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An Assessment of regional hydrogeological framework
of the Mesozoic aquifer system of Jordan
by
Kamal Moh'd Khdier
A thesis submitted to the Faculty of Science and Engineering of
the University of Birmingham for
the degree of Doctor of Philosophy
School of Earth Sciences University of Birmingham Birmingham B15 2TT United Kingdom
December 1997
SYNOPSIS
The groundwater flow systems of the carbonate aquifer systems of the Western Highlands and Central Plateau of Jordan are complex. They reflect the changes in climate and geology of the study area.
The aquifer systems of the Western Highlands and Central Plateau of Jordan are developed in a thick sequence of Upper Cretaceous-Cainozoic carbonate rocks that dip gently east and northeastward. The sequence exhibits vertical and lateral variation in lithology; there is a general lateral transition from marine deposits (mainly carbonates) in the north and west to continental deposits (sandy facies) in the south and southeast. Since deposition, however, compression, extension, intrusive and volcanic episodes, and erosion have greatly modified the distribution and thickness of the carbonate rocks.
Based on regional contrasts in hydraulic conductivity, the regional carbonate aquifer system is divided into three aquifers separated vertically by three intervening confining units. The aquifers from top to bottom are: the Rijam Aquifer System (B4), the Amman-Wadi Sir Aquifer System (B2/A7), and the Hummar Aquifer System (A4). The B2/A7 is the most extensive and continuous aquifer system in Jordan. It is the main source of water in the country. The regional carbonate aquifer systems are underlain by a thick Precambrian- Lower Cretaceous, mostly arenaceous, sequence which comprises the deep sandstone aquifer system.
The hydraulic parameters of the aquifer systems were inferred from aquifer tests, groundwater flow modelling, and the inherent relation between the stratigraphy and hydraulic parameters. The areal distribution of hydraulic parameters generally reflects the characteristics of the sedimentary sequence; in the northern parts of the study area, where the carbonate rocks dominate and the effect of tectonics and the degree of karstification are high, a wide range and erratic distribution of hydraulic parameters are expected. In the southeast, the hydraulic parameters are more uniform due to the increase in sand content in the sedimentary sequence. Furthermore, the increase in sand content in the Lower Ajlun Group (Al-6) results in there units becoming aquifers which are then in hydraulic continuity with the overlying B21 A 7 aquifer system. However, even in the northern and western parts of the study area, the hydraulic conductivity of the A 1-6 Group is interpreted to be higher than thought before, and the Group should hence be considered an aquitard which transmitting water downwards into the deep sandstone aquifer system. Vertical hydraulic conductivity of the confining units is the most important factor affecting the regional groundwater flow system.
The flow within the regional aquifer system, in general, is controlled by the altitude of major recharge areas, major discharge areas, and major structural features. Thus topography provides the major control for the regional aquifer system.
Recharge occurs by direct infiltration of rainfall in outcrop areas, indirect recharge through the transmission losses of flood flow via wadi beds, vertical leakage through underlying and overlying strata, water transfer from adjacent aquifer systems, or by lateral boundary flow from outside the study area. The areal distribution of recharge to the regional groundwater flow regime of the carbonate aquifer system was calculated for the water budget from what is known about precipitation, total runoff, and evapotranspiration, and analysed by groundwater flow model simulation.
Most of the recharge enters the aquifer in the structurally high outcrop areas (recharge mound). Much of the groundwater from the carbonate aquifer system is discharged to the land surface by numerous springs. The locations of these springs are controlled by permeability variations in the rocks and water levels related to landsurface altitude which cause the water to discharge at the surface. The springs have been classified as local, intermediate, and regional on the basis of the surface catchment area of the spring.
Groundwater levels in the B2/ A 7 aquifer system mostly vary in response to short-term fluctuations in recharge and long-term variations in discharge. Most of the fluctuation in recharge results from cyclic patterns in precipitation, and most of the variation in discharge results from pump age trends. Water levels have declined where and when changes in the rates of recharge and natural discharge have not compensated for increasing rates of groundwater abstraction.
Groundwater flow in the study area was conceptualised as relatively shallow. intermediate, and regional flows primarily through the carbonate sediments of the Western Highlands and the Central Plateau superimposed over deeper flow through primarily sandstone sediments. Three-dimensional groundwater f10w models were used to simulate the concept of groundwater flow in the area. The area was subdivided into five subregions that approximately cover, individually, Upper Zerqa, Wadi Wala, Wadi Mujib, Wadi Hasa, and ]afr basins. Each subregion was modelled separately and then compiled in one regional model. Six model layers were used to simulate relatively shallow and deep flow. The upper five layers were used to simulate the flow in the carbonate aquifer systems. The lowest model layer was used to simulate the concept of deep flow in the sandstone aquifer system.
The results of the model calibration and sensitivity analysis show that the calibrated values of the model input are, for the most part, consistent and within the range of reasonable possibilities.
Definition of the f10w system was accomplished through examination of the following results derived from the calibrated model: (1) regional water budget, (2) potentiometric surfaces. (3) vertical leakage between aquifers, and (4) lateral flow directions in the aquifers. The model simulations show that after development the system approaches a new state of equilibrium, in which the amount of abstractions was balanced by an increase in total recharge, a decrease in discharge to river valleys, a decrease in storage, and a decrease in downward leakage and water flowing out of the system outside the study area.
Although the aim was to describe the framework hydrogeology of the carbonate aquifer system in the Western Highlands and Central Plateau of Jordan and the model simulations were entirely conceptual, this study presents estimate of the direction and magnitude of now Crom recharge to discharge areas and discusses where the results agree and disagree with the hypotheses and hydrological estimates reported by other investigators.
To my parent for providing strength and comfort To my sisters and brothers for their unlimited support
To my nieces and nephews for believing in me when it mattered
ACKNOWLEDGEMENTS
I would like to thank the many people, without whom this work would not
have been possible.
First I would like to express my gratitude to Mrs Dorothy Williams of the
Student Support and Counselling Service for her advice, guidance and moral support.
I am grateful for her keen, humane treatment, and above all for the confidence and
motivation that she has created in me during my stay in Birmingham.
I would also like to thank Mrs Thelma Barron, the former Assistant registrar of
the Science Division, for her fairness and courage. Without her support it is unlikely
that I would have been able to complete my studies.
I would like to express my thanks to Dr. J. Tellam for his encouragement and
supervision throughout the final stages of this study. His assistance, suggestions and
advice were invaluable.
I am also indebted to numerous other colleagues in the School of Earth
Sciences of the University of Birmingham for their help in many ways, particularly Mr
Petros Handjis, Mr Stephen Buss, Miss Vivi Hatzichristodulu, and Mr Richard
Greswell. I am also grateful to my internal and external Examiners who will spend
their valuable time reading and evaluating this work.
I would also like to thank the many people in Jordan who have aided field
work and provided much necessary data. Many thanks go to the staff of the Water
Authority of Jordan and Water Research and Study Centre of Jordan University.
Finally, I am indebted to everyone who contributed to completing this work
whom I have failed to mention.
CONTENT
page
CHAPTER ONE INTRODUCTION 1
1.1 BACKGROUND ................................................................................................................... .2 1.2 GENERAL GEOMORPHOLOGy ....................................................................................... .4 1.3 THE STUDY AREA .............................................................................................................. 8 1.4 TOPOGRAPHy ..................................................................................................................... 9 1.5 SOIL AND VEGETATION ................................................................................................. 11 1.6 AGRICULTURE .................................................................................................................. 11 1.7 WATER DEMAND ............................................................................................................. 11 1.8 DRILLING ........................................................................................................................... 13 1.9 PREVIOUS WORK ............................................................................................................. 14 1.10 PURPOSE OF THE STUDy ............................................................................................... 17 1.11 SCOPE AND METHODOLOGY ........................................................................................ 18 1.12 STRUCTURE OF THESIS ................................................................................................... 21
CHAPTER TWO GEOLOGY 22 2.1 REGIONAL GEOLOGy ....................................................................................................... .22
2.1.1 OVERVIEW .......................................................................................................... 22 2.1.2 OUTLINE LITHOSTRA TIGRAPHy ................................................................... 24
2.2 GEOLOGY OF THE STUDY AREA ................................................................................... .26 2.3 STRA TIGRAPHy ................................................................................................................. .28
2.3.1 THE PRECAMBRIAN BASEMENT COMPLEX ............................................... 28 2.3.2 THE PALAEOZOIC SUCCESSION .................................................................... .29
2.3.2.1 THE DISI SANDSTONE GROUP ..................................................... .29 2.3.2.2 THE KHREIM SANDSTONE GROUP ............................................ .29
2.3.3 THE MESOZOIC SUCCESSION ........................................................................ .30 2.3.3.1 ZERQA GROUP ................................................................................. .30 2.3.3.2 THE KURNUB GROUP ..................................................................... .30 2.3.3.3 THE AJLUN GROUP ......................................................................... .31
NA'UR FORMATION (AII2) .......................................................... 38 FUHEIS FORMATION (A3) ........................................................... .40 HUMMARFORMATION (A4) ....................................................... .41 SHUE'IB FORMATION (A5/6) ....................................................... .41 WADI SIR FORMATION (A7) ....................................................... .42
2.3.3.4 THE BELQA GROUP ....................................................................... .44 AMMAN FORMATION (BII2) ....................................................... .47 MUWAQQAR FORMATION (B3) ................................................. .48
2.3.4 THE CAINOZOIC SUCCESSION ..................................................................... .49 RIJAM FORMATION (B4) ............................................................. .49
2.3.5 POST EOCENE SEDIMENTS: ........................................................................... 50 2.3.6 RECENT DEPOSITS ........................................................................................... 50
2.4 VOLCANICS .................................... · ................................................................................... 50 2.5 THE GEOLOGICAL STRUCTURE .................................................................................... .51
2.5.1 STRUCTURAL ELEMENTS .............................................................................. 53 2.5.2 MINOR STRUCTURE ......................................................................................... 57
CHAPTER THREE HYDROLOGY 58 3.1 CLIMATE ............................................................................................................................... 58 3.2 TEMPERATURE AND HUMIDITy ..................................................................................... 60 3.3 RAINFALL ............................................................................................................................. 62 3.4 EVAPORATION .................................................................................................................... 72 3.5 RUNOFF ................................................................................................................................ 78
INTERCEPTION ............................................................................................................ 80 EVAPOTRANSPIRATION ...................................... ........................................................ 80 BANK STORAGE ........................................................................................................... 81 SURFACE STORAGE AND DETENTION. .................................................................... 81 INFILTRA TION .............................................................................................................. 81
CURVE NUMBER APPROACH .................................................................... 84
APPLICATION OF THE CN METHOD ......................................................... 87
3.5.1 RUNOFF IN JORDAN ....................................................................................................... 88 3.5.2 RUNOFF IN THE STUDY AREA ..................................................................................... 92
3.5.2.1 UPPER ZERQA CATCHMENT ...................................................................... 92 3.5.2.2 WADI MUJIB CATCHMENT ......................................................................... 97 3.5.2.3 WADI HASA CATCHMENT ........................................................................ 102 3.5.2.4 JAFR CATCHMENT ..................................................................................... 104
3.6 CONCLUSION .................................................................................................................... 106 CHAPTER FOUR AQUIFER SYSTEM 109
4.1 AQUIFER SYSTEMS IN JORDAN .................................................................................. 1 09 4.2 AQUIFER SYSTEMS IN THE STUDY AREA ................................................................. .112
4.2.1 EXTENT AND LITHOLOGY ........................................................................... 114 4.2.1.1 THE NA'UR AQUIFER SYSTEM (A 112) ....................................... 114 4.2.1.2 THE HUMMAR AQUIFER SYSTEM (A4) ................................... .114 4.2.1.3 AMMAN - WADI SIR AQUIFER SYSTEM (B2/A7) ..................... 116 4.2.1.4 THE RIJAM AQUIFER SYSTEM (B4) ............................................ 120 4.2.1.5 LOWER AJLUN GROUP AQUIFER SYSTEM (A 1-6) ................. 122
CHAPTER FIVE AQUIFER PROPERTIES 125 5.1 INTRODUCTION ................................................................................................................. 125 5.2 ROCK FABRIC AND STRUCTURE ................................................................................... 125 5.3 PUMPING TESTS ............................................................................................................... 128
5.3.1 PUMPING TESTS IN THE STUDY AREA ..................................................... 128 5.3.2 RESULTS OF PUMPING TEST ANAL YSIS .................................................... 131
5.3.2.1 SPECIFIC CAPACITY ...................................................................... 131 5.3.2.2 TRANSMISSIVITY AND PERMEABILITY ................................... 133 5.3.2.3 VERTICAL HYDRAULIC CONDUCTIVITY ................................ 135 5.3.2.4 STORAGE COEFFICIENT.. ............................................................ 136
CALCULATING CONFINED STORAGE COEFFICIENTS ...................... 136 ESTIMATING STORAGE COEFFICIENTS FROM PUMPING
TESTS. .............................................................................................. 137 5.3.2.5 DISCUSSIONS ................................................................................... 141
5.4 ESTIMATION OF T FROM SC .......................................................................................... 143 5.4.1 APPLICATION OF THE METHOD ................................................................... 153 5.4.2 AREAL DISTRIBUTION OF PERMEABILITY ............................................... 158 5.4.3 DISCUSSION ...................................................................................................... 161
5.5 HUM MAR (A4) AQUIFER SySTEM ................................................................................. 163 5.6 RIJAM (B4) AQUIFER SySTEM ........................................................................................ 167 5.7 LOWER AJLUN GROUP (AI-6) AQUIFER SYSTEM ..................................................... 168
CHAPTER SIX RECHARGE 170 6.1 INTRODUCTION ................................................................................................................. 170 6.2 RECHARGE MECHANISMS .............................................................................................. 170 6.3 RECHARGE ESTIMATION ................................................................................................ 172 6.4 DIRECT RECHARGE .......................................................................................................... 172
6.4.1 SOIL-WATER BALANCE ................................................................................. 174 ACTUAL AND POTENTIAL EV APOTRANSPIRA TION .......................... 176 SOIL MOISTURE DEFICIT .......................................................................... 177 RECHARGE CALCULATION AND RESULTS .......................................... 180 DISCUSSION .............................................................. : .................................. 185
6.4.2 WATER BUDGET .............................................................................................. 187 6.5 INDIRECT RECHARGE ..................................................................................................... 190
6.5.1 RECHARGE THROUGH WADI BEDS ............................................................ 190 6.5.2 LATERAL BOUNDARY FLOW ....................................................................... 194 6.5.3 WATER TRANSFER ......................................................................................... 195
6.6 TOTAL RECHARGE .......................................................................................................... 196 6.6.1 GROUNDWATER BALANCE .......................................................................... 196 6.6.2 THE RESPONSE OF GROUNDWATER TO THE TOTAL RECHARGE ....... 198
6.6.2.1 WATER LEVEL FLUCTUA TIONS ................................................. 198 6.6.2.2 SPRING DISCHARGES .................................................................... 199
HYDROGRAPH ANAL YSIS ....................................................... .203 RECESSION HYDROGRAPH ANAL YSIS ................................. .205 HYDROCHEMICAL ANAL YSIS ................................................ .21 0 ENVIRONMENTAL ISOTOPES ANAL YSIS ............................. .212
CONCLUSION OF THE METHOD ............................................................ .212 APPLICATION OF THE METHOD ............................................... 215 SPRING CATCHMENTS AND RECHARGE COEFFICIENT ...... 216
6.7 RECHARGE MOUNDS ...................................................................................................... 222 6.8 RECHARGE TO RlJAM AQUIFER SySTEM ................................................................. .224 6.9 RECHARGE TO HUMMAR AQUIFER SYSTEM .......................................................... .226 6.10 RECHARGE TO LOWERAJLUN GROUP (AI-6) AQUIFER SYSTEM ........................ .226 6.11 CONCLUSION ................................... : ................................................................................. 228
CHAPTER SEVEN GROUNDWATER FLOW 231 7.1 GENERAL ............................................................................................................................... 231 7.2 FLOW MECHANISMS .......................................................................................................... .233
7.2.1 FLOW DISTRIBUTION ....................................................................................... .234 7.2.2 GROUNDWATER STRA TIFICA TION .............................................................. .235 7.2.3 CONCEPTUAL FLOW MODEL ........................................................................... 237 7.2.4 GEOLOGICAL STRUCTURES AND GROUNDWATER MOVEMENT .......... .239
7.3 REGIONAL GROUNDWATER FLOW ................................................................................ .241 7.3.1 UPPER ZERQA BASIN ........................................................................................ .241
HUMMAR AQUIFER SYSTEM (A4) ............................................................. .245 7.3.2 WADI WALA BASIN ........................................................................................... .248 7.3.3 WADI MUJIB BASIN ........................................................................................... .249 7.3.4 WADI HASA BASIN ............................................................................................ .253 7.3.5 JAFR BASIN .......................................................................................................... 254
7.3.5.1 INTRODUCTION ................................................................................ .254 7.3.5.2 AMMAN-WAD! SIR AQUIFER SYSTEM (B2/ A 7) ........................... 254
SHIDIY A AREA ................................................................................ 256 7.3.5.3 LOWER AJLUN GROUP AQUIFER SYSTEM (AI-6) ..................... .259 7.3.5.4 RlJAM AQUIFER SYSTEM (B4) ....................................................... .261
7.4 HYDRAULIC GRADIENTS .................................................................................................. 261 7.5 WATER LEVEL FLUCTUATIONS ...................................................................................... 267 7.6 AQUIFER INTERRELATION .............................................................................................. .271
7.6.1 LEAKAGE BETWEEN B2/A7 AND A4 AQUIFER SYSTEMS ......................... 271 7.6.2 LEAKAGE BETWEEN B2/A7 AND AI-6 .......................................................... .272 7.6.3 LEAKAGE BETWEEN B2/A7 AND K-D AQUIFER SYSTEMS ...................... .273
CHAPTER EIGHT GROUNDWATER MODELLING 274 8.1 INTRODUCTION .................................................................................................................. .274 8.2 MODEL DEVELOPMENT ................................................................................................... .274 8.3 GENERAL ASSUMPTIONS AND LIMIT A TIONS .............................................................. 279 8.4 APPROACH ............................................................................................................................ 281 8.5 MODEL GRID AND LA YERS .............................................................................................. 285 8.6 BOUNDARY CONDITIONS ................................................................................................. 297 8.7 INPUT DATA ......................................................................................................................... 299 8.8 MODEL SIMULATIONS ...................................................................................................... 304
8.8.1 STRATEGy .......................................................................................................... 304 8.8.2 STEADY STATE CALIBRA TION ..................................................................... .305
8.8.2.1 SIMULATION RESULTS .................................................................. .308 8.8.2.1.1 FLOW SUBREGIONS ...................................................... .309
AMMAN-ZERQA AREA .................................................. .309 WADI WALA BASIN ......................................................... 312 WAD! MUJIB AND WAD! HASA BASINS ..................... .313 JAFR BASINS .................................................................... .314
8.8.2.1.2 SIMULATED HYDRAULIC PROPER TIES .................................... 316 HYDRAULIC CONDUCTIVITy ..................................................... .316 LEAKANCE ................................................................................... .319 STREAMBED CONDUCT ANCE ................................................... 321 RECHARGE .................................................................................... 323
8.8.3 TRANSIENT CALIBRA TION .......................................................................... .324 8.8.3.1 STORAGE COEFFICIENT .............................................................. .332
8.8.4 REGIONAL GROUNDWATER BUDGET.. ...................................................... 334 8.8.5 SENSITIVITY ANALySIS ................................................................................ .338
8.8.5.1 HYDRAULIC CONDUCTIVITY ..................................................... .339 8.8.5.2 VERTICAL HYDRAULIC CONDUCTIVITY .................................. 34I 8.8.5.3 RECHARGE ........................................................................................ 343 8.8.5.4 STORAGE COEFFICIENT ................................................................ .344 8.8.5.5 SUMMARY AND DISCUSSION ...................................................... .345
8.9 MODEL RELIABILITy ....................................................................................................... .346 8.10 DISCUSSION ...................................................................................................................... .349
CHAPTER NINE SUMMARY AND DISCUSSION 353 CHAPTER TEN CONCLUSIONS AND RECOMMENDATIONS 365
10.1 CONCLUSIONS ................................................................................................................... 365 1 0.2 RECOMMENDATIONS ...................................................................................................... 361
REFERENCES 363 APPENDICES 373
Appendix (AI) Well list in the study area ...................................................................................... 374 Appendix (B1) Definition ofSCS Hydrologic Soil Groups (HSG) ................................................ 388 Appendix (B2) Runoff curve numbers ......................................................................................... .389
Appendix (B2.1) Runoff curve number for Urban Areas ........................................................ 390 Appendix (B2.2) Runoff curve number for cultivated Agricultural Lands .............................. 391 Appendix (B2.3) Runoff curve number for other Agricultural Lands ...................................... 392 Appendix (B2.4) Runoff curve numbers for Arid and Semiarid Rangelands .......................... .393
Appendix (B3) Surface Water in Jordan ......................................................................................... 394 Appendix (B4) Runoffmeasurments in the study area .................................................................. .395
Appendix (B4.1) Runoff measurments for Zerqa River at Sukhna Gauging Station in MCM ........................................................................................................................ 396 Appendix (B4.2) Runoffmeasurments for Wadi Wala at Karak Road in MCM ...................... .397 Appendix (B4.3) Runoffmeasurments for Wadi Wala at weir in MCM .................................. .398 Appendix (B4.4) Runoffmeasurments for Wadi Swaqa in MCM ............................................ .398 Appendix (B4.5) Runoffmeasurments for Wadi Mujib at Karak Road in MCM ...................... 399 Appendix (B4.6) mean annual observed flood flow of Hasa River at Tannur in MCM ............ .400 Appendix (B4.7) Observed runoff discharge of Has a River at Ghor Safi in MCM .................... 401 Appendix (B4.8) Mean annual observed flood flow of Wadi Jurdhan in MCM ........................ .402
Appendix (CI) Results of pump in test analysis in the B2/A7 aquifer system ................................ .403 Appendix (DI) Soil moisture balance (mm) for West Amman sub-catchment for the
water year 1982/1983 ................................................................................................................ .406
LIST OF TABLES
page Table (2.1) Geological succession in Jordan and occurrences in the study area ........................ 28 Table (2.2) Correlation oflitho-stratigraphic units of the Ajlun Group recognised by
various authors ......................................................................................................... .36 Table (2.3) Occurrences of Ajlun Group Formations ................................................................. .38 Table (2.4) Correlation of litho-stratigraphic units of the Belqa Group recognised by
various authors .......................................................................................................... 46 Table (2.5) Occurrences ofBelqa Group Formations ................................................................. .47 Table (3.1) Mean monthly temperature for selected stations for the period
1937-1985 in (oC) ....................................................................................................... 61 Table (3.2) Seasonal ranges and annual mean of relative humidity (%) at selected
stations ........................................................................................................................ 62 Table (3.3) Probability of various daily rainfall amounts in 100 years at selected
stations ........................................................................................................................ 69 Table (3.4) Relationship between Eo and PET ............................................................................. 75 Table (3.5) Mean monthly values for the Class-A-Pan (Eo) and potential
evapotranspiration (PET) for selected stations in (mm) ........................... ; ................ 77 Table (3.6) Classification of the hydrological soil groups (HSG) ............................................... 85 Table (3.7) Mean annual runoff coefficient (%) for the different groups of catchments ............ 92 Table (3.8) Spring discharge data for the main springs in the Upper Zerqa Basin ...................... 95 Table (3.9) Estimated flood flows (MCM/a) in the Upper Zerqa Basin obtained by using
the CN method .......................................................................................................... 96 Table (3.10) Spring discharge data for the main springs in the Wadi Mujib Basin .................... 99 Table (3.11) Estimated flood flows (MCM/a) in the Wadi Mujib Basin obtained by using
the CN method ........................................................................................................ 10 I Table (3.12) Spring discharge data for the main springs in the Wadi Hasa Basin ..................... l 04 Table (3.13) Estimated flood flows (MCM/a) in the Wadi Hasa Basin obtained by using
the CN method ........................................................................................................ 1 05 Table (3.14) Spring discharge data for the main springs in the J afr Basin ................................ l 06 Table (3.15) Estimated flood flows (MCM/a) in the Jafr Basin obtained by using
the CN method ...................................................................................................... 107 Table (4.1) The hydrogeological units of the study area .......................................................... .112 Table (5.1) The average limestone-marl ratio for the different formations ............................... 127 Table (5.2 ) Summary of pumping tests results in the study area .............................................. 131 Table (5.3) Frequency distribution of specific capacity from pump tests (%) .......................... 133 Table (5.4) Frequency distribution of transmissivity from pump tests (%) ............................... 134 Table (5.5) Frequency distribution of permeability from pumping tests (%) ............................ 135 Table (5.6) Storage coefficient and specific yield from pumping tests .................................... .138 Table (5.7) Storage coefficient and specific yield calculated by Ramsahoye and Lang
method .................................................................................................................... 140 Table (5.8) Statistical results of estimates of transmissivity and hydraulic conductivity from
pump tests and specific capacity for the data used in the analysis .......................... I 55 Table (5.9) Statistical results of calculated transmissivity and hydraulic conductivity
from specific capacity .............................................................................................. 155 Table (5.10) Results of pump test analysis of Hummar Aquifer System in Amman -Zerqa
area ......................................................................................................................... 164 Table (5.11) Results of pumping tests in the Rijam Aquifer System in Jafr Basin ................... 168 Table (6.1) Calculation of Soil Moisture Deficit (mm) at selected stations .............................. 179
Table (6.2 ) C and D values (mm) for various soil types in Jordan .......................................... .181 Table (6.3 ) Field capacities (%) values by weight for various soil types in Jordan ................ .181 Table (6.4) Results of direct recharge calculation to the B2/A7 aquifer system (in MCM)
by using soil moisture balance method ................................................................... 184 Table (6.5) Results of direct recharge calculation (in MCM) for the B2/A7 aquifer system
by using the water budget method .......................................................................... .188 Table (6.6) Results of indirect recharge calculation (in MCM) ................................................ 194 Table (6.7) Recharge estimation from lateral boundary flow .................................................... 195 Table (6.8) Groundwater balance of the Amman/Wadi Sir aquifer system ............................... 198 Table (6.9) The hydrological parameters of the recession curve model for Ras el Ain spring .. 209 Table (6.10) Results of recession hydrograph analysis for some springs .................................. 216 Table (7.1) Long term groundwater level fluctuations of the Rijam (B4) aquifer system ......... 261 Table (7.2) Estimated groundwater velocities and transient times along flow lines from the
recharge mounds to discharge areas .......................................................................... 266 Table (8.1) Observed and simulated water levels for selected observation wells .................... .309 Table (8.2) Abstraction used in simulations, by aquifer, area, and time period ....................... .325 Table (8.3) Stress periods and time steps used in the simulations (days) ................................ .327 Table (8.4) Simulated steady-state groundwater flow budget... ............................................... .337 Table (8.5) Maximum drawdown (m) for selected observation wells due to different
storage coefficients at the end of transient simulations .......................................... .345
LIST OF FIGURES
page Figure (1.1) Location map of Jordan ............................................................................................ .3 Figure (1.2) Location map of the study area .................................................................................. 5 Figure (1.3) The physiographic-geologic provinces in Jordan ....................................................... 7 Figure (1.4) The soil zones in Jordan ........................................................................................... 12 Figure (2.1) Palaeogeography of Jordan ..................................................................................... .23 Figure (2.2) General geological map of Jordan ............................................................................ 25 Figure (2.3) Geological map of the study area ............................................................................. 27 Figure (2.4) Change in lithology of the Ajlun Group in Southern Jordan ................................... .32 Figure (2.5) Isopachyte map of the Ajlun Group ........................................................................ .33 Figure (2.6) Generalised cross-section through the Mesozoic marine succession from
the northwest to the southeast. ................................................................................. .34 Figure (2.7) Percentage of sand in the Ajlun Group ................................................................... .37 Figure (2.8) Geological cross-sections in the study area ............................................................. .39 Figure (2.9) Type sections of the Amman-Wadi Sir Formation .................................................. .43 Figure (2.10) Isopachyte map of the Belqa Group ...................................................................... .45 Figure (2.11) The structural pattern of Jordan ............................................................................. 54 Figure (3.1) Distribution of Mediterranean bioclimatic stages in Jordan .................................... 59 Figure (3.2) Mean annual rainfall (mm/a) for the period 1938-1985 .......................................... 63 Figure (3.3) Annual rainfall variation around the mean at selected stations ............................... 67 Figure (3.4) Accumulated departure from the mean annual rainfall for selected stations .......... 68 Figure (3.5) Coefficient of variation of annual rainfall as a function of mean annual
rainfall for selected stations in the Near East and North Africa ............................... 69 Figure (3.6) Monthly rainfall at selected stations ........................................................................ 70 Figure (3.7) Morton's (1985) model of evaporation .................................................................... 73 Figure (3.8) Monthly potential evapotranspiration (PET) vs. Monthly Class-A Pan
evaporation(Eo) ........................................................................................................ 76 Figure (3.9) Monthly potential evapotranspiration (PET) at selected stations ............................ 78 Figure (3.10) Mean annual potential evapotranspiration (PET) .................................................. 79 Figure (3.11) Graphical solution of SCS curve Number Method ................................................ 86 Figure (3.12) Surface water catchments in Jordan ....................................................................... 89 Figure (3.13) Upper Zerqa Catchment. ........................................................................................ 93 Figure (3.14) Wadi Mujib Catchment... ....................................................................................... 98 Figure (3.15) Wadi Hasa and Jafr Catchments ........................................................................... 103 Figure (4.1) Generalised hydrogeological section in the study area .......................................... 113 Figure (4.2) Hydrogeological setting of the A4 aquifer system in the Amman-Zerqa area ...... .l15 Figure (4.3) Hydrogeological setting of the B2/A7 aquifer system .......................................... 118 Figure (4.4) Hydrogeological setting of the B4 aquifer system in the Jafr Basin ...................... 121 Figure (4.5) Hydrogeological setting of the AI-6 aquifer system in the Hasa and Jafr basins . .l23 Figure (5.1) Examples of pumping test data analysis ................................................................. 132 Figure (5.2) Relation between transmissivity and specific capacity .......................................... 146 Figure (5.3) Effect of varying storage coefficient on theoretical relations between specific
capacity and transmissivity .................................................................................... 147 Figure (5.4) Effect of aquifer anisotropy on theoretical relations between transmissivity
and well specific capacity ...................................................................................... 149 Figure (5.5) Effect of partial penetration on theoretical relations between aquifer
transmissivity and well specific capacity ............................................................... 150 Figure (5.6) Effect of vertical fracture on theoretical relations between aquifer
transmissivity and well specific capacity ............................................................... 152 Figure (5.7) Log~log relation between transmissivity and specific capacity ............................. 154 Figure (5.8) Frequency distribution of the calculated transmissivity ......................................... 156
Figure (5.9) Frequency distribution of the calculated permeability .......................................... .156 Figure (5.10) Ranges of hydraulic conductivity and permeability for various geological
materials, showing ranges determined from specific capacity estimates for the Amman-Wadi Sir aquifer system ................................................................... 157
Figure (5.11) The relationship between permeability index and permeability for the B2/A7 aquifer system ........................................................................................... 159
Figure (5.12) Areal distribution of permeability in the B2/A7 aquifer system .......................... 160 Figure (5.13) The relationship between permeability index and permeability for the
Hummar aquifer system ....................................................................................... 165 Figure (5.14) Areal distribution of permeability in the A4 aquifer system ............................... 166 Figure (6.1) The outcrop area of the B2/A 7 aquifer system ...................................................... 173 Figure (6.2) Schematic diagram showing the conceptual model of the soil-water balance
method .................................................................................................................... 175 Figure (6.3) Field capacity determination for basalt soil (after Lloyd et at., 1966) .................. .182 Figure (6.4) Soil moisture content (SMC) variation with time at selected sites for depth
range between 0-50 cm ........................................................................................... 183 Figure (6.5) Relation between rainfall and recharge coefficient (%) ........................................ .189 Figure (6.6) Relation between recharge and recharge coefficient (%) ....................................... 189 Figure (6.7) Groundwater level fluctuations due to rainfall in the year 198511986 ................... 200 Figure (6.8) Location map of the Ras el Ain spring showing the Wadi Abdoun
sub-catchment. ........................................................................................................ 202 Figure (6.9) Geological cross section ofRas el Ain spring ........................................................ 204 Figure (6.10) Average monthly values ofRas el Ain spring discharges and rainfall in
Wadi Abdoun Basin ............................................................................................... 204 Figure (6.11) Peak flow versus accumulated rainfall ................................................................. 205 Figure (6.12) Accumulated infiltration calculated after subtracting the threshold value .......... .206 Figure (6.13) Analysis of recession curve ofRas el Ain spring ................................................. 207 Figure (6.14) Graph of qt versus time for Ras el Ain spring ...................................................... 208 Figure (6.15) Chemical analysis ofRas el Ain spring ............................................................... .211 Figure (6.16) Environmental isotopes analysis ofRas el Ain spring ........................................ .213 Figure (6.17) Recharge versus spring catchment areas .............................................................. 219 Figure (6.18) Distribution of recharge to the B2/A7 aquifer system ......................................... 223 Figure (6.19) Outcrop of Hummar Formation showing local drainage pattern ......................... 227 Figure (6.20) Geological cross-section in Hummar Formation to the NW of Zerqa ................. 227 Figure (7.1) Generalised hydrogeological profile of the regional aquifer systems .................. 232 Figure (7.2) Conceptual model of groundwater flow in the regional aquifer systems .............. 238 Figure (7.3) The potentiometric surface map of the B2/A7 aquifer system in the
study area ............................................................................................................... 242 Figure (7.4) The potentiometric surface map of the Amman-Wadi Sir aquifer system in
Amman-Zerqa area ................................................................................................. 244 Figure (7.5) Hydrogeological profile along the Zerqa River .................................................... .246 Figure (7.6) The potentiometric surface map of the Hummar aquifer system in
Amman-Zerqa area ................................................................................................ 247 Figure (7.7) Hydrogeological profile in Wadi Mujib Basin ...................................................... 252 Figure (7.8) Hydrogeological profile in Jafr Basin .................................................................. .257 Figure (7.9) The potentiometric surface map of the AI-6 aquifer system in Jafr Basin .......... .260 Figure (7.10) The potentiometric surface map of the Rijam aquifer system in Jafr Basin ....... .262 Figure (7.11) Estimated predevelopment hydraulic gradients along selected flow lines
from the recharge mounds to the discharge areas ................................................ 264 Figure (7.12) Location map of the observation wells in the study area ................................... .268 Figure (7.13) Observation well hydrographs in the study area ................................................. 269 Figure (8.1) Conceptualisation of the regional groundwater flow model... ............................. .282 Figure (8.2) Regional and subregional model areas ................................................................. 286
Figure (8.3) Conceptualisation of the subregional groundwater flow models .......................... 287 Figure (8.4) Finite-difference grid for the Amman-Zerqa subregional model.. ........................ 290 Figure (8.5) Finite-difference grid for the Wadi Wala subregional model.. ............................. 291 Figure (8.6) Finite-difference grid for the Amman-Zerqa and Wadi Wala subregional
model ..................................................................................................................... 292 Figure (8.7) Finite-difference grid for the Upper Wadi Mujib and Wadi Hasa subregional
model ..................................................................................................................... 293 Figure (8.8) Finite-difference grid for the Wadi Mujib and Wadi Hasa subregional model.. .. .294 Figure (8.9) Finite-difference grid for the Jafr subregional model.. .......................................... 295 Figure (8.10) Finite-difference grid for the regional model.. .................................................... 296 Figure (8.11) Diagrams for calculation ofverticalleakance ..................................................... 302 Figure (8.12) Simulated steady state water levels for the B2/ A 7 aquifer system .................... .31 0 Figure (8.13) Areal distribution of calibrated hydraulic conductivity of the B2/ A 7 aquifer
system ................................................................................................................. .317 Figure (8.14) The major areas of groundwater abstractions .................................................... .326 Figure (8.15) Observed and simulated drawdown in observation wells .................................. .330 Figure (8.16) Simulated predevelopment water budget for the regional aquifer systems
(MCM/a) .............................................................................................................. .335 Figure (8.17) Sensitivity of the B2/ A 7 aquifer system to changes in hydraulic
conductivity ....................................................................................................... .340 Figure (8.18) Sensitivity of the B2/ A 7 aquifer system to changes in vertical hydraulic
conductivity of the A 1-6 aquitard ....................................................................... .342 Figure (8.19) Sensitivity of the B2/A7 aquifer system to changes in recharge ........................ 343
CHAPTER ONE
INTRODUCTION
This thesis summanses a study that was conducted to describe the
hydrogeological framework and associated groundwater flow system of the carbonate
aquifer systems in the Western Highlands and Central Plateau of Jordan. The area studied
is geologicaly complex. Rocks range in age from Precambrian to Recent, and the history
of the area includes many episodes of sedimentation, volcanic activity, and tectonic
defonnation.
The source of groundwater is precipitation that falls primarily on the higher
mountain ranges, where part of the precipitation is estimated to recharge groundwater.
Most of the groundwater discharges to springs in the low parts of the many valleys and as
subsurface outflow into the Dead Sea.
The development of energy-related resources, power generation, industrial
development, increasing in·igation, and increased water demand for domestic and
municipal use in the country are dependent on the availability of water resources. Owing
to the climatic conditions in Jordan, surface water not only has been appropriated fully in
much of the country but also is limited, unpredictable, and not a dependable source of
water. Therefore, emphasis is placed upon groundwater as the main source of water.
However, locally some wadis contain low baseflows receiving water from spring
discharges in the highlands.
