstraction"point. The determination of the transmissivity and storage functions is

necessary for evaluating in hydrodynamic terms the relationship between groundwater and the

aquifer boundaries over long time periods.

To investigate the effect of groundwater abstraction upon river flows a term called the 2

response time (T/SL ) may be used, where L is the distance from the river to the basin divide

parallel to the river. When a new abstraction regime starts there is some delay before the

aquifer reaches a new state of equilibrium. This delay is directly related to the response

time. If the response time is relatively high and the well is near the river, then a new

state of groundwater equilibrium is rapidly reached (Downing et al., 1973).

81


4.2.1 Hydraulic Conductivity

The hydraulic conductivity as defined by Darcy's Law is expressed in terms of velocity, and

is similar to the coefficient of permeability. Since the aquifer may for structural reasons

have a preferred direction of flow, the hydraulic conductivity may also have a directional

variation. It is dependent also upon porosity, grain size and the specific surface of the

grains. The standard unit is given in terms of metres per second (m/s) at a temperature of o

20 C. Accurate estimates of hydraulic conductivity require temperatures to be considered

because it is influenced by viscosity which in turn is influenced by temperature. For example o o

a fall in temperature from 25 C to 5 C can lower the hydraulic conductivity by 40 per cent.

Other factors affecting hydraulic conductivity are the compaction or cementation of the

constituent grains of the aquifer and the nature of the water flow which is generally assumed

to be laminar. Typical values of hydraulic conductivity are given in Table 4.2.3.

Under natural conditions the variation in hydraulic conductivity is large, ranging from — 3 —8 —12

10 m/s for coarse sand down to 10 m/s or even 10 m/s for clay. However, samples taken from the same source can show large differences and the extent to which the hydraulic

conductivity varies is more significant than its precise value. Thus a variation of 20 per

cent may be considered negligible whilst one of 200 or 300 per cent would be highly

significant. The determination of hydraulic conductivity in the laboratory is likely to

produce results differing from those obtained in the field. This is due in the main to the

small volume of the samples and the comparatively short duration of the tests.

4.2.1.1 Laboratory Determination of Hydraulic Conductivity

The Hazen (1911) formula uses the relationship between grain size and permeability, and the

hydraulic conductivity is given by:

2 K = 116 (d _) in CGS units

where d is the effective grain size defined as the maximum grain size of the finest 10 per

cent of the sample.

Although the results obtained are only approximate, the method is useful since only a

short time is required for its determination. Alternatively, permeameters can be used for

direct measurements of hydraulic conductivity based either upon the constant head or falling

head type. However, undisturbed samples are necessary for realistic results and such samples

are often difficult to obtain.

4.2.1.2 Field Determination of Hydraulic Conductivity

The Le franc (1937) or Lugeon (1933) tests occupy only a very short period of time, but the

values obtained apply only to that part of the aquifer immediately adjacent to the well screen.

Pumping tests over a longer period of time permit the evaluation of hydraulic conductivity over

a zone extending some distance from the well. The results of such tests provide values which

reflect the average conditions throughout a considerable volume of aquifer and are thus more

82


in keeping with the generalisations made previously. For steady state conditions the methods

proposed by Dupuit (1863) can be used. For non-steady state conditions, the methods proposed

by Theis (1935) and their many derivatives may be employed.

4.2.2 Transmissivity

Interpretation of pumping test data by the Theis (1935), Hantush (1964), and Cooper-Jacob

(1946) solutions of the diffusivity equation is widely used, and permits the determination of

the transmissivity with a high degree of accuracy.

The reactions within a pumped aquifer depend essentially upon the local value of

transmissivity, 'local' being defined as that area within the influence of the cone of

depression of the pumped well. Within this area of influence the aquifer is considered to be

homogeneous, and it is further assumed that the areal extent of the aquifer is infinite and

that boundary conditions can therefore be ignored. However, it is unsafe to extrapolate the

transmissivity values thus obtained to points distant in an extensive and heterogeneous

aquifer.

The accuracy of transmissivity estimates is also affected by the well-loss factor and by

partial penetration of the aquifer, whilst early cessation of pumping may result in

insufficient data being obtained. The well-loss factor can be determined by the analysis of

data obtained during the recovery period after pumping has stopped. Using methods such as

those of Kozeny or Hantush (1961) permits the calculation of the real transmissivity in cases

where either the depth of the well is inadequate or the duration of the pumping test is

insufficient (Prickett and Ungemach, 1970) . A frequent problem, particularly in alluvial

deposits is that the thickness of the aquifer is difficult to determine.

A comparison between transmissivity values from short duration pumping tests and from

long-term tests was made in Tunisia (Besbes, 1971). In that study the ratio between the

two values was less than 2.0 in 74 per cent of the tests analysed, and less than 5.0 in 96 per

cent of the cases.

Transmissivity may also be estimated by using the general equation to express drawdown

(A) in relation to the rate of pumping (Q):

A = AQ + BQ 4.2(1)

where A and B are constants.

However, this equation does not take into account the effects of clogging. In fact the

constant A is dependent upon transmissivity (T), storativity (S), degree of well penetration

and the disturbance of the aquifer adjacent to-the well. B is the well-loss constant, and

when B = 0 it is possible to express A as follows (for a fully penetrating, homogeneous and

isotropic aquifer):

83


Q 4TTT

where u = ^

and W is the well function (Theis, 1935)

r is the radius of the well, and

t is the time from start of pumping.

The specific capacity j- is related to transmissivity, and Pollack (1967) attempted to

quantify this relationship. From an examination of results from 274 wells in the Paris basin,

Levassor and Talbot (1976) derived the following relationship:

T - 1.25 Ç-A

A geostatistical approach together with the application of Kriging techniques (Delhomme

and Delfiner, 1973) is useful for estimating transmissivities throughout an aquifer using all

available data. By using the fictive point method, it is possible to make an a priori

quantification of the reduction in uncertainty corresponding to any improvement in the

observation network (Delhomme et al., 1977).

4.2.3 Specific Yield

The volume of water that drains from an aquifer under the influence of gravity may be

characterised by the term 'specific yield' which is frequently expressed as a percentage by

volume of the drained formation.

The specific yield (effective porosity) may be determined in the field by the analysis

of pumping tests data, in the laboratory using undisturbed samples or in the absence of

experimental evidence by reference to tables of generalised values. Estimates of the specific

yield and hydraulic conductivity are of considerable value in the estimation of changes in

groundwater storage and the flow of groundwater into rivers. In the absence of experimental

verification approximate values of these coefficients may be used such as those in Table 4.2.3.

However they vary considerably with local conditions. For example experimental data has

been analysed by the Valdai Research Hydrological Laboratory (USSR) for river basins with

glacial deposits and a landscape featuring terminal moraines. In these conditions values of

the specific yield may vary within the limits shown in Table 4.2.4 (Kapotova, 1978).

4.2.4 Coefficient of Storage

The coefficient of storage, or storativity, of an aquifer may be defined as the volume of water

which an aquifer releases per unit surface area per unit decline in stage. In unconfined

conditions this coefficient approximates to the specific yield provided that drainage by

gravity is complete. For many confined aquifers, the coefficient of storage lies between

0.0001 and 0.001. The most precise method of measuring this parameter is by analysing pumping

84


test data where observation boreholes (piezometres) have been used. When such data are not

available, there are two alternative methods for estimating the coefficient of storage.

The first method is based upon aquifer compressibility, and the equation for the

coefficient of storage is:

S = d . n - P - ( ß ^ + 3 ) water pores

where d is the thickness of the aquifer

n is the porosity

H is the specific weight of water

$ is the compressibility.

Table 4.2.3 Mean values of specific yield and hydraulic conductivity for selected

lithologies (glacial terrain)

4.2(3)

Range of mean values of coefficients

Lithology Specific

minimum

Q.10

0.15

0.20

0.25

0.02

0.O1

yield

maximum

0.15

0.20

0.25

0.35

0.03

0.10

Hydraulic (

minimum

1 X

6 x

6 x

2 x

1 x

2 x

lo"4

lO"4

lO"3

lO"2

lO"5

lO"3

con m/s)

ductivity

maximum

6

6

2

6

1

6

-4 x 10

x 10"3

-2 x 10

x 10"2

x 10"3

x 10"2

Very fine sand with loam

Fine sand with clay

Medium sand

Coarse sand and gravel

Argillaceous sandstone

Fissured limestone

There is also a relationship demonstrated by Fatt (in Craft and Hawkins, 1959) between

the reduction in pore space {$ ) and the compression of the aquifer. The latter parameter

is a function of the difference between the weight of the overburden and the hydrostatic

pressure within the aquifer. However, results obtained by this method are very approximate

and the degree of error is uncertain.

The second method is based upon the barometric effect (BE) which is defined as the ratio

between a change in piezometric head tAH) in the confined aquifer and the corresponding change

in atmospheric pressure expressed as metres of water. Using the same symbols as in equation

4.2(3) the coefficient of storage is calculated as follows:

(d . n . ß water

)/BE 4.2(4)

85

Aceuraoy of Methods of Assessment

Table 4.2.4 Range of values of the specific yield for selected lithologies

(after Kapotova, 1978)

T . ̂ , , Range of coefficient Lithology

Stiff loamy clay

Stiff loamy clay with le

Clayey loam

Clayey loam with a high

Stiff loam

Medium loam

Quicksand

Fine sand with silt

Fine sand

Medium sand with clay

Medium sand

Coarse sand and gravel

mses of

organic

sand

content

minimum

<0.001

0.01

0.02

0.03

0.04

0.06

0.03

0.10

0.12

0.15

0.20

0.25

maximum

0.005

0.025

0.03

0.04

0.05

0.07

0.08

0.15

0.25

0.20

0.30

0.35

4.2.5 Infiltration

The infiltration characteristics may be described in terms of the time taken for water to

infiltrate the soil and underlying aquifer. Water on the surface of the ground infiltrates

the soil first by filling the pore spaces, then by moving downwards under the influence of

gravity. The size of the pore spaces together with their total volume is therefore

important, and equally important is the degree to which the pores are interconnected.

Infiltration rates may be determined by using various methods such as:

Porchet method

Muntz-Laine method

Double cylinder method.

Under favourable geological conditions infiltration rates of approximately 5 to 20 mm/hr

may be observed.

4.3 Relative Accuracy of the Methods of Assessment

A review of contemporary methods of estimating the interaction between surface water and

groundwater shows that the quantitative solution of local problems is based extensively on the

use of hydrometric and hydrometeorological information (Brown et al., 1972; De Wiest, 1965;

Popov, 1969a; Toebes and Ouryvaev, 1971), The use of river flow information and the channel

water balance method are described in Chapter 3, and comprehensive information is given in

various publications concerning the methods of estimating the errors with their systematic

86


and random components (Anon, 1977a; Popov, 1975; Sokolov and Chapman, 1974) . In addition

sections 4.1 and 4.2 discuss the problems associated with the accuracy of measuring runoff and

aquifer characteristics.

When examining the interaction process with the aid of available information and

conceptual models, some systematic errors.are likely to occur when comparing estimated and

observed groundwater levels and river flows. In most cases the determination of quantitative

estimates of the interaction is a very complicated problem.

4.3.1 Channel Water Balance

The availability and accuracy of field data has a strong influence on the reliability of the

channel water balance (CWB) calculation. The accuracy of the CWB calculation can be given in

terms of a mean quadratic value (the standard deviation) that incorporates the rather complex

errors arising from the computation of each element, ç., as shown in equation 4.3(1).

/i=n 2 fi™ C = J Ï C : = A 2 (d, Q H ) 2 4.3(1)

where d is the relative mean quadratic error (coefficient of variation) of discharge Q, and

n is the total number of the CWB components.

The relative errors, d, are determined for the elements of the CWB appropriate to the

methods adopted (Anon, 1977a; Anon, 1977c; WMO, 1965).