Long-term, large scale water needs will reqUire development of groundwater
resources. Without knowledge of the hydrogeologic characteristics of the groundwater
system and its response to abstractions, large sustained yields of groundwater cannot
produced efficiently, and sound management plans cannot be fonnulated. Proper
development, use, and conservation of groundwater can be achieved only through an
understanding of the regional geologic framework and its effect on the response of the
hydrologic system to climate and to water supply development.
The study area was divided into five subdivisions, which allowed evaluation of
the different aspects of the framework hydrogeology of the aquifer system within each
subregion. The subregional studies address local subdivisions of the aquifer system in
greater detail than the regional study. Data collected and compiled in these subregional
studies were then integrated to define the regional framework of the geology, hydrology,
and groundwater flow regime for the Western Highlands and Central Plateau of Jordan
study area.
A groundwater flow model was designed and calibrated to improve understanding
of the groundwater flow regime. However, unlike many studies, which emphasised the
predictive capabilities of groundwater flow models, the groundwater flow model used in
this study primarily for analysis of the regional groundwater flow system, to evaluate the
effect of regional geological structure on water levels and inferred groundwater flow, and
to provide general estimation of the direction and magnitude of groundwater flow within
the study area.
1.1 BACKGROUND
The Hashemite Kingdom of Jordan is a country of about 96500 km2 in the
northern part of the Arabian Peninsula. It lies between latitudes 29.5° N and 33° N and
longitudes 35° E and 39.5° E. The part ofthe Kingdom which lies to the east of the Jordan
Rift Valley is bordered to the north by Syria, to the north-east by Iraq, and to the east and
south-east by Saudi Arabia (Figure 1.1). According to the 1988 statistics, the population
of Jordan is about 4 millions, with an average annual growth rate of about 3.7%.
About 80% of the total area of Jordan is classified to have semi-arid to arid
climate with less than 200 mmla of rainfall and high potential evaporation rate exceeding
2000 mmla. In the high rainfall area - in the Western Highlands - the climate is
Mediterranean type with rainfall reaching 650 mmla in some places. Moving eastward,
the climate rapidly changes to semi-arid and arid, with lower rainfall and higher
temperature in the south and east. Rain in Jordan mostly falls during winter (October-
2
Figure (1.1) Location map of Jordan
May) on the Western Highlands of limestones, marls, sandstones, and to a lesser extent
over areas covered by igneous rocks in the south and south-east of Jordan.
The basement complex is unconformably overlain by variable thicknesses of
sandstones and shale of Cambrian, Ordovician, and Silurian ages, of continental and
marine origin. The rock units, gently dipping towards the north and north-east become
overlain by a succession of younger marine sediments which are mostly made of
carbonate of Upper Cretaceous to Eocene in age.
The National Water Master Plan of Jordan (1977) in a resource appraisal defines
twelve surface water basins in Jordan (Figure 1.2), and divides the groundwater systems
in the country into three major aquifer systems or complexes: the deep sandstone aquifer
complex, the Upper Cretaceous carbonate aquifer, and the shallow aquifer complex.
The Upper Cretaceous carbonate aquifer system, the subject of this study, forms
the major regional aquifer system of Jordan. It is essentially continuous and contains very
productive aquifers throughout the country. Four aquifer system have been recognised,
the first which has regional importance is the Amman-Wadi Sir (B2/A7) aquifer system
which extends throughout much of the entire country and varies considerably in
lithology, depth of occurrence, hydraulic properties and resource development. The other
three aquifers are the Na'ur (AI/2), the Hummar (A4) and the Rijam (B4) aquifer system
are of importance locally in limited areas.
The Upper Cretaceous carbonate aquifer system is the main source of water for
municipal and industrial water supplies. It is used extensively for self-supplied industrial,
rural and domestic water supplies. Development of the system began as early as 1960. By
mid 1970's and early 1980's many wells had been drilled especially where flowing wells
or shallow water were obtainable. Abstraction of groundwater from the aquifer system
has increased almost steadily in relation to the growth in population, attaining rate of
about 93 MCMla in 1985.
1.2 GENERAL GEOMORPHOLOGY
Jordan may be divided into seven distinctive physiographic proVInces which
coincide with the geological provinces (Figure 1.3):
4
100
000
900
• • • • • • •
. .
,,' • .i :. I
• I .
• •
." .. ' f
;~
.' .'
200
" , -->'ann "I.
, ... °I.!V - ... -..", -c.. ~ . '" - '1/", ••••• "., ,- .. I!'I'
'. \-: \ '.-r ~ __ --. "
Ya"t0uk Basin',
Azraq Basin
300
Figure (1.2) Location map of the study area
'\ \.
'\ '\
\. '\.
Basin
'\. \.
\. \.
r I
I I
I
\. \.
\. '\
,> ."
.,,"
I -----
.S The study area
400
---
1. Highland West of the Rift: this area includes the hilly regions of structural
upwards of folded and faulted, mainly Upper Cretaceous-Lower Tertiary rock sequence
and drainage systems eastwards to the Jordan Rift Valley and westwards to
Mediterranean Sea.
2. Wadi Araba-Jordan Rift (the Rift Province) is a narrow depression that extends
from the Gulf of Aqaba for approximately 360 km north to Lake Tiberias. It represents
but a small fraction of the East African-Asian Minor Rift System. The floor of the rift
rises gradually from the Gulf of Aqaba to altitudes of 250 m above sea level (mas I) at the
watershed in the Central Wadi Araba, then the floor falls gently northward to the surface
of the Dead Sea, 392 m below sea level (mbsl). To the north of the Dead Sea, the Jordan
River Valley rises to 212 mbsl at Lake Tiberias.
3. Highland East of the Rift: this stretches north-northeast to north for about 370
km from the Gulf of Aqaba to Lake Tiberias. In general it slopes gently toward the
Central Plateau in the east, whereas it slopes very steeply toward the Rift Province in the
west. The highest altitudes in the country (about 1850 masl) are in the southern part of the
Mountain Ridge Province in the Jebal ash Sharah.
4. Southern Mountain Desert: this occupies the area south of the Ras en Naqb
Escarpment, and extends southward into Saudi Arabia. The Precambrian Basement
Complex in the west between Aqaba and Quweira, and the overlying Palaeozoic and
Mesozoic sandstone to the east and south-east, form rough and steep mountains rising up
to 1754 mas!. The sandstone also forms steep and bizarre cliffs and mountains. Farther to
the east, an inselberg landscape of table mountains and large depressions are the
dominant topographic features.
5. The Central Plateau includes Al Jafr and the Al Azraq-Wadi Sirhan basins to
the east of the Mountain Ridge Province In the north and in the east the Central Plateau
falls to the flat, wide southeast-striking Azraq-Wadi Sirhan Basin. This basin consists of
low undulating hills with eroded valleys which are partly covered or surrounded by the
debris of arid weathering. The outliers reach altitude of about 1000 mas!. To the south,
the altitude of the Central Plateau gradually drops to about 850 masl and is lowest in the
6
31"
29"
\ \
~\ 02' :z: , Z (oJ A.
( .... --' , ,
I , , l I
I -'" 1
III S ... o
37"
.... .. /' \ .. / , .. ,
.... / \ .. / \
,,/ \ ,'''''' ,
.. / '. '\ .. \ .. / ~ , /" , Northeastern Plateau ,
,. .. I \ ~ '\ /" !.,
~ --_""'-___ ,1 I "
,-) \ ...... \~ (,_ -r-', Northern Plateau Basalt / .. ,..,.... "/
I ,;" .... , ,'\..... ' """,,-,
/ \ " (" ..... _, : ,.,.",-- ef \96! .. ,............, ......... ~. --_.&......... ':: .".,/-".\~ft:,...;-
,/ \..... 1-~ ) .... ..-;:..:~:,. .. ~~ .. : .... "s> ....... .J.::::... .. .. I ...' .-:::~~~ 1 ..... ~r- '. f .. ~~', .... I '. '\ I . \" I .•.• "'. " .... I ... ~ '\
I ... \~ ',"" : .... 's,'''''. I ....,~ '\ I ".. ,\," •••.
J Central Plateau .... '. .... '\~" , '. \~ .. (Inc:lud .. AI Jat, and AI Az,aq· ". \~.; '~""
AI Jafr Basin
" WJdf .s Sir';:lln a.sins) ...... \\ '<,. ......... ' ...... ". \ ,~
:.' .......................... . .... \ ........ """ "\ /".'\ ..•...
~':. . \ . I ... \ .... I ". ..
/ '. \ '. \
......... ................ ..f~;::=~\ .... \ .. .. - .... -----, \ . "" •••.•• ,... ,_~ ___ ' \ I
, I?, ,I : \------1 , __ $ an Naqb , \ rt }
--------... / ",oese J I Mountainous: , ) _-----
... -.......... .... ---- .... ..-.---~ __ ~ /1 t:..-----.. -- ' -'--"", --~t'" -......... _?~ ,_
"'-~.!'"_I!!!. ---:" ---..' 0 .
~I-'..,..., I., --'" , ...... ~r'-...,.._..,... __ IOO ... , :..... ___ ....:I~fO::....:.;·'LO"rnoES o 50 lJo MILES
After Bender (1974)
Figure (1.3) The physiographic - geologic provinces in Jordan
wide Al Jafr Basin. The central part of this basin is an extensive mudflat. Farther south,
the altitudes of the Upper Cretaceous and Lower Tertiary sequence of the Central Plateau
is more than 1500 masl at the escarpment ofRas Naqb, the boundary between the Central
Plateau and the Southern Mountain Desert Province.
6. The Northern Plateau Basalt Province lies to the east of the northern part of the
Al Azraq-Wadi Sirhan Basin. The Plateau Basalt forms a shield of almost inaccessible
flows, fissure effusions and isolated volcanoes, gradually dropping from an altitude of
about 1,100 mas I at the Syrian-Jordanian border on the north to approximately 550 masl
in the south, close to the Wadi Sirhan Basin.
7. The Northeastern Plateau to the east of the Plateau Basalt extends as a
monotonous quasi-peniplained landscape eastward across the Iraqi border, north into
Syria and south into Saudi Arabia.
1.3 THE STUDY AREA
The study area (Figure 1.2) includes the Western Highlands to the east of the
escarpment to Jordan Valley from Zerqa River in the north to the southern limits of the
Jafr Catchment in the south. This includes the southern desert of the Central Plateau and
covers a total of approximately 23,350 km2• The study area stretches over the following
surface water basins (Figure 1.2):
1. Upper Amman-Zerqa Basin
This basin covers an area of approximately 850 km2 and incorporates the upper
most part of the Zerqa River drainage system and two major groundwater aquifers.
2. Wadi Mujih Basin
The Wadi Mujib Basin extends to the south of the Upper Amman-Zerqa Basin.
The Mujib River consists of both Wadi Wala and Wadi Mujib which have a combined
catchment area of about 6800 km2
• The basin is fairly rich in terms of groundwater
resources, together with a flood flow water source which appears only in short periods in
the rainy seasons.
8
3. Wadi Hasa Basin
The Wadi Hasa Basin has a catchment area of about 2200 km2 and incorporates
the drainage system of the Wadi Hasa and its tributaries. The basin is fairly poor in
groundwater resources, however, numerous springs issue in the western part ofthe basin.
4. JaJr Basin
The Jafr Basin is located in the southern part of the Central Jordan Plain and lies
to the east of the Western Highlands. The basin has an area of 13500 km2, most of which
classified as arid desert with mean annual rainfall of about 50 mm. Surface water is
limited to the few spring discharges along the western part of the basin.
1.4 TOPOGRAPHY
The topography and landforms that make up each of the previous physiographic
provinces or surface catchment areas have an effect on the areal distribution of recharge
and discharge and the occurrence and movement of groundwater. They reflect rock type,
geologic structure, degree of weathering and other features, knowledge which aids
understanding ofthe groundwater system.
The fault escarpment on the eastern side of the Jordan Valley Graben forms the
natural western boundary of the study area. Maximum elevations along the crest of the
escarpment are 1700 mas I to the west ofMa'an (Figure 1.2), 1250 mas I near Mazar, 1100
mas I east of Karak, 1000 mas I near Amman, and 800 masl east of Dhiban. The
escarpment is breached by a number of westward draining valleys within the study area,
the largest of these are; Wadi Zerqa, Wadi Mujib, and Wadi Hasa. Erosion in these
valleys has been rejuvenated by successive lowering of base level in the Jordan Valley
and the Dead Sea and as a result, deep gorges have been cut in the Mesozoic sediments
which underlie the Western Highlands. The headwaters of these drainages extend far into
the Central Plateau (Figure 1.2).
East of the escarpment there is gradual decline in elevation, and a gently
dissected plateau formed from flat lying sediments which have been eroded to form a
cuesta landscape. The elevation range from 1000 masl in the foothills of the Western
Highlands to less than 600 masl in the Wadi Sirhan depression near the Saudi Arabia
9
border. Topographic depressions which form the foci of internal drainage basins on the
plateau culminate near Jafr at less than 850 masI.
In Amman-Zerqa area the topography IS dominated by the Amman-Zerqa
Syncline structure, which forms along depression starting in Wadi Abdoun west of
Amman and runs towards the northeast. The elevation of the ground level falls from
about 800 masl to 550 masl along the syncline. The principal wadi in the area is Amman
Zerqa wadi which becomes the River Zerqa at the northwestern part where it leaves the
area at altitude of about 450 masI. To the west and northwest of the syncline structure a
mountainous area with marked topographic relief reaches more than 1000 mas I with a
maximum of 1086 masl; east of this, there is a gradual decline in elevation where hilly
country extends eastwards to the desert borders.
To the south of Amman-Zerqa Basin is the Wadi Mujib Basin which is mainly a
plateau land to the east of the Dead Sea and defined by the surface water catchment of the
Wadi Mujib and its principal tributary Wadi Wala. The majority of the catchment is at an
elevation of between 700 and 900 masl to the east of the hills which mark the edge of the
Jordan Valley Escarpment. The wadis Mujib and Wala have each cut gorges through the
hills to where they join some 3 km upstream of the Dead Sea. Both wadis in their lower
reaches have cut down to the saturated sections of water bearing formations so that
perennial flow is maintained by spring discharges. In the southeastern part of the Mujib
Basin, there lies a flat muddy swamp of Qa el Hafira with an area of 30 km2•
The Hasa Basin lies at elevations between 400 mas I at the basin outlet near
Tannour and 1250 masl in the Eastern Highlands. The wadis in the South-western
Highlands are characteristically narrow and moderately incised, while wadis are flat in
the eastern part of the basin where the elevation is about 900 masI. All the wadis in the
upstream reaches drain flushing floods to the central playa named Qa EI Jinz.
The Jafr Basin displays a classic centripetal drainage pattern with all wadis
draining from encircling highlands to the central EI J afr Playa, the largest concave in
Jordan. The catchments lies at an elevation of between 850 masl in EI Jafr Playa and 1750
masl in the Western Highlands.
10
1.5 SOIL AND VEGETATION
Soils and vegetation cover are indicators of the quantity of precipitation,
temperature and altitude. They change from grey lowland desertic soils with perennial
shrubs developed in areas with less than 150 mm mean annual rainfall, to brown soils
with a fairly complete cover of perennial shrubs and grasses in areas having mean annual
rainfall of 150-300 mm (Figure 1.4). Further to the west and along the Western
Highlands, as the altitude and precipitation increases and temperature decreases, red and
yellow Mediterranean soils with mountain forest are developed in areas where the mean
annual rainfall exceeds 300 mm. Other smaller biotic communities grow where
hydrologic conditions are favourable. The most prevalent is the dense growth of
phreaphytes commonly found along perennial and intermittent stream courses.
In some areas azonal soils are developed such as the weathered basalt in the
northeast, the saline soils in the topographic depressions (in Azraq, Hasa and Jafr),
alluvial soils and regosols formed from recently deposited detrital materials, and
litho sols-thinly covered consolidated rocks- such as basalt flows.
1.6 AGRICULTURE
As the country has a semi-arid to arid climate, different types of agriculture and
land use are found. In the Western Highlands, extensive agricultural development is being
carried out to take advantage of the higher rainfall of more than 200-600 mmJa and the
cool winter. Along the perennial streams and wadis, scattered farm projects are to be
found depending on surface water for irrigation; in some areas surface water supply is
supplemented by tube wells.
Several scattered irrigated farm projects are found to the east of the Western
Highlands from the Qastal area to the southeast of Amman, to Ma'an area in the south. In
the Jafr Basin, some oasis agriculture is being practised near Jafr town.
1.7 WATERDEMAND
Urbanisation and changes in regional development in Jordan are increasing the
need for water. However, many ofthe water resources are probably overdeveloped, such
11
200 300
: : : ::: :: ! : :::: . :: ; ~ : ..,. i : : : : : :..:. : ~ . -- ! .
I ~ 100:-' --hri:~~~~~7.~,.c777"7-T-T-.T-;A-------c::o - 100
~ c::: ~
o ,
I 30· ...:
I I I
i ,
e5c~--~--~~~~---~---~~----~--~-~B50
200 300
, I , I : Brown Soils (Yellow Soils ot Jordan)
Yellow Med iterraneon Soils I Grey Oese r' Soils
• Red Medi terraneon So ils :: : : Basalt Fields
After Bender (1974)
3~0
Figure (1.4) The soil zones in Jordan
that future growth may be seriously restricted unless a supplemental supply is made
available.
Of particular concern is the shortage of fresh water for domestic use, and this is
becoming increasingly a very serious problem. Approximately 96% of the population is
now supplied with drinking water from springs and groundwater, but according to the
Jordanian Ministry of Water and Irrigation's estimates (1988) water consumption is
expected to increase fairly rapidly in the future. Jordan is expected to require nearly 266
million cubic metres (MCM) of water for annual consumption by the year 2005. This
indicates that a shortage of75 MCM/a will exist, and that intensive efforts will have to be
made to find new water resources to meet the growing demand on water for different
purposes.
In the last decade, a series of water resources development projects have been
carried out all over the country.
1.8 DRILLING
Groundwater supplies most of the water demand for the entire country; therefore,
drilling started as early as 1916 when the first well was drilled in Amman-Zerqa area. The
Water Authority of Jordan (W AJ) inventory shows that in 1972 there were 737 producing
wells in Jordan. At present, the number of boreholes is unavailable, but it is expected to
exceed double this figure.
In the study area, significant withdrawal of water from wells began in the early
1960s in Amman-Zerqa area, where the alluvial aquifer systems were probably the first to
be developed for water supply because of the shallow depth to the water table and the
ease of drilling. In late 1964, groundwater development for irrigation purposes
commenced in the Jafr Basin; here the shallow Rijam aquifer system was used. The
development of water supplies in the Wadi Mujib Basin from the Amman/Wadi Sir
aquifer system did not begin until the early 1970s. Since the mid 1970s onward, many
wells have been drilled in the Wadi Mujib Basin by the private sector for irrigation
farming of vegetables.
13
Groundwater development and drilling increased steadily to a maximum in the
1980s. A huge number of wells have been drilled in the study area. Most of the well
names and numbers used by previous authors were adopted during this study. Only in the
Amman-Zerqa area, the well names and numbering systems used by previous authors are
found to be confusing, so for the purpose of this study, they have been given new
numbers which begin with the prefixed "A" for Amman followed by a serial number.
Similarly the private wells in Wadi Mujib Basin were given the prefix "PV" for private
followed by a serial number. Apart from the wells tapping the Rijam aquifer system in the
Jafr area and the Hummar aquifer system in the Amman-Zerqa area, most of the wells
through out the study area, are penetrating the regional AmmanIW adi Sir aquifer system.
However, number of wells were drilled into the deep sandstone aquifer system in the
Baqa'a area to the northwest of Amman.
The number of wells which have been found to tap the main aquifer systems
through the study area is about 701. These wells are listed in Appendix (AI), the list
including the basic information about the wells such as co-ordinate (Palestine grid),
aquifer system, ground surface elevation, total depth, water depth, groundwater level,
yield of pumping test, drawdown, specific capacity, and the permeability as calculated
from the specific capacity data during the course of this study. The well numbers used in
calculations and referred to in the text appear in the piezometric surface map (Figure 7.3).
1.9 PREVIOUS WORK
Groundwater investigations in Jordan began in the early part of the 20th century.
In the period since the Second World War there have been several studies on a regional
basis which to some degree have involved the evaluation of water resources. For the most
part, the scope of these studies is local or subregional. Geological and hydrogeological
data and interpretations from many of these reports have been used in this study.
The first assessment of the water resources of Jordan was made by Ionides and
Blake (1939). The water supply of Jordan was then described by Shaw (1947). The
geology, structural geology, stratigraphy and natural history of Jordan including some
14
groundwater data were described and summarised by Quennell (1951, 1956), Burdon and
Quennell (1959), Wentzel and Morton (1959), and Bender (1968).
Masri (1963) mapped and described the geology of the Amman -Zerqa area. The
water resources of most of the Western Highlands and part of the adj acent Plateau were
investigated in the period 1962-1965 by Sir M. MacDonald and Partners in association
with Hunting Technical Services Ltd. and the Jordan Office for Geological and
Engineering Services. These investigation included geological mapping, hydrogeological
studies, estimates of recharge and compilation of an inventory of springs.
The earliest and most important and comprehensive hydrogeological study of
Jordan was made by Parker (1970), who described the physiography, geology, structural
geology, stratigraphy, aquifer systems, extensions and characteristics, occurrence and
movement of groundwater and the chemical quality, environmental isotope analysis, use,
and availability of water in the study area.
Barber (1975) carried out an appraisal of the water resources and domestic
demands of east Jordan. His resources estimates were based on previous work and upon
data in the files of the Natural Resources Authority of Jordan (NRA). Agrar-und
Hydrotechnick (1977) compiled all the available data about geology, hydrology, and
groundwater for all the country for the construction of the National Water Master Plan of
Jordan (WMP). The work undertaken for the WMP provides the most comprehensive
regional surface and groundwater study ever carried out in Jordan. Because of the recent
increase in the use of groundwater as a source for irrigation and public supply, more
comprehensive, detailed investigations of the water resources including mathematical
model evaluation were undertaken by various authors.
The Amman-Zerqa area has been included III a large number of regional
geological and hydrogeological investigations. In addition, a number of specific studies
pertaining to this particular basin have been undertaken. These range in scope from
reports on groundwater exploration to mathematical model studies for evaluation and
management of the groundwater resources. For example, Mudallal (1973) carried out a
hydrogeological investigation in the Amman-Zerqa area. His study included
mathematical modelling and recharge estimates. Hemud (1973) investigated the declining
15
piezometric head of the artesian Hummar Aquifer System. The VBB
VATTENBYGGNADSBYRAN (VBB) in 1977 undertook a study to identify potential
water resources inside and outside Amman-Zerqa area to provide for the immediate and
long-term needs of the area and to formulate a water resources master plan to meet
Amman's water requirements up to the year 2005. Their study included drilling new
borehole, pump test analysis, mathematical model evaluation, recharge estimates, and
water quality.
Howard Humphreys and Sons (1977), by using the existing studies, assessed the
water resources in the northern part of Jordan. Carr and Barber (1972) undertook a model
study to predict the water level decline due to pumping in Qatrana area in Wadi Mujib.
In recent years great efforts have been made in conducting comprehensive
hydrogeological studies in the southern part of the country in order to facilitate the supply
for the increasing water demands in the area. Howard Humphreys Ltd. (1986) studied the
hydrogeology and the hydrochemistry of the Mesozoic-Cainozoic aquifer of the Ma'an
Shidiya-EI Jafr region in Southern Jordan. In addition, their study includes a
reconnaissance study of the Palaeozoic (Disi) sandstone aquifer system south and
southeast of Shidiya.
The German Federal Institute for Geosciences and Natural Resources (BGR) with
the co-operation of the Central Water Authority of Jordan (WAJ) in 1987 studied the
possibilities of, and constraints to, groundwater development for the water supply of the
envisaged oilshale processing plant in the Lajun area in Wadi Mujib; the study included
drilling, pumping test analysis, mathematical modelling and groundwater quality
investigations.
The Japan International Co-operation Agency (JICA), with the co-operation of
W AJ, has conducted a series of water studies in the central and southern part of the
country. In 1987, JICNWAJ studied the hydrogeology and water use of the Mujib
Watershed and the possibility of supply to Amman. In 1990, they conducted a
comprehensive study for the Wadi Hasa and Jafr basins, the study including drilling new
observation boreholes and groundwater mathematical modelling.
16
Numerous reports, papers and theses have also been published by many authors,
covering various aspects of geology, hydrogeology and water chemistry ofthe study area.
The data analyses and interpretation in these reports have provided the background and
detailed information about the aquifer systems, the geology, and the water chemistry
throughout the study area.
1.10 PURPOSE OF THE STUDY
Most of the sedimentary rocks overlying the Palaeozoic sequences and covering
most of the country, are the carbonate sequences of the Upper Cretaceous and Lower
Tertiary. These rocks form the major aquifer systems that supply a major part of the water
needs in the country. The Mesozoic sediments form a series of aquifers and
aquic1udes/aquitards in which five aquifer systems can be recognised. Four of these
aquifer systems are of carbonate origin. The fifth aquifer system, a sandstone aquifer
system, is present at the base of the sequence and has limited potential in the study area.
Many areas depend on the carbonate aquifer systems for all or part of their water
supplies.
The depositional thickness and lithology of the Mesozoic carbonate rocks are
extremely variable. The possible effect of major structures and change in rock type and
lithology on groundwater flow is the subject of this study.
A considerable amount of work has been done previously on evaluation of these
aquifer systems, but a comprehensive understanding of the hydrodynamics of these
groundwater systems is not available or has not been attempted.
The general purpose of this study is therefore to produce a conceptual evaluation
of groundwater flow and to better define the relationships between recharge, discharge,
water level, and aquifer characteristics. The study area includes the Jafr Basin, with its
differences in lithofacies and aquifer thickness, in order to allow the effects of geological
variation on hydrology and hydrodynamic pattern to be investigated.
The other objectives of this study are to analyse the changes that have occurred
between predevelopment times and present flow system, integrate the results of previous
studies that address either individual aspects of the aquifer system or local geographic
17
areas, and to provide some capability for evaluating the effects of future groundwater
development on the system.
These objectives can best be met by constructing a regional scale digital model of
the aquifer systems, supplemented by more detailed subregional models. Such models
will provide a framework for the interpretation and evaluation of the distributions of
observed aquifer characteristics and their relation to present and past patterns of
groundwater flow. They should also allow estimation of the yields available for each unit
as well as the impact of obtaining such yields.
In carrying out the study, the geometry and aquifer interrelationships as well as
the effect of large geological features (e.g. large intrusive bodies, regional lineaments and
faults) on water levels and groundwater flows will be inferred. Computer simulation will
be used to evaluate and help determine the regional distribution of such hydrogeologic
properties as hydraulic conductivity and leakance, especially in the downgradient parts of
the flow system where the data are sparse, and will also be used to help assess the
consistency of hypotheses, concepts, estimates and observations.
The results of this study intend to fully document and demonstrate the different
aquifer parameters and the description of the hydrogeologic framework and associated
flow systems of the carbonate aquifers, so that can be used by others to evaluate specific
groundwater management for the principal aquifer systems.
1.11 SCOPE AND METHODOLOGY
Literature search and data collection, involving compilation of geologic,
hydrologic, hydraulic and chemical data from published reports and from the files of
W AJ, dominated the early part of the study. These data were used to prepare a series of
maps and to describe the geology, hydrology and water quality of the aquifer systems.
Additional data were collected to fill major gaps in information.
No previous study has considered the carbonate aquifer system in the central and
southern part of Jordan as a continuous single hydrogeological system. However, a
regional approach is required if such problems as water level declines or the effects of
lateral changes in aquifer characteristics and regional structural systems on groundwater
18
flow are to be properly addressed. A major advantage of such an approach is that the
effect of such conditions as severe drought or widespread intensive pump age can be
analysed for the entire system, not just for a small part of it. As the analysis is regional in
scope, it does not address site-specific problems caused by intricate localised quality,
hydrogeologic, lithologic or structural discontinuities. The study is intended to answer
questions about the lateral flow of groundwater from recharge to discharge areas, its
vertical movement, and the general water yielding properties ofthe aquifer system.
Available hydrologic data provide most of the necessary information for the
interpretation and conceptualisation of the aquifer system. The physical boundaries of the
aquifers and the confining units are presented by different previous studies. The early
stages of the study comprised the compilation of a geological data inventory. A structural
map of the study area was then constructed using this data base: emphasis were placed on
describing the structure and lithological changes in the carbonate sediments that
constitute the aquifer systems. Climatic data published by W AJ for the period 1937-1985
were used to define rainfall, temperature and evapotranspiration to estimate the effective
rainfall and recharge to the aquifers. Soil characteristics, soil moisture deficits (SMD),
and infiltration properties were collected from previous studies.
Historical surface water data including baseflow, flood flow and spring discharges
were compiled and interpreted along with the calculated runoff by using the CN method.
These data were used for recharge estimation, water budget analysis and to find the
interrelationships between surface and ground water. Recession hydro graph analysis for
some springs was carried out where data allowed.
Hydraulic characteristics of aquifers and confining units were initially estimated
from analysis from geophysical and lithological logs of boreholes, data on specific
capacity of wells, flow net analysis, and from the available pumping test data. Hydraulic
conductivity and specific capacity maps were constructed for the area, and the relation
between the different aquifer characteristics was assessed.
Groundwater modelling was selected as the best approach to accomplish the
overall objective owing to its capability to integrate the many aspects of groundwater
19
systems. A groundwater flow model takes into account interaction between and
interdependence of aquifer geometry, properties, recharge, discharge, boundary
conditions, groundwater abstraction and groundwater-surface water exchanges.
The three-dimensional groundwater flow model code MODFLOW (McDonald
and Harbough, 1984) with the processing software PM (Wen-Hsing Chiang and
Wolfgang Kinzelbach, 1991) was used in this study. It was calibrated under steady-state
and transient-state conditions. The most important source of calibration data was the
inventory of water wells that is maintained by W AJ. Data from more than 700 wells in
the study area were used in this study.
Computer simulation was used extensively to evaluate and help determine the
regional distribution of aquifer hydrogeological properties and the effect of groundwater
development. The sensitivity of the model-generated water levels to selected variation in
hydraulic characteristics was tested. Simulation was accomplished with a coarse-mesh
regional model and six subregional models having smaller grid spacing. Since the study
area was divided into five subregional areas according to the location of major
groundwater divides and barriers, investigations within each subregion focused on local
flow systems and water problems in more detailed. Because of the physical and hydraulic
interconnection of these aquifer systems, computer simulation of adjoining parts of the
aquifer system have been compared to ensure that there are no conspicuous anomalies in
hydraulic heads and that reasonable amounts of water are simulated as passing between
the systems, in the direction indicated by field observations.
Regional groundwater chemistry and environmental isotope analysis, included in
previous studies, have been used to relate the observed variations in water chemistry to
changes in flow conditions.
Most discussions in this thesis are concerned with the B2/ A 7 aquifer system, the
most extensive exploitable aquifer system in Jordan, it is of primary interest for this
study. And it is been referred to as the "aquifer system".
20
1.12 STRUCTURE OF THESIS
In this thesis the research work is divided into three major categories. The first
part ( Chapter 2 to 4) deals with the geological, structural and climatic evolution of the
study area, and the framework hydrogeology of the different aquifer systems.
The second part (Chapter 5 and 6) concentrates on the different hydraulic
parameters of, and the estimation of recharge and water balance for, the different aquifer
systems.
The third part (Chapter 7 and 8) describes the different aspects of groundwater
flow hydrodynamics and mathematical model simulations. Discussions and conclusions
of the research are presented in Chapter 9.
21
2.1 REGIONAL GEOLOGY
2.1.1 OVERVIEW
CHAPTER TWO
GEOLOGY
Jordan lies on the northern edge of the Nubo-Arabian Precambrian Shield. This
shield is exposed in south-western Jordan and extends undermost of Africa and the
Arabian Peninsula. It is characterised by Precambrian plutonic and metamorphic rocks,
and by some minor occurrences of Upper Proterozoic sedimentary rocks, which is known
as the Precambrian basement complex.
The Precambrian basement complex has repeatedly moved up and down during
epierogenic activities ranging in age from Cambrian to early Tertiary. These movements
resulted in several marine transgressions and regressions of the Tethys Sea, which lay to
the west and north-west, over part of, or all of Jordan. The basement complex produced
the material from which, during certain periods, continental sediments were deposited in
the Tethys Sea.
During the transgressions, marine sediments of considerable thickness were laid
down. Inland of the transgression coastlines, and during intervals of regression, terrestrial
deposits accumulated: these consists mainly of sandstone of the Nubian facies, with no or
few fossils. This pattern of regressions and transgressions explains the pattern of the
different lithofacies - marine calcareous, marine sandy and continental sandy - of
Cambrian, Ordovician and Silurian sandstone and shale of continental and marine origin
which unconformably overlie the rocks ofthe Precambrian basement complex.
Regionally, the marine influence on the deposition increases toward the north and
west during the transgressive intervals of the Middle Cambrian, Early Ordovician, Early
and Middle Triassic, Middle Jurassic and Middle Cretaceous to Oligocene times.
Different shorelines have been formed due to these successive transgressions (Figure2.1).
Th~ total thickness of all post-Proterozoic sedimentary rock is generally 2000-3000 m;
38'
100 ISO KILOMETRES r-"~-T~-r--'---TI--~I~--------~I-,
33'
After Bender (1974)
Figure (2.1) Paleogeography of Jordan
SO
r:r::1 ~
~
D
EXPLANATION
Nubo-Arabian Shield
Stable shelf
Transition to unstable shelf. Unstable shelf in the northwest
Approximate east and southeast border of marine facIes
Direction of marine ingression. Border of marine facies beyond mapped area
however, post-Proterozoic rocks exceed 4000 m in thickness in the Jafr Basin in south
central Jordan, and 5000 m in the Al Azraq-Wadi Sirhan Basin in north-central Jordan.
It is believed that the epierogenic movements have resulted, in addition to the up
down movements, a horizontal drift. This drift is shown by the northward slide of the
East Jordan block with respect to the West Jordan block. Burdon and Quennell(1959)
stated that the displacement of this slide is estimated to be in the order of 107 km. It took
place along the Jordan Dead Sea-Wadi Araba Rift Valley.
2.1.2 OUTLINE LITHOSTRATIGRAPHY
The Precambrian basement complex consists mainly of grandiorite granite, with
minor acidic and basic intrusive dykes. In addition, the Sannuj Conglomerate, a molasse
type deposit, crops out at the south-eastern end ofthe Dead Sea.
The Palaeozoic sediments directly overlie the Precambrian rock complex. Olexcon
(1967) indicated that the Palaeozoic sedimentation persisted into the Carboniferous
period. During the Pennian, Palaeozoic rocks were tilted eastwards and eroded exposing
the older parts ofthe succession.
During the Middle Triassic, marine sedimentation resumed, but was restricted to
the north-west of Jordan. Transgression of the sea took place toward the end of Triassic
and the whole area was subjected to erosion. The third transgression occurred in the
Middle of Jurassic resulting in a further marine sedimentation of sandstone, dolomite and
mudstone in north-western Jordan. The regression of the Upper Jurassic was followed by
the extensive terrestrial deposition of the Kurnub Sandstone Group up until the end of the
Lower Cretaceous when a major transgression commenced eventually covering most of
the country. The main marine sediments were limestones, dolomites and marls. These
sediments are fonn the "Ajlun Group". However, the type of sediments changed into
chalks, cherts and marls towards the end of the Turonian without stratigraphic break, and
this sequence tenned the "Belqa Group".
It is believed that the final withdrawal of the sea from East Jordan took place
throughout Upper Oligocene- Lower Miocene. After that all the sediments are believed to
24
100
000
900
N
-\rE .;,-;':
S Y R I A ... ::::r:::::~
Q.
,::,
(.)
(.)
o
s
JERUSALEM
CI
UJ
•
--------------------------- -----------===-====-==-=---==-----------
-------=-------------=--
200 300
Compiled/rom: Burden and Quennell, (1959), Parker (1970), and Bender (1974).
(Figure 2.2) General geological map of Jordan.
50 km
I
LEGEND
II,\SALT
HECENT
CENOZOIC
MESOZOIC
PALEOZOIC
PRE-CAMOIUAN
400
be of terrestrial or lacustrine origin (Burdon and Quennell, 1959). Figure (2.2) shows the
general geology of Jordan.
The basaltic eruptions in the north-eastern part of Jordan took place at intervals
during the Middle Eocene to Recent. The major lava flows emanate from Jebel Druze in
the southern part of Syria. However minor flows originated from vents east and south
east ofthe Dead Sea.
The volcanic activities were contemporaneous with the major tectonic movements
which tilted the Mesozoic-Tertiary sediments towards the east and north-east, and faulted
and folded them adjacent to the Rift Valley (Parker, 1970).
During the Pleistocene, lacustrine sedimentation of marl, sandy limestone,
limestone, gypsum and clay occurred in Jordan Valley, Azraq and Jafr areas.
Recent sediments consist of sands and gravel in Jordan Valley and in the closed
drainage areas.
2.2 GEOLOGY OF THE STUDY AREA
The geology of the study area is mainly of sedimentary origin ranging in age from
Cambrian to Recent, overlying unconformably the Precambrian basement complex.
However, in some areas, volcanics of Quaternary age do occur and very locally the
Precambrian outcrops.
The sedimentary succession which reaches up to 5500 m in thickness is mainly
the result of a series of regional transgressions and regressions of the Tethys Sea. The
lower part of the succession is mainly of sandstone of Palaeozoic and Lower Mesozoic
age, while the upper part is mainly composed of limestone, marls and cherts of upper
Mesozoic and Cainozoic age. The geology is illustrated by Figure (2.3) and the
stratigraphical succession is summarised by Table (2.1).
Whilst the aim was to define the geology of the area for the hydrogeological
study, it was found that the stratigraphic nomenclature established by the early workers in
Jordan ( Quennell; 1951, Burdon and Quennell; 1959, Masri; 1963, MacDonald et al;
1965, Wiesemann; 1966, Parker; 1970 and others) was convenient to use and, therefore,
will be retained and adopted for this study.
26
200~_~~I] 180~
160
140
120
100
080
111\.["," o
060 1----+-4--
040
020
980
960~_
920
900
From the WMP, 1977 .... r........L....:.. __ ~_~ ____ ~ __ .:i
Figure (2.3) Geological map of the study area --"_.- - Kh
~~4"ID ... :.:.:;+;:: Be
LEGEND
Sand and gravel Quaternary marl and gypsum Basalt Rijam Formation Muwaqqar Formation Amman-Wadi Sir Formation lower Ajlun Group Kurnub-Zerqa Group Khreim Group Disi Group Basement Complex
Faull
TERNARY Pliocene Miocene
TERTIARY Oligocene Volcanic
Eocene
Palaeocene Belqa Group
UPPER CRETACEOUS
Ajlun Cenomanian Group
PRECAMBRIAN
shaded area indicates
Table (2.1) Geological succession in Jordan and occurrences in the study area
2.3 STRATIGRAPHY
2.3.1 THE PRECAMBRIAN BASEMENT COMPLEX
Precambrian rocks including the crystalline basement complex and the overlying
Sarmuj Conglomerate (Blanckenhorn, 1912) and the slate graywacke series (Bender,
1968) underlie the area, although they do not outcrop. These units are restricted to local
exposure in south Jordan, the east side of southern Wadi Arab a, and to the south-east
shore of the Dead Sea. Hornblendite, hornblende gabbro, diorite, quartz diorite,
grandiorite, mica aplite granite, and quartz porphyry are the most common plutonic rocks
in Jordan.
28
In the study area the Precambrian rocks occur at variable depth beneath a variable
thickness of Palaeozoic and later sediments. They are nearest to the surface in the extreme
south-west, where they are at an altitude of 500 m below sea level. In the Jafr trough, the
basement is up to 2000 m below sea level; up to 5000 and 3800 m below sea level in the
Azraq trough and in the north-east of Jordan respectively (Bender, 1975). A line of swells
in the basement extends north-westwards from Bayir, with contour closures at Safra,
Amman, and Ajlun.
2.3.2 THE PALAEOZOIC SUCCESSION
The Palaeozoic succession comprises a thick accumulation of primarily
aranaceous sediments. It is known to underlie the study area but outcrops are restricted to
the lower slopes of the rift escarpment and the escarpment face in the vicinity of Ras en
Naqb. These rocks are best exposed in the Southern Desert where two groups have been
recognised (Lloyd, 1969)- the Disi Sandstone Group and the Khreim Sandstone Group.
2.3.2.1 THE DISI SANDSTONE GROUP
This Group, unconformably overlying the Precambrian basement complex, was
exclusively continental in the Cambrian, but became mixed marine and deltaic and finally
fully marine in the early Ordovician, when the first major transgression of the Tethys Sea
across Jordan occurred.
The Disi Group forms the most extensive arenaceous deposits in the Arabian
Peninsula. It consists of 1000 m of medium-to coarse grained sandstone and siltstone of
Lower Cambrian to Middle Ordovician age.
2.3.2.2 THE KHREIM SANDSTONE GROUP
The Khreim Sandstone Group comprises mainly 600-800 m of fine-grained
micaceous sandstone and siltstone which were deposited in an unstable shallow marine
environment at the unstable shelf edge of the Tethys Sea during the middle and late
Ordovician and early Silurian.
29
The Palaeozoic succession was tilted to the east and eroded to expose
progressively older beds from east to west before the deposition of the Mesozoic
sediments, which therefore overlie the Palaeozoic with angular unconformity.
In the study area the Palaeozoic succession has only been proved by drilling. At
Safra well (PP3) south-east of Amman, the Palaeozoic succession was encountered at a
depth of 890 m below the ground surface; it consists of 1660 m of sandstones and
limestones. In an oil test well at Baqa'a (PP6), sandstones and subsidiary limestones
which are believed to be of Palaeozoic age were found between 911 and 2329 m below
the ground surface. In the centre of El Jafr, the top of the Palaeozoic succession was
encountered at 600-800 m below the ground surface, and seismic data suggest that the top
of basement may be at about 4000 m below the ground surface.
2.3.3 THE MESOZOIC SUCCESSION
2.3.3.1 ZERQA GROUP
The Triassic/Jurassic Zerqa Group unconformably succeeds the Palaeozoic rocks
and occurs only in the northern part of the study area. Rocks of this Group outcrop in the
lower slopes of the Jordan Valley and in the lower valleys of the deep dissecting wadis
between Wadi Zerqa and Wadi Mujib. The Zerqa Group has been subdivided into the
Ma'in Formation (Zl) of Triassic age, comprised principally of limestone, shales and
marls, and the Azab Formation (Z2) of Jurassic age comprised principally of dolomitic
limestone with marls and sandstone. They were probably deposited in epineritic to
lagoonal or shallow neritic environments. Both formations wedge out from north to south,
but the Azab Formation extends southwards for only 20 km from Wadi Zerqa (Parker,
1970). South of Wadi Hisban, the Azab Formation has completely wedged out and the
younger Kurnub Group directly overlies the Ma'in Formation. In Wadi Mujib the Zerqa
Group is completely absent.
2.3.3.2 THE KURNUB GROUP
. The Kurnub Group is predominantly comprised of sandstones and marls of lower
Cretaceous age and underlies the whole of the study area. It outcrops in the side wadis as
30
well as along the escarpments, and in eroded anticlinal structures ofBaqa'a, Wadi Sir and
Na'ur. The Group is thickest to the west of Amman, where 350 m of sandstone is
recorded, and it thins toward the south and south-east; its thickness ranges between 180
and 230 m in the Dead Sea area, and between 50 and 200 m in the Jafr area, and it is
absent in the extreme south.
The Group has been subdivided into two formations. The Arda Formation (Kl),
the lower Formation, comprises a thick sequence of massive light grey cross-bedded
loosely cemented sandstones which in places, particularly in the south, contains minor
clay horizons and occasional marls. The Subeihi Formation (K2), the upper Formation,
comprises interbedded sandstones, silts, clay and marls. It is readily distinguishable by
the distinctly varicoloured nature of the friable, crossbedded sandstones which make up
the major portion of the Formation. The upper boundary of the Formation is only readily
distinguishable where it is overlain by the carbonate facies of the overlying Ajlun Group.
Where it outcrops along the southern escarpment it cannot be separated with certainty
from the overlying sandy facies of the Ajlun Group.
2.3.3.3 THE AJLUN GROUP
Upper Cretaceous strata underlie most of the northern part of the Mountain Ridge
Provinces and they cover almost the entire Central Plateau. The beginning of the late
Cretaceous sedimentation in west, north and central Jordan is marked by marine
carbonates, whereas in south Jordan, deposition of sandy continental sediments
continued. The marine calcareous marly and siliceous rock units of the Cenomanian,
Turonian, Santonian, Campanian, and Maestrichtian overlie the continental sandstone
with onlap toward the south-east. Therefore, the marine carbonates in the north and west
become continental sandy facies in the south and south-east (Figure 2.4).
The Ajlun Group embraces all the marine sediments of Cenomanian-Turonian age
which overlie the Kurnub Sandstone, and consists of limestone, dolomite, marl, shale,
chalk and sometimes sandstone. The maximum thickness of the Group is 700 m in the
Azraq trough and between 500 to 550 m in the Amman-Zerqa area; it thins southwards to
reach about 50 m at Batn el Ghul (Figure 2.5 & 2.6).
31
L C :I
Z 0
~ c u 0 oJ
I ;;
!. ... .. •
. • -• I : i .
! I
~I gl "'I cl 31 :;:,
· · · ·
. . ~
c 0
! .
'i ~ :: E .. .; .. ... 1 c " D . D . .. en :I ", .
L = 0 e .. z = oJ ... c
.. • .. 1 · :! . .. · en E .. ~ · en
.. .. i . i >
i 0
c ns '
"C '-0 -, C '-Q) .c -:l 0 en c a. :l 0 '-
C)
C :l ......
. ~
CI) .c -.... 0 >-C) 0 0 .c ~
c CI)
6' C) c "- ns 0) .c
~ ""- 0
1 ,L I~ ,e I"
i~ I Z 'E
I~
~ -Q) "It 1: N ~ -CI) ~ '-Jg :l
C) <:( u.
200 250 300 350 400 450
.. ».",.
~0~----~~------~~~~~~r-------~r-----~--r-~~"'~---4~
150
~ I·. CC 100 ~ c:: ~
C:)
~ ,,-\C:l \. '. '\
~ 00
J /' "". /,. ,. i
"0: ,.,.
9001~----~--~~~~ __ ~~~~rr~------~--------+---------~0
.--.-.... I .--.-.-. ~O 850~----~--~~~~~~~.L-------~------~=----------r------~~~-
I
200 250 300 450
____ Isopachy" ( at 50 and 100 m intervals 1 • Location of measured thickness
After Parker (1970)
Figure (2.5) Isopachyte map of the Ajlun Group
EOCENE TO
PALEOCENE
MAESTRI CHTlAN
Irbld Aria Jemal Abyad
SENONIAN
TURONIAN
,, 0. -1 r~-;-"~- =..r. ..&.... r-L:1 . • . • - ::c' ." ,-,£::t;4~I . " ' J "T", . ,_,,"'..i' ·;'·'·' · ·j,· :T,~I.":'T,,:t :-t'~· ·- .~-- . ~;~~~~i~w.;~~~4+~f.t~"*·;·'·'·-~~-~ ' cT.._1_ •. : ,-.J _ :-=S-~-~L: L f __ : I . Jon · r: ~_. :- . ~-:ct_ Xl~!~· ~~~~~~¥i:Ji~~;g;(t
CENOMANIAN
LOWER CRETACEOUS
Llmillon.
Dolomitic IImlllon.
Sand, IImlllon.
Marl, IImlllon.
g-ifi _=_1IfI~ 5" '" c"
=~.~,~.~;;2:.
Chalk
~;:> Sand
.. '. Phoaphar".
... CharI
Mati
Shale
Sandy cloy
After Parker (1970)
Figure (2.6) Generallised cross-section through the Mesozoic marine
succession from the northwest to the southeast.
Ajlun Group was named as Ajlun Series for the first time by Quennell (1951) and
Burdon and Quennell (1959), then as the Ajlun Group by Sir M. MacDonnald (1965).
Several systems of subdivision of the Group have been proposed by different authors
(Table 2.2). Wolfart (1959), was the first to attempt a lithological subdivision of the
Group. Masri (1963) subdivided the Group into five lithological units by using local
terms to indicate the different Formations of the Group and symbols Al to A 7 as
abbreviation of the Ajlun sequences as follows:
Wadi Sir Formation (A7) Turonian
Shu'eib Formation (A5/6) Upper Cenomanian
Hummar Formation (A4) Upper Cenomanian
Fuheis Formation (A3) Middle Cenomanian
N a'ur Formation (Al/2) Lower Cenomanian
MacDonnald (1965) subdivided the Group into three units: (Al-2), (A3-6) and
(A7). But they subdivided the A7 into three subunits as A7a, A7b and A7c (Table 2.2).
While The German Geological Mission to Jordan (1961-1966) subdivided the Ajlun
Group on stratigraphical basis into three main units:
Sandy Limestone and Massive Limestone Unit
Echinodal Limestone Unit
Nodular Limestone Unit
The Sandstone Aquifer Projects of the UNDP (1965-1968) used the same
subdivisions as Masri (1963).
The Ajlun Group represents the most laterally variable sedimentary sequence in
the entire succession above the Palaeozoic, particularly in the south and south-east.
Consequently it presents the greatest difficulty when correlating between boreholes. In
the Western Highlands and in the northern part it has been well described at outcrop by
various authors who have attempted stratigraphic correlation on the basis of observed
facies changes (Wiesemann, 1966 and Bender, 1968).
In the south and south-east, rapid facies changes occur, sometimes with virtual
elimination of the carbonate horizons and thinning of the Group to not more than about
35
Parker (1965-1968)
F ormation A 7 I Formation A7 I A 7 c Limestone stone Unit & I Formaqtion I Formation A7b Massive A7 A7 Limestone & marl Limestone A7a Unit
Shue'ib Marly Sandstone, Sandy Echinoidal Shue'ib Shue'ib Formation Limestone limestone, shales Limestone Formation Formation A5/6 Formation marls, marly Unit A5/6 A5/6
A5/6 limestone A3-6 A51 A 1-7
Hummar A4 Limestone A4 HummarA4 Fuheis Formation Formation Formation A3
Formation A3 I Marl A3 Nodular Fuheis A3
Formation Limestone Formation Marly nodular I Na'ur A 112 I Limestone Limestons, Unit Na'ur limestones A2 Formation Formation A2 sandy, marly, Formation
marls, shales A 112
Marl Al limestones Al I I Formation dolomite
Table (2.2) Correlation of litho-stratigraphic units of the Ajlun Group recognised by various authors
200 250 - 300 350 400 450
36 i , ,,.-- 'J 38" 33~., ~ ,,,' ,," -.'
2S0~----~-+-------+~~~~==+---------~--------~~~~"~--~2~
S Y R I A ! )< I J' .. "."."." . I -'"
200'1~ ____ ~ _____ ~ ____ ~~~ ____ ~'_I'~ ______________ ~200
i • ,,,,. l-·-/ 1°" It, I 0 Zerqa I I L,zo~: AMMAN ,zo..,
150,: -----+-----l2Io~i ----+/--O-A-Z-ra-q -+I-----'I----.::.:~' JISO
! oMadaba 1, .-.-'-' , i I I .J._._._._'-,.-- ~ I 100 ... · -+rl::::/-~-:-::-:--~-------!...' ----~----...L---;;--'IOO
i " ct 1 I .'. 0: I
'" ~ I i \ .,,', I
oso;...· ____ ~~--------\~'----------~---·~~~--~i--------"~~--~-----~"o i oTalila >iO~Q ! 0 :0:'/ !'<~~ "0
1\050.
I ,( ",0·'° I I \. I • . """ i """ ,0'" 1 \ , : • w. --- I """,/"'"' 0~10 i·
ooo! ,,- -- """ """ .cO I '. , .0 ShQ.ube~'. _-;' <"" /' .,0" 0 I ''. 1
000
, • / . --- .;..< ./' -- 1 01 . > I' ....... -- I./' "" &0 ° I ", I Wedi Mu~a. / -.-----;-- ,/"'"/ -": 10°10 I ".",
I, 0 ,' •• i . //I~ "~afr----;-__ --~....c;_ &0'/0 ,I" I I / • !--~ ------ I .., I i ! / 0 • .,07 / 11~-.:::::-=----1_,!00/0 l
950'--....: •. ~-7-/-/"'7:"'7"-.+ //-T.I-+-/+-r /;:"'/:--~""-::::::-:=~=-"ilr----A----~950 ! ,oOfo'""Cfifa;'Jnyaq"ob "/ / I· ,I :"'00 10°/0 -'::- ....... /' /) / _._.J 300
I 0"1..;r........ /' . / / _._.-' ! ~ 4'00/0"-:. ./ /. ~/ --' . , oOuwerra..,o"lo;/ / • .1 / I ,r' i I I .c!lo.//' VI .i' I
900 I 0°10 A --------+--f--------7----------.j.....;"-------~ i I 1 --;co,. -polO [/ I 900
"j I /t; , I I;' j
r·-·-,L._._ ,I ./// i' 850' , '-. I ~o; .
200 250
Isopercenti Ie of sand content
• location of observed section or borehole samples
After Khdier (1965)
'7°1
300 350
o 50 tefft, ... ' ............. ' --"--'--'I
I
400
Figure (2.7) Percentage of sand in the Ajlun Group
850 ,so
450
50 m and sometimes it becomes diachronous with the Subeihi Formation of the Upper
Kurnub Group. The general pattern in lithofacies changes is that of an increase in sand
content from the north and west toward the south-east (Figure 2.7).
In the southern area of the Central Plateau, a formation name of the "Fassu'a" for
the entire Ajlun Group was proposed by Weisman (1966), because of the difficulties of
distinguishing the formation units in the Ajlun Group. The Fassu'a Formation which it
has its type section on the escarpment near Batn Ghul in fact represents only the lower Al
to A6 formations of the Ajlun Group, since the Wadi Sir Formation (A7) is of importance
to the study and can be identified over most of the area. The occurrence of the different
formations of the Ajlun Group are shown in Table (2.3) and Figure (2.8).
Formation Area Thickness( m) Rock Type Wadi Sir Amman 85-90 limestone and chert
A7 Mujib 70-128 limestone, chert, & marly limestone
Jafr 60-100 sandy limestone & sandstone Shu'eib Amman 75-100 marl, limestone, & shale A5/6 Mujib 127 marly limestone & shale
Jafr 75-100 marly sandstone, sandy marl, & marl Hummar Amman 40-65 dolomitic limestone
A4 Mujib 7-12 dolomitic limestone Jafr 20 sandstone
Fuheis Amman 80 limestone, marl, & chalk A3 Mujib 70 shale, marl, & limestone
Jafr 90 shales, clays, sandstone, & limestone Na'ur Amman 160-270 limestone, dolomite, & marl
A 112 Mujib 120-170 marl & limestone Jafr 50 shales, sandstone & limestone
Table (2.3) Occurrence of Ajlun Group Formations
NA'UR FORMATION (A1/2)
The Na'ur Formation comprises the basal member of the earliest late Cretaceous
sedimentation in west, north and central Jordan. It forms the lower part of the Nodular
Limestone Unit of Bender (1968). It was named "Na'ur" by Masri (1963) for the rocks
exposed near Na'ur village in the vicinity of Amman. It consists of a lower marly unit
"AI" and an upper limestone unit "A2".
38
r=-'" .. g .. ." oS ;l :;(
r=-'" .. g .. ."
.€ ~
1100 Upper Zerqa Basin
1000 900 800 700 600 500 400 300 200 AV2
100 NW
0 SE
1300 Wadi Mujib Basin
1200
Wadi el Abiad 1100 1000
Karak-Wadi Fiha fault line
900 800 700 600 500 400 300 200 100
0
The bOlmdary between Hasa and Jaft basins 1500r--z~ __________________________________________________________ ~
1400 1300 1200 1100 1000 900 800 700 600 500 400 300 200 100 W
Karak-Wadi el Fiha Fault Line
E OL-________________________________________________________ ~
Figure (2.8) Geological cross-sections in the study area
39
The type section of the Na'ur Formation is marked by a thick ledge of limestone.
This limestone is grey and sometimes pink in colour, is hard crystalline, coarse grained,
fractured, and in some places contains chert nodules. This bed is underlain by alternating
beds of shales, limestone and marl with occasional sand and sandy marl.
The Formation outcrops extensively along the rift escarpment and also inland along
the major wadi courses. The maximum thickness ofthe Formation was found to be that of
the type section at Na'ur area where it attains a thickness ranging from 200 to 230 m. To
the south, in the Madaba area, the Formation has a variable combined thickness ranging
from 120 to 250 m.
In the Dead Sea area the Formation consists of three strong light brown or grey
brownish dolomitic limestone units which individually may be up to 20-25 m in
thickness, alternating with weak or moderately weak yellow or greenish grey thinly
bedded silty mudstones.
The Formation thins out to the south and south-east where it comprises the lower
part of the Fassu'a Formation. It is composed of literal sandy facies almost entirely,
containing some limestone in the south-east, and becoming more shaley and marly in the
east.
FUHEIS FORMATION (A3)
The Fuheis Formation comprises the upper part of the Nodular Limestone unite. It
is composed of 70-80 m of marls intercalated with marly limestone. The Author in 1988
described a type section for the Formation to the west of Amman to be 70 m of marls and
chalk with occasional interbedding of thinly horizontal layers of limestone. At the top of
the section a 60 cm thick of very hard, brown, fine crystalline limestone bed is overlain
by a soft marly limestone layer of 1 m in thickness containing some salts. The marl
dominates the main part of the Formation, comprising more than 60% of the whole
section (Khdier, 1988).
To the south the Formation consists of moderately weak to moderately strong
grey and light brown thinly to medium bedded silty limestone or calcareous siltstone
interbedded with weak to moderately weak yellowish brownish very thinly bedded
40
calcareous silty and very fine sandy mudstones which breakdown to sandy clays at the
surface. In places these mudstones contain gypsum. Further to the south and south-east
the Formation thins out and becomes more sandy.
HUMMAR FORMATION (A4)
The Hummar Formation comprises the lower part of the Echinoidal Limestone
Unit of Bender (1968). In the north, in the Amman area, the Formation is a distinctive
well recognised limestone, light to dark grey, occasionally pinkish, hard, coarse grained,
and dolomitic. The thickness of the Formation ranges between 40 and 65 m in the
Amman area, where it forms an important aquifer, and has been penetrated by many
wells. Toward the north and north-east the Formation thins out and becomes marly. The
Formation thins out also toward the south and south-east and becomes less important. In
the Madaba area it consists of about 30 m of massive well-jointed crystalline limestones.
This changes to marly limestone very similar to the underlying Fuheis Formation in Wadi
Hisban. While in the Wadi Mujib, it is described as the lower part of the Echinoidal
Limestone Unit, and contains a 7-12 m thick very dense dolomitic limestone.
SHU'EIB FORMATION (AS/6)
The Shu'eib Formation forms the upper part of the Echinoidal Limestone Unit as
well as the upper part of the Fassu'a Formation in the south and south~east. In the north
the Formation is principally composed of marls, marly limestones and shales with a total
thickness of about 65-100 m. It acts as an aquiclude between the upper and the lower
aquifers of the Amman-Zerqa groundwater basin.
Facies changes resulting in a variation in marl to limestone ratio make it difficult
to separate the Formation from the overlying Wadi Sir Formation in the north and from
the underlying Nodular Limestone Unit in the south. However, in some places, the
Echinoidal and the Nodular limestone Units have been mapped as one unit.
In Wadi Mujib, the Formation consists of two parts: the lower part consists of
moderately strong silty limestone interbedded with calcareous shaley mudstones; and the
41
upper part consists of generally moderately strong to strong, white, light brown, or light
grey, finely crystalline limestone with occasional chert nodules.
Southward from the Mujib the limestone becomes progressively more finely
sandy and dolomitic. There is also a tendency for the limestone in the upper most part of
the sequence to become sandy and grade into the overlying Formation.
WADI SIR FORMATION (A7)
The Wadi Sir Formation comprises the uppermost part of the Ajlun Group. It
corresponds to the Bender (1968) Massive Limestone Unit of north Jordan or the Sandy
Limestone Unit of South Jordan. The Formation has a wide outcrop in the Western
Highlands. Karstic weathering appears to be well developed, especially in the upper
section ofthe Formation.
The Formation considered to be the highest yielding aquifer in Jordan and has
been penetrated by numerous wells all over the country.
In the north the Formation consists of limestone, white to light grey, seml
crystalline, lithographic occasionally chalky and marly towards its bottom. The
Formation consists of mostly limestones with chert concretions forming discontinuous
bands. In some places the Formation contains white, fine-grained, quartzitic sands, and
the upper part is of sandy limestone. The thickness of the Formation in Amman-Zerqa
area is about 100 m. The author in 1988 described the type section of the formation in
Amman area (Figure 2.9) to be about 65 m thick.
The regional thickness of the formation varies widely from the north to the south
of the country. It has a thickness of 185 m in the north as penetrated in the Qumeim well
to the west of Irbid. To the south the thickness of the Formation decreases and the
lithology becomes more sandy. The complete sequence of the Formation is exposed in
Wadi Mujib; the Mujib area is a transitional zone between the Massive Limestone Unit of
north Jordan, represented by thick bedded massive limestones, and the sandy facies of the
southern Jordan. The Wadi Mujib sequence consists of massive, fine, sandy limestones;
the limestone contains chert and fine grained sandstone layers. Intercalations of
bituminous shales and mudstones frequently occur in the area of EI Lajun (Ruef and
42
350 r---------------------------------------------------------~
300
250
200
150
LEGEND
chert
100 sandstone
sandy marl
marly limestone
50 marl
phosphatic limestone
limestone with chert
limestone
A1TD11an W.Wala W.Mujib
Figure (2.9) Type sections of the Amman-Wadi Sir Fonnation
43
Jeresat, 1965). Sometimes the unit contains also massive marl layers. The thickness of the
Formation in the Wadi Mujib area ranges between 70 and 100 m. In Wadi Bin Hammad
the Formation forms a 100 m vertical cliff of alternating, generally moderately weak to
moderately strong thinly to medium bedded silty, sandy mudstones, locally dolomitic
limestones, and fine to coarse sandstone (Bender, 1974). Towards the south and the
south-east the sand content increases until fine sandy layers make up the dominant part of
the unit. Here the whole of Ajlun Group is sandy, and hence cannot be differentiated into
formations. At Ras en Naqb, Weisman (1968) described the Wadi Sir Formation as 39 m
of sandy limestone with chert nodules in the middle section and with marl bands toward
the base. To the south toward Batn Ghul, the Formation is no longer recognisable, the
carbonate horizons having been virtually eliminated and the entire Ajlun sequence
replaced by approximately 50 m of sandy deposits.
2.3.3.4 THE BELQA GROUP
The Belqa Group overlies the Ajlun Group throughout most of the country. The
Group includes all the sediments from the end of Turonian to the Oligocene. These
sediments began to be laid down when the Tethys Sea was transgressing inland towards
the south-east (Burdon and Quennell, 1959). The Group attained its name "Belqa" after
the Belqa district where they are well exposed. The Belqa Group sediments lie
conformably above Ajlun Group except in some places where the lowermost part of the
Group is missing. The Belqa Group mainly consists of chert and carbonate rocks. It
reaches a maximum thickness of more than 600 m in the Azraq area and north of Irbid
(Figure 2.10).
The Group has been subdivided into several Formations and units by various
authors. Wolfart (1959) subdivided the Belqa Group sediments into sequences. Masri
(1963), as with the Ajlun Group, used local terms to indicate the different formations of
the Group and symbols prefixed by "B" as abbreviations. Sir M. MacDonald (1965)
subdivided the group into five formations. The German Geological Mission (1961-1966)
subdivided the group into four units. The Sandstone Aquifer Project of the UNDP
44
200 250 350 400 I
13· 37"
250~----~-+------~~~~~~---------+--------~~~----~2~
SYRIA ,,-
.",'
.'" J ",.",.",.",
'OOlr3-z-.--~r-~------~----~~-------·-/~-+O--M5------~--------,-z.~j'CO
150'~-----+---f---"""'-~--~"-:'r-'-r-----------:-I----------:I-------.-_-· .-j 50
I .-.-, . A 9. .~'- i I .-' ·VC: ' I ','- i
H ~ : 100·~· ---+~b'----:--::-'--;-+-t~-->...--''''<-'>'''':::-'-.... .-;;---...... -----------co ---100
i \,: Ii " ~
\ I " c::;, 050~--~~~~~~~~---~~~--~-----:-------"~_~~----IC50
1.. 3,. i \ ~ I "c::, I , \
e tJ i", lOCI( I ,
: '\ .~~~~~~~~"-:~-----~-------~ ____ ~'\ __ coo
2CO 250
--- /sopochyte (at 50 m interyal )
• Location of measured or compiled
thickness
After Parker (1970)
>
o
300 350 400 450
Figure (2.10) Isopachyte map of the 8elqa Group
marly limestone of Yarmouk area & Nummulitic Limestone Unit of Sirhan &
B4 I Absent Chert B4 I Chalk, B4 I Chert-Limestone Unit & chert I & limestone
Formation
Chalk B3 I Muwaqqar B31 Chalk- B3 & marl Formation marl
Limestone I Amman B chalk, B2 Formation chert
Limestone BIb
marl, chalky marl
Formation
Limestone Formation
Chalky BI marl Formation
Limestone & Chert Unit Chalk B3 & Marl Formation
Limestone & Phosphate Formation
Chalk, BI Marl, Chert & Sand Formation
Chalk-Marl Unit
Lime stone Unit
Parker (1965-1968)
Formation
Rijam B4 Formation
Muwaqqar B3 Formation
Amman B2 Formation
Wadi BI Ghudran Formation
Rijam B4 Formation
Muwaqqar Formation
Amman B2 Formation
Table (2.4) Correlation oflitho-stratigraphic units ofthe Belqa Group recognised by various authors
(1965-1968) also subdivided the group into five formations. The different subdivisions
and the correlation between the litho-stratigraphic units of the Belqa Group recognised by
various authors are summarised in Table (2.4), While Table (2.5) shows the occurrences
of the Belqa Group formations in the Western Highlands and the Central Plateau.
Formation Area Thickness (m) Rock Type Rijam Amman ? limestone, chert, & chalk
B4 Mujib ? limestone, chert, & chalk Jafr 40-50 limestone, chert, & chalk
Muwaqqar Amman remnants chalk & marl B3 Mujib 70-150 marl, chalk, chert, marly limestone, bituminous
Jafr 20-450 marl,shale, marly limestone, bituminous, sandy Amman Amman 80-110 limestone, chert, marl, & phosphate
B2 Mujib 100-260 limestone, marl, chert,& phosphate Jafr 30-100 chert, limestone, marl, in part phosphatic, sandy
Ghudran Amman 15-35 chalk & marl Bl Mujib 0-15 marls tone, marl, in part silicified
Jafr 10-15 chalk occasionally silicified
Table (2.5) Occurrence of Belqa Group Formations
AMMAN FORMATION (B1I2)
The lower part of the Belqa Group is composed of two Formations: the Ruseifa or
Wadi Ghudran (B1) Formation and the Amman (B2) Formation. The Wadi Ghudran (B1)
Formation has been proposed by the sandstone aquifer project (UNDP, 1970) for a
sequence of chalk and marl which forms the lowermost part of the Belqa Group and
overlies directly the Wadi Sir Formation of the uppermost Ajlun Group.
The B1 Formation is a thin and discontinuous unit. It can only be recognised as a
Formation in northern Jordan where it its thickness may reach 50-60 m. In the Amman
Zerqa area, as in many other localities, this Formation thins out and in some places is
missing. Because of the difficulties in distinguishing the Wadi GhudranlRuseifa
Formation (B 1) from the Amman Formation (B2), they are referred to together as the (B2
or B 112) Formation.
The Amman Formation (B2) is the most lithostratigraphically continuous unit
occurring over the study area. It outcrops extensively in the upland plateau area east of
47
the Western Highlands. It comprises two Members; the Silicified Limestone Member of
Santonian age and the Phosphorite Member of Campanian age, which correspond to the
Silicified Limestone Unit and Phosphorite Unit of the German Geological Mission (1968)
respectively.
The Silicified Limestone Member consists of a suite of similar lithologies
comprising alternating thin - bedded silicified limestone with chert, marly limestone,
marl and limestone (Figure 2.9). The Phosphorite Member consists of alternating thin -
bedded limestone, more-or-Iess silicified or calcified phosphorite layers and coquina
beds. The Formation becomes sandy toward the south and south-east. A characteristics
feature of the Formation is the undulating structure of its beds, this undulation does not
affect the under or the overlying Formations.
The thickness of the Amman Formation ranges between 80 and 110m in the
Amman-Zerqa area, and attaining up to 200 m in thickness in the Madaba area. In Wadi
Mujib the thickness of the Silicified Limestone Member is approximately 135 m, but the
thickness decreases to the east and to the south, where at Mahattat al Hasa and Al Qatrana
it is only about 40-70 m thick, and thinning to less than 20 m at Zakhimat al Hasa. The
thickness of the Phosphorite Member is approximately 90 m in the Wadi Mujib area,
decreasing to the east and south. The thickness of the Amman Formation ranges between
50 and 100 m in the Jafr through to less than 30 m to the south in the Shidiya area.
The Amman Formation (B1I2) together with the underlying Wadi Sir Formation
(A7) form the most important groundwater aquifer system in the country.
MUW AQQAR FORMATION (B3)
The Muwaqqar Formation comprises the upper part of the Belqa Group and it is
equivalent to Chalk Marl Unit of Bender (1968). It consists of chalk, marl, chalky
limestone, bituminous marl, shales and chert nodules. The Formation outcrops
extensively in the eastern part of the study area, a long belt trending north-south
extending parallel to the highlands. The author recorded in 1988 a remnant of the
Formation in Ruseifa, Marka and Jabal Al Akhdar in the Amman area (Khdier, 1988).
The thickness of the Formation ranges from about 20 m in south-east Jordan to more than
48
450 m as drilled in the Jafr Basin; the thick sequence is restricted to basins that strike
north-west. The Formation reaches a maximum thickness of 146 m in El Lajun (Abu
Ajamieh, 1980) and a thickness of up to 107 m in Wadi El Moghar in the south-eastern
Mujib catchments. According to Masri (1963), the thickness of the Formation in the
south-east of Amman area ranges between 60 and 70 m , while drilling logs shows that
the thickness may exceed 100 m. But the sandstone aquifer project (UNDP, 1965-68)
noted a thickness of about 270 m to the south-east of the Muwaqqar village. In Irbid
district, Wolfart (1959) recorded a thickness of 200-240 m, while Sir M. MacDonald
(1965) recorded a thickness of up to 320 m in the northwest of Jordan.
2.3.4 THE CAINOZOIC SUCCESSION
RIJAM FORMATION (B4)
The Rijam Formation (B4) is the uppermost unit of the Belqa Group in the study
area, and hence the uppermost formation of the major conformable sedimentary sequence
extending from the Lower Cretaceous to the Early Cainozoic. However, in the
northernmost part of the country and in the Azraq and Sirhan basins, the Formation is
overlain by the Wadi Shallala (B5) Formation, the youngest of the five divisions of the
Belqa Group.
The name of the Formation was first introduced by the Sandstone Aquifer Project
(UNDP, 1968). However this Formation has been recognised by Wolfart (1959) in Irbid
district in the northern part of Jordan as the Chalk Chert Sequence, Sir M. MacDonald
(1965) and the German Geological Mission (1961-1968) recorded the Formation as Chert
Limestone Unit. The Formation crops out in a very limited region, such as in the south
eastern comer of the Amman-Zerqa Basin, the easternmost part of the Wadi Mujib Basin
and in the northern part of the Jafr Basin.
Heimbach (1965) described a type section at Jebel Rijam as 38.5 m of alternating
thin layers of cherts, limestones, marly limestones, and chalks, and conformably overlies
the Muwaqqar Formation, Khdier (1966) recorded a thickness of 41 m in the area east of
Jabel Rijam, and Sir M. MacDonald (1965) recorded a thickness of about 40 m in the
northern part of Jordan. The thickness of the Formation at Jafr Basin is around 50 m.
49
The Formation forms shallow aquifers in the central parts of the Azraq and Jafr
basins, and has been penetrated by numerous wells.
2.3.5 POST EOCENE SEDIMENTS
The Belqa Group is overlain unconformably by a number of sedimentary
formations. Local names have been proposed for these formations, which include the
Plateau Group, the Sirhan Formation, the Azraq Formation and the Jafr Formation. The
Plateau Group comprises sandstone and sandy limestones of Miocene-Pliocene age. The
Sirhan Formation is a sequence of sandy limestone, sandstone and marls of Miocene
Pliocene age, which are correlated by the UNDP (1966) with the Wadi Shallala
Formation. The Azraq Formation consists of limestone, sandy marl and gypsum of
Pleistocene age. The Jafr Formation comprises exclusively fine grained, light coloured,
compact and hard lacustrine limestone, and occurs in the central part of the Jafr Basin.
2.3.6 RECENT DEPOSITS
The recent superficial deposits within the study area are of variable thicknesses
and lithology. They consist mainly of gravels, aeolian sands and playa muds and silts, and
form a mantle that obscures the outcrops of the consolidated sediments. The playa
deposits sometimes contain gypsum and other evaporites in disseminated form. In some
places the wadi alluvial deposits form shallow aquifers and have been penetrated by
numerous wells.
2.4 VOLCANICS
During the late Neogene and Pleistocene, basaltic volcanism was widespread in
Jordan, the most important episodes of which occurred along the Mountain Ridge
Provinces, in central and south Jordan, and within the Plateau Basalt province in north
east Jordan (Figure 2.3). The basaltic rocks along the Mountain Ridge province are
restricted to an area approximately 20 km wide and 110 km long between Jabel Uneiza in
the south and Wadi Mujib in the north. Along the north-west striking Al Karak Wadi el
Fiha fracture zone which extends across the entire Central Plateau for more than 300 km,
50
several isolated basalt flows and basalt dykes are present (Figure 2.11). In Southern
Jordan between Quweira and Mudawwara, there is another striking fracture zone along
which basalt dykes extend for about 70 km through Lower Palaeozoic sandstone. In the
Northern Plateau Basalt province of Jebel Druze, a succession of six lava flows lies
unconformably on the sedimentary succession. These basalt flows range in age from
Oligocene to Recent, the most recent being dated to about 4000 years before present.
These basalts form an important aquifer system in Azraq - Wadi Dhuleil area.
Many rather small areas of basalt were observed by Bender (1968) and others
within the rift zone between the Dead Sea and Lake Tiberias.
2.5 THE GEOLOGICAL STRUCTURE
The geological structure of Jordan is fairly well-known as a result of work by
Bender (1975), Quennell (1956), Burdon (1959), Lloyd (1969), Parker (1970), and others.
However, recent investigations and detailed drilling have assisted in defining more
precisely and have revealed new significant structural trends which were not identified by
the earlier studies.
Regionally, the structure of the study area is affected by the presence of the Nubo
- Arabian Shield and the formation of the Wadi Araba-Jordan Rift. The Wadi Arab
Jordan Rift forms a 360 km long section of the East Africa-North Syria Fault System, a
system is recognisable over a distance of 6000 km. The structural pattern, as seen in
exposures on the east side of the rift, and the morphology of the surface of the
Precambrian Basement Complex suggest that a structural zone of weakness (geosuture)
already existed at the end ofthe Precambrian.