The results of the CWB computation are considered to be reliable if the absolute value

of the remainder term in the CWB equation is significantly different from the maximum error of

its determination.

I S> I » Œ Ç 4.3(2) 1 o ' p o

where Œ is the p% confidence level from the error computation in terms of its mean

quadratic value, and p is the confidence coefficient or probability p.

For reliable estimates of the CWB computation, the minimum acceptable value of the

probability p may be given as p = 0.95, for which « = 2 . Therefore the criterion for a P

reliable CWB computation can now be written in-the form of the following inequality:

I Q 0 I > 2 C0 4.3(3)

or dQ $ 0.5 | Qo | 4.3(4)

For the case where relatively small changes in runoff occur in a reach, the accuracy of

computations depends mainly on the accuracy of discharge measurements at the upstream and

downstream cross sections. Errors in the other elements of the CWB should be neglected because

of their relatively small values. In this case the mean quadratic error of discharge

measurement in the reach (AQ) is estimated by the formula of Karasev:

87


V 1 + Kc 2

a - ̂ — - 2 - 4.3(5) AQ Q 1 - KQ

where d is the relative mean quadratic error in the runoff computation, with d = Ç / Q, and

k is the relative change in discharge in the reach, with K = Q9 /Q. . Q y ¿ i

The relative error d rapidly increases with decreases in the value of AQ. Thus, when

K = 0.75 it is equal to 5 ç , but when K = 0.9 it increases to 13.5 ç .

In the case of small discharge differences the appropriate inequality is

| AQ | » 2.82 ÇQ

This shows that to reliably determine the value of small losses from or increments to a

river discharge, AQ, the error in the runoff estimate at the cross sections, ç , must be

satisfied by the inequality:

ÇQ « 0.35 | AQ | 4.3(6)

In other words, ç must be less than one third of the estimated difference in runoff at

the two cross sections.

When d = 5%, values of [ AQ | » 1/15Q, may be reliably estimated. When K Q = 0.85, the

error d from equation 4.3(5) will be less or equal to 45%. In practice reasonable results

may be obtained from the water balance computation when | AQ | >0.3Q , that limits the

relative error in Q to a value less than or equal to 20%.

If Ç J- 0.5 | Q | then other physical factors may be considered which induce changes in

runoff in addition to those included in the CWB computation.

If ç 5- I Q I the residual term has a similar value to its error of determination. Thus o i o i

the derived Q value may be considered as a random variable.

When 0.5 Q I < ç < I Q I the nature of the residual term is uncertain, and 1 o ' o ' o '

additional information from long-term data is required to conclude whether the derived values

of Q are random and due to chance or have a physical meaning.

4.3.2 Flow Separation

If the groundwater component of river flow is not estimated by the direct hydrometric method,

then the main errors depend upon the accuracy and suitability of the conceptual model used in

the flow separation and the elements of the computation. The latter include river discharge

values adopted as indices of the groundwater inflow and its dynamic coefficients which are

used to determine groundwater inflow to a river during flood periods.

Contemporary procedures in hydrological field work include taking discharge measurements that

later may be used to estimate the movement of water between rivers and aquifers. These

88


measurements have a specified accuracy (Ratner, 1970). However, to use the discharge values in

later computations, it is necssary to ascertain that river water is formed from groundwater

according to the assumptions inherent in the conceptual model. This should be done on the

basis of complementary information concerning the hydrometeorological and hydrogeological

regime.

In the absence of complete information about the groundwater regime, the quantitative

estimate of the local dynamic coefficients representing the groundwater inflow is rather

complicated. This is typically the case when hydrograph separation methods are used.

Improvements in the separation methods, and increased accuracy, must be based on studies of the

variations in space and time of the characteristics of the formation of groundwater inflow to

rivers. Thus correct allowances should be made for groundwater inflow from each major aquifer

and reach in the river basin under study.

4.3.3 Mathematical Models

There are several statistical measures that are used to describe the accuracy of mathematical

models and each of these measures can be influenced by physical and experimental factors.

In this section two measures of accuracy are discussed for rainfall-runoff models and their

relationship to various factors described based upon analyses for representative basins in

England and Wales (Wright, 1978). The examples are intended to give a general indication of

the factors that influence the accuracy of flow estimation and the important effect that basin

geology has on model accuracy.

The accuracy of equations relating river flows to weather data may be given in terms of

their residual error of estimate and the coefficient of multiple correlation. The residual

error is calculated from the difference between the measured and estimated flows for the

period of model.calibration and may be given in terms of a per cent of the flow. Thus an

equation with a residual error of 25 per cent would be expected to estimate flows within

50 per cent for 95 per cent of the time. Figure 26 shows the relationship between the residual

error, the coefficient of multiple correlation and flow variability for equations derived for

75 representative basins in England and Wales.

An equation relating rivers flows to weather data could have a residual error (E) of 25

per cent and a coefficient of multiple correlation (R) of 0.96 or 0.9 7 if the coefficient of

variation (C ) for monthly flows was between 0.3 and 0.4. The relationship between these

measures is given by equation 4.3(7).

E = C \1 - R2 4.3(7)

Table 3.5.7 indicates the rock types associated with specified values of flow

variability and weather conditions. Thus in areas with similar weather conditions and known

rock types it may be possible to estimate the probable accuracy of rainfall-runoff models

before equations have been derived. Rivers with a high proportion of groundwater flow will

tend to be characterised by relatively low values of flow variability because of the smoothing

89


Multiple correlation coefficient

Figure 26 Accuracy of equations relating river flow to weather data for representative

basins in England and Wales

effect of groundwater storage. In such cases the accuracy of equations relating runoff to

weather conditions may be typified by low values of the residual error, for example in the

range of 10 to 20 per cent. However, low values of the residual error may not necessarily

imply an accurate equation because the coefficient of multiple correlation may be

unacceptably low (Figure 26) .

When a rainfall-runoff model has been optimised for a large number of basins, then the

accuracy of the derived relationships may be related to specific factors using multiple

90


regression techniques. For example the residual error (E) for equations to estimate river

flows in 55 representative basins in England and Wales has been related to several factors as

shown in equation 4.3(8) .

16300 C - 9 5 O'11

E = 4.3(8) l A G - 0 7 * * ! - 1 7 ™ - 1 8

where C is the coefficient of variation of monthly flows (range 0.1 to 0.6, examples given

in Table 3.5.3)

C is the length of the model calibration period, in months (range 60 to 1000 months),

LAG is the maximum significant lag in the equation in months (range 1 to 18 months),

RAI is the number of rainfall stations used to estimate mean catchment rainfall

(range 3 to 27) ,

RIV is the type of flow measuring station, 1 represents a calibrated river section,

2 a structure in the river not designed specifically for flow measurement and 3 a

structure designed for flow measurement.

All the coefficients for the five independent variables were significant at the 5 per

cent level, and the equation accounted for 84 per cent of the variance in the residual error.

The single most important independent variable is flow variability (C ), which accounts for

75 per cent of the total variance in the residual error.

Although equation 4.3(8) is limited in application to a part of western Europe, the

accuracy of other types of model may be expected to be subject to similar factors. For

example accuracy may depend upon:

1. Variability of the dependent variable such as groundwater levels or river flows,

2. Length of the model calibration period,

3. The lag effect of storage at the surface, in the unsaturated zone or in the

saturated zone,

4. The sampling density or number of sampling points for variables such as rainfall

or groundwater levels,

5. Accuracy of measuring or estimating the variables such as rainfall, evaporation,

groundwater levels and flow.

The use of remote sensing from aircraft and satellites may be used to improve the

accuracy of estimating certain components of the hydrological cycle such as the areal extent

of aquifers, wetlands, saturated soils, irrigation and types of land use and vegetation cover

(Moore, 1979). Of particular interest for modelling purposes is the use of radar for

91


estimating the areal extent, duration and intensity of rainfall (Browning et al., 1977). The

extreme variability of thunderstorm rainfall is well known, but even widespread frontal rain

can include areas of substantially heavier rain. Thus radar is being developed to improve the

accuracy of areal rainfall estimates and this has some potential as an aid for assessing water

resource schemes in both temperate and arid regions. The areal coverage of rainfall stations

tends to be rather sparse in some parts of the world such as the tropical forest, desert and

mountainous regions of Asia, Africa and South America, although a better coverage often exists

along valleys (Anon, 1978).

In conclusion the accuracy of a mathematical model will depend upon how adequately the

components of the hydrological cycle are represented by mathematical equations. Thus

improvements in assessing the interaction between surface water and groundwater will depend

upon improving techniques for measuring the variables and improved modelling procedures

including a better understanding of the interaction process.

92

Case Studies

5. Case studies 5.1 Temperate Area: Great Ouse Pilot Scheme, UK

5.1.1 Introduction

Increasing effort has been directed in water resource development in England and Wales to make

more efficient use of groundwater stored in the Chalk and Triassic sandstones. One method

designed to achieve this objective is to regulate river flows by means of the intermittent

abstraction of groundwater to enable a relatively high river abstraction to be sustained at a

downstream point on the river system. This method of development maintains riverside amenities*

safeguards navigation interests and maximises the yield by utilising both surface runoff and

groundwater storage. The variability of river flows is reduced by the redistribution in time

of the groundwater component of river flow. In such schemes it is. essential to examine the

interaction of surface flow and groundwater to minimise both capital and operational costs.

The number and location of abstraction wells must be carefully designed to ensure that river

flows are adequately augmented in dry years to support the design yield.

Several experimental areas have been set up to examine the feasibility of regulating

river flows by the intermittent abstraction of groundwater. Groundwater abstraction losses in

pilot schemes have been estimated (Wright, 1974) and ascribed to either (a) pumped

intercepted base flow or (b) induced river bed infiltration (Kemp and Wright, 1977). This case

study describes some of the work carried out in the Thet pilot scheme in Norfolk. Figure 27

shows the mean annual rainfall less actual evaporation over a part of south-east England,

the Chalk and the Thet pilot scheme area.

Test pumping programmes for the 18 abstraction wells (Figure 28) were carried out

between 1968 and 1971 to determine the well and aquifer characteristics and provide some 2

practical experience in river regulation by groundwater. The pilot scheme area covers 71.5 km ,

varying in height from 15 m to 50 m above mean sea level. The land is mainly arable with some

forest and heath, and soils are free draining. A summary of the work carried out in the Great

Ouse pilot scheme area was described by Backshall et al. (1972), and due to the success of the

experiment proposals for development were outlined (Great Ouse River Authority, 1972).

93

Case Studies

Thet Pilot Area

Chalk ' 1 ' 1 ' i i i

Residual Rainfall - .100—'

(rainfall less evaporation)

Figure 27 Mean annual rainfall less evaporation and the Chalk outcrop in East Anglia

5.1.2 Description of the Pilot Scheme Area

The pilot scheme area is situated within the basin of the river Thet (Figure 28) where average

annual rainfall is approximately 600 mm and evaporation 450 mm. A drought duration of

94

Case Studies

Thet Pilot Area • •

Gauging station V

Abstraction wells ®

Met. station 0

Kilometres 0 1 2 . 3

/ V»

S

<

I

\

t I

/

J

Figure 28 Thet groundwater Pilot Scheme

approximately 9 months is that which tends to be critical to the maintenance of water supplies

(a) in eastern England from small and medium surface reservoirs, and (b) from groundwater

storage such as that required to maintain flows in the river Thet at 50 per cent of the annual

mean value throughout a 1 in 50 year drought event. If greater use is made of groundwater

storage by maintaining river flows at a higher proportion of the mean, then minimum ground

water levels during a 1 in 50 year drought would tend to occur after 18 or 30 months,

similar to the critical duration for larger surface reservoirs (Downing et dl., 1973).

There are five flow gauging stations on the river Thet and its tributaries, four of

which measure flows into and out of the pilot area. The fifth gauge, at Melford Bridge, is

downstream of the pilot area and has the longest record. It is a triangular profile weir

specially designed for flow measurement and constructed in 1962. The drainage area to this 2

gauge is 316 km . Table 5.1.1 contains the average rainfall, evaporation and river flow

components for the basin to Melford Bridge. Groundwater flow in the river Thet is commonly a

high proportion of the total flow during the summer six months, April to September.

95

Case Studies

Table 5.1.1 Mean rainfall, evaporation and river flow components : River Thet to Melford

Bridge (mm)

Rainfall

Evaporation

Total flow

Base flow

Jan

57

6

31

11

Feb

43

10

28

13

Mar

39

18

24

14

Apr

48

32

19

13

May

46

60

13

11

Jun

45

78

9

7

Jul

65

78

6

5

Aug

59

70

5

4

Sep

57

50

5

4

Oct

59

23

8

4

Nov

64

15

15

6

Dec

55

10

23

8

Annual Total

637

450

187

100

The bed rock in the pilot area is the Upper Chalk which is overlain in places by super

ficial deposits comprising mainly sands and gravels or Chalky Boulder Clay. The geological

succession is given in Table 5.1.2, and a brief description is given of the characteristics of

the alluvium, boulder clay and the Chalk.

Table 5.1.2 Geological Strata in the Thet Pilot Scheme area

Period Stratigraphie Unit Approximate maximum

thickness (m)

Recent Blown sands

Alluvium

Terrace deposits

Pleistocene Loam

Boulder clay

Sand and gravel

5

30

15

Cretaceous Upper Chalk

Middle Chalk

Lower Chalk

Gault

120

70

60

65

Alluvium deposits occur along the flood plain of the rivers. The alluvium has a

variable lithology reflecting the properties of the nearby Pleistocene deposits of Boulder

Clay and sand and gravel. Much of the higher ground and parts of the valleys are covered by

Boulder Clay, which is a stiff clay with a variable lithology. Infiltration to the Chalk is

reduced beneath the Boulder Clay due to its relatively impermeable nature, and groundwater in

the Chalk tends to be confined in these areas.

Areas of bare ChdUi crop out along the sides of valleys between the Boulder Clay on the

hills and the valley alluvium. The Chalk is a soft white fine grained limestone composed

almost entirely of organic remains. Although it has a porosity of between 25 and 40 per cent

the pore spaces are very small and the mean flow of water occurs along joints, fissures.

96

Case Studies

bedding planes and flint surfaces. The size of fissures generally reduces with depth and

most groundwater is obtained from the top 30 m of the aquifer, although individual hard beds

of rock below this may contribute significant supplies. The transmissivity varies from less 2 2 2

than 10 m /day below the high ground to between 750 m /day and 1500 m /day or more at favourable localities in the valleys. Values for the specific yield are generally between

1.0 and 3.0 per cent.

The Chalk is subdivided into three divisions, the Lower, Middle and Upper Chalk.

Groundwater levels fluctuate within the Upper Chalk or in the overlying drift. At the base of

the Upper Chalk there are hard bands of rock up to 8 m thick which are an important water

bearing horizon- Likewise, a hard band of rock exists at the base of the Middle Chalk. The

Lowev Chalk is generally less favourable for groundwater development.

5.1.3 Measurements

Sub-surface geophysical work in connection with the Pilot Scheme was begun in 1967 with the

objective to determine the physical properties of the Chalk, the correlation of the strata

between the wells and the study of groundwater quality. A resistivity survey was run to a

depth of 180 m in the deepest available borehole with measurements at 1 m intervals. The log

obtained was later used to interpret resistivity logs from the 18 abstraction wells.

Measurements of electrical conductivity and temperature confirmed that the quality of the water

did not change with depth. In addition the following logs were run:

spontaneous potential continuously recorded

single electrode resistance continuously recorded

multiple electrode resistivity 0.5 m readings

Gamma ray continuously recorded

Caliper continuously recorded

Spinner flowmeter selected intervals

In the western part of the pilot area the final depths of the abstraction wells are

related to the base of the Upper Chalk, and the final depths of the wells were controlled by

the use of electric logs.

The hydrometric network was established in two stages. The first stage was

substantially completed by the end of 1967 and involved the drilling of 24 observation bore

holes of 153 mm diameter, the construction of five river flow gauges, and the installation of

eight rain gauges. The second stage included drilling a further 56 observation boreholes,

constructing a further six flow measuring points on springs and tributaries and building a

climate station near the centre of the area. The flow was measured from each abstraction well.

97

Case Studies

5.1.4 Natural River Flow and Groundwater Level Relationship

Measurements were made of river flows and groundwater levels prior to the start of groundwater

abstraction so that flow and level relationships could be derived for natural conditions.

Ideally at least two years of such records should be available to enable accurate

relationships to be derived for use in assessing the effects of groundwater abstraction.

Methods used to assess the effects of pumping included:

1. cross correlation of river flows,

2. rainfall-runoff models relating river flows to weather data,

3. aquifer characteristics, and

4. analogue model of groundwater flow.

The most accurate method of assessing the change in river flow due to groundwater

abstraction was found to be that based upon the cross correlation of river flows (Wright, 1974).

Measured and estimated natural river flows based upon this method are shown in Figure 29;

based upon five day means.

River flows may also be estimated from weather data. A monthly rainfall-runoff

regression model has been optimised for many basins in England and Wales (Wright, 1978), and

one of the basins examined was the river Thet to Melford Bridge. The derived rainfall-runoff

equation for this basin has a coefficient of multiple correlation of 0.984 and a residual error

of 15 per cent of the flow. During average weather conditions a positive increment of

rainfall less evaporation falling over the basin will continue to significantly influence river

flows for 14 months after the event due to the large groundwater storage in the Chalk. A

typical response of river flow to 10 mm of rainfall less evaporation is shown in Figure 30.

Rather more than 30 per cent of the basin comprises semi-permeable boulder clay overlying

the Chalk. This accounts for the rapid response of river flow to rainfall during wet weather.

When the soil is below field capacity a positive value of rainfall less evaporation has a

maximum effect upon river flows approximately one month after the rainfall event.

The analysis of river hydrographs and groundwater levels provided estimates of the

coefficient of storage using the following method. During long periods of dry weather the

river flow recession curve may be written as equation 3.2(5) which is:

Q t = Ô 0e - Œ t 5.1(1)

and after integration becomes:

S t = — 5.1(2)

The relationship between mean basin groundwater levels and base flow may be written as:

W = a + bF 5.1(3)

98

Case Studie

a) G R O U N D W A T E R ABSTRACTION

_ 0.5 -

5 o

1.5

1.0

5 o

0.5

b) R IVER F L O W

I I

I ' 1_J target

regulation

flow

f low

augmentation

estimated natural flow

T T Sept May July Nov

1 971

Figure 29 Regulation of river flow at Bridgham by groundwater, 1971

By definition the coefficient of storage is the change in volume of stored groundwater

divided by the change of groundwater levels during the same period, thus:

st - St Cl 2 S = x 100%

1 2 5.1(4)

Combining equations 5.1(3), 5.1(4) and 5.1(5) we obtain:

Case Studies

2A 1 0 m m of residual rainfall (rainfall less evaporation)

Figure 30 Response of river flow to rainfall less evaporation at Melford Bridge

If the water levels are in metres and the base flow is in millimetres of runoff for a

period of one month/ and the dry weather flow approximates to the base flow, then equation

5.1(6) reduces to:

S = 0.00329

°<b 5.1(6)

where Q and Q are the flows at time t and when t = o respectively,

S is the volume of groundwater in active storage at time t,

W is the mean basin groundwater level

S is the coefficient of storage

F is the mean monthly base flow

and e, Œ , a and b are constants.

An analysis of the base flow and groundwater levels in the Thet catchment gave a value

of b = 0.19 in equation 5.1(3) , and a dry weather flow recession constant of = = 0.0113 in

equation 5.1(1). Using these values in equation 5.1(6) the storage coefficient was

calculated to be 1.5% which agreed with values derived by other methods.

An electrical analogue model was constructed of the Pilot Area and the surrounding

region based upon the analogue between Ohm's Law and Darcy's Law. The system was modelled as 2

a slow-time resistance-capacitance network covering an area of 1050 km . Resistor elements

simulated transmissivity and capacitative elements simulated the storage properties of the

Chalk. The average groundwater levels were modelled by simulating the annual average values

of infiltration and the variation in transmissivity within the aquifer. The response of the

model suggested that the storage coefficient varied from 3% where the Chalk crops out to

100

Case Studies

0.05% where the groundwater is confined by extensive boulder clay cover, giving an average

value for the area of 1.5%.

The analogue model enabled many pumping programmes to be investigated, and examined the

effect of inflow of groundwater into the Pilot Area during periods of maximum drawdown. In

regional developments this inflow would tend to reduce due to drawdown also occurring in

adjacent areas.

5.1.5 Analysis of Group Pumping Tests, 1971

During 1971 the groundwater abstraction rate was varied to regulate river flows at Bridgham to

predetermined rates, as shown in Figure 29. This method of river regulation increases low

flows and reduces flood peaks. Early in May the measured flows were similar to the estimated

flows indicating that the depressions in the groundwater table made by abstractions the

previous year had been substantially filled by the natural winter recharge. Later on in May

and June the groundwater abstraction was increased to keep the river flow at Bridgham near to

the initial regulation rate of 1000 1/sec. Minor floods in June and early August enabled

groundwater abstraction to cease from most wells for a day or two. At the beginning of each

day the flows at Redbridge and Bridgham were examined and the number of pumping wells

determined. The travel time of pumped groundwater discharges reaching Bridgham varied from

1 hour for nearby wells to 7 hours for those most distant, representing on average an effective

river flow velocity of 2 km/hr.

The pumping pattern in 1971 can be considered to comprise three stages separated by

prolonged periods of rain. During the first period which extended up to 9 June, the average

pumping rate was 300 1/sec and the average gain in flow 150 1/sec. The difference comprised

groundwater that was pumped which would naturally have found its way to the river as base flow.

Pumping in this period was mainly from wells at some distance from the main river channels

therefore induced infiltration through the river bed would have been negligible (Kemp and

Wright, 1977).

The second period of pumping, from 22 June to 27 July involved pumping 15 of the 18 wells

for much of the time. The net gain in flow generally varied between 70 to 80% -of the

abstraction rate which was similar to that for the third period which ended on 12 October.

Induced river bed infiltration probably did not exceed 10% of the groundwater abstraction.

Higher losses could have occurred during the 1970 tests when groundwater levels were lowered

below the river bed along several reaches.

Extensive tests in a number of pilot schemes have shown the economic advantages of care

ful selection of sites for abstraction wells. In areas where the river bed is relatively

impermeable wells may be sited close to the river to take advantage of the higher coefficients

of storage in such areas and reduced capital costs due to short pipelines to the river. Where

the river bed is permeable the wells have to be sited some distance from watercourses to avoid

induced river bed infiltration.

101

Case Studies

5.2 Temperate Area: The Moscow Artesian Basin, USSR

5.2.1 Introduc tion

This case study describes the interaction between river water and groundwater in the Moscow

Artesian Basin which is an area with intensive groundwater development. The groundwater regime

of the basin represents a complicated hydrodynamic system with the considerable groundwater

abstraction reducing groundwater runoff and modifying the natural interaction between ground

water and surface water.

In temperate climates a typical characteristic of the interaction is the groundwater

inflow to rivers. However high levels of groundwater development generally deplete both the

groundwater inflow (base flow) and the total runoff. Where the groundwater abstraction has

created a cone of depression that extends to the river channel runoff may be reduced also by

the loss of river water to groundwater. However the return of groundwater abstractions to

rivers will increase runoff in those areas where the groundwater would not otherwise have

contributed to runoff. Estimates are required of such changes in the natural interaction

between groundwater and river water to solve problems associated with hydrological aspects of

multi-purpose water resource schemes that include groundwater development.