The occurrences of late Proterozoic to Cambrian quartz porphyry volcanism in the
southern Wadi Araba, and the thickness and facies changes in the sedimentary
successions from Cambrian to Lower Tertiary indicate the continued tectonic activity of
the geosuture. However, the Nubo - Arabian Shield in Southern Jordan plunges regionally
to the north and north-east. Epierogenic movements affected the Palaeozoic strata in
Southern Jordan resulting in the gentle regional dip of these strata to the north and north
east. The Palaeozoic formations were in part eroded before the deposition of Lower
51
Cretaceous clastic rocks. Therefore, from west to east in South Jordan, the Lower
Cretaceous rocks overlie, with angular unconformity, progressively younger Palaeozoic
rock units that range in age from Cambrian in the west to Upper Silurian in the east.
The taphrogenic structural movements that initiated the formation of the present
rift apparently occurred along the pre-existing geosuture and started during the late
Eocene - Oligocene. In the late Oligocene - Miocene, the Jordan block was subjected to
uplifting movements resulting in continental erosion and locally continental deposition
of syntectonic conglomerates in some places in the southern part of the rift.
Major taphrogenic movements restarted in the Pliocene - Pleistocene and
continued during several intra - Pleistocene phases associated with the wide - spread
basalt volcanism of the Middle Pleistocene. The post Oligocene taphrogenic structural
movements were mostly of dip - slip type. Only local minor movements of tangential
compression and lateral displacement have been observed. Quennell (1959) and Freund
(1965) believed that major strike - slip displacement had occurred along the rift of the
order of 70 kIn to more than 100 km, but this idea was not supported by Bender (1968).
Taphrogenic movements in the rift strongly affected the area bordering the rift chiefly
along north-west-, north-, and north north-east - striking normal faults, antithetic and
synthetic fault systems, and narrow long grabens and horsts paralleling the rift.
In Central and Southern Jordan the geological structure is characterised by block
faulting reSUlting in broad epeirogenic swells and basins which dominantly strike north
west and west-north-west (e.g. Al Jafr Basin, Bayir - Kilwa Swell, and Azraq - Wadi as
Sirhan Basin). A pattern of approximately north-west, north north-east and east - striking
normal faults and flexures of minor displacements occur in the area. A few small
anticlines in Central and Southern Jordan such as at Jebel at Tahunah north-west of Ma'an
can only be explained by tangential compression.
The pattern of dominant block faults in Central and Southern Jordan gradually
changes northwards into another structural pattern in north Jordan where upwarping and
tilting become a common feature with faulting. However, the relatively thin and
dominantly competent beds in the south reacted to structural stresses by fracturing and
faulting, whereas the thicker and more incompetent beds in the north reacted to the same
52
stresses by arching, tilting and flexuring (e.g. the north-west - striking anticlinal trend of
Jebel Safra south-east of Amman, the uplift of Suweileh north-west of Amman, and the
upwarp of Ajlun).
2.5.1 STRUCTURAL ELEMENTS
Bender (1974) divided Jordan into five structural provinces. The study area is
situated within the Upwarping, Tilting and Block Faulting in East Jordan provinces. The
major structural features of the provinces which apply to the study area are illustrated in
Figure (2.11). They comprise the following:
1. Amman-Zerqa Syncline.
The Amman-Zerqa syncline IS a major structure. Its axis trends southwest
northeast and extends from about 12 km south of Amman and continues to the north
easterly along the Zerqa River through Ain Ghazal, Ruseifa, and Awajan. The northern
limb is located at the north-east of the Upper Amman - Zerqa Basin where it merges into
the Suweileh anticline in the Baqa'a Valley. The southern end of the syncline begins as a
fault striking north-south.
The transition between the syncline and anticline is known as the Amman-Zerqa
Flexure which runs parallel to the syncline axis. Burdon (1959) considered the flexure as
a part of a monocline. However, Masri (1963) considered the flexure as a major syncline
which extends more than 37 km, although, in the flexure zone, fault systems are
observed. MacDonald (1965) stated that "Amman - Zerqa area is bounded by two main
flexures. These flexures, Amman - Zerqa and Suweileh - Salt, bound the area in the south
and west. while to the north and east the area is bounded by positive tectonic areas of
Aj/un Dome and north-eastern block which is open to the east and cover by basalt of the
north-eastern Plateau". The same author also recorded that the Amman - Zerqa trough is
developed as double plunging syncline where greatest depth lies just south of Zerqa.
There are numerous smaller anticlinal - synclinal folds which are generally aligned a long
similar direction.
53
33'
32'
3"
30'
---------~ ----- ---------- -------
36' 37'
3S'
37' 38'
36'
Modified from Bender (1969)
Figure (2.11) The structural pattern of Jordan
EXPLANATION
RIL"alt ,Iikp and .. rru!lion alon .. r.ulla
• • VAjor bJI!,"ll lor lull}
\'okano
)IRjor r ... lt zon~
)hjnr ("ult lJrf, ..... -.I ,. .. ,...,,,It".'I"I" M,tluwl •
.... ., ... " ....... 11"'" rtf tAr" ..
---t--A"iaoelllwpil
.... ~ ...... ./,,.,..r.-", "'''~II/'''''''.t-""'
--t--A:cil or !M"di~ntary bRAin
S,." ... ..,"'r".h_tJ.! pi .. "., ... A,", hto_
AntidiM Synclinf! .~ ... " ... d"".'''''' .If .... It" ... ,.. .",....."" • • of pl ........ "",... i_ .. "" pI ...... IT"'" 1 .. " .....
----+----AntidiM Syndinfl
[)r"",,i"'" Io. ,,.,pt.,,WrI' _,#Md.; ,Mll'ltl. " ....... 11_ ot/ phl"fl*
vlw,..Ir_ .. .,.
_""00-Strudur. rontoW"ll .• pproxim"~
Slrwh''''' """'t",," _ I." ttl ,It. ,...... ........ 6" ...... Dot"" .. ,. ,\(,ddrrmlU'll" S,III ",.,. C~.r i"'"",,,' .... ,;.&1 ... "I ,"rltf"
.. RuiM
Bordn or d.maf('aUon line-
31'
30'
As a result of Amman -Zerqa synclinal structure, the regional dip is generally to
the south-east, e.g. towards the Zerqa River. The regional dip is 5-6 degrees, although
locally minor folds may have much greater angles of dip.
2. Amman-Zerqa Fault Systems.
The main fault systems observed in Amman - Zerqa area are consist chiefly of two
main groups. The major fault trending northeast - southwest parallel to and between the
Amman - Zerqa syncline and anticline. This fault has a displacement ranging between 50-
60 meters. the second group striking northwest - southeast with small vertical
displacement.
Burdon (1959) classified these faults and numerous of minor fault pattern striking
in different direction mainly northeast - southwest to be as tension faults.
3- Siwaqa Fault.
The prominent Siwaqa structure striking west - east extends for more than 60
kilometres between Jebel Siwaqa and the Dead Sea. This fault disappears under the basalt
plateau and change it's strike into west south-west direction west of Jebel Sirhan. The
vertical displacement range from 100meters in Siwaqa area, 120 meters at the western
end, and it's maximum of almost 200 meters 7 kilometres east of Wadi Nukheila. This
fault was later affected by tangential compression which resulted in steep overthrust,
steep anticlines and piecing of older through younger Upper Cretaceous sediments
(Bender, 1968).
4- Karak-Wadi EI Fiha Fault System.
The most prominent fault system in the area trending northwest - southeast and
extending to more than 300 kilometres fro Karak in the north-west to Saudi Arabia in the
south-east. This system consists of series of discrete faults, which are composed of series
of graben - horst structure with vertical displacement of about 100 meters. Bender (1968)
55
believed that this fault is apparently caused by deep tensional forces. In many places,
Pleistocene basalts intrude this fault zone.
There is a group of faults and lineaments in the area to the east of the Karak -
Wadi EI Fiha structure have predominant north-north-west - south-south-east strike.
Many of these faults changes it's strike from near northwest - southeast to south - east (EI
Hiyari, 1985).
5- EI Hasa Fault.
This fault system has westnorthwest - east southeast trend and extends for more
than 50 kilometres. the southern block is downthrown with vertical displacement of up to
1000 meters (Weisman, 1969).
6- Salwan Fault.
The Salwan fault zone striking west - east direction is an extensive, not
continuous structure which is frequently cut by a series of north - south trending discrete
fault systems. The fault is obvious to the north of Shaubak and disappears north of
Husseiniya. To the north-east of the Jafr Basin, it can be traced as "horse tail" fault
aligned westnorthwest - eastsoutheast, which parallel to the axis of the Jafr trough. These
lineament crossing the Karak - Wadi EI Fiha fault zone into Bayir Block. The Maximum
vertical displacement is estimated to be 200 meters in the Jafr trough.
7- Arja-Uweina Flexure / Fault Zone.
These are sub - parallel to the Wadi Arab - Jordan Rift and form small graben and
horst structures. This fault zone extends from the west at Jebel Uneiza in the north to
Jebel EI Batra in the south.
8- Jafr Trough.
The Jafr trough is the most significant geological feature in the Jafr Basin. It's
bounded by two parallel faulting zones with approximate westnorthwest - eastsoutheast
56
strike. The formations in the trough are controlled by the post Palaeozoic sedimentation,
the thickness of the sediments reach its maximum at the centre of the trough where
Muwaqqar Formation (B3) of the Maestrichtian age exceeds a thickness of 450 meters
(Figures 2.8 and 2.10).
9-Bayir-Kilwan Swell.
The only major swell structure in the region extending from Saudi Arabia in the
south northwards to Jebel Safra area to the south-east of Amman. The swell is suggested
by the configuration of the Basement Complex (Figure 2.11 ). Bender (1968) indicated
that this structure is characterised by a decrease in the thickness of the Cretaceous and
Palaeozoic sediments on either side of the swell. It thought that the Bayir - Kilwa swell
was active only in the Palaeozoic.
2.5.2 MINOR STRUCTURE
In the addition to the above major structures, there are other zones of minor
structures which are, however, also of significance, including the undulation feature in
the Amman Formation (B 112) and the minor fault and joint systems associated with the
major structures in the area. Furthermore, toward the south, local structural features are
observed such as the Qihati fault line to the south-east of Amman, the Mazar mound in
the Karak area, the anticlinal structures of Jebel Tahunah 8 km north-west ofMa'an, Jebel
Ruweifi, and Jebel Mutaramil - Sagrat areas, the synclinal structures of EI Lajun and
Jebel Batra areas, and the Sultani - Qatrana graben.
57
CHAPTER THREE
HYDROLOGY
3.1 CLIMATE
Jordan lies within the Mediterranean bioclimatic region of semi-arid to arid type
(Long, 1957). The essential feature of this climate is to have a dry, hot summer and cool
winter.
The climate regime is determined by the interaction of two major atmospheric circulation
patterns. During the winter, the temperature latitude climatic belt prevails and moist cool
air moves eastward from the Mediterranean. In the summer, the subtropical high pressure
belt of dry air causes relatively high temperature and no rainfall.
The climate features of the area can be described by considering north-south and
west-east trends (Figure 3.1). The climate in the northern and western mountainous area
is Mediterranean but, moving eastward, there is a rapid change to semi - arid and arid
types, as the influence of the Mediterranean Sea is replaced by that of the continental land
mass causing a decrease in rainfall and an increase in the ranges oftemperature. Farther to
the southeast in the EI Jafr Basin the climate has been classified as arid (Miller, 1951) or
as a Mediterranean Saharian climate of the warm variety (Long, 1957). Additionally there
is a marked secondary influence of topography upon the climatic parameters throughout
the country.
Rainfall in Jordan is produced by Eastern Europe and Western Mediterranean cold
fronts which are drawn by the Eastern Mediterranean low pressure system (Cyprus Low).
The Cyprus Low, therefore, is a dominant feature of the rainfall production in Jordan. The
incidence and strength of the cold fronts decreases southwards. Moreover, Jordan is
subject to several air-mass lifting mechanisms that commonly act together to bring about
the necessary air cooling required for rainfall to occur. This includes orographic, cyclonic
and convective lifting. The orographic lifting is always in progress over Jordan
throughout the year, and it plays a relatively major part in the rainfall production. The
cyclonic lifting is of mmor importance smce the centre of low
:: ::'\. ::: : : : '\ ::::::::\. : :: :::: ::: :\
~ 0:: <t'
050~--~~~~~~--r----------+------~--~------~~--~~------1
Boy,r::::::: : :::::::::::::: ............... ::::~.
LJEll]S :::::::::::::: ::::::::::::::: ::::::\. :::::::::::::: ::::::::::::::: ::::::::,
000 --------r-.. -.-.. -.-.. -.-.. -.,-.. ~ ... -.-.-.. -.-.. -.-.. -.-.~.-.-.. -.-.-.. -.-.. ~.~~
9 50~-+t-:-t7¥.\
: : : 0 Jofr : : : : : :
•••••••••••••••••••• 0 •••••••••••••••••••
••• "0 •••• 0 ••••••••••••••••••••••••••••••
•••••• •••• •••• ••••••• •••••• 0 •
•••••• •••••••• ••••••• •••• 0 ••• ...... ........ . ..... .... .... . ...... ........ ..... ......... . ••••••••••••••••• 0 •••••••••••
0' ••••••••••••
::::::::: :,;.. .. ....... " . ..... ~ . :::;--
: : : : : :...:..:.;..:: .'-!. • .;..: •• '-7.4:-7.-:-• .;..: •• ..;-:. . : : : : : : : :
LEGEND
I~:)<J SUb-humid
~? :} Semi-arid 30°
900f-------"d' ................ ..... ... ....... . ... ........... . .............. :: :::::::::: :;, :: ::::::::;,.;'
.............. ......... ;..,. ...... :::::;r' :::/ ::1
~ Arid-cool variety
m Arid-warm variety
o Saharian-cool variety : : : :::::;;. :: : ::: ",' 0 50.",. I::: :1 Saharian-warm variety : : : ;/ __ *==,,;,.,.""'==d-....
. 850~----~------~~~~~-------+--------~r---------~----~~·
200 300 . 350 400 45
Modified from Moorman (1959)
Figure (3.1) Distribution of Mediterranean bioclimatic stages in Jordan
pressure ( over Cyprus) is too far from Jordan to be really effective. Convective lifting is
typically present in thunderstorms (lbbitt, 1969).
The cold fronts during winter are in general restricted to the period October to
April or early May, and the greatest activity occurs in December to March. The
distribution of the rainfall is reflected to the orographic effect of the Western Highlands.
The high rainfall zones coincide with the high mountain range east of the Jordan Valley,
followed by a pronounced rain shadow in the lee of the mountains. Thus rainfall
decreases gradually from north to south and rapidly from west to east.
The average wind velocity is strongly affected by elevation, ranging between 4.2
mls in the highlands to less than 2 mls in areas of lower elevation: however, this low land
average excludes the period of sandstorms in the dry season. The sunshine period per day
varies from 7 to 10 hours in the dry season, and from 4 to 6 hours in the wet season.
All the climatic factors such as atmospheric pressure, temperature, relative
humidity and sunshine hours are essentially influenced by the climate macrotic trends;
therefore, the hydrological year has been defined as the period from October to
September, which is then called the water year. All the meteorological and hydrological
data have been then considered according to the water year.
There are more than 270 hydrometeorological stations distributed over the country
maintained by the W AJ, the Ministry of Communication, and the Ministry of Agriculture.
Many of the stations have records over 50 years, while some have been established in
recent years. The location of available stations and observation points in the study area is
shown in subsequent hydrological maps.
3.2 TEMPERATURE AND HUMIDITY
The study area experiences a marked annual and diurnal variation in temperature.
The monthly and annual mean air temperature for a selected stations shows that January
is the coldest and August the hottest month throughout the study area (Table 3.1).
Monthly mean air temperature varies between 5 and 25 °C. Some snowfall is often
recorded between January and March in the highlands.
60
Station Jan. Feb. Mar Apr. May Jun Jul. Aug. Sep. Oct. Nov. Dec. Annual
Amman 8.1 9.0 11.8 16.0 20.7 23.7 25.1 25.6 23.5 20.6 15.3 10.0 17.5 Na'ur 7.0 8.1 10.3 14.2 18.1 21.3 22.7 22.4 2Ll 18.9 13.0 9.2 15.5 Madaba 7.7 9.0 11.3 15.2 19.0 22.2 24.1 23.5 22.3 19.7 14.1 9.7 16.5 W.Wala 10.5 ILl 13.5 18.1 21.3 24.1 25.8 25.3 24.2 21.8 16.5 11.7 18.7 Rabba 7.6 8.9 11.2 15.3 19.5 22. 23.3 22.9 21.5 19.6 14.0 9.6 16.3 Qatrana 8.1 9.2 12.2 16.6 18.1 21.9 23.7 21.3 20.0 20.0 13.9 9.8 16.2 Shaubak 4.7 5.2 8.5 12.7 14.5 18.1 20.5 20.7 17.7 15.3 ILl 6.8 13.0 Udruh 7.5 9.0 12.1 16.3 20.5 23.8 25 25.4 23.1 19.5 13.7 9.3 17.1 Ma'an 5.8 8.2 10.6 14.3 18.9 21.8 24 23 22 17.6 11.6 7.6 15.5 El Jafr 7.9 9.4 13.6 17.7 21.6 25.3 25.9 26.6 24.6 20.3 14.3 9.2 18
Source: Meteorological office records, Hashemite Kingdom of Jordan.
Table(3.1) Mean monthly temperature for selected stations for the period 1937-1985 in (oC).
The mean temperature at each station is affected by its elevation; it has been
shown that the mean temperature falls by 0.6-0.9 °C for each 100 m of increased
elevation (Ministry of Transport, 1966). Consequently, the Jordan Valley experiences the
highest temperature in Jordan with a maximum mean monthly temperature of 31 °C in
August and a minimum of 14 0C in January. Typical diurnal variations are from 24 to 38
0C in August and from 9 to 18 0C in January. Exceptionally frost has been recorded on
the lower slopes of the escarpment but not on the valley floor.
The Western Highlands experience a climate considerably cooler than the Jordan
Valley in the west. At Amman Airport the mean temperature ranges from 8 to 25 °C in
August. The typical diurnal temperature range is from 4 to 12 0C in January and 18 to
33 0C in August. To the south in the Wadi Mujib Basin the monthly mean air temperature
range between 5 to 25 0C. Frost occurs in most years in January and February.
The desert area to the east, being under continental climatic influence, experiences
greater extremes of temperature. Night frosts are more common than elsewhere, mean
minimum January temperature is about 3 °C and mean maximum is 14 0C. In August,
mean extreme temperatures are 20 and 38 °C
Temperature data obtained from several stations in the Jafr Basin in the southeast
show the very wide range in temperature characteristics of the desert area. The mean
monthly values in the coldest month (January) are between 1 and 3 oC, and the maximum
61
mean monthly temperature in July and August lies between 30 and 42 0C according to
location.
Relative humidity varies with location and season and ranges between 30
and 75 %. Annual mean and ranges of relative humidity in the wet season (December -
March) and in the dry season (May - November) for selected stations are shown in Table
(3.2).
Station Wet Season Dry Season Annual
Dier Alla 55-69 30-47 46 Amman 60-70 36-55 51 W.Mujib 60-75 40-55 54 Shaubak 65-74 46-64 58 Udruh 60-69 44-63 55 Maan 53-64 35-56 48 Jafr 51-62 37-59 49 H5 43-57 23-41 38 Source: Meteorological office records, Hashemite Kingdom of Jordan.
Table (3.2) Seasonal ranges and annual mean of relative humidity (%) at selected
stations.
3.3 RAINFALL
Rainfall in Jordan occurs in the wet season which begins in October and ends in
May. During autumn and spring, thunderstonns occur over very limited areas for periods
of about an hour. These thunderstonns tend to occur in succession, therefore sporadic
rainfall occurs for a period of several hours, and exceptionally, for a day or two. During
the winter more widespread stonns occur at intervals of 2 to 3 weeks, each lasting for 2 to
3 days. Occasionally more persistent, continuous and unifonn rain is experienced which
lasts for a period of one day at a time.
Figure (3.2) shows the mean annual rainfall map for the study area for the period
from 1938 - 1985. The map was compiled from data provided by the available rainfall
stations in the study area. A complete record was available for only some of the stations,
the record for many of the stations being either discontinuous or too short. The
discontinuous records have been completed by estimating the missing values by using
least square method which is applied by finding a good linear correlation in monthly total
62
200 r--------------,~~~~~rr--_n~------------.---------------,
100
000
I , , I
I
s
lit I ~ , ~ , s'
£/ I
I , \
I I
I I
I I
I
, I
I I
I ,
_Rallf
~Naqb
-Quwelra
_Ohaba
-Urn Jirnal
, , , ,
-Khan Zabeeb
atti..MuJib / .. .. \ .........
.. '.. )S~aaa
)I~:~ 8 ...
_Ma'an
-.... 300-
LEGEND
Rainfall station
Isohyet of 300 mm
• \ 25 0 25 50
900U' ____________ ~~~--------------~~-----E3----E3~-E3--~----------------J 200 300
Figure (3.2) Mean annual rainfall (mm/a) for the period 1938-1985.
rainfall between the. station with the missing values and the neighbouring stations. For
stations having short records of data, the long term mean annual rainfall (1938-1985) was
estimated by comparing the available records at each station with those of adjacent
stations which have complete records, provided that these stations are in good
correlation. The following formula was used:
where:
LM = LM,(:~) .......................................... : ................................................. (3.1)
LM = computed long term mean annual rainfall (mm)
LMr = long term mean annual rainfall at reference station (mm)
SM = short term mean annual rainfall from available data (mm)
SMr = short term mean annual rainfall at reference station for the same
period as SM (mm).
The distribution of the mean annual rainfall (Figure 3.2) shows that the highest
rainfall zones correspond to the major mountain block of the Western Highlands.
However, the isohyetal lines are almost parallel with the elevation contour lines in the
Western Highlands, and the mean rainfall decreases eastwards to the inland desert. Thus,
the average annual rainfall varies from more than 500 mm in the west to less than 50 mm
in the east. The deeply incised wadis which separate the blocks are marked by narrow,
east-west regions of lower rainfall. The Central Plateau is an area of generally low
rainfall.
The average rainfall decreases rapidly westwards from the escarpment highs into
the Jordan Valley, Dead Sea and Wadi Araba. From the northern end of the Dead Sea
southwards the mean annual rainfall decreases from less than 100 mm to less than
50 mm in most of Wadi Araba. North of the Dead Sea, up the Jordan Valley the rainfall
increases up to 400 mm near Lake Tiberias.
64
The 200 mm isohyet approximates to the eastern limit of the Western Highlands,
and in most of the Central Plateau the mean annual rainfall is less than 100 mm. The
Eastern Plateau and the Eastern Desert have mean annual rainfall of around 50 mm.
In the northern part of the study area, in the Upper Amman-Zerqa Basin, the mean
annual rainfall reaches 600-650 mm to the west of the basin between Suweileh and Salt.
The rainfall decreases rapidly eastward from more than 500 mm near the western
watershed to less than 100 mm east ofZerqa.
Rainfall in the Wadi Mujib Basin is poor, as much ofthe catchment lies in an area
with a mean annual rainfall of 50-150 mm. However, in the western and northern part of
the basin, the mean annual rainfall locally reaches 500 mm. In the Wadi Wala Sub-basin,
in the northern part of the Wadi Mujib Basin (Figure 3.14), the annual rainfall ranges
between less than 100 mm in the southeastern part to more than 500 mm in the
northwestern part with an average of 189 mrnJa. Whilst in the southern part of the Wadi
Mujib Basin, the mean annual rainfall ranges between less than 50 mm in the
southeastern part to more than 350 mm in the western part with an average of 128 mrnJa.
The average annual rainfall for the whole Wadi Mujib Basin is about 130 mrnJa.
In the Wadi Hasa and J afr basins, the isohyetal map shows that the annual rainfall
decreases from 300 mm in the western and northwestern highlands to less than 40 mm in
the east. Rainfall decreases dramatically in the western highlands from 250 mm northwest
ofUdruh to 40 mm at Ma'an, which illustrates the rain shadow of the western highlands.
The average annual rainfall of the Upper Wadi Hasa Basin, the Wadi Jurdhan Catchment,
and the Jafr Basin is estimated at 89, 129 and 52 mm respectively.
The annual rainfall varies widely about the mean. Figures (3.3) and (3.4) show the
annual rainfall variations about the means and the accumulated departure of the annual
rainfall from the mean for selected stations. They indicate that, since 1945 till the end of
the record, the annual rainfall usually varies around the mean over 1 to 5 year cycles, but
it is more common to have drought periods for continuous 3-5 year periods than to have
wet periods of more than three years. It shows also that, in general, a continuous decrease
in the mean annual rainfall for long period up to 20 years is quite common. After the
65
drought period of 1958 to 1963, the southern part of the country experienced relatively
wetter conditions than the north.
In general, the coefficient of variation (Cv) of annual rainfall varies between
0.4 and 0.7, which again indicates that annual rainfall can vary substantially from year to
year. Inter-annual variations appear to be less in the Western Highlands (coefficient of
variation generally less than 0.5 ) but greater in the central and eastern desert areas (0.5-
0.7). Cv increases with decrease of the amount of rainfall. Figure (3.5) shows the relation
between Cv and the mean annual rainfall at selected stations in the study area, together
with data from other stations in the Near East and North Africa. This figure indicates that
the relation between rainfall and Cv in the study area is relatively weak and the Cv value
in the area with less rainfall is higher.
The monthly rainfall varies much more widely than the annual rainfall, especially
in the south and southeast. Figure (3.6) shows the monthly distribution of rainfall for
selected stations. It shows that December to March is, on average, the wettest period in
the rainy season. Monthly maximum values may exceed the monthly mean by four to five
times. The spread of the minimum values is even greater as almost any month may show
little or no rain on some occasions during the measured periods. Traces of rain have been
recorded in the period from June to September.
The probability of occurrence of a series of daily rainfall amounts has been
calculated by Parker (1970) for selected stations in Jordan. The return period is expressed
as the probable number of occurrences of a given daily rainfall in 100 years. The results
of the calculations are shown in Table (3.3) which indicates magnitudes of daily
rainfall which have return periods ranging from one or two years to once or twice each
100 years.
In the Western Highlands a single day having a rainfall of about 20 mm may be
expected to occur once each year, while in the Central Plateau, the same daily rainfall has
a return period of two to three years. High daily intensities are fairly common on the
Plateau and days with rainfall in excess of the annual mean for the station have a return
period of several times each 100 years.
66
900
E 800 .s 700 " -I
600 -I « u. 500 z ~ 400 -I 300 « => 200 z z 100 «
0
900 , ••••••• Naur
E 800 " " ___ Jiza " .s 700 " . ____ fv1azar " , , , , " , ,
-I ' . , " ' . -I 600
, "
LE 500 z ~ 400 -I 300 « => 200 z z 100 «
0
800
E 700 _ •• _. Qatrana
.s •••• Rabba -I
600 ___ Shaubak -I _____ Tafila
" LE 500 z I, f.
~ 400
-I 300 « => 200 z z 100 «
0
140 ___ fv1aan E .s 120 ____ Jafr -I 100 -I « u. 80 z ~ 60 -I « 40 => z z 20 «
0 <Xl 0 N '<I' <0 <Xl 0 N ;1i <0 <Xl 0 N <1i <0 <Xl 0 N '<I' <0 <Xl 0 N ~ C') -a; '<I' -a; '<I' '<I' LO LO LO LO <0 <0 <0 <0 ...... ...... ...... ...... ...... <Xl <Xl O'l O'l O'l O'l O'l O'l O'l O'l O'l O'l O'l O'l O'l O'l O'l O'l O'l O'l O'l O'l O'l O'l
YEAR
Figure (3.3) Annual rainfall variation around the mean at selected stations (the horizontal lines indicate the means).
67
E .s ILl a: :::> I-a: q; a.. ILl D D ILl I-q; -' :::> ::E :::> u u q;
E .s ILl a: :::> I-a: q; a.. ILl D D ILl I-q; -' :::> ::E :::> u u q;
800 JIZA
600 t, ,
400
200
0 .... ,\
\ .r"O\ ' ~ tt\ -: ... ",,: \ ) \ I: , , ....... J., .. \woo .' ,
\ I: , " " Pi,' , -200 " ''\ I JUBBHA '.J , " V:
\ : .. ' " , ... '.. I , I, ... '" " -400 A ...
-600
1400
1200
1000
800
600
400
200 MA'AN
........ ."'_ .... -- ...... -'" _ .... ; -'"" , .. '. 0 - - ............ -
-200 B -400
.... .... .... .... .... .... .... .... .... .... .... .... .... .... co '£ co co co co co co co co co co co co .j:o. .j:o. .j:o. 01 01 01 ~ ~ 0> -...j -...j -...j -...j
0 W 0> co I\) 01 (Xl -...j 0 W 0> co
YEAR
Figure (3.4 A,B) Accumulated departure from the mean annual rainfall for selected stations
68
.... .... co co (Xl (Xl I\) 01
200
180 ...... 160 ;,R ~ > 140 u c: 0 120 :; .~
> 100
'0 80 1: Q) 60 ·0 IE Q) 40 0 u
20
0
Station
Irbid
Amman
Shaubak
Rabba
Maan
Jafr
Azraq
H5
H4
0 100
o Stations in the Near East & North Africa
y = 834.4x·O.5791
Shaubak Ka.rak Naur
BJ.Xi"O~O(t°~·2&I~o~d=¥i:~~~~~~.JJerUSalem 00
200 300 400 500 Mean Annual Rainfall (nm)
600
Figure (3.5) Coefficient of variation of annual rainfall as a function of mean annual rainfall for selected stations in the Near East and
North Africa (Modified from FAD Irrigation and drainage Paper 37).
Daily Rainfall Amount in (mm)
110 100 90 80 70 60 50 40 30 20 10
1.5 2.9 5.3 10.0 18.2 30.3 50.0 72.9 92.6
1.2 2.5 5.4 11.4 23.8 43.4 69.4 93.5
1.3 2.9 5.9 11.8 23.8 43.5 71.4 90.0
1.4 2.8 5.7 10.9 20.8 36.4 58.8 84.7 98.8
1.5 4.4 12.5 33.3 70.4
1.1 3.4 10.5 28.6 65.4
1.5 4.4 12.5 33.3 70.7
1.8 5.3 14.7 37.0 74.1
1.1 3.0 8.3 21.3 47.6 84.0
Source: Parker (1970).
5
88.5
84.7
88.5
90.9
95.2
Table (3.3) Probability of various daily rainfall amounts in 100 years at selected
weather stations.
69
350
300
E 250 .§.
200 ...J
<i! 150 li-z
~ 100
50
0
180 160
E 140
.§. 120
...J 100
...J
It 80 z 60 ~ 40
20 0
250
200
E .§. 150 ...J
<i! li- 100 z ~
50
0
300
250
E 200 .§. ...J 150 <i! li-Z 100 ~
50
0
. Mn JLEB-IA
.Ave Olvlax
nJ J .I .. 1 n o N D J F M A M
MJNTH
. Mn SLJ<HNA
.Ave Olvlax
lJl I J J J 11 o N D J F M A M
M:)NTH
. Mn SAHAB
.Ave Olvlax
JlI I ,J J I I
ON D J F MA M
MJNTH
. Mn W.WALA
.Ave Olvlax
Jl. I I I I I ~
o N D J F M A M
M:)NTH
Figure(3.6) Monthly rainfall at selected stations
E .§. ...J
~ z ~
E .§. ...J
<i! li-Z
~
E .§. ...J
<i! li-Z
~
E .§. ...J
~ z
~
300 .....
250 . Mn A~NAIRFORT .Ave Olvlax
200
150
100
50
0 JlI I ,J J J il ONDJ FMAM
M:)NTH
400 . Mn NALR 359 .Ave 300 oWa 250
200
150
100
50 _
0 0 N D J F M A M
MJNTH
180
160 . Mn JIZA .Ave
140 Olvlax 120
100
80
60
40
20
0 0 N D J F M A M
M:)NTH
160 140 _ . Mn DHABA
.Ave 120 _ Olvlax 100
80
60 40
20
0 0 N D J F M A M
MJNTH
120 _
100 ElMn .Ave
E OWex
E 80
:r -' 60 <t: u-z ~
40
20
0 a N 0
450
400
350
III Mn .Ave o Wax
E 300 .s -' 250 ;a! 200 u-z 150 ~ 100
50
0 -n Jl I a N 0
80
70 li Mn . Ave
60 o Wax E E :r
50 _
;a! 40 _ u-z 30 ~ 20
10
0 a N 0
140
120
E 100
liMn .Ave o Wax
.s 80 -'
-' <t: 60 u.. z ~ 40
20
0 - il I
a N o
Figure (3,6) Continued
SrNAOA
J F M A M
MONTH
Mll.ZAR
I I I I n
F M A M
MONTH
HASA
F M A M
MJNTH
SHA.L8AK
I jl F M A M
MJNTH
400
350
E 300
IIIMn KARAK - .Ave o Wax
.s 250 -' -' 200 ii z 150
~ 100
50
0 ] ~ I I I I • -ONDJ FMAM
MONTH
120
100 .Mn OAlRANA .Ave
E 80 .s oWe
-' -' 60 <t: u-z 40 ~
20
0 11 n I I I I II n aND F M A M
MONTH
200 180 - .Mn TARLA
160
E 140
_ .Ave OWex
.s 120 -' 100 ;a! u- 80 z ~ 60
40 20 0 :Jl Jl I I ...,
a N 0 J F M A M
MONTH
50 45 .Mn MII.'AN 40 _ .Ave
E o Wax
35 _ .s 30 _ -' 25 ~ z 20
~ 15 10 5 __
0 a N 0 J F M A M
MONTH
3.4 EVAPORATION
Evaporation is the direct transfer of water from the ground to the atmosphere,
through the phase change of water from liquid to vapour. Evaporation of water from plant
surfaces that has traversed from the soil through the plant is termed transpiration. All
evaporation from plant surfaces is not transpiration since intercepted water by vegetation
before it reaches the ground is also evaporated. The combined evaporation from the
surface of the ground and the transpiration from the vegetation is termed
evapotranspiration (ET) to represent the net water loss from the total ground surface. ET
generally involves a large fraction of the total rainfall. In arid climates most of the
rainfall, 90% or more, may be lost through ET. In more humid climates, ET may account
for 40-70% of the annual rainfall. ET is more complex than evaporation since plant and
soil factors affect the process. The rate at which ET occurs from well-watered, actively
growing, completely vegetated surfaces is termed potential evapotranspiration (PET).
The definition of PET is based on a continuous water supply, with the relationship
between ET and PET being dependent mainly on soil moisture content. Such relationship
fails to take into account the existence of stomatal controls of transpiration that do not
depend on soil moisture content and of existence of a feedback mechanism whereby
changes in ET alter the temperature and humidity of the overpassing air which in turn
changes PET.
Instead of using PET as a causal agent for estimating ET, Morton (1985) used
PET as an effect of changes in ET caused by changes in the availability of water for
evaporation from a large area (Figure 3.7). ET increases from zero when there is no water
available for evaporation from the surrounding area to a constant rate of wet
environmental areal evapotranspiration (Ew) when there are no limitations on the
availability of water. In contrast, PET decreases from 2Ew, when ET=O and the air is hot
and dry, down to a constant rate of Ew, when ET=Ew and the air is cool and humid. Thus
Morton postulates that PET, in steady state conditions, is a negative index ofET and that
changes in ET and PET might be equal and opposite (Nash, 1989).
72
PEd (2Ew ) = Dry environment potential evapotranspiration
+-PE = Potential evapotranspiration
Ew = Wet environmental areal evapotranspiration
+-Et = Areal evapotranspiration
OL--------------------. Water supply to soil-plant surface of area
Figure (3.7) Morton's (1985) model of evaporation.
Evaporation is a very important component of the hydrologic cycle. Reliable
information on evaporation losses is required in water balance studies and in planning,
designing and operation of existing or proposed reservoirs and irrigation projects.
The problem of estimating actual or potential evaporation and evapotranspiration
losses is particularly difficult. These losses could be estimated directly by using the pans
and lysimeters, or indirectly by using the different formulas available, or a combination of
both. For example, one approach to estimating PET is to set it equal to a coefficient times
pan evaporation, the coefficient used may range from 0.6-1 (Haan et.al, 1994).
Theoretical approaches are based on energy budgets, mass transfer relationships, or a
combination ofthese approaches.
For evaporation to take place water must be present and energy must be available
to provide the latent heat of evaporation and the increase in temperature needed to bring
the water to the evaporation point. In a semi-arid country, like Jordan, the presence of
water must be stressed. In the early stage of the rainy period each year, when there is very
little pre-existing soil water available, the amount of evaporation will be determined
73
entirely by the amount of rainfall. It is also important that these two requirements are
present simultaneously. Also the presence of a layer of dried soil between the energy
source (atmosphere) and a deeper soil layer containing water will provide some resistance
to evaporation and reduce its rate and amount.
For regional studies, experience has shown that, in determining potential
evaporation, the methods embracing the measurement of the climatic elements involved
and the assessment of their rational relationships, have proved far more reliable than
direct evaporation measurements, or indirect empirical methods involving the limited use
of only certain specified climatic elements.
MacDonald (1965) and Lloyd (1966) have shown that the semi-empirical
combination method of Penman proved to be the most reliable in climatic conditions
similar to those in Jordan. Therefore, Penman's method has been employed by NRA and
W AJ in Jordan to determine open water evaporation and potential evapotranspiration
from other surfaces. The method is based on a combination of aerodynamic and energy
balance approaches and requires meteorological observations of air temperature, relative
humidity, solar radiation and wind velocity.
Class-A Pans are used in all of the evaporation stations in Jordan since the 1960s.