5.2.2 The Moscow Basin

An analytical solution has been described by Minkin and Kontsebovsky (1979) to estimate the

decrease in runoff due to the effect of specified major groundwater abstractions. However the

solution for the Moscow basin is complicated by several factors such as the regional nature

of the problem caused by the extensive network of abstraction wells and the availability of

detailed hydrogeological information. Moreover an adequate account has to be taken of the

variation in runoff caused by the return of groundwater to rivers after for example domestic

or industrial use.

The regional assessment of the interaction between surface water and groundwater, and the

prediction of variations in runoff due to the effect of intensive groundwater development,

should be obtained by the analysis of hydrometric data (Anon, 1973b; Anon, 1974). Thus

artificial changes in runoff may be assessed by taking the difference between the measured

river flow and the estimated natural flow. To illustrate this method of investigating the

interaction some analyses are presented for the Moscow Artesian Basin.

The region under investigation has a temperate climate and lies near the centre of the

Eastern European plain. The countryside is typically undulating with hills, valleys, moraine

ridges and lowland areas, and is at the centre of the Moscow Artesian Basin. This basin is

hydrogeologically complex with aquifers that are drained by a network of surface channels that

have developed in the zone of intensive water exchange associated with the Quaternary and

Mesozoic deposits (Lebedeva, 1972).

102

Case Studies

5.2.3 Analyses

Base flow in this area is typically associated with a backwater type of regime as described in

section 3.2.1 and Figure 2. There is a relatively small variation in base flow during a year.

The effect of man's activities, in particular the intensive and substantial groundwater

development, has changed the hydrological regime of the river basins. Thus there now exists

the problem of determining the natural groundwater resources of the region and substantiating

the future safe groundwater yield based upon the available data. To solve this problem

required the estimation of the groundwater inflow to the rivers, together with the associated

inflow parameters and the determination of the difference between the measured river flows and

the estimated natural runoff.

Groundwater inflow values were obtained for successive reaches of the drainage basins by

analysing the measured long-term runoff data, supplemented by low flow measurements taken as

part of hydrometric surveys during notable historic dry periods. From this information the

long-term annual groundwater inflow values also were obtained (Ratner, 1978). The careful

location of river gauging stations, based upon hydrometric surveys, enables due account to be

taken of hydrogeological features that influence the formation and characteristics of ground

water inflow.

Annual values of the groundwater inflow were obtained using a simplified method based

upon low flow data recorded during the summer and winter low flow periods. This method is

suitable for those cases where there is a relatively small variability in groundwater runoff.

The groundwater runoff characteristics of the zone under consideration can be determined from

an examination of the relevant hydrogeological data (Popov, 1972).

Values of the natural groundwater inflow were estimated for those reaches with a

disturbed river flow regime caused by the effect of groundwater abstraction. Methods used to

reconstruct the river flow hydrograph included correlation, multiple correlation, channel water

balance and hydrological analogy. The method used for specific cases depended upon the

available data and the extent to which the natural regime, relating to the interaction between

groundwater and river water, was changed by man's actions (Ustiuzhanin, 1974). All the above

methods were used in deriving the probability distribution of the long-term groundwater inflow.

The deviations of the estimated actual groundwater inflow from the reconstructed natural

values for each water year were assumed to be equal to the change in the groundwater inflow

caused by groundwater abstraction.

This method may be applied to assess the natural groundwater resources in areas where

man has changed the natural interaction processes between river water and groundwater. Also

due account may be taken of the variations in groundwater inflow caused by intensive ground

water development. Thus up-to-date estimates can be made of variations in river flow caused

by reduced groundwater inflow or by the infiltration of river water to cones of depression in

the water table aaused by groundwater abstraction. The latter is applicable to specific

regions near industrial cities as shown in Table 5.2.1. The generalised regional situation

illustrating the disturbance to natural groundwater inflow to rivers in the central part of

103

Case Studies

the Moscow Artesian Basin is shown in Figure 31.

Table 5.2.1 Reduction in runoff caused by depressed groundwater levels in selected reaches

of rivers near industrial cities in the Moscow Artesian Basin

River

1

Kliazjma

Kliazjma

Pakhra

River Reach

2

near Noginsk

near Orekhovo-

zouevo

near Podolsk

Year

3

1970

1971

1970

1971

1970

1971

Basin Area

(km2)

4

210

210

2360

2360

1310

1310

Length of

reach

(km)

5

10.0

10.0

37.2

37.2

25.2

25 .2

reduced ground

water inflow

(m3/s)

6

0.6

0.45

0.80

2.50

1.40

0.70

Reduction in r

river water

infiltration

(m3/s)

7

0.2

0.65

0.00

0.00

0.44

1.18

unof f

total reduction

(m3/s)

6+7=8

0.80

1.10

0.80

2.50

1.84

1.88

(1/s.km2)

8-4=9

3.8

5.2

0.3

1.1

1.4

1.4

(1/s.km)

8*5=10

80.0

110.0

21.5

67.0

73.0

75.0

Moscow near 1971 4300 Lytkarino and Bronnitsy

54.0 10.0 2.3 185.0

Moscow near Kolomna

1971 2100 18.0 (8.3) (4.0) (460.0)

Kliazjma near Kovrovo

1972 3800 16.0 2.0 0.0 2.0 0.5 125.0

Within the area under consideration two main regions may be identified where there are

substantial disturbances to the natural interaction between groundwater and river water. These

two regions are within the influence of the so called 'Moscow' and 'Meshcherskij' depressions

in the water table caused by intensive groundwater development. At the present time in the

Moscow region the groundwater inflow to rivers is reduced by 20-35% of the annual natural

value, and the annual total runoff is reduced by 5-25%. The corresponding values for the

second region are 25-60% and 10-25% respectively. A comparison of these changes in runoff with

the rates of groundwater abstraction in the respective basins proves the decisive influence

that groundwater development has on these changes in runoff.

5.2.4 Future Situation

By the year 2000 the two regions are predicted to be within the area of a major depression in

the groundwater table caused by continued intensive groundwater development (according to

V.S. Plotnikov) . Therefore it is necessary to estimate a safe level of groundwater development

104

Case Studies

Key 1 Regional decrease in groundwater inflow 2 Increased runoff due to effluent discharges 3 Predicted regional lowering of groundwater table from the year 1980 to 2000 4 River reaches with reduction of groundwater inflow 5 River control sections 6 Reduction of groundwater flow as percentage of annual mean

Figure 31 Disturbance of groundwater inflow to rivers in the central part of the Moscow

Artesian Basin (after B.S. Ustiuzhanin)

for the whole of the Moscow Artesian Basin. Because the methods described are reasonably

accurate, they are used to estimate possible changes in runoff and regional changes in the

interaction between groundwater and river water. Thus the methods should be the basis for

organising future hydrological and hydrogeological investigations to enable an adequate assess

ment to be made of the impact of various water resource projects.

105

Case Studies

5.3 Arid Area with Irrigation: Chu Valley, USSR

5.3.1 . Introduction

The development of irrigation can change significantly the components of the water balance of

a basin and modify the river flow regime, including for example changing the interaction

between river water and groundwater. Changes in this interaction may be quantified using the

channel water balance method described in section 3.1. Even in areas with limited hydrological

data this method may be used to assess generalised changes that occur in the flow regime. To

illustrate this type of situation an example is given based upon information for the Chu river

basin in Middle Asia (Sumarokova, 1976} Kharchenko and Tsytsenko, 1976).

5.3.2 The River Chu

From its mountainous upper basin the river Chu flows through the Boam gorge, after which it

flows along the Chu valley and eventually disappears in the deserts of Mujunkumy and Betpak-

Daly. The climate of the Chu valley has typical continental features with a dry summer and

low annual precipitation which in the flood plain is 170-180 mm. Evaporation from surface

water is 850-900 mm per annum.

5.3.3 Description of the Study Reaches

The channel water balance has been computed for the zone of intensive irrigation development

which has been divided into six typical reaches. These reaches were chosen to ensure that

homogeneous conditions obtained relating to both the interaction between river water and

groundwater and the location of the water body inducing changes in the flow regime. A brief

description is given of each reach.

Reach 1 is from the Ortotoisk reservoir to the confluence of the river Chon-Kemin. Here the

river flows along the narrow Boam gorge and has insignificant tributaries.

Reach 2 is from the confluence of the river Chon-Kemin to the Buruldaiskii bridge which is at

the point where the river flows into the Chu valley.

Reach 3 is from the Buruldaiskii bridge to the town of Tokmak. In this zone there are losses

of river water through the permeable detritus which comprises relatively large rock

fragments. Groundwater is located at a depth of 25-100 m.

Reach 4 is from the town of Tokmak to the downstream side of the Chumysh barrage. This is

characterised by an intensive wedge-shaped stream of groundwater flowing into the

channel in addition to subchannel flow which occurs because of a change in the

composition of the valley deposits. The gravel and pebble deposits of reach three have

been replaced by a sandy clay loam.

Reach 5 is from the downstream side of the Chumysh barrage to the Tashutkul barrage. This

reach differs from the last because of the considerable volume of drainage water that

is collected. There is a small flood plain with a typical flat channel.

106

Case Studies

Reach 6 is from the downstream side of the Tashutkul barrage to the village of Furmanovka.

The characteristics are similar to those of reach five. Below Furmanovka the river

flows in the direction of the Aral Sea and disappears in sand.

5.3.4 The Channel Water Balance

The channel water balance of the separate reaches is computed by analysing the river flow data

for the upstream and downstream sections of the reach. Hydrometrie data and abstractions are

available for the period 1911-1973. Precipitation on the water surface and evaporation have

been considered to be practically equal. To reduce the effect of random errors in estimates

of some of the water balance components, the computation has been completed to show the main

characteristics based upon average values for successive five year periods. The results are

shown in Table 5.3.1. The channel water balance for the study period enables the following

conclusions to be made for each reach concerning the characteristics of the interaction

between river water and groundwater.

In reaches one and two there are alternating periods of groundwater inflow and outflow,

generally within the limits ..of 1—3 m / s . These flows are relatively small and close to the

errors associated with the discharge measurements for the reaches.

In the third reach there are relatively substantial losses to groundwater of 17-33 m /s .

The data from recent years have shown losses to be generally lower than in earlier years. This

is due to augmented river abstractions that have decreased the mean discharge of the reach.

There is a sustained groundwater inflow to the river in reach four with maximum values

coinciding with periods of maximum river water abstraction and the relatively wet ten year

period 1950-1960. The augmentation of groundwater inflow in such periods apparently reflects

the increasing volume of return water from irrigated areas.

In reach five there is an overall tendency for groundwater inflow to be augmented with

increasing irrigation. A considerable increase of groundwater inflow has been observed with

the opening of the West Great Chu canal. Although there has been an increase in irrigated

areas over the last two decades this has coincided with a more rational use of water for

irrigation. Thus from 1960 there has been a tendency for groundwater inflow to decrease even

during relatively wet years such as 1969 and 1970 (Sumarokova, 1976).

During the period under consideration stream flow losses in reach six have been modified

by increments of groundwater inflow from returned irrigation water.

5.3.5 Summary

By examining the information in Table 5.3.1 it is apparent that during the period of study

certain trends have occurred in the data that quantify the interaction process. The total

groundwater losses in all reaches have decreased from 40 to 22 m /s. During the same 63-year

period the total groundwater inflow, including returned irrigation water, increased from 3 3 3

23 m /s in 1911-1915 to a peak of 67 m /s in 1956-1960 after which it decreased to 33 m /s by

107

Case Studies

1971-1973. The errors associated with these estimates as 25-35% according to Karasseff

(In Anon, 1977a).

Table 5.3.1 Tabulated values of groundwater inflow and losses for selected reaches of the

river Chu, 1911-73

Reach

1

2

3

4

5

6

Losses

Gains

1911-15

3.42

1.30

-32.1

18.4

-0.50

-7.98

40.6

23.1

Mean gain (+) or Loss (-)

1916-20 1921-25 1926-30

0.60 1.30

-12.4

in period

1931-35

3.22

1.30

-33.7

15.7

4.05

-8.10

41.8

24.3

(m3/s)

1936-40

2.47

-0.70

-26.2

20.4

-3.37

-8.27

38.5

22.9

1941-45

2.10

0.42

-28.2

18.2

7.56

-7.93

36.1

28.2

Reach

1

2

3

4

5

6

Losses

Gains

1946-50

-1.00

0.10

-27.5

20.4

13.2

-10.5

39.0

33.7

Mean gain

1951-55

-0.67

1.39

-30.5

42.9

16.5

4.61

31.2

65.4

(+) or Loss

1956-60

1.02

-1.54

-24.4

33.2

23.7

9.73

25.9

67.6

(-) in

1961-65

-1.45

-2.52

-17.0

24.3

12.1

9.80

21.0

46.2

period (m /s)

1966-70

-3.67

-0.25

-17.6

29.4

10.3

15.0

21.5

54.7

i

1971-73

-3.97

-0.50

-17.7

14.6

11.0

6.53

22.2

32.1

Although parts of the channel water balance compilation described in this study are

approximate, nevertheless they are of considerable practical importance. In particular they

have become the basis of more detailed water balance studies for the Chu river basin to

establish the influence of irrigation schemes and the effect upon water resources of changes

in the flow regime. Moreover, the analyses relating to the interaction process have already

enabled a number of practical measures to be adopted for the more rational use of water

resources in the Chu valley. For example in areas where a significant loss of water has been

identified through the permeable detritus, these losses may be reduced to zero by facing the

channel with concrete. This case study shows that the channel water balance method may be used

with standard hydrometric data to quantify the interaction between groundwater and surface

water, without necessarily any special field investigations.

108

Case Studies

5.4 Arid Area: Groundwater Replenishment by Surface Water, Tunisia

5.4.1 Introduction

This case study is an example of aquifer replenishment from surface runoff in an arid region

and is for the plain of Kairouan in Tunisia (Besbes et al., 1978). The plain of Kairouan 2

has the largest groundwater storage in central Tunisia, with a surface area of 3 000 km . It

is replenished mainly by flood water from two ephemeral streams the wadis Zeroud and

Merguellil. Flood waters in this internal drainage basin spread out and infiltrate to the

groundwater table then flow to the sebkhas which are natural depressions with brackish water

and evaporate. The annual potential evaporation of 1 500 mm substantially exceeds the mean

annual rainfall of 300 mm.

5.4.2 Aquifer Recharge in the Kairouan Plain

The upper aquifer comprises a concave depression that has been filled with lenticular

Pliocene - quaternary detritus which is more than 500 m thick, in which the phreatic ground

water surface covers the entire plain. However, there exists a second aquifer at depth which

is contained within a semi-permeable Quaternary formation.

Recharge areas around the boundary of the plain provide a common intake area to both

aquifers. Figure 32 shows the groundwater levels in the Kairouan plain, with the higher levels

associated with the wadis Zeroud and Merguellil. This confirms that recharge occurs

preferentially in these areas. Although the low flows of these two wadis completely infiltrate

on reaching the plain, an examination of the movement of groundwater levels and surface runoff

events over a period of several years shows that the floodwaters play a major role in

replenishing the groundwater (Figure 33) .

Annual runoff from the Zeroud and Merguellil wadis that infiltrates to the groundwater 6 3 6 3

table in the Kairouan plain is estimated to average 12 x 10 m and 7 x 10 m respectively, 6 3

whereas direct recharge from rainfall is estimated as 6 x 10 m per annum over an area of

1 000 km .

5.4.3 Recharge by Surface Runoff from the Zeroud Wadi

The determination of effective increases in groundwater levels resulting from floods will be

described for the period of 18 months from January 1968 to June 1969. Most of the 23

piezometers in the network recorded changes in the groundwater level corresponding to the

passage of four flood events. These floods were measured at the river gauging station at Sidi

Saad during the following months:

6 3 1 February to April 1968, 15 x 10 m

6 3 2 June 1968, 54 x 10 m

6 3 3 September 1968, 4 x 10 m

6 3 4 April to May 1968, 7 x 10 m

Case Studies

Figure 32 Groundwater level contours in the Kairouan plain

Groundwater levels at location Z3 began to decline on 10 July 1968 following the flood

event of June that year. The recession in groundwater levels after 10 July was analysed and

was found by trial and error to be asymptotic to a depth of 27.8 m. With this information two

experimental points are sufficient to calculate the recession constant. Then other points can

be calculated and the groundwater recession curve drawn as shown in Figure 34. All the

recession curves may now be deduced for each flood event by means of a simple translation.

110

Case Studies

FLOW OF ZEROUD AT SIDI SAAD (1/9/69-31/8/74).

"E_

w^ time (days x10).

oo 36.50 73.00 109.50 146.00 182.5

Figure 33 Hydrograph of river flows and groundwater levels in the Zeroud wadi

The same analysis can be repeated on readings from each piezometer to determine the effective

recharge for each flood event, and from this information maps may be drawn showing the depth

of recharge (iso-recharge) for these periods.

Ill

Case Studies

Figure 34 Groundwater levels and increments of recharge at piezometer Z.., 1967 to 1969

The maps show that in the upper part of the plain the groundwater levels reach a maximum

level approximately 140 days after the flood event. In the downstrean part of the plain an

event such as that of June 1968 caused a significant increase in levels only by the winter of

1969. This lateral spread of the flood wave in the saturated zone is dependent upon the

aquifer characteristics. For example the porosity is 0.1 and 0.05 in the upstream and down

stream parts of the plain respectively. Calculations show that 140 days after the flood 6 3

recharge event, the volume of saturated rock increased by 150 x 10 m which was caused by 6 3

13.5 x 10 m of water reaching the water table. In addition flood water that has infiltrated

the surface may take four months to flow through 60 m of the unsaturated zone to reach the

aquifer below.

112

Concluding Remarks and Recommendations

6. Concluding remarks and recommendations 6 .1 Concluding Remarks

The traditional division between surface water and groundwater disciplines has tended to be

reduced in recent years with the result that some useful advances have been made in under

standing the interaction process. This report has emphasised the importance of understanding

the interaction between surface water and groundwater in different climatic conditions. In

different climates certain aspects of the interaction are dominant. For example, in arid

regions groundwater recharge may be derived from mountain rivers or from intermittent surface

runoff that is generated by intense storms, in temperate regions recharge is derived mainly

from precipitation and in colder regions recharge is associated with prevailing temperatures

and snow melt.

In many regions an understanding of the interaction process is necessary for the

satisfactory operation and long-term planning of water resource schemes. This may include a

study of the characteristics of groundwater recharge, aquifer properties, groundwater flow and

river flow. In temperate regions the groundwater component is frequently the main component of

low river flows, and in arid regions schemes may be designed to increase groundwater recharge

from surface runoff. Water resource schemes are beina developed to take advantage of the

differing storage, recharge and flow characteristics of surface water and groundwater. For

example some schemes are designed to utilise groundwater storage to artificially regulate river

flows. High flows are reduced and low flows are increased, thus reducing the natural

variability of river flows. The efficient development of water resources in all these cases

depends in part upon a study of the interaction between surface water and groundwater.

In arid areas techniques to fully utilise available water resources may include dams

strategically placed to induce a higher proportion of groundwater recharge during times of

flood flow. Thus a higher proportion of recharge occurs at a favourable location for

resource development and flows to the oceans, seas or other saline areas are reduced. Where

irrigation is practiced careful supervision is necessary to utilise fully the available water

and at the same time to minimise the build up of harmful chemicals in the soil. The continued

re-cycling of irrigation water through soils, aquifers and river channels may lead to salinity

problems and reduced crop yields.

113


The development of groundwater resources in temperate regions is likely to have some

influence upon the quantity and quality of river water especially during periods of low river

flow. In addition, the direct abstraction of river water for supply purposes is limited by

quantity and quality considerations. During periods of low river flow the main component of

flow may be from groundwater sources. Thus, although the quality of groundwater generally is

very good, this can be maintained only by the careful control of possible sources of pollution.

Once a groundwater source has become polluted it may take several years and a considerable

cost to restore the aquifer to its original state. The quality of groundwater and base flow

may be affected seriously by fertilisers applied to farmland, waste disposal tips and

accidents such as oil or chemical spillage.

Various improvements have been made in recent years in the methods for assessing the

interaction between surface water and groundwater. These improvements are in several fields

such as instrumentation, the use of tracers and improved conceptual and mathematical models of

the system. Improvements have been made in the accuracy of estimating all elements of the

hydrological cycle. A method that has been used widely to assess the interaction process is

that based upon the channel water balance. Measurements of river flow are made at carefully

selected locations and appropriate meteorological and hydrogeological information is obtained.

In addition the analysis of river flow hydrographs including flow separation and recession

curve analysis can provide useful information.

A particularly useful aid in understanding the interaction process has been the use of

tracers. For example environmental isotope techniques now provide the hydrologist and hydro-

geologist with a method to study the actual mass transfer of water. The method has the

advantage that radioactive tracers are not introduced into the system so that problems of

health and safety do not arise. Furthermore the scale of the investigation is not limited

since studies may be carried out in different climatic regions far removed from the laboratory

where the analyses are made.

In many situations the use of a mathematical model is advisable and often essential to

investigate the interaction between surface water and groundwater. The models may comprise

mainly the groundwater aspects, the combined surface water and groundwater system, or mainly

the surface water components. In addition the model may include water quality considerations.

A complete knowledge of the hydrogeology and the geometry of a given problem area is rarely if

ever available. Thus estimates have to be made of suitable parameter values to be used in

models. Sensitivity analyses may be carried out over a range of parameter values to give an

indication of the reliability of a simulation exercise. Such analyses may be used also to

indicate what further field work is necessary so that each component of the hydrological

cycle can be represented adequately by the modelling technique.

In addition models may be used with long-term weather records, where available, to

synthesise long sequences of groundwater recharge and river flow data. By such means the

severity of historic droughts can be assessed and the synthesised data may be used to assist in

the management and design of water resource schemes.

114


The further understanding of the interaction process will be of considerable benefit to

mankind. The management of water resources may be improved, in arid regions a scarce commodity

may be utilised more efficiently, crop production may be improved and safeguarded, and in

temperate regions local amenities and nevigation may be protected. The characteristics of the

interaction should be understood so that harmful developments are avoided and positive benefits

ensue. The international dissemination of knowledge as described in this report is intended to

contribute towards this objective.

6.2 Recommendations and Further Research

International co-operation should be maintained to disseminate information concerning the

development of new instruments and techniques that contribute to the understanding of the

interaction between surface water and groundwater.

Some elements of the hydrological cycle are not easy to measure. Those that require

further effort to improve the accuracy of measurement include, the estimation of areal rainfall

in arid areas, the estimation of evaporation for specific types of land use at various

latitudes, the estimation of flow characteristics within the unsaturated zone and the

estimation of river flows in arid areas.

The use of remote sensing for estimating areal rainfall and soil conditions should be

investigated further.

Changes in land use may directly or indirectly influence the quantity and quality of both

groundwater and river flow. Therefore the effect of changes in land use should continue to be

studied especially its effect upon evaporation, groundwater recharge and quality.

Water quality problems that may arise due to human interference includes those associated

with heavy metals such as cadmium, lead and mercury. These should be monitored together with

the presence of bacteria and viruses in groundwater and base flow.

The movement of water in the lower levels of aquifers and semi-permeable rocks should be

studied to assist in the location of waste disposal sites. In particular the location of

nuclear waste sites should take into account the movement of groundwater and its interaction

with surface water.

Techniques for estimating the age of groundwater may be developed further as an aid to

understanding the movement of groundwater at depth.

Many types of mathematical model are available for investigating groundwater and surface

water problems and these should be used to investigate the characteristics of the elements of

the interaction process.