They provide data on open water evaporation (Eo) for more than 20 stations distributed all
over the country. Attempts have been made to study the relationship between the monthly
values of Eo and PET (Figure 3.8). The relationship found to be linear, and it does not
seem to vary very much from year to year, but varies within the year with higher values in
cold and wet months and lower values in hot and dry months. It also shows some
variation between stations in different locations. Thus the relationship may only appertain
to the station of derivation, although stations established in similar climatic zones may
have a similar relationship.
For the purposes of this study three relationships between Eo and PET were
obtained; for the wet period ( October-April), for the dry period (May -September), and
for the whole year. The linear regressions as listed in Table (3.4) are in the usual form of
Y = aX + b , where Y and X refer to PET and Eo, respectively. The regressions appear
reliable under average climatic conditions but do not hold for exceptional conditions
74
whereby abnormally high temperature and windspeeds give high and erratic results. Such
correlations will facilitate the use of incomplete records for the determination of PET.
Period Linear Regression R~ Corr. Coef. st. Dev.
Wet Y = 0.5822(X) + 13.023 0.7313 0.86 51.04
Dry Y = 0.5221(X) + 27.342 0.5824 0.76 84.58
Annual Y = 0.5537(X) + 17.046 0.8553 0.93 99.70
Table (3.4) Relationship between Eo and PET.
The great climatic and topographic variation in Jordan results in a wide variation
In evaporation. The mean annual values of open water evaporation and potential
evapotranspiration for selected stations are shown in Table (3.5). The mean annual Class
A Pan evaporation ranges between 1768 and 4186 mm, which is more than 3-10 times as
high as annual rainfall. The mean annual potential evaporation ranges between 1153 and
2490 mm. Figure (3.9) shows the monthly potential evapotranspiration at selected stations
in the study area. As would be expected, for all the stations, about 70 % of the annual
evaporation is recorded between April and September during the hottest months of the
year. During the period when rainfall may be expected ( October-April ), the average
monthly PET ranges between 26-248 mm.
The areal distribution map of annual potential evaporation (Figure 3.10) shows
that the distributions are almost parallel to the isohyets; however, PET increases in the
opposite direction to the rainfall, reflecting mainly the variation in temperature.
As discussed above, evapotranspiration is the net result of various climatic
factors. Thus, comparisons between these factors and the evapotranspiration estimates for
the various stations are quite difficult. However, the general feature of the
evapotranspiration estimates is that, for most of the stations, regardless of the
location and the climatic regime prevailing at the station, there is no great
75
E .§. I-w a..
E .§. I-w a..
E .§. I-w a..
Dry Period
350
300 Y = 0.5221 x + 27.342 R2 =0.5824
250
200
150
100
50
0 0 100 200 300 400
Eo (mn)
Dry Period 350
300 Y = 0.5221 x + 27.342 R2 =0.5824
250
200
150
100
50
0 0 100 200 300 400
Eo(mn)
Annual 350
300 Y = 0.5537x + 17.046 R2 = 0.8553
250
200
150
100
50
0 0 100 200 300 400
Eo (mn)
Figure (3.8) Monthly Potential evapotranspiration (PET) vs. monthly Class-A-Pan evaporation (Eo)
76
500
500
•
500
Station Type Jan Feb Mar Apr May Jun Ju1 Aug Sep Oct Nov
Amman Eo 86 98 153 207 299 350 370 341 278 222 136
PET 55 72 111 146 193 229 240 211 153 109 67
Baqa'a Eo 71 94 140 205 279 336 362 335 272 207 127
PET 58 75 107 148 173 196 203 193 151 108 67
Zeituneh Eo 59 77 130 163 228 260 270 260 210 140 80
PET 49 65 113 128 180 208 218 205 165 105 63
Dhaba PET 62 70 109 165 225 270 285 270 225 150 105
Rabba PET 35 49 88 121 - - - - - 90 54
Qatrana PET 62 84 124 210 240 255 270 240 210 150 105
Udruh Eo 121 137 198 287 390 488 488 503 391 287 159
PET ~ 2J ill l8.Q UQ m ill ru UQ ill 1M
Shaubak Eo 65 63 98 153 200 227 250 227 199 144 86
PET 26 48 84 103 ill ill. ill ill ill ill 49
Tafila PET 58 64 101 119 - - - - - - 49
Abur Eo 91 89 132 193 261 288 329 303 255 198 118
PET 22 ~ 2Q ill ill ill 12.8. ill lQQ ill Rl
Hasa Eo 103 134 190 276 377 455 488 451 346 236 149
PET 33 60 95 108 ill ~ lli ill 2.Q1 UQ 47
Ma'an Eo 117 138 210 315 390 370 377 425 284 275 143
PET 45 58 99 155 173 193 213 227 153 108 61
Jafr Eo 180 213 285 405 515 612 741 632 246 140 120
PET ill ill ill ill ill ill ill ll2 ill 21 ~
Source: Meteorological office records, Hashemite Kmgdom of Jordan.
NB.: Underlined values are calculatedfrom the relations between Eo and PET.
Table (3.5) Mean monthly values for the Class-A Pan (Eo) and potential
evapotranspiration (PET) for selected stations in (mm).
Dec Ann.
87 2626
52 1683
81 2506
55 1534
58 1935
48 1547
60 2011
36 -60 2010
115 3564
au lli2
55 1768
33 1153
33 -81 2338
@ lill.
103 3306
32 1761
107 3150
40 1525
98 4186
1Q lliQ
difference between the evapotranspiration estimates at the different stations during
winter. Greater differences occur in summer, when the air temperature plays an even
more significant role in detennining evapotranspiration, and as the other factors (e.g. the
humidity and the wind velocity) become similar all over the country. For example, during
winter, the evaporation at Amman is higher than at Maan and Rasa despite the fact that
the latter stations are located in the southern desert and consequently experience higher
air temperatures. This could be explained by the high wind velocity at Amman during
winter.
77
250 300
200 Amman 250 Qatrana
E150 i' 200
.s e 150 tu 100
I-
~ 100 c.. 50 50
0 0
0 c u. « ""') « 0 c u. « ""') «
200 250 Shaubak 200
150 Maan E E150 .5.100 .5. I- tu 100 w c.. c..
I 50 50
0 0 0 c u. « ""') « 0 c u. « ""') «
Month Month
Figure (3.9) Monthly potential evapotranspiration (PET) at selected stations.
Even in similar regions, for example the relative difference in winter
evapotranspiration between Hasa and Ma'an, could be related to the location of the Hasa
Station within an irrigated area. For the stations in the Western Highlands, the Tafila
evapotranspiration is higher than those for Shaubak and Rabba, demonstrating the
variable nature of climatic conditions close to the edge of the escarpment.
In most of the study area, moisture supply to the evaporating surface is very
limited by the shortage of rainfall, so actual evaporation from the area is much less than
potential evaporation.
3.5 RUNOFF
Runoff is the process by which water, in addition to the base flow, reaches streams
by travelling through the soil surface or falling directly into the stream channels. It results
from the excess rainfall after abstraction. And it depends on the climate and on the basin
drainage characteristics.
78
,-IL 7 L.~''''''''''
Sarna{' r------ -"SYRIA ./ ./
'" ./
• ./ (1856) ./ -.- . ./
150
V , ./
0 f-o--- Urn Jimal .""'1'
.--\ . . ""-,-,/' I DierAlla L-ZJ7) H5
~ ·k~·-8aqa'a, • (1983) ~ ('1391) • ~ (1534) I 32° " 32°_ «tj Amman Air Port Azraq
I . • ~f • I I (1683) -.-
Zeituneh (1942) -'
J I • Dhaba -.- . .",.,........ .
-' (1547) ... -' -' w.way-i2011)........." .~.
<t' ""
, '" -/' ctJ·-
~~~~~. \ <t , c:
• Qatrana
" <t'
~ " .. __ I· (2Q1,~) -, ~ " ~31°
~ ,
~ 31°
i Abur Hasa '.~ , I • , , , (1498) (1761} i
I 1\ " , I , Shaubak ! \
f---- I. I , - '. I > (1153) I
I .;'
Udruh r Jafr ./.'/'
I 1" , .. , • ~2159)Ma·an I (2490) I
I • I (1525)
i I , I
20
100
050
000
950
I .~ . ',0" I ..".,,.,,,.. , 30°
! .-' ~.-' ! 'r'-I • Evaporation s~tion
I --Disi ,
(1683) PET (mm) I / • ')(,)("1
/ (2168)
I /
/'
'" 0 50 Kms
t--._._. '" I I I ! I , t-._. '" \
--.-.... -36' 37· I 38·
I
85(
200 250 300 350 400
Figure (3.10) Mean annual potential evapotranspiration (PET).
Abstraction from rainfall are losses from rainfall that do not show up as storm
runoff. The volume of runoff thereafter, is the volume of rainfall minus the volume of
abstraction which is also known as excess rainfall or effective rainfall. Abstraction
include interception, evapotranspiration, surface storage and surface detention, bank
storage, and infiltration.
INTERCEPTION
Interception is the amount of rainfall that is intercepted by vegetation before it
reaches the ground. It varies with the type, density and stage of growth of the vegetation,
intensity of the rainfall, and wind speeds. On an annual basis, interception may involve a
significant percentage of the total rainfall especially in dense forest (Dunne and Leopold,
1978). In a semi-arid environment with less vegetation cover such as Jordan, the amount
of interception is less significant. Thus for the purposes of this study, since the amount of
rainfall satisfying interception storage is generally a small percentage of the total storm
rainfall, interception assumed negligible. However, it is included in other hydrological
cycle parameters, as intercepted water, eventually evaporated or may later fall to the
ground.
EVAPOTRANSPIRATION
Evapotranspiration as discussed earlier, is a very important component of the
hydrological cycle. It involves a large fraction ofthe total annual rainfall, as much as 90%
or more in semi-arid and arid climates. In spite of the high total fraction of rainfall
involved in the evaporation process on an annual basis, for individual rainstorms it is less
significant. Infiltration, as it will be discussed later, is the more significant component of
abstraction during storm events: eventually infiltrated water may be evapotranspirated
after the rainstorm ceases, but at a relatively slower rate. Therefore abstraction by
evapotranspiration is more significant during the times between rainstorm events, not
during the storm events themselves. A higher percentage, however, is expected in the
urban part of a basin where it is covered by concrete, asphalt, roof, etc.
80
BANK STORAGE
Bank storage represents losses from streamflow into the bank of the stream. As
this water seeps back into the stream in later stages, thus it is not actually a loss from
runoffbut a storage and delay in the runoff process.
SURFACE STORAGE AND DETENTION
Surface storage is the volume of water required to fill depressions and other
surface storages before surface runoff begins. Detention storage is the build-up of small
depths of water required to support the runoff process. Some workers define the surface
storage to include the interception. Actual measurements of surface storage and detention
are extremely difficult to make and consequently are practically non-existent. Many
authors (Wright-McLaughlin Engineers, 1969, Terstriep and Stall, 1974, Linsley et at.,
1949, Tholin and Keifer; 1960, and Viessman, 1967) propose different values for surface
storage and detention depending on the size and slope of depression, the land cover, and
the duration of the rainfall event. However, it is recognisable that a watershed surface is
made up depressions of various sizes and that as some of the smaller depressions are
filled, surface runoff can begin even though the larger depressions are still filling.
Furthermore, for long duration rainfall, the values of surface storage will not appreciably
affect estimated runoff rates since the early part of the storm would fill this storage prior
to the occurrence of major runoff producing part of the rainfall. As might be expected,
surface storage is of greater importance on flat surfaces than on steep surfaces. For the
purposes of this study as far as long duration rainfall is concerned, surface storage and
detention, which are believed ultimately to evapotranspirate or infiltrate, are grouped into
and considered to be analogous with soil moisture deficit as far as the soil capacity to
hold water is concerned.
INFILTRATION
The major abstraction from rainfall during significant runoff-producing storms is
infiltration of water into pervious soil. The processes of infiltration of water and
subsequent water movement is exceedingly complex. In general, the infiltration rate is
81
dependent on soil physical properties, vegetation cover, antecedent soil water conditions,
rainfall intensity, and the slope ofthe infiltration surface.
Soil has a finite capacity to absorb water. The infiltration rate for any soil depends
on the permeability of that soil; it is not uncommon to find a soil with variable
permeability with depth, which results in great alteration in the pattern of infiltration. The
infiltration capacity varies not only from soil to soil but is also different for dry versus
moist condition in the same soil. If a soil is initially dry, the infiltration capacity is high.
Surface effects between the soil particles and the water exert a tension that draws the
moisture downward into the soil through labyrinthine capillary passages. These capillary
forces decrease with increased soil moisture content causing a drop in the infiltration
capacity. Eventually, the infiltration capacity reaches a more or less constant, or
equilibrium value, after which the runoff begins. Thus rain falling on a wet soil will
produce more runoff at a higher rate than the same rain on a dry soil.
Bare soils tend to have lower infiltration rates than soil protected by a vegetation
cover, since the impacting rain drops breaks down soil aggregates and small particles are
carried into the soil pores. The net result is a lowering of the infiltration rate.
Light rainfalls are easily absorbed, but heavier rains soon saturate the surface soil
layer, consequently decreases the infiltration capacity, which results in runoff rates being
near the rainfall rates. Furthermore, high intensity rains are more effective in sealing the
soil surface as they have more energy to breaks down the soil aggregates.
The time available for infiltration is a function of the slope of the infiltration
surface. On a steep slope, the water tends to run off rapidly and thus it takes less
opportunity for infiltration than on a gentler slope. Moreover, the soil type and thickness
found on steeper slopes is generally not the same as on flatter slopes.
In Jordan, observation shows that the first rainfall events in the early part of the
ramy season produce relatively higher runoff than the later rainfall. This could be
attributed to the sealing process produced by the break down of the soil aggregates or the
precipitation of secondary minerals in the pores of the soil as these water ultimately leave
the soil profile by evaporation. Precipitation of secondary minerals probably takes place
earlier during the previous dry season. This phenomenon is more severe and lasts for
82
longer in the Plateau and south-eastern desert where the potential evapotranspiration is
high.
The combination of all the factors governmg the infiltration throughout a
watershed interact in such a fashion as to result in a very complex spatial and temporal
distribution of infiltration. At some locations, the infiltration capacity may be so high as
to practically never produce surface runoff, whereas other areas may have low infiltration
capacities and produce surface flow from light rainfalls.
A great deal of effort has been expended in developing the mathematical theory of
the infiltration of water into soils and subsequent movement of this water within the soil.
Theoretically based equations, such those based on the continuity equation and Darcy's
Law applied to unsaturated flow, have been found to be difficult to use and of limited
application. Therefore, a great many empirical relationships have been proposed. The
formulas and their derivations need not be considered here. Horton (1940) introduced an
equation which fitted the experimental data on decreasing infiltration rates as a function
of time. The difficulty with this equation is that it does not account for variations in
rainfall intensity and thus has no provision for a recovery of infiltration capacity during
periods of low or no rainfall. Holtan (1961) has advanced an empirical infiltration
equation based on the concept that the infiltration rate is proportional to the unfilled
capacity of the soil to hold water. It has the advantage over the Horton equation in that it
has a more physical basis and can describe infiltration and the recovery of infiltration
capacity during periods of low or no rainfall.
Over the years many other empirical infiltration models have been proposed.
Because of the general lack of values for the parameters for these various models and the
nonhomogeneity of soils, these models have not been widely applied. Instead, a steady
infiltration loss rate from the rainfall rate has been defined to obtain the effective rainfall
rate: this infiltration loss rate is equivalent to the storm water runoff rate. Wright
McLaughlin Engineers (1969) proposed values of different constant infiltration loss rates
for different storm frequencies based on field tested. Often the constant infiltration loss
rate is termed the <l> index.
83
CURVE NUMBER APPROACH
The Soil Conservation Services (SCS) of the US Department of Agriculture
(1972, 1985) introduced a fonnula which combines infiltration losses with initial
abstractions and estimates rainfall excess or equivalently the runoffvolume:
where:
(p -O.2SY Q=~--<--
P+0.8S for P ~ 0.2S ................................... (3.2)
25400 S = - 254 .................................................................................... (3.3)
CN
Ia = 0.2S ................................................................................................ (3.4)
Q = the accumulated runoff volume or rainfall excess (mm)
P = the accumulated rainfall (mm)
S = maximum soil water retention (mm)
CN = curve number
Ia = initial abstraction (mm)
The SCS has classified more than 4000 soils into four hydrologic soil groups
(HSG) according to their minimum infiltration rate obtained for bare soil after prolonged
wetting. The four HSG are denoted by the letters A, B, C, and D (Appendix Bl).Choice
of the HSG can also be made based on the texture of the exposed surface soil as shown in
Table (3.6) (Brakensiek and Rawls, 1983).
The eN values were assigned by plotting observed runoff versus measured
rainfall for a number of experimental plots scattered throughout the USA (Figure 3.11).
The CN values were then correlated with the land use. Appendix (B.2) gives a summary
of CN values for various land-use and treatment combinations. The curve number of an
area indicates the runoff potential of the area.
84
HSG Soil texture
A Sand, loamy sand, or sandy loam B Silt loam or loam C Sandy clay loam D Clay loam, silty clay loam, sandy clay, silty clay, or clay
After Brakenslek and Rawls (1983)
Table (3.6) Classification of the hydrological soil groups (H8G)
Recognising that abstractions from rainfall depend on the antecedent soil moisture
conditions that exist at the time a rainstorm occurs, three antecedent conditions have been
defined. The curve numbers given in Appendix (B.2) are for antecedent conditions of
type II, which are based on the median values for CN taken from sample rainfall and
runoff data. Antecedent condition type I is used when there has been little rainfall
preceding the rainfall in question, while condition III is used where there has been
considerable rainfall prior to the rain in question. Curve numbers for antecedent
conditions I or III can be estimated by (Chow et a/., 1988) :
4.2CN(II) CN(I)= 10-0.058CN(II) ..................................................................... (3.5)
23CN(II) CN(III)= 10 + O.l3CN(II) ..................................................................... (3.6)
where CN(I) , CN(II) , and CN(III) represent curve numbers for antecedent conditions I,
II, III, respectively.
Once the CN is obtained, Eq. (3.2) and Eq. (3.3) can be used to estimate the
accumulated rainfall excess as a function of total accumulated rainfall. Equation (3.2)
indicates that P must exceed 0.28 before any runoff is generated. Thus a rainfall volume
of 0.28 must fall before runoff is initiated.
85
8~mrnffi.ffi~mmffiffi~mmmm~ C u rv e 5 0 nth i 5 5 h ee tare for -t-1-+-t--T--f-if-,l<-+-t-FIH-:.A-H4H4-I-+..H-H.4-++-l
i- tile ca5e la = 0.2S, 50 that (p-02S)2 .!4-1Yl IXI I:YI U4 I I,..f I 1)4 I 1..vI I I
0= , P+O 8S I TTTT ~ 17fT17 . 17 D 1/ ~ po' C7
, , ,~.£i i/ 1/1 I.' -,t:f-o,'/'
",~,i. it"'J~~v v . ,,'<> -1/
I ('':;/j 'b""
~";'/1 ./ ~ I ~ ~ ~
. ~ V Ill/ "1-rrr-,,\~ ~
I
1=+ 6
If)
~ 5 u .S -Q.4 :t= 0 C :J l-
t) Q) I-
i:5 z
~
o z 3 4 5 6 7 8 9 10 11 12
Rainfall (P), inches
Figure (3.11) Graphical solution of the SCS Curve Number Method
It should be noted that the CN approach is a runoff approach and not an
infiltration approach. Using it as an infiltration approach can lead to errors. Certainly
infiltration is a factor affecting runoff, but so is quick return flow and initial abstraction.
Combining the CN approach with infiltration approaches such as minimum retention
rates carries the CN concept beyond its original intent and beyond the data on which the
CN values are based. Since the derivation of curve numbers includes factors in addition
to infiltration, it is, in fact, a non-Hortonian approach to runoff estimation.
APPLICATION OF THE CN METHOD
The CN method was applied during this study to estimate the volume of surface
runoff in the study area. The area was subdivided into different sub-catchments, each
covering main wadis and tributaries having the same hydrological characteristics. Storm
events and stream flow data for the different sub-catchments were analysed for
comparison and to find the relation between rainfall and runoff. A storm event is defined
here as the rainfall volume that occurs in a 24 hour period. All the storm events occurring
in the period 1980-1985 were used. Rainfall distribution and the volume of rainfall in
each of the sub-catchments were estimated by using the Theissen polygon technique. This
record is believed to be representative as it comprises two years of high rainfall above the
average (1981,1983), two years of low rainfall below the average (1982, 1984), and one
year of rainfall approximately close to the average rainfall in the area (1985).
The soil cover and vegetation were discussed in Chapter One. In determining the
hydrological soil groups, and for the purpose of runoff and infiltration estimates, in
addition to the soil texture, consideration was given to the soil thickness and the influence
of the rock type on the soil characteristics, especially in areas were the soil cover is thin.
High infiltration capacity gravel and alluvial fan areas were found along the major wadis
in the area. Impervious mudflat covers the bottom of some topographic depressions in the
eastern and south-eastern desert such as Qa Hafira, Qa Jinz, and Qa Jafr. Depending on
these elements, three hydrologic soil groups were assigned in the study area; Group A, B,
andC.
87
The vegetation cover is limited due to the low rainfall, deforestation, and thin soil
cover. Scrubs and trees are found only at a few places in the Western Highlands, and
during spring the highlands carry a light cover of grass. Seasonal cultivation is limited to
those places in the Western Highlands where the soil cover and the annual rainfall are
favourable.
The curve number of an area indicates the runoff potential of the area, which
depends on many factors as discussed above. Accordingly the CN values were assigned
for each sub-catchment. These ranges between CN70 - CN90. The only assumption
which has been introduced in choosing the CN value is that for areas with steep
topography a high curve number value is chosen, since these areas have less infiltration
capacity and high runoff potential. Furthermore, thin soil cover is always found along the
escarpments and in areas with steep topography.
The antecedent moisture condition is chosen to be CN(n) which is defined as
"the average case for annual floods, that is, an average of the conditions which have
preceded the occurrence of the maximum annual flood on numerous water sheds" (US
Department of Agriculture, 1972, 1985).
The results are discussed in detail in the following sections. The volume of
rainfall and runoff and the runoff coefficient results are obtained as the average of the
mean values for the sub-catchments.
3.5.1 RUNOFF IN JORDAN
The fault escarpment on the eastern side of the Jordan Valley-Wadi Araba graben
is breached by a number of westward draining wadis in the zone between the Syrian
border and the southern end of the Dead Sea (Figure 3.12). The largest of these are the
Yarmouk Valley, the Wadi Zerqa, the Wadi Mujib, and the Wadi Rasa. More than half of
the country drains westward directly to the Dead Sea or to the Jordan River Valley and
thence into the Dead Sea by these wadis. The headwaters of these drainage systems
extend eastwards into the Plateau. The Yarmouk River Valley catchment is mostly in
Syria: only a small part lies within Jordan, and this drains the extreme north-west part of
the country. The Wadi Zerqa drains much of the northern part of the area. Its most
88
100
000
900
• • • , , • , • ,
'-
~ ~ c:
" , Ya- /-
,,""0'!kl - - '" -., :..~~. ~j.,
......• _, ,' .... e,. .. -, \., : \ ----t ""-, , . " Yarrflouk Basm ,
!
-E Jo-...--
. • •
• •
° ...,
• tb' :-Q~ : ~ 0 . ~~
" ~ t: . u· ..... • ~tI) '::ii;tb' '. ll:j
" .. ... i
J
.~ .. ' t
;~
Wadi Araba Basin South
200
Jafr Basin
Azraq Basin
300
Figure(3.12) Surface water catchments in Jordan.
, , , , , , , , ,
- -
N
W-\rE s
400
--
easterly headwaters extend into southern Syria. The Wadis Mujib and Hasa drain a large
part of the high rainfall zone of the central part of East Jordan and discharge directly to
the Dead Sea. In addition, a large number of side-wadi catchments, eroded into the rift
escarpment, drain westward to the Jordan Valley and Dead Sea. To the south of the Dead
Sea, along the escarpment to Wadi Araba, many wadis breach the western highlands to
the east of the escarpment and drain westward to wadi Araba and thence to the Dead Sea
or, in the extreme south of Wadi Araba, to the Red Sea. Many of these westward draining
wadis have cut down to intersect the saturation zones of the aquifer systems underlying
the Western Highlands and the Plateau. Thus a perennial flow is maintained by spring
discharges along many ofthese wadis.
The rest of the country, mostly desert, is drained by five major closed catchments
(Figure 3.12): the Azraq Basin, the Jafr Basin, the Wadi Hamad Basin, the Wadi Sirhan
Basin, and the Southern Desert Basin. The greater part of these internal basins lie within a
rainfall zone of less than 100 mmla. Only the Wadi Jurdhan sub-catchment in the western
part of the Jafr Basin, lies in the 50-250 mmla rainfall zone.
For the purposes of a country-wide review and assessment of the surface water
resources under the scope of this study, the drainage areas of Jordan can be classified into
five groups. Each belongs to catchments having similar topography, vegetation cover, and
meteorology:
GROUPA:
Group A includes the large catchments which stretch between the Central Plateau
and the Jordan Rift Valley. It includes part of the Western Highlands where high
rainfall is expected. This group is sub-divided into two sub-groups according to the
area of discharge:
I: Where discharge is into the Jordan River Valley.
II : Where discharge is into the Dead Sea.
GROUPB:
Group B includes the small, steeply sloping, rift side catchments along the
escarpment which discharge to the Jordan River Valley and the Dead Sea. It
90
includes most of the highly rainfall western Highlands. On the same basis as in
Group A, it is sub-divided into two sub-groups: I and II.
GROUpe:
Group C includes the small catchments along the eastern side of Wadi Araba.
These catchments, particularly in the southern part of Wadi Araba lie in a low
rainfall zone. Thus this group is subdivided into two sub-groups:
I: In the northern part of Wadi Araba.
n : In the southern part of Wadi Araba.
GROUPD:
Group D includes the large flat desert catchments which occupy a substantial part
of the country to the east of the Central Plateau. Group D catchments lie wholly
within a low rainfall zone.
GROUPE:
Group E includes the flat desert catchments in the low rainfall zone of the Central
Plateau, but belong to watershed areas of Group A as far as the area of discharge is
concerned.
Meteorological and drainage characteristics data for the different catchments in
Jordan were analysed to find the relationship between rainfall and runoff (Appendix B.3).
As discussed before, there are many factors affecting runoff processes. These factors are
either due to the rainfall intensity, duration, and time distributions, or to catchment
drainage characteristics. The runoff in Jordan varies between the catchments according to
the above mentioned parameters. Generally the runoff coefficients are low with respect to
the rainfall, and range between 2 and 7% (Table 3.7). It is believed that for the type of
climate and superficial materials prevailing in Jordan, the threshold rainfall necessary for
runoff is high. Thus a considerable amount of the annual rainfall goes to satisfy the soil
moisture deficits before runoff will occur. As expected, the topography in the desert
areas is less favourable for runoff, and the average annual runoff is therefore low.
Although, in some areas, where the conditions are less favourable for infiltration to occur,
91
a high runoff coefficient was recorded. The playas collect stonn runoff for long periods
before drying through evaporation.
Among the other parameters, rainfall and catchment areas are the most important
factors affecting the runoff in Jordan: the runoff increases with increase in rainfall and in
catchment areas.
Group Minimum Maximum Average Weighted Average
A 4.6 8.3 6.94 6.94 B 0.6 6 2.95 2.84 C 0.4 4.5 1.81 2.07 D 1.2 3.3 2.25 2.03 E 3.4 5.3 3.97 4.05
Table (3.7) Mean annual runoff coefficient (%) for the different groups of
catchments.
3.5.2 RUNOFF IN THE STUDY AREA
The study area includes four surface catchment areas of groups A, D, and E. Apart
from the Jafr catchment which is a closed desert catchment draining the Jafr Basin into
the Jafr Playa, the rest of the study area is drained by numerous wadis into the Jordan
River Valley and the Dead Sea.
3.5.2.1 UPPER - ZERQA CATCHMENT
This catchment is considered to be a sub-catchment of Wadi Zerqa drainage area.
It has an area of approximately 850 km2, and is drained by the Upper Zerqa, which
comprises a number of wadis tributary to the Zerqa River (Figure 3.13).
BASE FLOW AND SPRING DISCHARGES
The only perennial stream in this catchment is the Zerqa River itself which begins
at the Ras el Ain spring in the western part of Amman. This stream is subsequently fed by
other springs either in or close by the river.
92
-....:
235 265
Figure (3.13) Upper Zerqa Catchment
N
w-<r E
s
LEGEND
,/ .. River Wadi
• Well
a Spring
~ Gauging Station
o S IOl.rn t-s-;--';;z-=-' ---.. ----.
During the rainy season the baseflow from springs is supplemented by periodic
flash surface runoff from side wadis which for the remainder of the time are dry. The
natural baseflow is greatly reduced by pumping from the springs and aquifers and at times
of low flow the river is reduced to almost nothing.
Within the Upper Zerqa catchment several springs are found. Their waters used to
be the main water supply for the Amman and Zerqa cities, but in recent years, due to
groundwater abstraction, the flows have reduced significantly, and some of them now are
almost dry. Table (3.8) lists the main springs and their flows in the area.
Though there are other springs in the area their contribution to runoff is
insignificant compared to that of the main springs which are listed in Table (3.8). Some
springs along the Zerqa River emerge only in high rainy season, whilst they are dry most
of the time.
The discharges of these springs show a close correlation with the rainfall and vary
greatly, not only between years but also within a year.
Baseflow in the Zerqa River is maintained by spring discharges in the area, but
also varies depending on the rainfall and the abstractions at the springs. The baseflow
record at the Sukhna gauging station is intermittent and of inadequate length to provide
an estimate of mean annual flow. However, a record which covers only nine years
in the period between 1971-1985 was obtained (Appendix B4.l). The mean annual
baseflow ranges between 2.19 and 25.13 MCMla, with an average of 9.76 MCMla.
However, it is worth understanding that the baseflow measurements at Sukhna
gauging station do not reflect the natural baseflow, since the springs which would
otherwise maintain the baseflow are depleted by abstraction. Furthermore, for the period
of the record, the baseflow measurements include a combination of springs discharge and
effluent discharge from sewage and industrial treatment plants along the Zerqa River.
Only recently have these effluents been diverted to the main treatment plant which again
discharges its effluent into the Zerqa River but outside the study area downstream of the
Sukhna gauging station. The mismatch between the mean annual baseflow (Appendix
B4.1) and the springs discharge (Table 3.8) is also due to the difference in the record
periods.
94
Spring Co-ordinates Number of Discharge (m'/h)
East I North Measurements Minimum Maximum Mean Ras el Ain 237.000 150.500 70 6.23 2820 412 Zarbi 243.400 158.80 62 0.0 948 67 Ruseifa 246.600 158.610 15 99.8 400 215 Zerqa 252.600 162.700 85 0.0 3240 818 Sukhna 250.500 171.000 185 0.0 1300 687 Nimra 248.000 172.800 110 31.3 511 194 Hussaya 251.600 173.400 15 3.06 134 42 Source: WAJ (1986), Sprmg flow data m Jordan, Techmcal Paper No. 51, WAJ, Jordan.
Table (3.8) Spring discharge data for the main springs in Upper Zerqa Catchment
FLOOD FLOW
There are numerous non-perennial wadis draining the Upper Zerqa Catchment,
The most important is Wadi Abdoun-Wadi Seil with adjoining smaller wadis from the
rainy escarpment area such as Wadi Hanutia, Wadi Ghubar, Wadi Saqra, Wadi Haddada,
and Wadi Zarbi. Other big wadis like Wadi Qatar, Wadi el Madhana, Wadi Khaja, Wadi
Hassor, Wadi Sa'ieda, and Wadi Hussaya lie within low rainfall zones and flow only
occasionally.
The reliable record of flood flows at the Sukhna gauging station is of inadequate
length to provide an estimate of mean annual flow. The record shows mean annual flood
flow ranges between 2.17 and 39.13 MCM and an average of 9.14MCMla. However
Agrar und Hydrotechnik (1977) have estimated 2.7 MCMla and 8.6 MCMla of flood
flows at Wadi Abdoun and Sukhna Gauging station respectively.
In this study the catchment area was subdivided into six sub-catchments, each
covering the surface drainage area of the main wadis in the catchment (Figure 3.13).
Then, the curve number approach of the Soil Conservation Services (SCS) of the US
Department of Agriculture (1972, 1985) was applied to estimate the rainfall excess or the
runoff for each subcatchment. Rainfall data for the period 1980-1985 were used to
estimate the mean annual flood flow and the relation between flood flow and the rainfall
in the area.
The Upper Zerqa catchment includes two major cities, Amman and Zerqa. The
urbanised zone occupies a substantial percentage of the total area. Thus, an increase of
impervious cover is expected, which will increase the runoff. Although, in recent high
95
rainfall years, springs along the Zerqa River within the urbanised zone which have been
dry for many years, emerged again indicating that, despite the water proofing introduced
by the urbanisation, infiltration still occurs. It is believed that steep topography IS
analogous to imperviosity as far as the initial abstraction or infiltration capacity IS
concerned. Hence high curve numbers were chosen also for the areas with steep
topography. The results are shown in Table (3.9).
The estimated annual flood flow varies widely between the subcatchments and
from year to year within the same catchment, depending on the subcatchment drainage
characteristics and the rainfall intensities and distributions. It ranges between 0.7-11.4 %,
with lower values in the eastern part where the rainfall and urbanisation is less. The
average runoff coefficient for the whole basin is estimated to be 6.2 % (Table 3.9).
The mean annual flood flow volumes range between 3.73 and 22.49 MCMla with
an average of 12.43 MCMla, which is higher than the measured volume at Sukhna
gauging station. It is possible that the estimated volume is correct, and that part of the
flow either infiltrates and/or evaporates before it reaches the measuring point at Sukhna.
Catchments Year Ave. 1981 1982 1983 1984 1985
AZI Rainfall 34.9 23.2 33.5 24.6 23.7 28.0 (64 km2
) Runoff 2.6 0.11 1.2 1.0 1.2 1.2 RO/P (%) 7.4 0.5 3.5 4.1 4.9 4.2
AZ2 Rainfall 68.9 64.2 93.7 55.4 63 69.1 (167 km2
) Runoff 11.6 3.3 11.6 5.5 8.1 8.0 RO/P(%) 16.8 5.1 12.4 9.9 12.9 11.4
Upper Zerqa AZ3 Rainfall 29.6 25.4 44.4 18.7 30.1 29.6 Basin (121 km2
) Runoff 3.7 0.32 3.4 0.1 2.4 2.0 RO/P (%) 12.6 I.3 7.6 0.4 8.0 6.0
AZ4 Rainfall 21.8 14.8 21.2 13.2 17.1 17.6 (107 km2
) Runoff 2.3 0.D3 0.3 0.14 0.60 0.7 RO/P(%) 10.5 0.2 1.4 1.1 3.4 3.3
AZ5 Rainfall 29.0 19.2 23.5 14.4 23.3 21.9 (156 km2
) Runoff 0.8 0.0 0.0 0.0 0.02 0.16 RO/P (%) 2.7 0.0 0.0 0.0 0.1 0.7
AZ6 Rainfall 41.5 29.2 41.2 25.2 42.9 36.0
(235 km2) Runoff 1.6 0.0 0.0 0.0 0.54 0.42
RO/P(%) 3.8 0.0 0.0 0.0 I.3 1.2
TOTAL Rainfall 226 176 258 151 200 202 (850 km2
) Runoff 22.5 3.7 16.4 6.7 12.9 12.44 RO/P (%) 10.0 2.1 6.4 4.4 6.5 6.2
Table (3.9) Estimated flood flows (MCMJa) in the Upper Zerqa Basin obtained by using the eN method.
96
3.5.2.2 WADI MUJIB CATCHMENT
This catchment covers about 6530 Ian2 of mainly plateau land to the east of the
Dead Sea and is defined by the surface water catchment of the Wadi Mujib (4500 Ian2)
and its principal tributary the Wadi Wala (2030 Ian2) (Figure 3.14).
Wadi Mujib originates at the southern end of the catchment and its mam
tributaries are Wadi Hafira, Wadi Sultani, and Wadi Sueida. The upper reaches of Wadi
Hafira and Wadi Sultani consist of muddy flat areas called "Qa". Wadi Wala originates at
the northern part of the area; its main tributaries are Wadi Zafaran, Wadi Halq, and Wadi
Shabik. The downstream reach ofthe main Wadi Wala is called Wadi Haidan.
Both Wadi Mujib and Wadi Wala drain the area westward directly into the Dead
Sea. The river bed slopes start gently in the east and become steep downstream in the
western part.
A series of flood and baseflow discharge measurements have been carried out in
both the Wadi Mujib and Wadi Wala since 1962 (Appendix B4.2-B4.5). Prior to this,
only occasional and intermittent measurements at Wadi Wala and for spring discharges
have been taken.
BASEFLOW AND SPRING DISCHARGES
The Wadis Mujib and Wala have each cut gorges through the hills to where they
join some three Ian upstream of the Dead Sea. Both rivers in their lower reaches have cut
down to the saturated sections of water bearing formations so that perennial flow is
maintained by spring discharges. Baseflow begins with small springs and seeps increase
gradually downstream depending upon the phreatic level within the aquifer which, in tum
depends upon the extent of recharge from rainfall during the former wet season. During
the wet season, a significant storm runoff derived by short-lived surface flow responses to
heavy rainfall, may add to the baseflow.