115

References

REFERENCES

Amusia, A.Z. 1974. Priemy otsenki podzemnoi sostavliaiushchei stoka gornykh rek pri razlichnoi gidrogeologicheskoi izuchennosti (Methods of estimating the groundwater flow component of mountain rivers with variable information). Trans, gos. gidrol. Inst., vol. 213, p. 44-52.

Anon. 1968. Izuchenie podruslovogo stoka pri vodnobalansovykh issledovaniakh (studies of sub-channel runoff in water balance investigations). Leningrad. Trans, gos. gidrol. Inst. 44 p.

Anon. 1973. Metody issledovanii vodnoga balansa territorii i kartirovanie ego elementov (Methods of studying the areal water balance and variations in its elements). Moscow, Edition of Inst, of Geography SA SSSR, 220 p.

Anon. 1974. Otsenka podzemnogo pritoka v reki i estestvennykh resoursov podzermykh bod b usloviakh naroushenia rezhima reahnogo stoka (Estimates of groundwater inflow to rivers and groundwater resources in conditions of changes in the river flow regime). Leningrad, Gidrometeoizdat, 53 p.

Anon. 1977a. Metodicheskie- ukazania upravleniem Gidrometsluzhby N90. Sostavlenie ruslovykh vodnykh balansov (Methods recommended by the Chief Administration of the Hydrometeorological Service N90. Compilation of the channel water balance). Leningrad, Gidrometeoizdat, 103 p.

Anon. 1977b. Gidrometricheskaia otsenka vzaimodeistvia rechnykh i podzemnykh vod (Hydrometrie estimate of the interaction between groundwater and river water). Leningrad, 77 p.

Anon. 1977c. Metodicheskie rekomendatsii po otsenke tochnosti i gidrologicheskomu kontroliu dannykh gosudarstennogo ucheta vod i ikh ispolzovania (Recommended methods for estimating the accuracy, hydrological control and use of the state water data). Leningrad, Gidrometeoizdat, 117 p.

Anon. 1978. World water balance and water resources of the earth. Paris Unesco, studies and Reports in Hydrology No. 25, 663 p.

Appolov, B.A.; Kalinin, G.P.; Komarov, V.D. 1974. Kurs gidrologicheskikh prognozov (Courses on hydrological prediction). Leningrad, Gidrometeoizdat, 419 p.

Aravin, V.l.; Numerov, S.N. 1965. Theory of fluid flow in undeformable porpus media. Israel program for scientific translation, 511 p.

Backshall, W.F.; Downing, R.A.; Law, F.M. 1972. Great Ouse groundwater study. Wat. and Wat. Engng., vol. 76, p. 215-233.

Balek, J. 1978. A storm rainfall pattern above the Central African Plateau. Hydrol. Sei. Bull. vol. 23, no. 1, p. 151-156.

Barnes, H.H., Jr. 1974. Measurement of flash floods in the United States - State of the art. Symposium on Flash Floods, Paris. Publ. IAHS 112. p. 48-51.

Benson, M.A. 1968. Measurement of peak discharge by indirect methods. WMO Technical note 90, Publ. WMO 225, T.P. 119, Geneva.

Besbes, M. 1971. Etude des pertes de charge dans les forages - Application à la determination des transmissivités par essai de pompage de court durée. Ressources en Eau de Tunisie.

Besbes, M.; Delhomme, J.P.; de Marsily, G. 1978. Estimating recharge from ephemeral streams in arid regions: a case study at Kairouan, Tunisia. Wat. Resour. Res., vol. 14, no. 2, p. 281-290.

Bochever, F.M. 1969. Osnovy gidrogeologicheskikh raschetov (The fundamentals of hydrogeological computations). 2nd edition. Moscow, Nedra, 326 p.

116

References

Bochever, F . M . 1979. Principles and methods of evaluation of safe groundwater yield. International Association of Hydrogeologists Vilnius Congress, 10-15 July, mémoires, vol. xv, p . 8-11.

Brown, R . H . ; Konoplyantsev, A . A . ; Ineson, J. (eds.) 1972. Ground water studies. An international guide for research and practice. Paris, Unesco, Studies and reports in hydrology No. 7 .

Brown, L.J . ; Taylor, C . B . 1974. Geohydrology of the Kaikoura Plain, Marlborough, New Zealand. Proc. Symp. on Isotope Techniques in Groundwater Hydrology 1974, IAEA STI/PUB/373, 1, p . 169-188.

Browning, K . A . ; Bussell, R . B . ; Cole, J.A. 1977. Radar for rain forecasting and river management. Wat. Power and Dam Const., vol. 29, (12) p . 38-42.

Bruch, Jr., J .C. 1979. Seepage streams out of canal and ditches overlying shallow water tables. J. Hydrol., vol. 41, p . 31-41.

Callede, J. 1977. Oscillations journalières du bébit des rivières en l'absence de précipitations. Cahiers ORSTOM, Série Hydrologie vol. XIV no. 3, Paris.

Carlin, F . ; Magri, G . ; Cervellati, A . ; Gonfiantini, R. 1975. Use of environmental isotopes to investigate the interconnections between the Reno river and groundwater (Northern Italy). Proc. Symp. Isotope ratios as pollutant source and behaviour indicators. IAEA STI/PUB/ 382, p . 179-194.

Cherkauer, D . S . 1977. Effects of urban lakes on quantity and quality of baseflow. Wat. Resour. Bull., vol. 13, No. 6, p . 1119-1130.

Clarke, R . T . 1973. A review of some mathematical models used in hydrology, with observations on their calibration and use. J. Hydrol., vol. 19, p . 1-20.

Cooper, H . H . ; Jacob, C . E . 1946. A generalised graphical method for evaluating formation constants and summarizing well field history. Trans. Am. Geophys. Union, vol. 27, p . 526-534.

Craft, B . C . and Hawkins, M . F . 1959. Applied Petroleum Reservoir Engineering. New Jersey, Pentice Hall, 437 p .

Cruette, J.; Rodier, J.A. 1971. Mesures de débits de L'Oued Zeroud pendant les crues exceptionnelles de L'Automne 1969, Cahiers Orstom, série Hydrologie, vol. VIII, No. 1, Paris.

Cunningham, A . B . ; Sinclair, P.J. 1979. Application and analysis of a coupled surface and groundwater model. In: W Back and D A Stephenson (Guest-Editors) , Contemporary Hydrogeology - The George Burke Maxey Memorial Volume. J. Hydrol., vol. 43, p. 129-148.

Di Wiest, R . J . M . 1965. Geohydrology. New York, Wiley 366 pp.

Delhomme, J .P. ; Delfiner, P . 1973. Application du krigeage à l'optimisation d'une campagne pluviomátrique en zone aride. Actes du colloque Elaboration des projects d'utilisation des resources en eau sous données suffisantes, Unesco-AIHS-OMM. Madrid-Tome 1 p . 171-190.

Delhomme, J .P. ; Levassor, A . ; Frasnetti, J .C. 1977. Reconnaissance optimale du champ des transmissivités. International Association of Hydrogeologists Birmingham Congress, 24-30 July, mémoires, vol. xiii, part 1, p . E25-E39.

Dobroumov, B . M . ; Popov, O . V . ; Sokolov, B . L . 1976. Metody Kolichestvennoi otsenki podzemnogo pritoka v reki i ihk prakticheskoe primenenie. Trudy IV Vsesoiuznogo gldrologicheskogo siezda (Quantitative methods of estimating groundwater inflow to rivers and their practical application). Trans. IV All Union hydrological meeting. Leningrad, Gidrometeoizdat, p . 299-310.

117

References

Downing, R . A . ; Williams, B.P.J . 1969. The groundwater hydrology of the Lincolnshire Limestone. United Kingdom. Water Resources Board, Reading, 160 p .

Downing, R . A . ; Oakes, D . B . ; Wilkinson, W . B . j Wright, C E . 1973. Regional development of groundwater resources in combination with surface water. J. Hydrol., vol. 22, p . 153-177.

Dupuit, J. 1863. Etudes théoriques et pratiques sur le mouvement des eaux dans les canaux découverts et à travers les terrains permeables (2nd ed.) Paris, Dunod, 304 p .

Fontes, J . C . ; Gonfiantini, R . ; Roche, M . A . 1970. Deuterium et Oxygène-18 dans les eaux du Lac Tchad. Proa. Symp. in Isotope Hydrology 1970, IAEA STI/PUB/255, p . 387-402.

Fox, I .A.; Rushton, K . R . 1976. Rapid recharge in a limestone aquifer. Groundwater, vol. 14, p . 21-27.

Friedrich, W . 1954. Lysimetrmessung und andere gewa'sserkundliche Verfahren zur Ermittlung der Grundwassererneuerung. Z.dt. geol. Ges. Bd. 106, Teil 1, S. 41-48.

Fritz, P . ; Drimmie, R . J . ; Rendev, F . W . 1974. Stable isotope contents of a major Prairie aquifer in central Manitoba, Canada. Proa. Symp. on Isotope Techniques in Groundwater Hydrology, IAEA STI/PUB/373, vol. 1, p . 379-396.

Gonfiantini, R . ; Togliatti, V . ; Tongiorgi, E . 1962. II Rapporto 18Q/160 nell 'Acqua del Lago di Bracciano e delle Falde a Sud-Est del Lago, Noticiario del C.N.E.N., anno 8, 6, 3-9.

Great Ouse River Authority. 1972. Great Ouse groundwater pilot scheme. Final Report. Cambridge, 103 p .

Hantush, M . S . 1962. Drawdown around a partially penetrating well. Am. Soc. Civ. Eng. Trans., vol. 127, part I, p . 268-283.

Harpaz, Y . 1970. Practical experiences of well recharge. Water Research Centre, Medmenham. Artificial Groundwater Recharge Conference, paper 12.

Hazen, A . 1911. Discussion of: Dams on sand foundations by Koenig, A . C . Trans. Am. Soc. Civ. Eng. vol. 73, p . 199.

Herbert, R. 1970. Modelling partially penetrating rivers on aquifer models. Groundwater vol. 8, p . 29-36.

Herschy, R . W . (ed.) 1978. Hydrometry: principles and practices. Chichester, Wiley, 511 p .

Ineson, J.; Downing, R . A . 1964. The groundwater component of river discharge and its relationship to hydrogeology. J. Inst. Wat. Engrs., vol. 18, p . 519-541.

Institution of Civil Engineers. 1978. Thames groundwater scheme. Conference proceedings 12-13 April. Inst. Civil Eng., London 241 p .

International Atomic Energy Agency. 1969, 1970, 1971, 1973, 1975, 1979. Environmental isotope data nos. 1-6: world survey of isotope concentration in precipitation.

International Organisation for Standardisation. 19 70, 1973. Liquid flow measurement in open channels, slope-area method. I . S . O . , Geneva.

Jacob, C . E . 1950. Flow of groundwater. In: Engineering Hydraulics (Ed. Rouse, H . ) . New York, Wiley, p . 321-386.

Kafri, U . ; Arad, A . 19 79. Current subsurface intrusion cf Mediterranean seawater - a possible source of groundwater salinity in the Rift Valley system, Israel. J. Hydrol., vol. 44, p . 267-287.

Kalinin, G . P . ; Abalian, T . S . 1957. Ob opredelenii podzemnogo pitatia rek (On the determination of groundwater inflow to a river). Meteorología i gidrologia, No. 5, p . 12-20.

118

References

Kapotova, N . I . 1978. Izuchenie regima podzemnykh vod primenitelno k raschetu vodnogo balansa vodosborov. (Groundwater regime studies for the purpose of computation of the water balance of river basins). In Izuchenie: raschet elementov vodnogo balansa (Studies and computation of water balance elements), part II, edition GGI, Valday, USSR.

Kemp, J .B . ; Wright, C E . 1977. The assessment of river regulation losses. International Association of Hydrogeologists Birmingham Congress, 24-30 July, mémoires, vol. xiii, part 1, p . D1-D19.

Kharchenko, C.I.j Tsytsenko, K . V . 1976. Otsenka vliania irrigatsionnykh meropriatii na rechnoi stok : na primere r. Chu (Estimate of irrigation influence upon streamflow : the river Chu example). Trans, gos. gidrol. Inst., vol. 230, p . 6-24.

Konikow, L . F . ; Bredehoeft, J .D. 1974. Modelling flow and chemical quality changes in an irrigated stream-aquifer system. Wat. Resour. Res., vol. 10, no. 3, p. 545-562.

Konoplyantsev, A . A . 19 70. Evaluation of the groundwater balance of large territories. Symposium on world water balance. Vol. 1 IAHS-Unesco, Publ. 92. p . 71-77.

Kudelin, B. i . 1969. Printsipy regionalnoi otsenki estestvennykh resursov podzemnykh vod (Principle of the regional estimate of natural groundwater resources). Moscow, MGU 344p¿

Lebedeva, N . A . 1972. Estestvennye resoursy podzemnykh vod Moskovskogo artezianskogo basseina (Natural groundwater resources in the Moscow Artesian Basin). Moscow, Nauka, 147 p .

Levassor, A . ; Talbot, A. 1976. Sur le débit spécifique l'interprétation des essais de puits et les pertes de charge dans les forages d'eau. Ecole des Mines LHM/RD/76/8 .

Lugeon, M . 1933. Barrages et Geologie, Dunod, Paris.

Minkin, E . L . ; Kontsebovsky, S.J. 1979. Vlianie eksplouatatsii podzemnykh vod na poverkhnostny stok (Effect of groundwater development on surface runoff). International Association of Hydrogeologists Vilnius Congress, 10-15 July, mémoires, vol. xv, p . 22-26.

Moore, G . K . 1979. What is a picture worth? A history of remote sensing. Hydrol. Sei. Bull. vol. 24, no. 4, p . 477-485.

Morel, E . H . ; Wright, C . E . 1978. Methods of estimating natural groundwater recharge. Reading, United Kingdom, Central Water Planning Unit Technical Note No. 25, 40 p .

Morgante, S . ; Mosetti, F . ; Tongiorgi, E . 1966. Moderne indagini Idrologiche nella zona di Gorizia, Boll. Geofis. Teor. Appl., vol. 8, p . 114-138.

Natermann, E . 1951. Die Linie des landfristigen Grundwassere (Auz) und die Trockenwetterabflusslinie (TWL). Die Wasserwirtschaft, (Sonder H . ) .

Nordberg, L . ; Soveri, J. 1978. Seasonal and multiannual changes of groundwater storage in Norden and the adjacent part of the USSR. Nordic Hydrological Conference, Hanasaari.

Nutbrown, D . A . ; Downing, R . A . 1976. Normal-mode analysis of the structure of baseflow-recession curves. J. Hydrol., vol. 30, no. 4, p . 327-340.

Oakes, D . B . 1977. The movement of water and solutes through the unsaturated zone of the Chalk in the United Kingdom. 3rd International Hydrology Symposium, Fort Collins.

Payne, B . R . 1970. Water balance of lake Chala and its relation to groundwater from tritium and stable isotope data. J. Hydrol., vol. 11, p . 47-58.

Payne, B . R . ; Shroeter, P. 1979. The importance of infiltration from the Chimbo river in Ecuador to groundwater using environmental isotope variations. Proa. Symp. Isotope Hydrology 1978, I . A . E . A . , vol. I, p . 145-156.

Peck, A . J . ; Hurle, D . H . 1973. Chloride balance of some farmed and forested catchments in southwestern Australia. Wat. Resour. Res., vol. 9, No. 3, p. 648-657.

119

References

Penman, H . L . 1949. The dependence of transpiration on weather and soil conditions. J. Soil Science, Oxford, vol. 1, p . 74-89.

Pollack, Z . 1967. Possibility to use simplified method of estimation of transmissibility and coefficient of permeability according to unit specific yield. Mémoires de I.A.I.E. Tome VIII congrès d'Istanbul p . 481-485.

Popov, O . V . 1969a. Razvitie issledovanii podzemnogo pritoka v reki (Study of the development of groundwater inflow to a river). Trudy gos. gidrol. Inst., vol. 116, p . 5-28.

Popov, O . V . 1969b. Podzemnoe pitanie rek. (Groundwater feeding of the river). Leningrad/ Gidrometeoizdat.

Popov, o.v. 1970. Nekotorye elementy analiza vzaimodeistvia rechnykh i podzemnykh vod (Particular elements of the analysis of the interaction between groundwater and surface water). Trans, gos. gidrol. Inst., vol. 182, p . 5-25.

Popov, O . V . 1972. Prakticheskie zadachi v problème izouchenia vzaimodeistvia poverkhnostnykh i podzemnykh vod (Practical problems in investigation of the interaction between 'surface water and groundwater). Trans, gos. gidrol. Inst., vol. 188, p . 5-22.

Popov, O . V . 1975. Kteorii podzemnogo pritoka (On groundwater inflow). Trans, gos. gidrol. Inst., vol. 226, p . 5-18.

Popov, O . V . 1978. Raionirovanie prirodnykh regionov po usloviam formirovania podzemnogo pritoka (Natural regions based on groundwater inflow conditions). Trans, gos. gidrol. Inst., vol. 253, p . 3-15.

Prickett, T . A . ; Ungemach, P. 1970. Séminaire sur les applications des modèles analogiques a l'hydrogeologie, Strasbourg.

Prickett, T . A . ; Lonnquist, C G . 1971. Selected digital computer techniques for groundwater resources evaluation. Illinois State Water Survey, Bull. 55.

Prickett, T . A . 1975. Modelling techniques for groundwater evaluation. Advances in Hydroscience. Vol. 10. Academic Press, New York, p. 1-143.

Rae, G . 19 78. Mine drainage from coalfields in England and Wales a summary of its distribution and relationship to water resources. United Kingdom, Reading, Central Water Planning Unit Technical Note No. 24, 21 p .

Ratner, H . C . 1970. Tochnost ekspedistionnykh gidrometricheskikh rabot pri izuchenii podzemnogo stoka (Accuracy of hydrometric works in groundwater flow studies). Trans, gos. gidrol. Inst., vol. 182, p . 51-72.

Ratner, N . C . 1978. Raschet mnogoletnikh kharakteristik podzemnogo pritoka v reki po dannym gidrometricheskikh ciemok (Account of long-term data of groundwater inflow to rivers from hydrometric surveys). Trans, gos. gidrol. Inst., vol. 253, p . 31-41.

Ratner, N . S . 1972. Orazmechshenii seti pri gidrometricheskoi otsenke podzemnogo stoka (On the network location for hydrometric estimates of groundwater flow). Trans, gos. gidrol. Inst., vol. 188, p . 153-170.

Riggs, H . S . 1953. A method of forecasting low flow of streams. Trans. Am. Geophys. Un., vol. 34, no. 3.

Roche, M . 1966. Recherche d'un hydrogramme standard. Cahier Orstom d'Hydrologie. NO. G. Paris.

Rodda, J .C . ; Downing, R .A. j Law, F . M . 1976. Systematic hydrology. Newnes-Butterworth, 399 p .

Rushton, K . R . ; Herbert, R. 1966. Groundwater flow studies by resistence network. Geotechnlque, vol. xvi, p . 264-267.

Rushton, K . R . 1976. Rapid recharge in a limestone aquifer. Ground water, vol. 14(1).

120

References

Rushton, K . R . ; Tomlinson, L . M . 1979. Possible mechanisms for leakage between aquifers and rivers. J. Hydrol., vol. 40, p . 49-65.

Schneider, W . J . 1965. Areal variability of low flows in a basin of diverse geologic units. Wat. Resour. Res., vol. 1, no. 4 , p . 509-515.

Shestakov, V . M . 1975. Dinanrika podzemnykh Vod. (Groundwater dynamics). Moscow MGU, 327 p .

Smiles, D . E . ; Knight, J .H. 1979. The transient water table beneath a leaking canal. J. Hydrol., vol. 44, p. 149-162.

Snyder, F .F . 1939. A conception of runoff phenomena. Trans. Am. Geophys. Un., 20th annual meeting, p . 725-738.

Sokolov, A . A . ; Chapman, T . G . (eds.) 1974. Methods for water balance computations. An international guide for research and practice. Paris, Unesco. 127 p . (Studies and reports in hydrology, 17).

Sokolov, B . L . 1974. Ispolzovanie krivykh spada pri analize uslovii formirovania i raschetakh podzemnogo stoka v reki (Use of recession curves to analyse formation conditions and estimation of groundwater flow). Trans, gos. gidrol. Inst., vol. 213, p . 152-170.

Sonuga, J .O. 1977. Hydrological aspects of the drought event in Nigeria-1972/73. Hydrol. Sei. Bull. vol. 22, no. 4, p . 487-502.

Soveri, J. 1973. Pohjaveden korkeuden valtakunnallisesta havainnoinnista ja sen uudelleen jarjestelyist'á vesihallituksessa. Rakennusgeologisen yhdlstyksen julkaisuja. Vol. 8, Helsinki.

Spink, A . E . F . ; Rushton, K . R . 1979. The use of aquifer models in assessment of groundwater recharge. International Assoc, for Hydraulic Research. Proc. XVIII Congress, Sept. 1979. Cagliari, Italy.

Subramanian, V. 1979. Chemical and suspended-sediment characteristics of rivers in India. J. Hydrol., vol. 44, p . 37-55.

Sumarokova, V . V . 1976. Izmenenia vodoobmena rusia r. Chu s prilegalushchei territorii v sviazi s razvitiem oroshaemogo zemledelia (Water exchange fluctuation in the Chu river with adjacent areas in connection with irrigation development). Trans, gos. gidrol. Inst., vol. 230, p . 25-33.

Theis, C . V . 19 35. The relation between the lowering of the piezometric surface and the rate and duration of discharge of a well using groundwater storage. Trans. Am. geophys. Un., vol. 16, p . 519-524.

Todd, D . K . 1960. Groundwater hydrology. Wiley, 336 p .

Toebes, C . ; Ouryvaev, V . (eds.) 1970. Representative and experimental hasins. An international guide for research and practice. Paris, Unesco. 348 p . (Studies and reports in hydrology, 4) .

United Kingdom, Central Water Planning Unit. 1977. Nitrate and water resources with particular reference to groundwater, C W P U , Reading, 64 p .

United Kingdom, Department of the Environment and Welsh Office, 1970. Report of a river pollution survey of England and Wales. London, HMSO, 270 p .

United Kingdom, Institute of Hydrology, 1980. Low flow studies report. NERC, Wallingford.

United Kingdom, Natural Environment Research Council, 1975. Flood studies report, vol. 1 Hydrological studies, ÍERC, UK. 550 p .

References

Ustiuzhanin, B.S. 1974. Otsenka izmenenii stoka rek tsentralnoi chasti Moskovskogo artezianskogo basseina pod vlianiem kroupnykh vodozaborov podzemnykh vod (Estimate of runoff variation in the central part of the Moscow Artesian Basin under the effect of major groundwater abstractions). Trans, gos. gidral. Inst., vol. 213, p. 127-151.

Van Keulen, H. 1975. The use of simulation in the study of soil moisture transport processes. Modelling and simulation of water resource systems. North Holland Publishing Co. p. 291-298.

Vladimirov, A.M. 1966. Characteristics of formation and computation of minimum flow in the small rivers of the European USSR. Trans, gos. gidrol. Inst., No. 133.

Williams, A.F.; Holmes, J.W. 1978. A novel method of estimating the discharge of water from mound springs of the Great Artesian Basin, central Australia, J. Hydrol., vol. 38, p. 263-272.

Winter, T.C. 1976. Numerical simulation analysis of the interaction of lakes and groundwater. U.S. Geological Survey, Professional paper 1001, 45 p.

World Meteorological Organisation. 1965. Guide to hydrometerological practices. Geneva. (WMO Publication 168-TP82).