In Wadi Wala the baseflow maintained by springs and groundwater runoff starts at
just upstream of the Wala bridge in Wadi Haidan at an elevation of 450 mas I. At 5 km
downstream from the Wala bridge at an elevation 350 masl, the baseflow suddenly
increases up to a mean annual discharge of 15 MCM, and up to 23 MCM downstream at
97
N
140 W~E s
120
100
080
040 LEGEND
"'=.:.- River Wadi
• Government Well
0 Private Well
~ Test Well
C1' Spring
020 ~ Gauging Station
~~~::2 Muddy Swamp Area
200 220
Modified from JICA (1988)
AMMAN
•
240
•
o 0 o
••
Figure (3.14) Wadi Mujib Catchment.
o ~ 10 km , ,
260 280 300
the confluence with Wadi Mujib. Most of the baseflow in the Wadi Wala is derived from
the Wadi Haidan which is dependent on the groundwater runoff from the B2/A7 aquifer
system, the major exploited aquifer in the study area. Table (3.10) shows discharge data
for the main springs in the Wadi Mujib catchment.
In the Wadi Mujib, the baseflow springs start from just upstream of the Mujib
bridge at an elevation of about 150 masI. The baseflow gradually increases downwards
collecting the spring water from the sandstone aquifer system. The flow discharge just
upstream of the confluence is about 12 MCMla.
The mean annual baseflow for the whole Mujib catchment is about 35 MCM. It
varies between years depending on the annual rainfall over the catchment.
Spring Co-ordinates Number of Discharge (m~/h) East I North Measurements Minimum Maximum Mean
W. Haidan 219.300 108.000 54 893 4540 1723 Lajjun 232.200 072.100 147 17 119 45 Muztawiyya 216.400 086.200 8 5 15 9 Er- Rashash 215.300 090.800 9 2 12 6 El - Khajajah 216.00 088.000 9 1 12 5 Arafat 215.300 091.200 13 3 7 5 Magbouleh 215.600 091.500 8 1 6 3 Um-Ma'ual 214.600 090.300 7 1 2 1 Source: WAJ (1986), Sprmg flow data m Jordan, Techmcal Paper No. 51, WAJ, Jordan.
Table (3.10) Spring discharge data for the main springs in Wadi Mujib Catchment.
FLOOD FLOW
Flood water in the wadis depends on the capricious nature of the storms in the
catchments, which occur during the wet season from October to May. These flash floods
discharge directly into the Dead Sea within a few days after the rain storm. Only two
small dams on the upper tributaries of the Wadi Mujib exist in the area.
Direct measurements of flood flows have been carried out at four gauging stations
for more than 25 years. These gauging stations are: Wadi Wala at Karak Road, Wadi
Wala at weir, Wadi Siwaqa at Desert Highway, and Wadi Mujib at Karak Road (Figure
3.14). A summary of monthly runoff at each station is given in Appendix (B4.2-B4.5).
99
According to the observed data, annual flood flow varies from almost negligible
to as high as to be considered the major component of the total flow. The mean annual
flows at Wadi Wala, Wadi Mujib, and Wadi Siwaqa are 19.1, 20.42, and 4.3 MCM
respectively. The difference is the result of differences in rainfall pattern and the
topography of the catchments.
To evaluate the relation between the flood flow and the rainfall, the curve number
method was used. The area is divided into 12 sub-basins (Figure 3.14), each sub-basin
covering a main wadi with its tributaries having similar geological, hydrological, and
topographical characteristics. The flood flow depth is assumed to be uniform over a sub
basin. Area W6 at the lower reaches of Wadi Mujib and Wadi Wala, which extends from
the gauging stations of the Wadis Wala and Mujib at Karak Road downstream into the
Dead Sea, is comprised of 230 km2 of the Wadi Wala catchment and 160 km2 of the
Wadi Mujib catchment. Table (3.11) shows the estimated flood flow volume in the study
area.
In general the calculated flood flow correlated well with the observed at the
gauging stations. The calculated annual mean volume for the whole Wadi Mujib
catchment is about 55 MCM. For the Wadi Wala and Wadi Mujib catchments prior to the
gauging stations, the mean annual flows are 22.4 and 23.3 MCM respectively. Area W6,
and according to the steep topography and the concentration of rainfall in the surrounding
mountains, estimated to discharge 9 MCM of flood flow every year.
However, despite the good correlation between the estimated and observed flood
flow, there seemed to be large differences for some floods. In the Wadi Siwaqa sub-basin,
for example, flows were estimated to be only 1 MCMla against 4.3 MCMla as observed
at Siwaqa gauging station. This could be explained by the rainfall intensities and the
condition of the soil cover during the storm, since the record shows that the area receives
flash flood produced by heavy short-duration rain storms. Most of the high anomalies in
the record happened in one day or sometimes in a few hours. In such conditions most of
the rainfall runs off Such circumstances, however, are beyond the limits of the curve
number method of calculation. Furthermore, the record shows unrealistically high values
of flood flow, such as 35.2 MCM in April 1982 corresponding to 20.9 MCM at
100
Catchments Year Ave. 1981 1982 1983 1984 1985
WI Rainfall 123 88.8 163 78.8 134 117.4 (490 km2) Runoff 16.1 4.4 13.4 3.4 15.9 10.6
ROIP (%) 13.1 4.9 8.2 4.3 11.9 9.0 W2 Rainfall 50.8 38.5 53.5 31.9 58.8 46.7 (380 km2) Runoff 3.9 0.13 0.61 0.34 4.10 1.8
RO/P (%) 7.6 0.33 1.1 1.1 7.0 3.9
W3 Rainfall 41.2 32.9 44.7 22.9 41.7 36.7 (340 km2) Runoff 1.2 0.0 0.0 0.06 1.4 0.53
RO/P (%) 3.0 0.27 3.3 1.5 Wadi Wala W4 Rainfall 30.4 18.8 29.6 15.4 22.6 23.34
(240 km2) Runoff 3.2 0.43 0.12 0.94 1.27 1.2 RO/P (%) 10.4 2.3 0.4 6.1 5.6 5.1
W5 Rainfall 56.7 46.7 87.5 40.0 80.7 64.1 (350 km2) Runoff 15.0 3.4 8.0 1.8 12.9 8.2
RO/P (%) 22.9 7.3 9.1 4.5 16.0 12.8
W6 Rainfall 78.1 61.9 88.4 54.4 73.8 71.3 (390 km2) Runoff 18.1 2.1 9.2 4.9 11.0 9.00
RO/P (%) 23.2 3.3 10.4 8.9 14.9 12.7
Total Rainfall 357 263 231 221 381 330 W.Wala Runoff 50.0 9.5 27.4 9.4 42 27.7 (2030 km2) RO/P (%) 14.8 3.6 6.7 4.7 11.3 8.7
M7 Rainfall 53.0 28.3 49.0 25.6 32.7 37.7 (460 km2) Runoff 3.5 0.4 0.01 0.7 0.2 1.0
RO/P (%) 6.6 1.5 0.02 2.7 0.6 2.5
M8 Rainfall 36.8 34.4 61.2 39.2 59.8 46.3 (320 km2) Runoff 2.6 0.0 0.4 0.9 3.4 1.5
RO/P (%) 6.9 0.7 2.3 5.7 3.1
M9 Rainfall 157 123 194 103 129 141.2
(640 km2) Runoff 16.9 0.5 13.6 5.3 12.0 9.7
RO/P (%) 10.8 0.4 7.0 5.1 9.3 6.8 Wadi Mujib MIO Rainfall 105 72.7 106 53.7 74.9 82.6
Basin (420 km2) Runoff 12.9 0.2 6.0 1.8 6.4 5.4 RO/P (%) 12.2 0.2 5.6 3.3 8.6 6.6
Mil Rainfall 144 79.7 175 93.1 110 120.4 (1490 km2) Runoff 10.9 1.3 0.3 3.0 2.8 3.6
RO/P (%) 7.6 1.6 0.2 3.2 2.5 3.0
M12 Rainfall 78.7 81.1 175 99.4 104 107.8
(1010 km2) Runoff 0.14 0.00 3.9 1.96 4.61 2.12 RO/P (%) 0.2 2.23 1.97 4.41 1.97
Total Rainfall 607 445.4 797 436 541 565
W.Mujib Runoff 54.4 3.16 28.1 15.6 33.9 27 (4500 km2) RO/P (%) 8.2 0.55 3.2 3.3 5.7 4.3
TOTAL WADI Rainfall 963 707 1027 658 922 895 MUJIBBASIN Runoff 104 13 56 25.0 76.0 55
(6530km2) RO/P (%) 10.8 0.18 4.6 3.8 8.2 6.2
Table (3.11) Estimated flood flows (MCM/a) in the Wadi Mujih Basin obtained by
using the CN method.
101
Wadi Wala and 18.1 MCM at Wadi Mujib for the same period. Eliminating such
erroneous values will reduce the mean annual flood flow at Siwaqa gauging station to less
than 2 MCMla.
The runoff coefficient is generally small since most of the rainfall is evaporated.
Its average ranges between 1.5 and 12.8 %, with lower values in the south and east and
higher in the west where the rainfall is high and the topography is steep. It increases with
increasing rainfall. The average runoff coefficient for the whole Wadi Mujib catchment is
estimated to be 6.2 % (Table 3.11).
3.5.2.3 WADI HAS A CATCHMENT
The Rasa catchment occupies 2198 km2 in the southern part of the Central Plateau
(Figure 3.15). The wadis in the Western Highlands are narrow and moderately incised,
while those in the eastern part of the catchment are flat. All the wadis in the upstream
reaches drain flash floods to the Qa EI Jinz central playa then into the Dead Sea by Wadi
Rasa.
At Rasa River there are two gauging stations, one at Hasa Tannur on the upper
part of the Rasa catchment, and the other one downstream at Ghour Safi. Runoff records
for both gauging stations show absence of data for considerable periods (Appendix B4.6
and B4.7). The essential feature of the record at Ghour Safi is the baseflow discharges
maintained by spring discharges emerging from the carbonate aquifer system in the upper
part of the basin and the sandstone aquifers in the lower part. It shows consistency
through the years of the record with a mean annual flow of about 27.67 MCMla. Spring
discharge data for the main springs emerging from the carbonate aquifer system are
shown in Table (3.12).
Observed flood flow records show mean annual discharge of about 8 MCM for
the Upper Rasa catchment at the Tannur gauging station and about 10.9 MCM for the
whole Rasa catchment at Ghour Safi.
The estimated flood flow volume as estimated using the curve number method is
summarised in Table (3.13). It indicates mean annual flood flows of 9.7 MCMla for the
whole catchment which is comparable with the observed volume at Ghour Safi. But the
102
000
900
Sea
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200 300
Figure (3.15) Wadi Hasa and Jafr catchments
N
w-\rE s
I •• •
t-:rz1:r=r=t= r==1
LEGEND
--.- River Wadi
. ." ~ . .. -.::. Muddy Swamp Area
d Spri.ng
~ Gauging Station
Spring Co-ordinate Number of Discharge (m'/h) East I North Measurements Minimum Maximum Mean
Irhab 211.000 035.500 13 2.8 10.3 4.8 Bir El-Harir 214.600 020.600 15 16.6 95 43.9 ElAweel 221.200 038.800 7 7.1 19.1 11.5 Mezrab 216.200 030.800 6 2.5 7.1 4.6 Mugheisel Kabeira 224.300 042.600 143 27 95.4 56.7 Yahoudiyya Kabira 224.600 042.500 127 7 27.1 15.8 Yahoudiyya Sagheira 224.700 042.600 81 0.8 9.9 2.9 Iflat 224.300 042.000 72 4 36.5 15.2 Mosalla Fouga 224.200 041.800 74 0 12.4 1.6 Mosalla Tehta 224.100 041.700 74 2.1 13.9 6.1 Irteija 224.100 041.800 10 0.5 3.6 1.2 Abu Shattal 223.300 042.000 9 2.6 9.3 6
Source: WAJ (1986), Spring flow data in Jordan, Techmcal Paper No. 51, WAJ, Jordan. Table (3.12) Spring discharge data for the main springs in Wadi Hasa Catchment.
estimated volume for the Upper Hasa catchment of about 6 MCM/a is less than the
observed volume at Tannur by 2 MCM. Given the low rainfall and the flat topography in
the eastern part, it is believed that the observed volume is too high to be justified:
estimated runoff coefficients for the area range from less than I % in the eastern part to
just over 4 % in the west with an average of 2.5 %. In the lower part of the catchment
the observed mean annual flood flow calculated from the difference between the values at
the two gauging stations is about 2.9 MCM/a, slightly low for the amount of annual
rainfall the area receives and the kind of topography. The estimated volume of 3.68
MCM/a and runoff coefficient of 6.5% are thought to be more realistic. It is possibly that
the time required for the flood flow to gather and reach the gauging station allowed part
of the flood water to evaporate before reaching the gauging station. The runoff coefficient
for the whole Hasa Basin estimated to be 3.3 %.
3.5.2.4 JAFR CATCHMENT
This catchment located in the southern part of the Central Plateau to the east of the
Western Highlands (Figure 3.15). It stretches over an area of 13427 km2, most of which is
classified as arid desert with mean annual rainfall of about 50 mm. The
catchment displays a typical centripetal drainage pattern with all wadis draining from
104
Catchments Year Ave. 1981 1982 1 1983 1 1984 1 1985
HI Rainfall 69.1 46.8 76.0 36.6 55.1 56.7 (335 km2
) Runoff 8.3 0.13 3.5 1.6 4.9 3.7 RO/P (%) 12.0 0.3 4.6 4.3 8.9 6.5
H2 Rainfall 124 91 168 107 94 117 (785 km2
) Runoff 13 0.05 6.6 3.9 1.5 4.9 Hasa Basin RO/P (%) 10.2 0.05 3.9 3.6 1.6 4.2
H3 Rainfall 105 69.8 199 113 123 122 (1400 km2
) Runoff 4.1 0.0 1.3 0.00 0.15 1.1 RO/P (%) 3.9 0.64 0.0 0.13 0.91 Rainfall 298 207 443 256 273 295
TOTAL Runoff 25.1 0.18 11.4 5.43 6.53 9.7 (2520 km2
) RO/P (%) 8.4 0.1 2.6 2.1 2.4 3.3
Table (3.13) Estimated flood flows (MCM/a) in the Wadi Hasa Basin obtained by
using the eN method.
the encircling highlands to the central playa, which is an extensive mudflat of about 240
km2.
The wadis draining into the central playa of the Jafr Basin are of two types: in the
western area, where the wadis have their headwaters in the Western Highlands where the
annual rainfall exceeds 150 mm; and in the eastern area, the wadis rise on the edge of the
depression where the annual rainfall is around 40 mm.
The baseflow within the Jafr Basin is limited to spring discharges found
exclusively in the Western Highlands. This baseflow has been developed as source of
potable water supply. Discharge data for the main springs in the area are tabulated in
Table (3.14). The mean annual spring discharge is 1.3 MCM. The long term records
shows there has been a general reduction in spring flow in recent decades.
In the Jafr Basin there is only one runoff gauging station, in the Wadi Jurdhan in
the Western Highlands. The station does not represent the basin as it is located in a high
rainfall area with different topographical regime; however, it does provide some idea
about runoff coefficients in the area. Appendix (B4.8) shows the observed annual runoff
and the runoff coefficients for the period of the record (1963-1982). The annual runoff
was found to range from none to 1.52 MCMla, with an average of about 0.50 MCMla,
while the observed runoff coefficient ranged between 0.2-5.8 % with an average of 1.9%.
105
Spring Co-ordinate Number of Discharge (m'/h) East I North Measurements Minimum Maximum Mean
Suweilem 220.900 955.800 132 0.0 22.9 7.0 Mayyet-Khewarah 220.300 956.900 74 0.5 12.1 5.6 Mayyet-Nakhleh 220.500 956.700 79 0.4 11.3 5.4 Udruh 207.300 971.300 201 0.0 120.0 32.7 Jerba-EI-Kabira 207.900 975.700 163 1.8 41.0 10.1 Jerba-El-Saghirah 207.500 974.200 95 1.2 12.1 5.8 Tumeiah (North) 207.300 972.700 15 2.0 11.0 5.2 Nijil-Shaubak 202.200 992.200 140 2.0 40.3 14.5 Basta 201.000 959.800 222 1.4 60.1 9.5 Ail (Janoubeyyeh) 200.800 958.000 185 3.1 18.0 7.6 Abu-Iea-Itham 200.800 956.400 97 0.5 15.3 5.9 Derbas 200.100 954.800 97 0.8 6.0 2.9 Uniq 199.300 955.000 95 1.6 8.5 5.9 Farthkh 198.800 956.100 96 0.9 20.8 8.0 EI-Dhaur 195.700 948.300 90 2.3 7.1 3.9 Mureigha 200.400 946.600 74 0.7 9.7 3.4 Sadaqa 197.500 952.600 97 0.0 5.6 1.7 Source: WAJ (1986), Sprmg flow data m Jordan, Techmcal Paper No. 51, WAJ, Jordan.
Table (3.14) Spring discharge data for the main springs in the Jafr Basin.
The estimated runoff volumes in the basin are summarised in Table (3.15). They
vary between the sub-catchments, with high values in the western part and lower values
in the east and southeast. The estimated mean annual volume for the Wadi Jurdhan sub
catchment, which covers approximately 350 km2, is found to be about 0.7 MCMla. It
compares well with the observed volume at the gauging station of 0.50 MCMla for an
area of only 222 km2 covered by the station. The mean annual runoff for the whole Jafr
Basin ranges between 0.4-73 MCMla and averages 16.9 MCMla. The runoff coefficient
ranges between 0.11-7% and average 3.15 %.
3.6 CONCLUSION
The climate in the study area can be divided into two major types: the
Mediterranean type on the Western Highlands and the semi-arid to arid type on most of
the Central Plateau and eastern desert. The climate characterised by cold winters and hot
dry summers. January is the coldest month, and August is the hottest. Average annual
temperature ranges from about 13 °c in some high mountainous areas to about 18.7 °c in
the low lands and in the extreme southeastern area. Temperatures are subject to large
106
Catchments Year Ave. 1981 1982 1983 1984 1985
Jl Rainfall 202 158 199 72.0 151 156.5 (1250 km2) Runoff 40.3 1.2 2.3 0.2 5.4 9.9
RO/P (%) 20.0 0.74 1.13 0.26 3.6 6.3 J2 Rainfall 45.8 31.5 27.7 12.1 21.6 27.8 (350 km2) Runoff 3.44 0.04 0.0 0.0 om 0.70
RO/P (%) 7.5 0.11 0.02 2.5
13 Rainfall 30.8 14.6 24.5 13.0 21.5 20.9 (340 km2) Runoff 2.9 0.00 0.0 0.0 0.11 0.61
Jafr Basin RO/P (%) 9.5 0.01 0.5 2.9 J4 Rainfall 34.8 30.0 21.6 13.6 32.5 26.5 (263 Knl) Runoff 1.1 0.35 0.0 0.17 1.6 0.62
% 3.0 1.2 1.22 4.8 2.4
J5 Rainfall 726 115 220 221 244 305.3 (11224km2) Runoff 25.3 0.0 0.0 0.0 0.0 5.1
RO/P (%) 3.5 1.7
Rainfall 1039 349 493 332 471 537 TOTAL Runoff 73 1.6 2.3 0.4 7.0 16.9 (13427 km2) RO/P (%) 7.0 0.5 0.5 0.11 1.5 3.15
Table (3.15) Estimated flood flows (MCM/a) in the Jafr Basin obtained by using the
eN method.
daily and seasonal fluctuations. Monthly mean temperatures vary between 5 and 25°C.
Large variations in temperature also occur within short distance due to topography.
Rainfall is primarily controlled by the Eastern Europe and Western Mediterranean
cold fronts which are drawn by the Eastern Mediterranean low pressure system. Rainfall
in the study is seasonal, occurring in the period October to May with the highest fall in
December and January. Rainfall outside this period would be an extremely rare event.
Precipitation generally decreases from west to east and from north to south. However, this
pattern changes locally in some areas owing to orographic effects over the higher
elevations of the Western Highlands. The mean annual precipitation decreases from
about 600 mmla in the northern Western Highlands to less than 50 mmla in the
southeastern desert. However, in the eastern and southeastern deserts, extended periods of
no rain and periods of flooding are not unusual. The mean annual volume of precipitation
over the whole study area is mounted to about 1929 MCMla.
Most of the precipitation is evaporated from the land surface, is transpired by
vegetations, or moves directly to nearby streams and wadis as overland flow. Depending
107
on the amount, duration, and intensity of the precipitation, as well as on the nature of the
terrain, soil, and topography, part of the precipitation infiltrates the land surface; some of
the infiltrated water may eventually recharge the groundwater system.
Estimates of average annual evapotranspiration rate within the study area vary
from a maximum of about 2490 mmla in the southeastern desert to a minimum of nearly
1153 mmla in the Western Highlands. The annual evapotranspiration rate decreases to the
north and west, reflecting regional climatic trends.
Runoff is the second largest element of the water budget in the study area (after
evapotranspiration). The curve number method used in this study, generates estimates of
runoff depth, Q(mm), as a function of rainfall depth, P(mm), and a storage term, S, which
is a function of the curve number, CN. The CN are assigned based on soil type and land
use, and are modified depending on soil moisture content and the time of rainfall. The
runoff averages about 9.3 % (179 MCM/a) of the amount of precipitation: 4.9 % (94
MCMla) as surface runoff, and about 4.4 %( 85 MCMla) as baseflow. Surface runoff is
generally most important where the terrain is steep, the soil texture is fine, and there is
little plant cover. Estimated surface runoff coefficients vary between the different
subcatchments, it ranges from about 3.15 % in the desert areas to more than 6.2 % in the
west and north. Baseflow is controlled largely by the underlying geology, the degree of
stream entrenchment, and the head relations between groundwater levels and water levels
in the surface drains. Shallow headwater streams receive baseflow from locally occurring,
principally unconfined aquifers. The major, more deeply entrenched streams-such as the
lower parts of the Wadi Mujib and Wadi Hasa-receive baseflow from the deep,
principally confined aquifers. Although over the long term the shallow streams drain off a
significant a mount of groundwater, many dry up during extended periods of little
precipitation. Because the major streams tap flow paths deeper in the regional
groundwater flow regime, they are less affected by either droughts or periods of above
average rainfall.
The areal pattern of runoff is similar to that of precipitation. Runoff generally
increases from east to west and in the northern part of the study area. However, this
pattern reflects the changes in climate, physiography, and geology of the study area.
108
CHAPTER FOUR
AQUIFER SYSTEM
An aquifer system is a group of lithological units which together have a
significantly greater permeability than major underlying or overlying units or
composites of units. Variation in permeability within a system gives rise to local
changes in aquifer characteristics, both laterally and vertically, and thus to complex flow
dynamics. Nevertheless, on the regional scale an aquifer system may be considered as a
single hydraulic unit. Such characteristics as hydrostatic head, flow directions and water
quality are regionally related in an aquifer system. Under certain geological conditions
two or more systems may interact to form a composite system.
4.1 AQUIFER SYSTEMS IN JORDAN
The Disi Group Aquifer underlies the entire country and sandstones within the
Khreim Group form low-yielding aquifers in an extensive region in the southern part of
the country. However the piezometric surface of these aquifers are in excess of 200 m
below the ground surface on the Plateau and in the Western Highlands. Moreover, the
aquifer is more than 900 m below ground surface except in some deeply incised wadis
along the rift, and the depth to this aquifer increases eastwards. These sediments outcrop
on the lower slopes of the rift escarpment and in the Southern Desert. They occur in a
zone of low rainfall, and receive little recharge. Groundwater movement in the Palaeozoic
aquifer is generally northeastwards in accord with the regional dip.
The Kurnub and Zerqa Groups which underlie almost the entire country form an
aquifer system which is continuously saturated. The two Groups are hydraulically
connected on a regional scale and they are regarded as a single system called the Kurnub
Zerqa Aquifer System (Parker, 1970). However, clayey strata of variable thickness and
lateral extent are present in both groups. The system has a deep piezometric surface, and
thus receives little direct recharge, except for limited outcrops in the core of breached
anticlines in the high rainfall zone. In the southeast the system is hydraulically connected
with the Amman-Wadi Sir system by the permeable zone in the Fassu'a Formation.
The thick limestones ofthe Na'ur Formation (A1I2) contain some water. However,
due to the limited outcrop and to the deep slopes where they occur, the aquifers receive
only limited recharge. They have poor permeability due to the limited development of
karstification. The aquifer has been tested in some wells, all of which were low yielding.
It supplies water to springs along the rift and in other deeply eroded areas.
The Hummar Formation (A4), separated from the A1I2 aquifer by the thick marls
of the Fuheis Formation (A3), forms locally important aquifer in the Amman- Zerqa area.
In this area it is recharged from rainfall on outcrops along the western limb of the
Amman-Zerqa syncline, and water moves eastward in accord with the dip and then moves
northwards to the Lower Zerqa Valley. In the northwest and in the north central part of
Jordan the aquifer receives little recharge and has low permeability. To the south of
Madaba, the Hummar Formation wedges out and the Fuheis and Shu'eib Formations form
a continuous aquiclude.
The Amman-Wadi Sir (B2/ A 7) aquifer system is the most important and extensive
in the study area. It outcrops in the high rainfall zone of the Western Highlands and
extends at depth beneath a cover of younger sediments on the Plateau. Groundwater
movement is generally eastwards in accord with the regional dip. However, many of the
larger side wadis of the Jordan Valley and Dead Sea cut back into the aquifer and cause
the flow to move westwards. The base flow of most of the perennial streams which
discharge to the Jordan Valley and Dead Sea are maintained, at least in part, by springs
rising from this system. It is estimated that more than half of the mean annual recharge is
discharged in this manner.
Where the aquifer is exposed in the Western Highlands the groundwater is
unconfined. Eastwards in the foothills and in the Plateau the system is confined by the
impervious marls of the Lower Muwaqqar Formation (B3). In limited areas to the
southeast of Shaubak the heads are sufficient to produce flowing wells. The marls of the
Shu'eib Formation (A5/6) form the lower confining beds to the system except in the
southeast where the marls and limestone of the Ajlun Group are laterally replaced by
110
arenaceous facies-the Fassu'a Formation. The Amman-Wadi Sir system IS then
hydraulically connected with and discharges into the Kurnub Group system.
The variable permeability ofthe Amman-Wadi Sir system is largely due to joints
and karstification of the limestone. However zones of high permeability may be
recognised which appear to relate more to concentration of flow than to structure.
Vertical changes in permeability occur within the system, and very less permeability
argillaceous layers are present. Except in the north of the area, the clayey units do not
appear to be of great lateral extent and their confining effects are only local. In the north,
the chalk and chalky marls of Wadi Ghudran Formation (Bl) are present between the
Amman and the Wadi Sir Formations and may form a more extensive aquiclude.
However their effects cannot be recognised in the regional pattern of the aquifer
characteristics and no marked hydraulic head differences between water bearing zones in
the Amman and Wadi Sir Formations are recorded.
The Rijam Formation (B4) forms an aquifer in the central part of the Jafr and
Azraq basins. In some parts of the Jafr Basin it forms a composite system with the
overlying limestones of the Jafr Formation and locally with the alluvium. The lower
bounding surface of the system is formed by the chalky marls and chalk of the Upper
Muwaqqar Formation. There is no extensive upper confining layer, though low
permeability alluvial materials may produce local artesian conditions. Groundwater
movement in the system is in an easterly direction. The water does not discharge at the
surface within the Jafr Basin so it must be transmitted laterally, probably through the
chalks of the Upper Muwaqqar Formation. Within the Jafr Basin the saturated zone of the
Rijam Formation occurs in an area of very low rainfall and direct recharge is probably
negligible. It is believed that flash floods in wadis provide most of the recharge. The
permeability of the aquifer is extremely variable and the water appears to move in well
developed solution channels along preferred paths separated by zones of low
permeability.
The extensive basalt rock sequence to the north of Azraq contains aquifers which
discharge into the closed groundwater system of Azraq Basin and into the Amman - Wadi
Sir system in the vicinity of Zerqa and Mafraq. The aquifers are recharged in the high
111
rainfall zone of the Jebel Druze mountains in Syria and the groundwater moves radically
southwards.
An upper shallow aquifer system occurs under water table conditions in the
alluvial deposits which are in hydraulic continuity with the Amman - Wadi Sir aquifer
system along Zerqa River in Amman - Zerqa area.
4.2 AQUIFER SYSTEMS IN THE STUDY AREA
The Mesozoic-Cainozoic carbonate sediments form a sequence of aquifers and
aquic1udes. Four aquifer systems have been recognised, the first which has regional
importance, is the Amman-Wadi Sir (B2/A7) aquifer system which extends throughout
much of the entire country and varies considerably in lithology, depth of occurrence,
hydraulic properties and resource development. The other three aquifers aare the Na'ur
(A1I2), the Hummar (A4) and the Rijam (B4) aquifer systems, and they are of importance
locally in limited areas.
The relation between the geology and hydrology that provides a basis for the
study of a regional flow system is shown in Table (4.1). This relation provides a simpler
system for study and is the foundation for the conceptual model for describing
groundwater flow in the carbonate aquifer systems. The conceptual model of groundwater
flow in the carbonate aquifer systems is shown on Figure (4.1).
Throughout the study area the Mesozoic-Cainozoic carbonate sediments are
underlain by arenaceous deposits of Lower Cretaceous and Palaeozoic age which contain
aquifer systems that have not been considered in detail in this study.
Formation
112
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4.2.1 EXTENT AND LITHOLOGY
4.2.1.1 THE NA'UR AQUIFER SYSTEM (A1I2)
The thick limestones of the Na'ur Formation contain some water. They outcrop on
the lower slopes of the rift escarpment and in the deeply incised wadis from near Wadi
Zerqa southward to the vicinity of Ras en Naqb. They are also exposed on the flanks of
several eroded anticlines to the northwest of Amman. The limestones are present in the
upper part of the Formation, the lower part consisting mainly of marls. This lower
Member forms the confining layer which separates the Na'ur aquifer from the underlying
sandstone aquifers in the Kurnub Group. The marls of the Fuheis Formation (A3) are the
major upper confining beds for the Na'ur aquifer.
The upper Member of the Na'ur Formation is not a continuous limestone
sequence, but contains marl units which locally have a thickness of 40-50 m. Thus it is
thought that the thicker limestones within the Member may form individual aquifers.
However the flow dynamics are insufficiently well known to establish the relationship
between the limestone units. In the northwestern part of the study area, there are three
main limestone beds in the upper member of the Na'ur Formation. The upper and lower
beds are the largest and have thicknesses of 30 - 40 m.
From the Wadi Mujib southwards only two thick limestones are present and these
are 30-40 m thick. The limestones are often dolomitised. East of Ras en Naqb the Na'ur
Formation is laterally replaced by the sandy facies of the Fassu'a Formation. It is believed
that there is lateral continuity between the sands and the limestones.
Due to the limited area of their outcrop and the steep slope where they occur, the
aquifer receives only limited recharge. They have poor permeability due to the limited
development of karstification.
4.2.1.2 THE HUMMAR AQUIFER SYSTEM (A4)
The Hummar Formation is a limestone unit in the Ajlun Group to the north of
Wadi Mujib and extends eastward at depth into the Azraq Basin. However the main area
in which the Formation is known to provide aquifer potential is in the Amman-Zerqa
114
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t-: 1l:
I II ~//
.. , // ...... • . . ~// ..... I , ... •
..:..' 'j/ / .. • //
..... • ...
.... (I) -~ en (I) 0 ~ L-
~ "C
0 >.
M E J: N e -00;;: N
"t:l -.:t (I)
I;:: -i3 (I)
~ L-::I en u:
0 It) .....
syncline (Figure 4.2). Elsewhere limited recharge, poor permeability and deep static water
levels limit the productivity of the aquifer.
The aquifer outcrops as a narrow band high on the northwestern flank of the
Amman-Zerqa synclinal structure. The Amman-Zerqa flexure and associated faulting are
believed to form a hydraulic barrier to the southeastward movement of water within the
aquifer. Within the syncline the aquifer is about 45 m thick and consists of limestone and
dolomitic limestone with thin shale and marl bands.
The Hummar Formation is known to be saturated in low amplitude, synclinal
structure in the vicinity of Salt. The structure is tilted to the southwest and the Formation
outcrops on its periphery and in the deeply cut valley of the Wadi Shue'ib and its
tributaries. Only the Salt well (S21) has penetrated the aquifer in this area, and here it
consists of 43 m of limestone and dolomitic limestone with thin marl bands.
North of Wadi Zerqa in the Irbid - Mafraq area, the Hummar Formation is water
bearing but the permeability is low and the water levels are deep. The aquifer outcrops in
the Ajlun Mountains, to the north of Jerash, and along the Zerqa Valley. It dips regionally
northwards towards the river Yarmouk beneath a cover of younger sediments. In the
Jerash area the aquifer is about 40 m thick and consists of dolomitic limestone and rudist
reefs.
The lower confining beds are marls and shaley marls of the Fuheis Formation.
Where the aquifer is artesian it is confined by marls and marly limestone of the Shue'ib
Formation.
The Hummar Formation forms a locally important aquifer in the Amman-Zerqa
area. The aquifer is confined by the overlying marls of the A5/6 Formation. To the south
of Mad aba, the Hummar Formation wedges out and the A3 and the A5/6 Formations form
a continuous aquiclude (Figure 4.1).
4.2.1.3 AMMAN - WADI SIR AQUIFER SYSTEM (B2/A7)
This is the most important and extensive aquifer system in the study area; it
outcrops in a large area of the Western Highlands, on the western edge of the Plateau,
116
and a long the top of the Ras en Naqb erosion escarpment (Figure 4.3). It is present at
depth beneath younger sediments on much of the Plateau, and is known to extend beneath
the basalt of the Wadi Dhuleil - Mafraq area. Much of the outcrop in the Western
Highlands is in an area of relatively high rainfall; elsewhere the system is exposed in a
zone of low rainfall.
Except in the eastern part of the Jafr Basin, the lower confining beds of the system
are the marls of the Shu'eib Fonnation (AS/6). In the Amman-Zerqa area the AS/6
Fonnation separates the B2/A7 aquifer system from the lower aquifer (A4) and is
considered to have low intergranular penneability, and any flow can only occurs through
fractures. Furthennore water flow from one aquifer to the other presupposes a difference
in piezometric level between them. An exception to this is where the AS/6 Fonnation is
not found and the two aquifer are in direct contact through the alluvial deposits as in the
north-west of Zerqa. The groundwater flow model (Chapter 8) indicates a leakage of
approximately 2.16 MCM/a from the A4 aquifer to the upper B2/A7 aquifer existed
probably before extraction started. To the south of Madaba, the A4 Fonnation wedges
out and the AS/6, A3, and the A1I2 Fonnations fonn a composite aquitard separates
between the principal B2/ A 7 and the deep sandstone aquifer systems (Figure 4.1).
To the east and south east of Ma'an the marls and limestones which fonn the
Ajlun Group in the west of the study area, are laterally replaced by a sandy facies. At Jafr
and eastwards the AS/6 Fonnation can no longer be recognised and the lower confining
strata are ill-defined, and then the B2/A7 aquifer system and the lower Ajlun Group
merge to form a single aquifer which is hydraulically connected with the underlying
Kurnub aquifer system.
The system as recognised in the study area includes the aquifer of Wadi Sir (A7)
and Amman (B2) Formations. These consist predominantly of limestone, sandy limestone
and silicified limestone. Sandstones are present in the southern part of the area and
become increasingly important until they form a large part of the system in the eastern
part of the Jafr Basin. Beds of chalk, marl and shale occur within the system, and they are
of limited thickness and lateral extent and form only minor aquicludes.
117
200
ISO f--- •.
100
050
000
gsa
gOO
ISO
] ..
. . ... '. 0.. ~
• . C) :
uJ :
ISO
After WMP (1977) 200 250
LEGEND
·l····· 1': ....
,... "
,
: ~l" -""'" ~ . _. :,.
,.of·
'/,........ ..---
'"I ... - .---
!
iii :1:;:1:11: Outcrop of the B2/A7 aquifer syst~m I I! 1,1,
Area of confined conditions of the B2/A7
~_--I Main groundwater divide Direction of groundwater flow
1---1 • Main spring discharge
)00 lSO '00
Figure (4.3) Hydrogeological setting of the B2/A7 aquifer system
In the north and northwest the chalks and marls of the Wadi Ghudran Formation
(Bl) are present between the Wadi Sir and the Amman Formations and are thought to
form an extensive zone of low permeability which separates the upper and lower parts of
the aquifer. Where the marls are thick the Bl Formation forms an aquiclude; where the
Formation consists mainly of chalk it probably acts as an aquitard to movement of water
between the aquifers of the Amman and Wadi Sir Formations. In the study area the Bl
Formation is thin and of limited lateral extent. In some areas as revealed by borehole JTl
in the western part of the Hasa Basin, the Formation forms a local aquiclude. However in
regional terms the A 7 and B2 aquifer systems can be considered to be hydraulically
connected and behave as single layer aquifer system.
Carbonate rocks are the predominant aquifer material of the Wadi Sir Formation.
In the northern part of the area the aquifer consists mainly of limestones, some of which
are dolomitised. Thin beds of chert and chert nodules are present, particularly in the upper
part of the Formation. In the south-western part of the area the limestones are often sandy
and thin beds of calcareous sandstone are common. Most of the limestones, are thinly
bedded, jointed, and often contain solution channels. The intensity of jointing increases in
the vicinity of the fault system. However, there is thought to be a regional pattern of
joints which resulted from the large scale disturbance associated with the taphrogenic
movements which formed the rift. In some areas the joints have been sealed by deposits
of secondary calcite.
The aquifer material of the Amman Formation consists of limestone, silicified
limestone, chert, phosphatic chert, phosphate rocks and sandstones. The sandstones which
occur in the southern part of the area are present towards the base of the unit and are
often weakly cemented. The component rocks are jointed, the intensity of jointing
increasing in the vicinity of faults. The carbonate rocks have, in some areas, developed
solution channels.
The maximum recorded thickness of the system occurs outside the study area at
Azraq well (PPl) where it is about 350 m thick. In the Western Highlands and the
western part of the Plateau the system ranges in thickness from 200 to 350 m. There is
marked thinning of these beds towards Bayer. In the eastern part of the Jafr Basin the
119
base of the system is difficult to define as the Ajlun Group consists predominantly of
permeable sandstone and sandy limestone, and in this region the water bearing units of
the Ajlun and Kurnub Groups are thought to form a single aquifer system.
In the Western Highlands where the aquifer is exposed the groundwater is unconfined.
Eastward, in the foothills of the plateau, the system is confined by the impervious marls
of the lower B3 Formation. The marls of the A5/6 Formation form the lower confining
bed to the system, except in the southeast, where the marls and limestone of the Ajlun
Group are laterally replaced by an arenaceous facies, the Fassu'a Formation. Here the
B21 A 7 aquifer system is then hydraulically connected with and discharges into the
Kurnub Group system.
4.2.1.4 THE RIJAM AQUIFER SYSTEM (B4)
The Rijam Formation forms an aquifer in the central part of the Jafr and Azraq
Basins. In some parts of the Jafr Basin it forms a composite system with the overlying
limestone of the Jafr Formation and locally with the alluvium. There is no extensive
upper confining layer, though low permeability alluvial material may produce local
artesian conditions.
The Rijam Formation (B4) comprises a shallow aquifer system in the central part
of the Jafr Basin. The approximate limit of saturation in B4 in the basin is shown in
Figure (4.4). It encompasses an area of about 1,250 square kilometres which extends
north - eastwards from the Jafr Playa. The water table is known to extend beyond these
limits into the chalks of the Muwaqqar Formation. However, the permeability of the chalk
is very low so that the effective limit of the aquifer productivity is the limit of saturation
in the Rijam Formation. The Formation is unsaturated in the western part of the Jafr
Basin; wells drilled in Wadi Ishush (845), Uneiza (856 and 863), and Wadi Jurdhan
(839) found to be dry. In most of the area the aquifer is unconfined, but in the area of the
Jafr Playa, alluvial clays form an upper seal which gives rise to localised artesian
conditions. In the vicinity of the playa a aquifer may be traced laterally and vertically
from the Rijam Formation into the Jafr Formation limestone, and locally into the sands
and gravels of the alluvium. The maximum recorded thickness of the saturated
120
000
s
to;, .......... ' •••• '. • I, I' .' .,........ " ••
• 0
.~ Karak:~" '. ...." .: , .... o _ .' ~
" ,t . "
. . .. .. .. • 'Wu •. ( ~ • • :""asa ,'.# •
LEGEND
~ outcrop of 84
~ limit of saturated 84
J ........ r. ') .: _"~ basin boundary
O~~Z5~'" ( Tafilj" .... ~..... " . ~ • • LI··· "1 .. ' wadi
( '7.' 'J , ......
i1,ft ." . J ""iI e. .......... _- \ / ,J iI~il] (' ". _ \ . " ~ .. r v-t-
(
, ~ • J .,
" .:/,' " ." ... ,. -,/ '-..
/ ' -, ,~ ./ ~\ ( -... --. .' \ () .\ ..... ' "
J.~~... < ) f._ .. - 1:. . ()
/ \ ' j \\" ... ' '!o Jafr Basin ....• . J.
,.-;-: \
J .... '\ . ,..-./\ ..... ,- '-' ..'
" .' ~ j..' 1 " ,,: J ____ . \ ) .,.. '--'-..J'" .I ~ '\. \ , .•..• ' . . ,I ,.3) 'I Aqaba .J ~ \,~,
/ ./~~\, -. ~ -'-'- / ~ .-.J
-'-200 300
Modified from JICA (1990).
Figure (4.4) Hydrogeological setting of the 84 aquifer system in the Jafr basins.
Rijam Formation is 41m in Jafr Well No.1 (PPI5). The thickness of saturated aquifer
varies in accord with the shape of the top of Muwaqqar Formation and decreases
westwards from the centre of the basin towards the limit of saturation. Depths to water
are shallow and range from 15 to 35 m below the ground surface.
The aquifer materials ofthe Rijam Formation are crystalline and chalky limestone,
chalks, and chert bands. These sediments are interbedded with marl and chalky marl
which form zones of low permeability within the section.
The Rijam aquifer system is exploited mainly for irrigation. The aquifer suffers
from high salinities, probably due to irrigation returns, and has limited yield, less than 1
MCMla.
4.2.1.5 LOWER AJLUN GROUP AQUIFER SYSTEM (A1-6)
In general, the Lower Ajlun Group Aquifer is multi-layered and comprises shaley
and marly units separating discrete aquifers which, in the west, consist of limestone, and
in the south and east of sandy limestone and sandstone. In the southeast of the study area
it has been possible to delineate an uppermost arenaceous layer 20-50 m thick in a direct
hydraulic continuity with the overlying Wadi Sir Formation. This layer is separated from
the main aquifer by clays and silty sands. The underlying main aquifer is arenaceous but
generally impure and therefore poorly productive (Humphreys, 1986).
The extent of the AI-6 aquifer system within the Jafr Basin is not well known,
particularly in the north and west. In the Western Highlands the Group does not appear to
act as an aquifer but rather as a series of aquicludes which in faulted situations, serve as
barriers to flow to the overlying B2/ A 7 aquifer thus giving rise to springs. Immediately to
the east of the Arja-Uweina flexure and to the northwest ( Jebel Uneiza region) the
nature and extent of the aquifer are again incompletely known due to the absence of deep
boreholes. In Malan (Borehole S I) the aquifer occurs at depth of 319 m, is multilayered
and consists of 266 m of shales, shaley limestone and sandy limestone (Parker, 1970).
South of Malan at borehole PHOI the aquifer has been found to be very shaley and
mostly dry (Humphreys, 1986). At Jafr (Borehole SIS) it has been encountered at depth
of341 m and comprises some 306 m of calcareous sandstone, clay and dolomite. In the
122
000
s
'. ' ..
o .
Karak
.. . . . ' . . \Wad'H '., .. :
) ...... ' aSa
' ..
' ... ' .. -. ... , .
. .... " . '. , I
. . ' .. . - ..
' .. ' . . .
' ..
LEGEND
I I Outcrop of A1-6 Formation
Limit of saturated A 1-6
Fault
. ···r':.: '...... ~ Flexure
Oe.:,,;.~~~IC'" ( Taf~la ,I ... ~.~ -~ .... /' basin boundary
( ~ / Wadi Ha~ ... ",: - .-' wadi
) ~ l3aSin a .: .... ~'_. (' ; '- r, \. '. . .. \
, ~...... '/' . ':.J _ .... ---~ . l ....
! I .... ; ~\. \-""\",
I. ' . " ) ....0..' ~ "'\.
o •• ' ~ ....... :'" - "" •• \" '. :., •• , ", ...."
( '. :':;" .. \ .... '.~ ') . ...... \ ) ..... " ., . .... ......... ....... ':. . '"., ~ . .... "... { Jafr Basin A'1 ~l$(
I ,.' 0" : • iflliill III \ i '/!!fl1lll!11lJ /11!!1ilT!1l1I111!1!T11/l11!J!!I!!{
central and southern area the extent of the aquifer has been fairly well delineated by
Howard Humphries, 1986. In the general Shediya area it is 200 m thick and consists of
sandstone and marls. Toward the southern escarpment a limit of saturation has been
defined which is located a few kilometres south of that of the B2/A7 (Figure 4.5). In the
southwest the AI-6 is probably dry and in the east-southeast, where the sandy facies is
predominant, it is in hydraulic continuity with the overlying B21 A 7 and the underlying
Kurnub aquifer system.
124
CHAPTER FIVE
AQUIFER PROPERTIES'
5.1 INTRODUCTION
Hydraulic characteristics of any aquifer are the main factors affecting the amount
of water in storage, the rate at which water moves through the aquifer, and the rate and
the areal extent of water level declines caused by groundwater withdrawal.
The essential hydraulic parameters for an aquifer system evaluation include: the
saturated thickness, storage and transmissivity. These have been estimated mainly from
rock fabric, borehole drilling, geophysical and aquifer test data.
The aquifer systems are described in detail in order of their economic importance,
therefore, most of the discussions refer to the B2/ A 7 aquifer system. The hydraulic
characteristics of the other aquifers are discussed under separate titles.
5.2 ROCK FABRIC AND STRUCTURE
The carbonate rocks, in the form of limestone and dolomite, consist mainly of the
minerals calcite and dolomite, often with minor amounts of clay. Nearly all dolomite is
secondary in origin, formed by geochemical alteration of calcite. This mineralogical
transformation causes an increase in porosity and permeability because the crystal lattice
of dolomite occupies 13% less space than that of calcite (Matthess, 1982).
Geological younger carbonate rocks commonly have porosities ranging from 20%
for coarse, blocky limestone to more than 50% for poorly indurated chalk (Davis, 1969).
With increasing depth of burial, the matrix of soft carbonate minerals is normally
compressed and recrystallized into more dense, less porous rock mass. The primary
permeability of old unfractured limestone and dolomite is commonly very low. Many
carbonate strata have appreciable secondary permeability as a result of fractures or
openings along bedding planes. Secondary openings in carbonate rock caused by changes
in the stress conditions may be enlarged as a result of calcite or dolomite dissolution by
circulating groundwater. Secondary permeability is more important than the primary,
since it provides avenues for movement of groundwater through otherwise virtually less
permeable rock.
Observation shows that the solution openings along vertical joints are generally
widely-spaced, especially in folded rocks where the fractures are associated with crests of
anticlines. But openings along bedding planes are more important from the point of view
of water yield from wells. However, concentrated vertical fractures together with the bed
plane openings provide high permeability in many areas.
It is found that permeability is higher in the area where the limestone is covered
by alluvial deposits: this is due to fracture enlargement by the water infiltrating through
the alluvium to the underlying carbonate rocks, since this water is usually undersaturated
with respect to calcite. In other areas the water becomes more saturated in calcite prior to
entry into the fracture zones in the carbonate rocks.
The B2/ A 7 aquifer system is extremely heterogeneous aquifer unit that transmits
water through fractures that commonly constitute a considerable percentage of the
thickness of an individual bed. These beds are separated by less transmissive marls and
marly limestones, in which the fractures are more or less vertical. Lateral groundwater
movement in these marls and marly limestone interbeds is probably negligible when
compared with the volume of water that moves laterally through the limestone beds. This
is because movement of groundwater in the limestone beds is controlled by fracture and
joint systems, whereas movement of groundwater in the interbed zones is controlled by
primary features.
It is concluded that the average permeability of the aquifer unit is a depth
integrated permeability for the limestone, marl, and marly limestone beds. The average
relative total thickness of limestones and marls for the different Formations of the Ajlun
and Belqa Groups have been calculated from the compiled geological type sections and
well logs (Table 5.1). The table shows a high percentage of limestone in the recognised
aquifer units, and a high percentage of marls in the confining units. The limestone-marl
ratio for the Na'ur Formation is 50%, but due to its deep burial and hence lack of
sufficient water percolation to improve permeability, the Formation has local, very
126
limited groundwater potential with poor yields. The same is true for the Shue'ib
Formation; although it contains a substantial percentage of limestones, the limestones are
in the form of thin intercalations within the thick marls of the Formation and thick layers
or sequences of marls are more common than limestone.
The saturated thickness of the B2/ A 7 varies locally and regionally within a wide
range between 2 and 365 m with a mean and median of about 99 and 94 m, respectively.
Local variations from the regional patterns of saturated thickness result from structural
troughs and ridge on the base of the aquifer system. Subregional increases in saturated
thickness resulted from deep sedimentary basins are present in the Wadi Mujib and Jafr
basins. As the topographic highs and lows produce highs and lows in the potentiometric
surface, the relief in the potentiometric surface affects the distribution of saturated
thickness. Areas of lesser saturated thickness associated with areas of lower hydraulic
heads are present throughput the study area, however, such areas are especially prominent
along the eastern slopes of the Western Highlands, along the western margin of the
Central Plateau. In the eastern parts of the Central Plateau, the B2/ A 7 aquifer system
confined by the Muwaqqar Formation, thus the saturated thickness of the regional
groundwater flow system might be considered from the total thickness of the B 112 and
A 7 formations.
Formation limestone:marl Notes
Rijam (B4) 80% aquifer in J afr area
Muwaqqar (B3) 20% confining unit
Amman (B2) 70% (B2/ A 7) extensive aquifer system all over
Wadi Sir (A7) 60% the country
Shue'ib (AS/6) 40% confining unit
Hummar (A4) 90% aquifer in Amman -Zerqa area
Fuheis (A3) 20% confining unit
Na'ur (A 112) 50% poor yielding aquifer in places
Table (5.1) The average limestone:marl percentage for the different formations
127
5.3 PUMPING TESTS
The basic analysis of well aquifer hydraulics by pumping test depends on an
idealised representation of the aquifer, its boundaries, and the nature of the applied stress.
Several assumptions are made whenever analytical methods are applied to the analysis of
aquifer-test data. Despite these restrictive assumptions, analytical methods such as Theis
(1935), and Hantush and Jacob (1955) solutions have been shown to produce
representative aquifer parameters for confined and leaky confined aquifers (Hantush,
1956, Walton, 1970, Lohman, 1972, and Kruseman and de Ridder, 1979). The Theis
solution can be seen as special case of the Hantush-Jacob solution in the limit when
leakage factor, B[L2], approaches infinity, i.e., when the leakage from the confining layer
is very small. The related aquifer parameters evaluated are transmissivity, T[L2/t] and
storage coefficient, S.
Preliminary aquifer investigation usually entails measuring the drawdown in wells
under the influence of pumping. These measurements help in choosing an appropriate
aquifer model, and in estimating aquifer transmissivity and storativity. If pumping tests
are conducted so that drawdown is measured at several locations, spatial variability in the
aquifer parameters can also be estimated. The aquifer model and parameter estimates can
then be used to estimate groundwater movement in the area of investigation.
5.3.1 PUMPING TESTS IN THE STUDY AREA
A large number of aquifer pumping tests have been carried out in the study area,
the majority conducted and supervised by the Natural Resources Authority (NRA) and
lately by the Water Authority of Jordan (WAJ). The analyses of the pumping test data
have been based on various formulae developed and modified by different authors. The
results of the pumping tests analyses are presented in Appendix (Cl). The main series of
pumping tests conducted in the study area are:
Parker (1970 ) during the sandstone project carried out 54 pumping tests in the
study area. They were constant discharge tests and both draw down and recovery data
128
were recorded. Well efficiency tests were also carried out. Most of the tested wells do not
completely penetrate the aquifer, and observation wells were available for only three of
the tests. The rate of pumping varied between the tests according to the capacity of the
wells.
The pumping test data were analysed for estimates of transmissivity (T) using the
Cooper and Jacob (1946) "straight line" approximation of the Theis (1935) non
equilibrium formula.
Mudallal (1973) reported 33 pumping test analyses for wells yielding water from
the B2/ A 7 aquifer system in Amman - Zerqa Basin. These wells were tested at constant
discharge, and drawdown and recovery data were recorded. In some cases it was
impossible to maintain a steady pumping rate throughout the test, so weighted averages
were taken when fluctuations in the water level were observed. The rate of pumping
varied between tests according to the capacity of the wells. Most of the tested wells
penetrated the aquifer. Observation wells were occasionally available. The durations of
the tests were in the order of 72 hours. In many cases a period of 24 hours was found to
be satisfactory, after which a steady pumping water level was approached. Recovery data
were recorded until the original (static) water level was attained.
The pumping test data were analysed using the Theis (1935) non - equilibrium
formula and the Cooper and Jacob (1946) straight line method.
The VBB (1977), in their study of the water resources in Amman-Zerqa Basin,
conducted a number of pumping tests, reporting 34 values of transmissivity for wells
tapping the B2/ A 7 aquifer system. Their efforts were mainly concentrated on recovery
tests. In a few cases, they used adjacent wells as observation wells. Prior to recovery, the
wells were pumped at constant rate long enough to reach steady state. The duration of the
tests ranged between 10 and 120 hours, and in several cases up to two or more weeks.
They used the Cooper and Jacob (1946) method for estimating the transmissivity of the
aquifer. The reported transmissivity refers to the analysis of measurements in the
observation wells.
Howard Humphreys (1986) examined the hydraulic parameters of the B2/A7
aquifer system in Jafr Basin by aquifer tests in 8 fully penetrating exploratory boreholes,
129
7 of which had observation boreholes constructed at approximately 20 m from the
pumped well. The tests comprised step-recovery tests and constant yield tests lasting
from 2-8 days.
The pumping test data were analysed by Theis, Jacob, Boulton and Hantush type
curve methods for confined, water table and leaky conditions, respectively.
The BGR (1987), during their study of the hydrogeology of the EI Lajun area in
Wadi Mujib Basin as phase No.1 of the EI Lajun Oilshale Feasibility study, conducted 15
pumping tests from which they reported 12 values of transmissivity and 5 values of
storage coefficient (since only 5 boreholes were equipped with observation wells). Each
test included a step-drawdown test and constant yield test for 48 and 72 hours
, respectively; one borehole was pumped for a 7 day period and another borehole was
selected to conduct a long-term pumping test with a duration of 18 days.
JICA (1987) selected four areas in Wadi Mujib, including Rumeil (Tl), Khan EI
Zabeeb «T2), Siwaqa (T3) and Qatrana (T4), for carrying out pumping tests. Pumping
tests were performed to estimate the aquifer parameters of the B21 A 7 aquifer system. To
assess the storage coefficient (S), two observation holes were installed at 20 m from the
Tl and T3 wells. Data analysis was performed using the modified equilibrium equations
for the pumping tests without observation holes, and the conventional non-equilibrium
equations for the pumping tests with observation holes.
JICA (1990), in their water resources study of the Jafr Basin, conducted 4
pumping tests in 4 selected areas in the study area, including JT 1, JT2, JT3 and JT 4 to
estimate the aquifer parameters in the major aquifer systems ofB2/A7 and AI-6. The JTl
and JT3 pumping tests were carried out in an area with static water levels of less than 200
m below ground surface. A step-drawdown test with five steps was performed in test well
JT3 to estimate the well efficiency, and a constant yield pumping tests were carried out
for 72 hours in both JT 1 and JT3 to estimate aquifer parameters. In JT2 and JT 4 where
the static water level is deeper than 200 m and the aquifer is poorly yielding, pump-in
tests were performed to estimate the permeability. The rate of constant injection was 36
litres/minute for two hours. Data analysis used the same method as in the Wadi Mujib
study (JICA, 1987).
130
5.3.2 RESULTS OF PUMPING TEST ANALYSIS
Due to the variation in lithology, diagenetic and structural phenomena, the aquifer
properties are extremely variable. As summarised above, a considerable number of wells
are reported to have been tested, but most of the tests are of short duration (usually
around 72h), and/or the test wells do not fully penetrate the aquifer, and/or no observation
wells are available . Furthermore, the pumping test data have been evaluated with the .
help of the Theis method, which assumes that the aquifer is confined, homogeneous,
isotropic, and receives no recharge from any source, therefore, the pumped well fully
penetrates the aquifer, and water is instantly released from storage with reduction of head.
It is clear that the aquifer departs radically from Theis and Jacob model.
Nevertheless it is believed that the non - equilibrium analysis provides a useful first
estimate of transmissivity, bearing in mind the regional nature of the study for which the
data were required.
A summary of aquifer properties produced by the tests is given in Table (5.2)
Examples of time-drawdown and time-recovery plots are shown in Figure (5.1).
Basin No.of SC(m'lh!m) T(m~lh) k(m/h)
Tests Range Mean Median Range Mean Median Range Mean Median
Amman 72 0.08-551 47.8 14.9 0.035-306 23.8 4.17 0.0004-36 1.38 0.1
W.Mujih 46 0.05-792 50.5 3.85 0.018-2346 103 4.6 0.0001-21.9 1.02 0.43
Jafr 33 0.03-138 26.5 8.18 0.01-1435 92.5 21.1 0.00005-45 2.01 0.17
All areas 151 0.03-792 43.9 9.1 0.01-2346 63 6.1 0.00005-45 1.41 0.09
Table (5.2 ) Summary of pumping tests results in the study area
5.3.2.1 SPECIFIC CAPACITY
Specific capacity is the term expressing the productivity of the wells; the larger
the specific capacity, the better the well. Specific capacity (SC) is defined as:
SC= QI s .............................................................................................. (5.1)
where Q= the pumping rate in m3/hour
s = the drawdown in the well in m.
131
II)
~ E :e B .5
~ 0
~ 0
UI e ., E .5
~ ~ 0
UI
i .5
~ l 0
II)
~ ., E .5
~ l 0
50
55
60
65
70
75
80
8
9
10
11
12
13
14
15
16
4
6
8
10
12
14
16
18
a 0.5
1
1.5·
2
2.5
3
3.5
4
4.5
5
I I 1111111 WeillNo.S60,1111 IQ=t04Iml'lhlll I I I I t I III I I I t I I I I I I I I I I I t I
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10 100 1000 10000 Time since pumping started in minutes
I I 1111 Well,Noj 565 11111 Q=Silm'i1h I I I III
- - --,- -1- ~ -:-~~~:~ - - - - ~ - -:- -:- ~~~~ ~~ - -S:~h.~~1rf1-:-:-~:~ - - I- +- 1-1-+ - - - - f- - -1- -1- + -i -t -i t- f- - - - -t - - -t - +- -1- J- t-H
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____ : __ ~~1:-15~~!~niiJ~~ ____ L _. I :_ ~ ~~~ ~~ ___ ~ __ ~ _~_:_:_~:~ I I I I I I I II I Itt I I I 1'1
____ 1 __ .1_ .l_I_L LLI.1 ____ L __ 1 __ I_.L...J...J __ ...J __ .l_.L_I_I_I_1..J
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: : : : : :::: :::: ::::: .. ~.~:::::: - - - -,- -1-T-,-r 1,-,1- ---1- -1- -'-I 11111--- - -,- ·T - 1 -'-1-1-'-1
I , I I I , , I I , I , I , 1 I II
10 100 1000 10000
Time since pumping started in minutes
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10 100 1000
Time since pumping started in minutes
0.1 10 100 1000 10000 Time since pumping started in minutes
Figure (5.1) Examples of pumping test data analyses
132
Specific capacity depends on, or reflects aquifer transmissivity and storage coefficient.
Hence, a large specific capacity indicates an aquifer to have a large transmissivity, and
vice versa.
Pumping test results (Table 5.2) show that the specific capacity varies widely
within the basin and between basins. It ranges between 0.03 and 792 m2/h1m with a mean
of about 44 m2/h1m. The frequency distribution of the specific capacity (Table 5.3) shows
that, in 51 % of the tests the specific capacity ranges between 0.03-10 m2/h1m, i.e. the
median is within this range. Only 21 % of the tests exceeded the mean. Specific capacity
is proportional to transmissivity for the wells tested; however, well losses owing to the
different types of well construction, methods of development and existing condition of
wells may have a significant effect on the range of specific capacities observed. In some
cases a decline in specific capacity during the pumping test was observed, which is
attributed either to a reduction in transmissivity due to lowering of the groundwater level
in the unconfined aquifer, or to an increase in well loss associated with the clogging or
deterioration of the well screen.
An attempt was made to relate specific capacities to transmissivities for the wells
tested, and then to estimate the aquifer transmissivity from the specific capacity.
Range Amman-Zerqa Wadi Mujib Hasa & Jafr All areas O-lO 41 61 55 51
10-20 19 9 15 15
20-30 6 4 6 5 30-40 3 9 6 5 40-50 4 3 0 3 50-lO0 14 3 9 9 >lOO 13 11 9 12
Table (5.3) Frequency distribution of specific capacity from pump tests (%)
5.3.2.2 TRANSMISSIVITY AND PERMEABILITY
Transmissivity (T) is a measure of the ability of an aquifer to transmit water. It
depends on the hydraulic conductivity (K) and the saturated thickness of water bearing
133
material. Various values of transmissivity and permeability were estimated by different
authors using different techniques (Appendix Cl).
The transmissivities calculated from the aquifer tests differ widely from well to
well with values ranging from 0.01-2346 m2/h (Table 5.2). Despite the great range in
transmissivity values, the frequency distributions, show that the majority of the tests give
transmissivities of less than 5 m2/h in the Amman-Zerqa and Wadi Mujib areas
(Table 5.4). In the Hasa and Jafr Basins, the transmissivity distribution is more uniform.
This may reflects the nature of the aquifer system in these areas: it has been shown in the
previous chapters that the aquifer is dominated by a sandy facies in the Jafr Basin, and
therefore the transmissivity depends more heavily on the primary permeability.
The variation in transmissivity may be due to differences in thickness of aquifer
penetrated by the wells, or to differences in permeability. However, discounting the factor
of saturated thickness by calculating the permeability for the section penetrated by the
wells, and the lack of any direct relationship between transmissivity and the saturation
thickness, suggests that the wide variation in transmissivities is mainly attributable to the
permeabilities. The calculations show a wide variation in permeability (0.00005-45 mIh),
with mean and median values of about 1.41 and 0.09 mIh respectively. The frequency
distribution of the permeability (Table 5.5) indicates that the percentage of values having
a permeability ofless than 0.01 mIh is higher in Wadi Mujib (30 %) and Amman-Zerqa
Range Amman-Zerqa Wadi Mujib Hasa & Jafr All areas 0-2.5 41.67 45.65 17.65 37.5 2.5-5 13.89 6.52 8.82 10.53 5-10 9.72 6.52 11.77 9.21 10-15 4.17 2.17 5.88 3.95 15-20 8.33 4.34 5.88 6.58 20-30 6.94 4.34 14.71 7.9 30-40 0.0 6.52 5.88 3.29 40-50 0.0 2.17 0.0 0.66 50-100 8.33 8.7 8.82 8.55 >100 6.95 12.51 20.59 11.84
Table (5.4) Frequency distribution of transmissivity from pump tests (%)
134
areas (11 %) than in the Jafr Basin (only 6 %). The higher penneabilities in the Amman
Zerqa area compared with the Wadi Mujib area may be due to the presence of the shallow
alluvial aquifers in direct hydraulic continuity with the underlying B2/ A 7 aquifer system.
The greater degree of karstification in the Amman-Zerqa area may also be important,
since the aquifer system in that area is mainly of the limestone of the Wadi Sir (A7)
Fonnation, while in the Wadi Mujib area the aquifer consists of the silicified Amman
(B1I2) and Wadi Sir (A7) Fonnations. The effects of sandstone on the penneability of the
aquifer system is obvious in the results from the Jafr Basin. 65 % of the penneability
samples in Jafr Basin lie in range ofpenneability between 0.01-0.5 m1h, compared with
61 % lying in a range of 0.0-0.1 mIh in Wadi Mujib.
Range Amman-Zerqa Wadi Mujib Rasa & Jafr All areas 0-0.01 10.61 30.43 5.88 15.75 0.01-0.05 22.73 23.9 17.65 21.92 0.05-0.1 16.67 6.52 17.65 13.7 0.1-0.5 25.76 19.57 29.41 24.66 0.5-1 7.58 2.17 8.82 6.16 1-1.5 3.03 8.70 0 4.11 1.5-2 1.52 2.17 2.94 2.05 2-5 7.58 2.17 14.71 7.53 >5 4.55 4.34 0 4.09
Table (5.5) Frequency distribution of permeability from pumping tests (%).
5.3.2.3 VERTICAL HYDRAULIC CONDUCTIVITY
The values for vertical hydraulic conductivity (VC) are largely unknown.
Application of aquifer testing methods for estimating VC is difficult. However, when
marl interbeds are present, the interbeds restrict vertical groundwater movement, and the
large scale VC can then be estimated on the basis of the hydrogeological characteristics
of the interbed lithology. Otherwise, VC may only be estimated by a numerical
groundwater flow model simulation. Such estimates represent the integrated effects of the
limestone beds and the interbeds. From previous numerical groundwater modelling
studies, only BGR (1987), reported values for vertical hydraulic conductivity: in the
B2/A7 aquifer system they present a value of about 0.00036 m1h; for the Al-6 aquitard
135
they suggest a range from 0.000016-0.00005 mIh, depending on the degree of fracturing
and geol~gic structure. The ratio of horizontal to vertical hydraulic conductivity has been
estimated to be about 0.01 for the B2/A7 system and a range of 0.012 to 0.047 for the A1-
6 system.
5.3.2.4 STORAGE COEFFICIENT
Storage coefficient is the amount of water that can be released from or added to
the groundwater reservoir. It is usually defined as the volume of water an aquifer system
releases from or takes into storage per unit surface area of aquifer per unit change in head
(Lohman et aI., 1972). In the zone of water table fluctuations, the storage coefficient is
virtually equal to the amount of water released from storage by gravity drainage, referred
to as specific yield. Below the zone of water table fluctuations, the storage coefficient is
the amount of water released by compression of the sediment and expansion of the water.
This amount is usually much less than the amount released by gravity drainage.
CALCULATING CONFINED STORAGE COEFFICIENTS
The specific storage (8s ) was described by Jacob (1940) on the basis of the
compressibility of the skeleton of the aquifer and the expandability of water, and can be
described by the following equation:
8 s = p g(a + n~ ) ...................................................................................................................... (5.2)
where p = the specific weight of water (kg/m3)
g = the acceleration due to gravity (N/m3
)
a = the compressibility of aquifer (m2/N)
n = the porosity
~ = the compressibility of water (m2/N)
Taking the value for porosity as 0.1 , the aquifer compressibility as 10'9 m2/N, the
compressibility of water as 4.4*10,10 m2/N (Freeze and Cherry, 1979), and the specific
136
weight of water as 1000 kglm3, the specific storage was calculated to be 1.0* 10-5 m-l
•
Original sources of compressibility data include Poland (1961), Domenico and Mufflin
(1965), Johnson et al. (1968), Riley and McClelland (1972), and Helm (1978). The
storage coefficient (S) of the B2/A7 aquifer system was then estimated by multiplying
specific storage by the estimates of aquifer thickness. The estimated minimum, median,
and maximum values of storage coefficient were found to be 2.05*10-5, 9.62*10-4
, and
3.73*10-3 respectively. These estimated values are within the range of storage coefficients
estimated from aquifer tests in the confined B2/ A 7 aquifer system, and consistence with
the range (0.00005-0.005) reported by Freeze and Cherry (1979) for confined aquifers.
ESTIMATING STORAGE COEFFICIENTS FROM PUMPING TESTS
Because few wells in the study area have nearby observation wells, reliably
estimated storage coefficient values are rare. Only 23 wells in the study area are reported
to have been tested for estimation of the storage coefficient of the B2/ A 7 aquifer system.
The drawdown and recovery data were analysed using the Theis (1963) and Cooper and
Jacob (1946) graphical methods. The results of these analyses are summarised in Table
(5.6).
The storage coefficients computed by Mudallal (1973) from data recorded from
three pumping tests (AI22-AI24) in the Amman-Zerqa area range from 0.004 to 0.7.
Parker (1970) -conducted three pumping tests with observation wells in Wadi
Mujib (S83) and Hasa Basin (S59 and S61A). The most comprehensive test was carried
out on well (S59). The well was pumped at an average rate of 180 m3/h for 72 hours, and
drawdown and recovery data were recorded from the pumping well and from three
observation wells (S40, S60, and S61A). The pumping well (S59) was located in the
confined part of the aquifer, while the three observation wells were located at distances in
the unconfined part. Thus a cone of pressure relief was developed during the early period
of the tests and water was provided from the confined storage. Later the cone intercepted
the free water table and the well then drew a proportion of the water from the unconfined
aquifer. This interaction between confined and unconfined conditions during the test
gives an underestimate of the specific yield of the aquifer. Two sets of results are also
137
obtained in 1987 from pumping tests analysis conducted by the BGR (LA series) and
fleA (T series) in the Wadi Mujib Basin (Table 5.6).
Howard Humphreys (1986) conducted comprehensive pumping test analyses in
the Jafr Basin to estimate the specific yield and storage coefficient of the B2/A7 aquifer
system. As shown in Table (5.6) the values of specific yield and storage coefficient
obtained from the tests found to range between 0.2 % and 13.33 % with an average value
of2.66 % and between 0.00001 and 0.03 with an average value of 0.006 respectively.
Pumping Observation Distance T S Sy Well Well (m) (m2/d) (%)
Al22(Zerqa 5) Z50bs. 150 660 0.014 A123(Khaw) Khaw Obs. 22 154 0.7 AI24(W.Rimam) W.Rimam Obs. 50 2640 0.004 S59 S40 783 7200 0.0085 1.3 S59 S60 1090 0.0055 0.3 S59 S61A 1100 1375 0.005 0.6 S61A S40 400 0.003 0.2 861A 861 4 718 3.9 883 844 49 1.9 LAI LAIA 68 41 0.00002 LA2 LA2A 30.2 1440 0.3 LA4 LA4A 41.1 588 0.03 LA9 LA9A 26.1 23 0.00034 LAB LA7 2600 88 0.00089 LA13 LA9 3800 124 0.00001 Tl TI0 20 373 0.0265 T3 T30 20 73 0.0009 SH5 PHOlO 20.22 863 0.00195 1.9 PHT5 PH05B 20.22 569 0.00235 4.8 PHT9 PH09B 20.34 301 0.000805 PHTII PHOIIB 19.92 1731 0.00914 PHT14 PH014 20.2 186 0.000885 PHT15 PH015 19.45 575 0.0024 13.33 PHT16 PH016 20.31 38 0.00076
Table (5.6) Storage coefficient and specific yield from pumping tests
138
It is considered that the estimates of specific yield obtained from the test are of the
right order of magnitude for the short-term storage of a highly permeable zone in this
aquifer system (Table 5.6). It is believed that the effects of delayed drainage were not
fully realised during the relatively short periods for which the wells were pumped. The
long-term specific yield of the aquifer would be somewhat higher than the results of the
pumping tests indicated. Bearing in mind the variability of the B2/ A 7 aquifer system, it
seems probable that zones of greater or less effective porosity occur at different levels
within the aquifer.
A simple method for determining specific yield (Sy) by the pumping test method
of Remson and Lang (1955) as modified by Ramsahoye and Lang (1961) to reduce the
time necessary to compute Sy was applied in this study. The method consists of
computing the volume of dewatered material in the cone of depression and comparing it
with the total volume of discharge water. The data from long-duration pumping tests for
sites where observation boreholes exist were used. The assumptions are made that all the
water is pumped from storage, and that the cone of depression has almost reached an
equilibrium shape.
Ramsahoye and Lang (1961) calculated the volume of dewatered material. (V)
within the cone of depression by:
Qr2 Ts logY = log-+5.45- ...................................................................... (5.3)
4T Q
The specific yield (Sy) is the water pumped during the test divided by the gross volume
of dewatered materials within the cone of depression:
S = Qt ................................................................................................. (5.4) y V
where Sy = specific yield
V = the volume of dewatered material in (m3)
139
Q= the discharge rate ofthe pumped well in (m3/day)
T = the transmissivity in (m2/day)
r = the horizontal distance from the pumped well to measuring
points in the cone of depression in (m)
s = the drawdown at distance r in (m)
t = the time since pumping began in (days)
The above formula was used to compute the specific yields in the study area
assuming that the drawdowns were stabilised after only one day of pumping. The
calculations were carried out using the observed and a hypothetical possible range of
transmissivity between 100 and 1000 m2/d. The results (Table 5.7) were found to be
broadly consistent with the results obtained from the pumping tests analysis (Table 5.6).
Pumping Observation Distance Time Ddawn Q T Sy Sy Sy Sy Well Well (m) (day) (m) (m3/d) (m2/d) (T=observ) (T=100m2/d) (T=500m2/d) (T=1000m2/d)
AI22 Z50bs. ISO 7 0.54 11440 660 0.005 0.011 0.008 0.002 AI23 KhawO. 22 3 1.2 1008 154 0.13 0.18 0.002 <IO'~
AI24 W.Rimam 50 4 0.275 2880 2640 0.18 0.14 0.44 0.5 S59 S40 783 3 0.11 4320 7200 0.005 0.0001 0.0005 0.005 S59 S60 1090 3 0.11 4320 0.00007 0.0002 0.003 S59 S61A 1100 3 0.11 4320 1375 0.003 0.00007 0.0003 0.002 S61A S40 400 2 0.14 2660 0.003 0.009 0.01 S83 S44 49 I 0.07 1986 56300 <10' 0.11 0.09 0.02 LAI LAIA 68 3 3.8 2160 41 0.014 0.01 <10' <10-' LA9 LA9A 26.1 7 5.7 1651 23 0.05 0.008 <10- <10-LA 13 LA7 2600 18 1.07 1058 88 0.00002 0.00002 <10- <10-LA 13 LA9 3800 18 0.68 1058 124 0.00001 0.00001 <10- <10-SH5 PHOIO 20.22 4 0.67 2134 863 0.18 0.60 0.50 0.10 PHT5B PH05B 20.22 5 1.4 2678 569 0.13 0.50 0.18 0.014 PHT5A PH05A 20.2 5 3.61 1693 319 0.0006 0.07 <10-' <IO-~
PHT5(A+B) PH05A 20.2 7.02 4.92 3473 444 0.002 0.2 0.0007 <10-
PHT5(A+B) PH05B 20.2 7.02 2.805 3473 444 0.05 0.4 0.03 0.0004 PHT9 PH09B 20.34 2 1.27 1564 301 0.14 0.4 0.03 0.0004 PHTl5 PHOl5 19.45 2.6 0.72 1296 575 0.11 0.50 0.2 om PHTl5 PH015 19.45 4 0.69 1201 575 0.10 0.50 0.15 0.008 PHTl6 PHOl6 20.31 2 6.71 473 38 0.0004 <10-> <10- <10-
Average 0.05 0.17 0.08 0.03
Table (5.7) Storage coefficient and specific yield calculated by Ramsaboye and Lang method.
140
It is believed that the regionallong-tenn specific yield of this system might range
from 1 to 10% depending on the degree of karstification. In the south and southeast, the
lithology is dominantly arenaceous which would suggest a higher value -perhaps between
10 and 15%. Because of the uncertainty, simulated values by groundwater flow model
(Chapter 8) may provide better estimation for the regional specific yield and storage
coefficient values.