World Meteorological Organisation. 1975. Guide to hydrological practices. 3rd edition. Geneva, p. 9-513 (WMO Publication 168).

Woudt van't, B.D.; Whittaker, J.; Nicolle, K. 1979. Ground water replenishment from river flow. Wat. Resour. Bull., vol. 15, No. 4, p. 1016-1027.

Wright, C.E. 1974. The assessment of regional groundwater schemes by river flow regression equations. J. Hydrol., vol. 26. p. 209-215.

Wright, C.E. 1978. Synthesis of river flows from weather data. United Kingdom, Reading, Central Water Planning Unit Technical Note No. 26, 100 p.

Zaltsberg, E.A. 1977. Mnogoletnij rezhim orovnya gryntovyh vod znony izbytochnogo yvlazheniya evropejskoj territorii CCCP. Izvesitiya vsesoy oznogo geograficheskogo obshchectva. Tom 109.

Zektser, i.S. 1977. Zakonomernosti formirovania podzemnogo stoka i nauehno-metodicheskaia osnova ego izuchenia (Groundwater flow formation regularities and the scientific and systematic basis of its study). Moscow, Nauka 173 p.

122

Selected Papers from 1979 Symposia

SELECTED PAPERS FROM 1979 SYMPOSIA

Artificial Groundwater Recharge : Dortmund, FRG, 14-18 May

Bibby, R. ; Brown, S .K. Research into the conjunctive use of surface and groundwater with artificial recharge in Sussex, England.

Bize, J. Artificial recharge in the regions of Varamin and Garmsar, Iran.

Blasy, L . Infiltration of drainage water to maintain the natural groundwater regime. Planning and test results in connection with the projected airport at Munich 11.

Dedk, J. Investigation of the supplying and draining process of a regional groundwater flow system, using environmental isotopes.

Gholamali, F. Experiences with artificial groundwater recharge at Djahrom, Southern Iran.

Houdaille, F. The groundwater of the Albien in the surroundings of Paris : New possibilities of exploitation based upon artificial recharge.

Peck, A.J . Groundwater recharge and loss : invited review paper.

Peters, G . Aspects of planned artificial groundwater recharge in the 'Fuhrberger Feld' area.

Robert, A . Artificial recharge of groundwater in Croissy : Systems of discharge.

Tanwar, B . S . Effects of irrigation on the groundwater system in the semi-arid zone of Haryana, India.

Wildschut, R.J. Practical applications of artificial recharge in North-Holland.

Methods for Evaluation of Groundwater Resources : Vilnius, USSR, 10-15 July

Bochever, F . M . 1979. Principles and methods of evaluation of safe ground-water yield, p . 8-11.

Gokhberg, L . K . ; Roshal, A . A . The consideration of surface run-off variation in simulation of ground-water withdrawal in river valleys, p . 128-130.

Minkin, E . L . ; Kontsebovsky, S.Ya. Estimating the ground-water development effect on surface run-off. p . 22-30.

Usenko, V . S . ; Altshul, A . K h . ; Zlotnik, V . A . ; Kalinin, M . Y u . Improving methods for safe groundwater yield evaluation taking account of the water development effect on the environment, p . 333-335.

Vallner, L . K . Ground-water discharge to streams as a check criterion in regional groundwater flow evaluation (summary). p . 66-69.

International Symposium on the Hydrology of Areas of Low Precipitation : Canberra, Australia, 10-13 December

Bogomolov, G . V . ; Stankevich, R . A . ; Chaban, M . O . Interactions between groundwater and surface water at sites of large groundwater withdrawals and methods of their estimation.

Cabrera, G . ; Iroumé, A . A finite element model applied to stream-aquifer relations during floods.

Emery, P . A . Geohydrology of the San Luis Valley, Colorado, USA.

Gelhar, L . W . ; Gross, G . W . ; Duffy, C.J. Stochastic methods of analysing groundwater recharge.

Morel-Seytoux, H . J . ; Illangasekare, T . ; Peters, G . Field verification of the concept of reach transmissivlty.

123

Tilles in this series

1. T h e use of analog and digital computers in hydrology. Proceedings of the Tucson Symposium, June 1966 / L'utilisation des calculatrices analogiques et des ordinateurs en hydrologie: Actes du colloque de Tucson, juin 1966. Vol. I & 2. Condition IASH-Unesco / Coédition AIHS-Unesco.

2. Water in the unsaturated zone. Proceedings of the Wageningen Sympos ium, August 1967 / L'eau dans la zone non saturée: Actes du symposium de Wageningen, août 1967. Edited by/Édité par P . E . Rijtema & H . Wassink. Vol. 1 & 2. Co-edition IASH-Unesco I Coédition AIHS-Unesco.

3. Floods and their computation. Proceedings of the Leningrad Sympos ium, August 1967 / Les crues et leur évaluation: Actes du colloque de Leningrad, août 1967. Vol. 1 & 2. Co-edition lASH-Unesco-WMO / Coédition AIHS-Unesco-OMM.

4. Representative and experimental basins. A n international guide for research and practice. Edited by C . Toebes and V . Ouryvaev. Published by Unesco. (Will also appear in Russian and Spanish.)

4. Les bassins représentatifs et expérimentaux: Guide international des pratiques en matière de recherche. Publié sous la direction de C . Toebes et V . Ouryvaev. Publié par l'Unesco. (A paraître également en espagnol et en russe.)

5. Discharge of selected rivers of the world / Débit de certains cours d'eau du m o n d e / Caudal de algunos ríos del m u n d o / P a c x o a u B O U M H36paHHbix p e u M H p a . Published by Unesco / Publié par ¡'Unesco.

Vol. I: General and régime characteristics of stations selected / Vol. I: Caractéristiques générales et caractéristiques du régime des stations choisies / Vol. I: Características generales y características del régimen de las estaciones seleccionadas/ T O M I: 06uine H - p e w H M H u e xapanTepucriiKit ii36pannux CTamiutt.

Vol. II: Monthly and annual discharges recorded at various selected stations (from start of observations up to 1964)/ Vol. II: Débits mensuels et annuels enregistrés en diverses stations sélectionnées (de l'origine des observations à l'année 1964) / Vol. II: Caudales mensuales y anuales registrados en diversas estaciones seleccionadas (desde el comienzo de las observaciones hasta el año 1964) / T O M II: MecHMHbie H ro^oubie p a c x o a u B O A L I , 3aperHCTpnpOBaHHbie paajiHM-HtiMii H36panHbiMn CTaminuMii (c itanajia naOJiioaeHHñ n o 1964 roaa).

Vol. Ill: M e a n monthly and extreme discharges (1965-1969) / Vol. H I : Débits mensuels moyens et débits extrêmes (1965-1969) / Vol. Ill: Caudales mensuales medianos y caudales extremos (1965-1969) / T O M III: CpeflHe-MecHHHue H

3KCTpeMa.ibHbie p a c x o a w (1965—1969 rr.).

Vol. Ill (part II): M e a n monthly and extreme discharges (1969-1972) / Vol. Ill (partie II): Débits mensuels moyens et débits extrêmes (1969-1972) / Vol. III (parte II): Caudales mensuales medianos y caudales extremos (1969-1972) / T O M III (nacTb II); C p e a H e - M e c H i H w e u aKCTpe.Majibiiuc p a c x o a u (1969—1972 rr.).

Vol. Ill (part III): Mean monthly and extreme discharges (1972-1975) (English, French, Spanish, Russian). 6. List of International Hydrological Decade Stations of the world / Liste des stations de la Décennie hydrologique internationale

existant dans le m o n d e / Lista de las estaciones del Decenio Hidrológico Internacional del m u n d o / C H H C O K CTaminíl M e w n y -HapoflHoro rnflpoJiornlieCKoro ÄecfiTHJieTHH 3eMHoro m a p a . / Published by Unesco / Publié par ¡'Unesco.

7. Ground-water studies. A n international guide for practice. Edited by R . B r o w n , J. Ineson, V . Konoplyantzev and V . Kovalevski. (Will also appear in French, Russian and Spanish / Paraîtra également en espagnol, en français et en russe.)

8. Land subsidence. Proceedings of the Tokyo Sympos ium, September 1969 / Affaissement du sol: Actes du colloque de T o k y o , septembre 1969. Vol. 1 & 2. Co-edition IASH-Unesco / Coédition AIHS-Unesco.

9. Hydrology of deltas. Proceedings of the Bucharest Symposium, M a y 1969 / Hydrologie des deltas: Actes du colloque de Bucarest, mai 1969. Vol. 1 & 2. Co-edition IASH-Unesco / Coédition AIHS-Unesco.

10. Status and trends of research in hydrology / Bilan et tendances de la recherche en hydrologie. Published by Unesco / Publié par ¡'Unesco.

11. World water balance. Proceedings of the Reading Symposium, July 1970 / Bilan hydrique mondial : Actes du colloque de Reading, juillet 1970. Vol. 1-3. Co-edition IAHS-Unesco-WMO / Coédition AIHS-Unesco-OMM.

12. Research on representative and experimental basins. Proceedings of the Wellington ( N e w Zealand) Sympos ium, December 1970 / Recherches sur les bassins représentatifs et expérimentaux: Actes du colloque de Wellington ( N . Z . ) , décembre 1970. Co-edition IASH-Unesco / Coédition AIHS-Unesco.

13. Hydrometry: Proceedings of the Koblenz Symposium, September 1970 / Hydrometrie : Actes du colloque de Coblence, September 1970. Coédition IAHS-Unesco-WMO.

14. Hydrologie information systems. Co-edition Unesco-WNO, 15. Mathematical models in hydrology: Proceedings of the Warsaw Symposium, July 1971 / Les modèles mathématiques en hydro

logie: Actes du colloque de Varsovie, juillet 1971. Vol. 1-3. Co-edition IAHS-Unesco-WMO. 16. Design of water resources projects with inadequate data: Proceedings of the Madrid Symposium, June 1973 / Elaboration des

projets d'utilisation des resources en eau sans données suffisantes: Actes du colloque de Madrid, juin 1973- Vol. 1-3. Coédition Unesco-WMO-IAHS.

17. Methods for water balance computations. A n international guide for research and practice. 18. Hydrological effects of urbanization. Report of the Sub-group on the Effects of Urbanization on the Hydrological Environment. 19. Hydrology of marsh-ridden areas. Proceedings of the Minsk Symposium, June 1972. 20. Hydrological maps. Co-edition Unesco-WMO.

21. World catalogue of very large floods/Répertoire mondial des très fortes crues/Catalogo mundial de grandes crecidas/ BceMHpHbiü KaTaJior fxuibiUHX HaBonKOB

22. Floodflow computation. Methods compiled from world experience. 23. Guidebook on water quality surveys. (In press.) 24. Effects of urbanization and industrialization on the hydrological regime and on water quality. Proceedings of the Amsterdam

Symposium, October 1977, convened by Unesco and organized by Unesco and the Netherlands National Committee for the IHP in co-operation with I A H S / Effets de l'urbanisation et de l'industrialisation sur le régime hydrologique et sur la qualité de l'eau. Actes du Colloque d'Amsterdam, Octobre 1977, convoqué par l'Unesco et organisé par l'Unesco et le Comité national des Pays-Bas pour le P H I en coopération avec l'AISH. (In press / Sous presse).

25. World water balance and water resources of the earth. 26. Impact of urbanization and industrialization on water resources planning and m a n a g e m e n t . 27. Socio-economic aspects of urban hydrology. 28- Casebook of methods of computation of quantitative changes in the hydrological régime of river basins due to h u m a n activities 29. Surface water and groundwater interaction.

(I.) SC.80 /XX.29 /A

Date post:	20-May-2020
Category:	Documents
Upload:	others
View:	3 times
Download:	0 times

Studies and reports in hydrology 29 · 2014-12-28 · 2.2 Recharge of Groundwater 15 2.2.1 Recharge...

Documents