5.3.2.5 DISCUSSIONS
Due to the heterogeneity of the B2/ A 7 aquifer system it is expected that
transmissivity and penneability values at certain locations depart widely from the
average. Penneability of the limestone ofB2/A7, as in many carbonate rocks, is provided
by solution enlargement of bedding planes and joints. Thus, original penneability was
controlled by the intensity and direction of the joint patterns and the degree to which the
joints are open. Highest penneabilities probably occurred in the vicinity of large jointing
and the zones of greatest tectonic disturbance in the area. The penneability provided by
the joint pattern has been enhanced by solution, and the present pattern of penneability
appears to be closely related to the degree of karstification, which is in tum related to the
quantity of water flowing through any part of the system. In general karstification
increases with the volume of flow, providing the water chemistry will allow the solution
of the calcium carbonate. The water must be acidic. Where the limestones tend to be
silicified or sandy, as in the upper part of the aquifer system and in the south and
southeast, the degree of karstification becomes low. In such cases, penneability is
affected primarily by the existing fracture pattern. Nevertheless, in some areas
particularly in the Amman-Zerqa area where the B2/ A 7 aquifer system is in hydraulic
continuity with the overlying alluvium, high transmissivity values are more related to the
alluvium than the underlying carbonate aquifer especially when the tested wells are
shallow and the saturated thickness do not exceed 25 meters.
141
Although aquifer homogeneity and isotropy generally do not exist under field
conditions, local aquifer test results can approximate regional conditions, particularly in
unconsolidated clastic aquifers where secondary permeability development is minor. In
carbonate systems, however, aquifer transmissivity is controlled largely by secondary
permeability, and regional trends in transmissivity may not exist. Examples of the degree
of areal variation of aquifer parameter values in the B2/ A 7 aquifer system are seen in the
long-term pumping test data in number of the tested wells, and extreme variations in
transmissivity over a short distance are observed. The anomalous results are site-specific
and mayor may not apply to other locations or situations. Therefore, odd high values of
aquifer parameters in certain areas are taken to indicate only local aquifer characteristics
and have been eliminated from further considerations in the assessment of the regional
distribution of transmissivities. For example, the average calculated transmissivity is
about 63 m2/h while the results show that 80% of the tests have transmissivities less than
50 m2/h and only 18 % of the tests have transmissivities exceeding 60 m2/h. In 76% of
the wells tested the mean permeability of the penetrated sections was in the range of
<0.01-0.5 mIh. This figure was exceeded in 24 % of the samples, from which only 4 %
of the tests indicate a permeability higher than 5 mIh.
Some of the test data showed evidence of lateral changes in permeability in the
vicinity of the wells. This is revealed by a change in slope of the log-normal plot of time
against the drawdown or recovery data. Changes in transmissivity by a factor of two or
three were not uncommon. The phenomenon is illustrated by the time-draw down graphs
ofWaheida well No.3 (S60) and Jarba well No.1 (S65) (Figure 5.1). Judeida well No.1
(S71) was tested when it had penetrated 76 m of saturated section and again when it had
been deepend a further 108 m: the mean permeability of the sections penetrated did not
change. Similar results were obtained from tests on Jafr well (SI5) (Parker, 1970). In
these cases the transmissivity of the section penetrated was almost directly proportional to
the thickness of aquifer penetrated. However, Sultani well No.2 (S66) was tested when it
had penetrated 34 m of aquifer and again when 96 m had been cut. In this case the
transmissivity increased from about 0.833 to 33.33 m2/h and the mean permeability from
about 0.025 to 0.33 mJh (Parker, 1970). The results of tests on S66 illustrate the range of
142
penneability which may occur within a section of the aquifer system and emphasise the
error which might be involved in using data from partially penetrating wells.
In many of the single well tests the drawdown stabilised within very short times
(few minutes) after pumping started and the recovery was almost instantaneous .. Such
drawdown and recovery readings, and because of the assumptions inherent in the
methods, made it impossible to calculate transmissivity by the non-equilibrium fonnula.
And since, the transmissivities derived from real aquifer test are minority in the data set
and represent only the test area, apply results from aquifer tests over a large area found to
be difficult. Therefore, attempts were made to calculate transmissivity from the
distinctive relationship between transmissivities and specific capacities.
5.4 ESTIMATION OF T FROM se. Transmissivity is often estimated from specific capacity data because of the
expense of conducting standard aquifer tests to obtain transmissivity and because of the
relative abundance of specific capacity data. Most often, analytic expressions relating
specific capacity to transmissivity derived by Thomasson and others (1960), Theis
(1963), Brown (1963), and Bradbury and Rothschild (1985) are used in this analysis.
Razack and Huntley (1991) demonstrate that turbulent well loss produces a poor
correlation between measured transmissivities and those estimated from specific capacity
from the above relations. This study focuses on a comparison between transmissivity and
specific capacity of wells completed mostly as open boreholes in fractured carbonate
aquifer systems, where turbulent well loss may be less important.
The theoretical relations between specific capacity and transmissivity have been
derived by solving the Dupuit-Thiem equation for transmissivity as a function of well
specific capacity (Q/s) (Thomasson and others, 1960):
(QI s) R T= In- ...................................................................................... {5.5)
21t r
where s = drawdown in the pumping well (m)
143
Q = discharge rate (m3/h)
T = aquifer transmissivity (m2/h)
R = radius of influence ofthe well (m)
r = radius of the well (m).
This approach results in a linear relation between transmissivity and specific
capacity of the form:
T = C(Q I s) ............................................................................................ (5.6)
Using a radius of influence ranging from 300-3000 ft, they noted that the constant, C in
equation (5.7) should range from 1460-1990. For pumping tests in valley fill sediments in
California, they noted that C varied from 1300-2200, and averaged about 1700,
corresponding to constants of 0.9, 1.5, and 1.2, respectively, for self consistent units of
specific capacity and transmissivity.
Theis (1963) and Brown (1963) used the Theis nonequilibrium equation to derive
similar relations between transmissivity and specific capacity for unconfined and
confined aquifers, respectively. Using 24 hour specific capacity data to illustrate the
approach, they arrived at similar range of the constant C, as Thomasson and others (1960)
with lower values corresponding to unconfined aquifers and higher values corresponding
to confined aquifers. Theis analysis is based on the Jacob equation:
T = 2~!; 10g( 2:;;t) ........................................................................... (5.7)
where T = transmissivity (L2/T)
Q = discharge (L3/T)
s = draw down in the well (L)
t = pumping time (T)
S = storage coefficient ( dimensionless)
144
r = radius of the well (L)
One complication not taken into account by the above analysis is the effect of
turbulent well loss in the well bore and gravel pack. The gravel pack and well screen
increase entrance velocities, which often produces turbulent flow. Jacob (1947) suggested
that total drawdown in a borehole is given by:
s = BQ + C t Q2 ....................................................................................... (5.8)
Where B = laminar head loss coefficient
C t = turbulent head loss coefficient
The introduction of this turbulent well loss term decreases the specific capacity of
a well for a given transmissivity. If this is not taken into account, transmissivities will be
underestimated from the corresponding specific capacity values.
Data from 116 wells completed in the B2/A7 aquifer system (Appendix Cl) were
used to explore the relation between transmissivity and specific capacity. Most of the
transmissivities were calculated using the slope of the time-drawdown and/or recovery
data. In some wells where more than one transmissivity value was calculated, the average
was used in the analysis. The drawdown figure applied to calculate specific capacity was
that recorded after prolonged pumping when the drawdown was assumed to have
stabilised. No account was taken of whether the data was recorded from a well tapping an
aquifer under confined or unconfined conditions.
Figure (5.2) is an arithmetic plot of transmissivity versus specific capacity for the
tested wells; it also shows the theoretical relationship between transmissivity and specific
capacity as predicted using the steady state approach of Thomasson et al. (1960). It shows
that the measured transmissivities are generally greater than the average of those
predicted by the theoretical relations. Unlike the wells completed in alluvium, the
deviation between the observed and theoretical relations cannot be explained by the effect
of turbulent well loss. Turbulent well loss would increase drawdown in the production
well for a given pumping rate, thereby decreasing the specific capacity of the well. For
145
the fractured-rock, however, the specific capacities are slightly less than would predicted
by the transmissivity. Possible explanations for the deviation between the observed and
theoretical relations might include:
1- The specific capacities were measured either after the water levels have
stabilised or at the end of the pumping test. This results in lower specific capacity for the
same transmissivity.
2- The results may be significantly sensitive to the storage coefficient. Figure
(5.3) shows the theoretical relations between transmissivity and specific capacity plotted
for storage coefficients of 0.0001, 0.001, and 0.01. In this case the Theis equation was
used to calculate the transmissivity. Because specific capacity varies with the logarithm
of liS, the solution is not very sensitive to variations in S.
3- Fractured-rock aquifers are often anisotropic, which modifies the response of
the well to aquifer testing. Neuman and others (1984) modified the Cooper and Jacob
(1947) equation to take into account anisotropic conditions:
s = 4" z;& log [(x,:;):T&T:;, )]s ................................................. (5.9)
• T(measured) 600 - - - - Linear (T=0.9*SC)
i 450
--- Linear (T=1.2*SC)
- - Linear (T=1.5*SC)
• •
o o 100 200 300 400 500
Specific capacity(m/h)
Figure(5.2) Relation between transmissivity and specific capacity
146
where x, y = distance in the x and y principal directions of anisotropy, respectively
T x = transmissivity in x direction
T y = transmissivity in y direction.
In an anisotropic aquifer, analysis of production well drawdown yields a value of
effective transmissivity, (Tx Ty )112, rather than the true transmissivity. However, if the
effective transmissivity and the ratio between the transmissivities in the two principal
directions are known, then Tx and Ty can be calculated:
Tx = (Tx / Ty)12 (Tx Ty f2 ....................................................................... (5.10)
This approach was applied to the range of measured transmissivities and
anisotropy ratios of one to 1000 (Figure 5.4), assuming a storage coefficient of 0.001.
Introducing anisotropy decreases the draw down measured in the production well,
resulting in an increase in specific capacity for a given value of transmissivity. However,
the measured transmissivities are reasonably well correlated with the transmissivities
given by the theoretical relation with specific capacity.
10000.-__________________________________________ ,
• T(measured) --_T(S=0.0001) • 1000 - - - - T(S=0.001)
100 - - T(S=0.01)
10
0.1
• 0.01 • • •
0.001 +-______ -...... ________ ------..--------.------~ 0.01 0.1 10 100 1000
Specific Capacity (m2/h)
Figure(5.3) Effect of varying storage coefficient on theoretical relations between specific capacity and transmissivity.
147
4- Most of the wells, particularly the private wells, penetrate less than the full
thickness of the aquifers. During the specific capacity tests, partially penetrating wells
may yield anomalously low values of specific capacity, depending on the ratio of
penetration ( L) to aquifer thickness (b ). In the study area, the L / b ratio is sometimes as
low as 0.3. Thus, a correction for partial penetration is necessary before estimating
transmissivity from specific capacity. For unsteady drawdown in a partially penetrating
well, Sternberg (1973) shows that:
T~ 2~!~[IO~ 2::;1) + 2sp ] ........................................................... (5.11)
where sp is a 'partial penetration factor' given by Brons and Marting (1961) as:
sp = 1~~~b( In~-G(L/b») ........................................................... (5.12)
where b = aquifer thickness
L = length of open interval
G = a function of the Lib ratio.
Brons and Marting evaluate G(L / b) for a various values of (b / r). Bradbury and
Rothschild (1985) found that the following equation:
G(L / b) = 2.948 - (7.363L / b) + 11.447(L / b)2 - 4.675(L / b)3 .................... (5.13)
fitted the data of Brons and Marting, with a correlation coefficient of 0.992.
148
:2 i;r
.s
.?;-.s: ·iii f/)
.!!l E f/) c: jg
10000
1000 • T(measured)
T(Tx/Ty=1)
100 __ T(Tx/Ty=1000)
10
0.1
• 0.01 •
0.001
0.01 0.1 10 100
Specific capacity (rrf/h)
Figure(S.4) Effect of aquifer anisotropy on theoretical relations between transmissivity and well speCific capacity.
•
1000
The effect of assumed partial penetrations ranging between 0.3 and 0.9 on the
theoretical relation between specific capacity and transmissivity for the B2/ A 7 aquifer
system is shown in Figure (5.5). The observed transmissivities are correlated well with
the transmissivities calculated by the theoretical relation with specific capacity for partial
penetration ratio of 0.9.
5- The presence of relatively open fractures may act to effectively increase the
radius of the production well, thereby decreasing drawdown and increasing specific
capacity. Gringarten and Witherspoon (1972) analysed the problem of drawdown in a
well which is drilled through a vertical fracture extending a distance Xf in both the
positive and negative direction along the x axis. The permeability of the fracture
intersecting the well is assumed to be sufficiently high that the drawdown is everywhere
the same along the fracture. Drawdown in the fracture, and therefore the well, is
controlled by the ability of the surrounding rock to transmit water to the fracture that
intersects the borehole. The early response to pumping is therefore controlled by the
storage coefficient of the surrounding fractured rock. Thus, a vertical fracture acts to
extend the radius storage capacity of the borehole.
149
10000
1000
:2 100 ;;r
.s
.?;-:~ 10 (f)
.!Q E (f) c: ~ I-
0.1
0.01
0.01
• T (measured) __ T(Ub=0.3)
- - T(Ub=0.6)
- - - - T(Ub=0.9)
• • 0.1
• 10
Specific capacity (n1/h)
100 1000
Figure{S.S) Effect of partial pentration on theoretical relation between aquifer transmissivity and well specific capacity
For a well intersecting a single, plane, vertical fracture in an otherwise
homogeneous, isotropic, confined aquifer, Gringarten and Witherspoon (1972) and
Gringarten and Ramey (1974) obtained the following general solutions for the drawdown
in the pumped well:
s= ~F(uvf ) ..................................................................................... (5.14) 47tT
where
F(u,,)= 2'/1tUvreJ kJ -Ei(-_I_J ................................................... {5.15) 112 uvf 4Uvf
and
. x -u
-Ei{-x) = J~du = the exponential integral ofx ............................. {5.16) o u
where
Tt uvf = -- .......................................................................................... {5.17)
Sx/
150
S = storage coefficient of the aquifer
T= transmissivity of the aquifer
x f = half length of the vertical fracture.
At early pumping times, when the drawdown in the well is governed by the horizontal
parallel flow from the aquifer into the fracture, the drawdown can be written as:
Q s= -F(uvf ) ..................................................................................... (5.18) 47tT
where
F(uvf ) = 2J7tUvf ................................................................................. (5.19)
or
log F(uvf ) = 0.5 log (uvf) + cons tan t ................................................ (5.20)
and consequently
s~ 2~1t~Xf2 Jt ................................................................................ (5.21)
or
log s = 0.5 log (t) + constant ............................................................. (5.22)
Gringarten and Ramey (1974) produced a log-log plot type curve F(Uvf) versus
uvf. They demonstrated the early-time parallel-flow period is characterised by a straight
line with a slope of 0.5. The parallel-flow period ends at approximately uvf = 1.6*10-1
(Gringarten and Ramey, 1975). If the aquifer has a low transmissivity and the fracture is
elongated, the parallel flow period may last relatively long. The pseudo-radial flow period
starts at uvf =2 (Gringarten et al. 1975). During this period, the drawdown in the well
varies according to the Theis equation for radial flow in a pumped, homogeneous,
isotropic, confined aquifer, plus a constant and can be approximated by the following
expression (Gringarten and Ramey, 1974):
151
2.3Q 16.59Tt s=--log 2 .......................................................................... (5.23)
41tT SX f
and hence
2.3Q 16.59Tt T=-log 2 ........................................................................... (5.24)
41ts SXf
This approach was applied to the range of B2/A7 transmissivities, with X f
ranging between 50 and 500 m (Figure 5.6), assuming a storage coefficient of 0.001. The
use of the Gringarten and Witherspoon equation results in a relatively good correlation
between the theoretical and observed relations between transmissivity and specific
capacity for fracture halflengths ofless than 50 m.
10000
• T (measured)
1000 - - T(xf=250m)
• • • • T(xf=500m)
2' 100 T(xf=50m) 0J-
.§. 10 Z.
:~ (J) (J)
'E (J) c: ~ 0.1 l-
• 0.01
0.001 ..
0.01 0.1 10 100
Specific Capacity (rrf/h)
Figure (5.6) Effect of a vertical fracture on theoretical relations between aquifer transmissivity and well specific capacity
152
•
1000
5.4.1 APPLICATION OF THE METHOD
The previous discussion provides an explanation for the relation between the
measured pairs of transmissivity and specific capacity and the relations predicted by the
theoretical relations.
Nevertheless, the scatter of data about the line is not excessive, it is considered
that an acceptable estimate of transmissivity can be obtained from specific capacity data
if the well loss is small and if there are no marked barrier conditions in the vicinity of the
well. The method was employed to compute the transmissivity of the B2/ A 7 aquifer
system.
For the data set obtained from the pumping tests results conducted in the B2/ A 7
aquifer system, the relationship was found to be reliable with a correlation coefficient of
about 0.95. Thus, the previous equations allow the determination of transmissivity from
specific capacity. In the more general case of unknown anisotropy, fracture length, and
storage coefficient, transmissivity could only be estimated from specific capacity using an
iterative approach. Although, the previous possible corrections do not appear to account
for discrepancies in the relations between specific capacity and measured transmissivity
using the theoretical relations of Theis (1963).
The log-log transformation of the data set (Figure 5.7), improves the correlation
coefficient on the one hand, and produces a normally distributed dependent variable
required by the normal regression on the other. The equation for the log-log regression
line is:
T = 1.0566(Q / s) 1.0655 ........................................................................... (5.25)
where T = estimated transmissivity (m2/h)
Q / s = specific capacity (m3/h1m)
It should be noted that the values of the regression coefficients are specific for the
units of transmissivity and specific capacity used in this analysis and for the data tested
from the B2/A7 aquifer system in Jordan.
153
Correlations between the estimates of transmissivity and hydraulic conductivity
from the specific capacity data and the values estimated using pumping tests are good,
with correlation coefficients of 0.96 and 0.98 for transmissivity and hydraulic
conductivity, respectively. Table (5.8) gives a statistical summary of transmissivity and
hydraulic conductivity estimates for a 95% confidence level for the mean. The table
shows that using many data points, the specific capacity estimates give a lower mean
hydraulic conductivity and less variation as indicated by the low standard deviation
values. However, the median values give better measures of the central tendency of the
data.
In spite of the well known difficulties in estimating transmissivity and hydraulic
conductivity from specific capacity data, the values obtained correlate well with the
results obtained from pumping tests. As noted by Winter (1981) the standard error in
estimating values of hydraulic conductivity is often close to 100% or even higher. Thus
the ranges of values shown in Table (5.8) are relatively narrow when compared to the
possible range of hydraulic conductivity.
:2 -;;;-
.s
1000
z. 10 ';:; 'Cij I/)
'E I/) c: ~ 0,1 I-
•
0,001 +----____ ,--______ ---.. _______ --1
0.Q1 100
Specific Capacity(n1/h1m)
Figure(5.7) Log-log relation between transmissivity and specific capacity
154
10000
Parameters SC (m'/h/m) T (m"/h) K(m/h) Pumping tests SC data Pumping tests SC data
No. Of Data 116 116 116 110 116 Range 0.03-792 0.01-2346 0.025-1296 0.00008-36 0.0002-35 Mean 37.18 67.6 54.31 l.33 1.18 median 5.60 7.90 6.88 0.093 0.078 Standard Dev. 95.25 235.01 151.26 4.44 4.08 Standard Error 8.88 21.82 14.04 0.42 0.39 Confidence 17.41 42.77 27.53 0.83 0.76
Table (5.8) Statistical results of estimates of transmissivity and hydraulic
conductivity from pump tests and specific capacity for the data used in the
analysis.
This approach is used for estimating transmissivity and hydraulic conductivity
from specific capacity data in the study area. The estimated values are given' in the well
inventory (Appendix AI) . Table (S.9) shows a statistical summary of the calculated
transmissivity and hydraulic conductivity data. Figures (S.8) and (S.9) show the
frequency distribution of calculated transmissivity and permeability.
Comparison between the results and the average hydraulic conductivity for
various materials reported by Freeze and Cherry (1979) shows that the range of values
obtained by this method lie within that of sandstone, silty sand, limestone and dolomite,
and karst limestone, with a mean and a median within the range of the silty sand and karst
limestone (Figure S.IO).
Parameters Saturated SC T K thickness (m) (m3/h/m) (m%) (m/h)
Count 563 405 405 390 Range 2-365 0.2-1500 0.02-2559 0.0002-36.4 Mean 99.18 54.62 84.10 1.47 Median 94.00 3.70 4.26 0.052 Standard Dev. 51.18 163.32 269.73 4.55 Standard Error 2.45 8.12 13.40 0.23 Confidence 4.8 15.91 26.27 0.45
Table (5.9) Statistical results of calculated transmissivity and hydraulic
conductivity from specific capacity.
155
100 100%
90 90%
80 80% ~ 0
70 70% 1)-c:
1)- 60 60% Q) :::l
c: tT
Q) 50 50% ~ :::l U. tT Q)
~ 40 40% ~ u. rn
30 - 30% :5 E :::l
20 20% u
10 10%
0 0% ..- LO LO 0 LO 0 LO 0 0 0 0 0 0 0 0 0 0 0
I N .b ~ ..- N ~ LO 0 0 0 0 0 0 0 0 0 0
0 I
I .b .b ..- N C") ..,. LO <0 ,... co (]) 0 N 0 0 0 0 I I 0 0 I I ..-
N N 0 0 0 0 0 0 LO 0 0 0 0 0 0 0 0 N C") V LO <0 ,... co 0
(])
Transrrissivity (rri/h)
Figure(5.8) Frequency distribution of the calculated transmissivity
70 100%
90% 60
80% ~
50 70% 1)-c:
1)- 60% Q)
40 :::l
c: ~ Q) 50% :::l
u. tT 30
Q)
~ 40% -~ u. rn
20 30% :5 E :::l
20% u 10 10%
0 0%
LO 0 LO LO LO ..- LO LO ~
N C") ..,. LO 0 LO 0 LO 0 N 0 ,... ci N ci I I I
~ ~ N N
0 ci 0 ci 0 9 N C") v .b 0
9 ci ~ .b .b ci .b I .b ,... ..- N
N
0 0 ..- N 0 ci ci 0 0 0 0 ci ci ci ci ci Perrreability (m'h)
Figure (5.9) Frequency distribution of the calculated permeability
156
Rocks Unconsolidated k k K K K deposits (dorcy) (cm2) (cm/s) (m/s) (goJldoyIft2
)
10.5 10.-3 10. 2
Q; 10. 4 10.-4 10. la-I 10.6
:>
II I~ 4 cu ~I c:= 00 - II' c: "'0 0 3 +-mean cu.Q '" ~~ I c: .Q~ 0
-:noc:", .!!! 2 +-median ~ cu o~ ~u oE",v
~I
'fn~'1 >-.-
~oo (I)
~E<II~ ~I :> 0 e:.- <II -Qj.2Ee: .2 gE~.2.2 ... EO", - 10-3 la-II 10-6 10.-8 ~I ~I~] (I)
~I 10-4 10-12 10-7 10.-9 10-2
~I I "0
10.-3 Q.I >:u "'00 10-5 10-13 10-8 10.-10 <11--.s=.v<.:)
~ 0<11 10.-4 <lie:
~§", ~.;: 10-6 10-14 10-9 10-11
~u~ c:o :>.- ° ::lE
la-S _..c~QJ
I 10.-7 10-15 10-10 10.-12 uo.cn-°O:J° ~ E g65 10.-6 ::l .2 c: I 10.-16 10-13 <IIO! 10-8 la-II E'-
I 10.-7
Figure (5.10) Ranges of hydraulic conductivity and permeability for various . geological materials, showing ranges determined from specific
capacity estimates for the Amman-Wadi Sir aquifer system (after Freeze and Cherry, 1979)
5.4.2 AREAL DISTRIBUTION OF PERMEABILITY
To map the regional distribution of the calculated permeability is extremely
difficult because of the wide range of permeability which contains small unmappable
values, and the sparcity of the data points in certain areas. Therefore, an attempt was
made to find an index in which a range of permeabilities can be related. A permeability
index (Pi) was calculated from the following formula:
Pi = 10 g (K x 10 6 ) .................................................................... (5.26)
where Pi = index of permeability,
K = permeability in m1h.
However, it should be noted that multiplication of permeability by 106 and
expression of the permeability index as a logarithmic number are simple devices to avoid
values in small fractions and to produce values which can be easily mapped and
compared. The permeability index for the B2/ A 7 aquifer system is found to range
between 2-8. The relationship between the permeability index (Pi) and permeability
expressed in mIh is shown in figure (5.11); for comparison with the aquifer test data, the
observed values from pump tests analysis were also plotted. Although, a unit difference
in permeability index represents a very wide range of permeability, it is believed that this
method provides a useful device for mapping regional changes in permeability in a
complex aquifer system.
In order to describe the different ranges of permeability according to the
permeability index, the following categories are used:
Pi K(mLh) D~scriptiQn
2-3 0.0001-0.001 very low 3-4 0.001-0.01 low 4-5 0.01-0.1 medium 5-6 0.1-1 high >6 >1 very high
158
8
• Observed value from purrp tests analysis
7
" !,..o' ~ "'~
~ ~'" ~ ... 2
I'"
0.0001 0.001
~ I.
,.; ~ •• •
41
0.01
~
~r. • ......
~'" • • ~4I~
0.1
Permeability (nih)
• ,.. ~~ ~
j:~4
•
Figure (5.11) The relationship between permeability index and permeability for the B2IA7 aquifer system
lJ; ~~
~
10 100
This classification is meant to apply only to the aquifer system in the study area,
and does not necessarily compare with other classifications. For example the range of
permeability between 0.00036-0.36 mIh is regarded by the U. S. Bureau of Reclamation
(1977), as being of medium permeability, while here it covers a wider range, from very
low to high permeability.
Data from 450 wells were used to map the areal distribution of the permeability
using a contour interval of one unit of permeability index (Figure 5.12). The map shows
that the pattern of permeability is complex. In general, low permeability zones are present
in the vicinity of the groundwater mounds; Pi increases downwards with the
groundwater flow gradients. Apparent increase in permeability is noted in areas of flow
convergence such as along the Amman-Zerqa syncline, Wadi Wala, Karak-Wadi Fiha
fault line, Wadi Rasa, around the western end of the Salwan fault line, and in the area to
the west of AIja-Uweina Flexure, where the structural barriers cause flow convergence
(this will be discussed in other parts of the thesis). In the main groundwater development
areas, along a south-north trend in the Central Plateau, the permeability is high (Pi = 5-6).
This pattern of permeability distribution can also be noted from the spacings of the
equipotential line.
159
rt kin
d 40 :2'0
:r- - Wadi
Fault
Pi(K)
7
6
5
4
3
2
160 180 300
Figure (6.12) Areal distribution of permeability in the B2IA7 aquifer system
In the northern area, the permeability is medium except along Amman-Zerqa
syncline, where high to very high permeability is developed. The permeability increases
towards the centre of the Amman-Zerqa areas in the Ruseifa region, while it decreases
eastwards towards the Amman flexure and westwards towards the unsaturated zone of the
aquifer. This can be explained by the increase of the degree of karstification and the
present of alluvial deposits in hydraulic continuity with the main aquifer system.
In the Wadi Mujib and Wadi Hasa basins, the permeability is highly affected by
the structures in the area. In addition to the very· high permeabilities in the water
convergence areas, a zone of high permeability extends in a south-north direction from
the Salwan fault through Hasa to Sultani, and continues as a high to very high
permeability to Swaqa. Another zone extends from Swaqa fault line to to the west of
Khan Zabeeb area northward along the central area ofthe Wadi Wala Basin.
In the Jafr Basin, regions with high permeability within a zone of medium
permeability are present to the west of the Arja-Uweina flexure. A zone of low
permeability is found to the south of the Salwan fault line and to the east of the Arja
Uweina flexure. Further to the east and southeast, as the lithology of the aquifer system
changes to a sandy facies, the permeability becomes more uniform and is regarded as
medium with few spots of high permeability in the southern part.
However, due to the heterogeneity of the B2/A7 aquifer system, it is expected that
the permeability figures at certain locations may depart from the general pattern.
5.4.3 DISCUSSION
Most of the lateral hydraulic conductivity values discussed in this study are depth
integrated average values that represent the general hydraulic characteristics of the entire
study-unit thickness.
The initial values of K for the aquifer were estimated from pumping test analysis
and specific capacity data using different methods. In these methods, the Theis equation
is used to calculate T, which is "equal to an integration of the hydraulic conductivity
values across the saturated thickness part of the aquifer perpendicular to the flow paths"
(Lohman et aI., 1972). However, the T calculated from specific capacity data was
161
assumed to be the product of the K and the thickness of the aquifer open to the wells.
This T was then divided by the length of the open section or screened interval in the
well to obtain a vertical averaged K of that interval. The estimates are therefore probably
too high because ofthe effect of the vertical flow within the interval.
Ideally, estimates of aquifer conductivities are made by conducting test on wells
that fully penetrate the aquifer of interest and that are cased off above and below that
aquifer zone. Because of the variability in hydraulic conductivities caused by folds,
faults, aquifer thickness, and intercalated sedimentary materials, many hundreds of such
aquifer tests would be necessary to delineate properly the hydraulic conductivity
distributions throughout the study area. Relatively few such aquifer tests have been made.
Indeed, cased wells completed in B2/ A 7 in the study area are rare. However, a large
number of specific capacity values for domestic, irrigation, and municipal wells are
available from which K can be estimated. Generally, these wells are neither cased
properly nor fully penetrate the aquifer system; thus, the estimated K values represent
vertically integrated values over the entire penetrated interval in th~ aquifer system.
Therefore, some K values may represent only a part of an aquifer at particular location,
whereas others may represent composite K values over different parts of the aquifer.
Because the B2/ A 7 aquifer system is relatively similar in structurally similar areas, for
the purpose of this study, this was not considered to be a serious problem.
The potential range in permeability values can be estimated by usmg the
distribution of thickness of the aquifer system and the estimated transmissivity. The
frequency distribution of the entire thickness of the B2/ A 7 aquifer system can be
determined from the well inventory data. Dividing the minimum, median and maximum
values of the transmissivity by the respective minimum, median and maximum values of
aquifer thickness gives a potential range of hydraulic conductivity from 0.01-7.01 mIh,
with an approximate median of about 0.046 mIh.
Estimates of K made from specific capacity data were supplemented with the data
from previous studies, which were based on specific capacity data, aquifer tests and
numerical groundwater flow modelling. The frequency distribution of K calculated for
the B2/A7 aquifer system is shown in Figure (5.9). The values range between 0.000163-
162
36.4 mIh, with mean and median K of about 1.5 and 0.05 mIh respectively. About 90 %
of the values are less than the mean, while 16 % of the values about the median lie in
range 0.025-0.075 mIh. The distribution of K for B2/A7 is presented in Figure (5.12).
This distribution is based on specific capacity data, and thus values might be high in areas
with no data.
The wide range in K reflects the heterogeneous nature of the B2/ A 7 aquifer
system. The largest estimated K values probably are due to local geological structure, to
changes in the thickness of the aquifer zone, to karstification, or to the presence of
alluvial deposits with hydraulic continuity with the main aquifer system. Many wells with
large K values are close to wells from which small values were estimated. This indicates
the high degree of heterogeneity. Observed physical variations in exposed sections of the
aquifer system indicate that, within a single bed, a wide range of lateral (and vertical)
hydraulic conductivity values exist.
The K in the interbeds is probably dependent on the primary features, whereas in
the limestone beds is mainly due to the fracture system. Lateral hydraulic conductivity
may be much larger than that estimated from specific capacity data for some zones
because these estimates represent the entire uncased penetrated intervals.
The hydraulic conductivity of the B2/ A 7 also may be affected by faults, since the
offsetting of the aquifer beds through faulting could produce low hydraulic conductivity
fault material, or might close the pore space through deposition of secondary minerals
along the fault plane.
Large variation in K and the lack of data in many areas preclude areal mapping of
K. Results from previous modelling studies indicate that small values of K are required to
describe the regional movement of groundwater in the carbonate aquifer system (VBB,
1977, Howard Humphreys, 1986, BGR, 1987, and JIeA, 1987, 1990).
5.5 HUMMAR (A4) AQUIFER SYSTEM
Within the study area, the A4 aquifer system is under development only in the
Amman-Zerqa area. Wells tapping the A4 aquifer have been pumping tested by Parker
163
(1970), Mudallal (1973), and VBB (1976). The results of these tests are listed in Table
(5.10).
Within the Amman-Zerqa syncline, all the wells are flowing. Specific capacity
ranges between 0.2 and 116 m3/h1m, this range reflecting variations in thickness of the
aquifer as well as changes in permeability. The permeability values vary between 0.0017
and 3.9 mIh, with a mean and median of about 0.41 and 0.062 mIh, respectively. 48% of
the values are less than the median value, and 83% are less than the mean value.
Nevertheless, only three locations have a permeability higher than 0.5 mIh.
Well Number S.Thick SC Transmissivity (m"Jh) Permeability
(m) (m%) Parker Mudallal VBB (m/h) (1970) (1973) (1976)
A11(S8) 25 2.2 2.583 0.1033
A13(PP278) 41 1.21 1.375 0.25 0.02
A14(PP469) 47 8.9 11.458 16.917 0.3021
A16 45 6.25 1.042 0.0233
A17 45 1.98 0.375 0.0083
A23 45 2.2 0.4167 1.333 0.0196
A24 45 12.92 2.0833 162 1.825
A26 45 1.24 0.25 0.00542
A181 45 0.083 0.00167
A182(S10) 45 0.2 very low A185 45 1.458 0.0325
A186(S17) 50 0.95 1.796 8.458 0.1025
A187 45 1.625 0.03625
A190(PP468) 13 0.82 1.042 0.4167 0.04
A193 45 4.09 0.667 5.417 0.0675
A194 45 0.26 0.04167 5.5 0.0617
A195 45 2.39 0.5 41.042 0.4617
A196 45 0.91 0.125 0.1 0.0025
A200(S14) 45 116 145.833 29.167 3.889
A211 45 4.4 0.708 6.5 0.08
A214(PP111) 47 0.09 0.1146 0.0025
A215(PP180) 45 73 93.75 2.0833
A216(PP221 ) 33 2.8 3.667 0.1113
A219(PP458) 41 0.74 2.567 0.0625
Table (5.10) Results of pump test analysis of Hummar Aquifer System in Amman
Zerqa area.
164
To map the regional distribution of the permeability for the Hummar aquifer
system in the Upper Zerqa Basin, the permeability indices were calculated in the same
way as for the B2/ A 7 system, but the values of permeability used were those obtained
from pumping tests analysis. It is believed that the number of pump tests conducted in the
system is sufficient to delineate the regional distribution of the permeability. The
relationship between the permeability index (Pi) and the permeability expressed in m/h is
shown in Figure (5.13).
It is believed that the permeability in the Hummar aquifer system arises from
primary permeability, dolomitization of the limestone, and karstification. The
permeability distribution map (Figure 5.14) shows an increase of permeability in the
direction of groundwater flow from the western limb of the Amman-Zerqa syncline
eastward then northwards. The groundwater flow rate increases from Amman northwards
towards the discharge area to the north of Zerqa. Consequently, the' degree of
karstification increases accordingly with the increase of groundwater flow. As the
saturated thickness of the aquifer remains more or less constant, the increase in
permeability reflects a comparable increase in the transmissivity. Outside the Amman
Zerqa syncline, particularly eastwards and southwards, the permeability is low: this is
attributed to the increased percentage of marl in these areas.
7 r------------r------------r-----------~----------_.
6~----------~------------~--------~~~--------~ ~
~ )(
~ 5~----------~----------~~~--------~----------~ .= ~ ~ 4~----------,~~----------~----------~----------~ Q)
E Q)
~ 3~----------~------------~----------~----------~
2~----------~------------~----------~----------~ 0.001 0.01 0.1
Permeability (m/h)
Figure (5.13) The relationship between permeability index and permeability for the Hum mar aq uifer system.
165
10
1~~_~---==------=~----~~----~~----~~----=-=---~~~--~~
Figure (6.14) Areal distribution of permeability Index In the A4 aquifer system
None of the pumping test sites was provided with observation wells so the storage
coefficient cannot be reliably calculated. However, Parker (1970) from water budget
calculations for the A4 aquifer in Arnman-Zerqa area, believed that a storage coefficient
of 0.01-0.1 might be appropriate. Groundwater modelling by VBB (1976) suggested
values of 0.001 and 0.05 for the confined and unconfined parts of the aquifer,
respectively.
5.6 RIJAM (B4) AQUIFER SYSTEM
The Rijam limestone. comprises a shallow aquifer system in the central part of the
Jafr Basin, where it is exploited mainly for irrigation. The aquifer characteristics have
been studied by Parker (1970) who conducted pumping tests on nine fully penetrating
wells: only for one test an observation well was available. The pumping test data were
analysed using the Cooper and Jacob (1946) straight line approximation of the Theis
(1935) non-equilibrium formula. The results are summarised in Table (5.11).
As shown in Table (5.11), the transmissivity of the Rijam aquifer system is
relatively high, it varies widely from well to well, ranging between 1.8 and 404 m2 Ih. The
variation in transmissivity does not show any relationship with saturated thickness. Well
(PP28) for example has the highest transmissivity and the shallowest saturated thickness.
The areal distribution of transmissivities shows a zone of high transmissivity to the
northwest of Jafr. The permeability of the wells tested ranges between 0.06 and 27.0 m1h.
This wide range in permeabilities is indicative of the karstic nature of the Rijam
limestone. Lateral changes in permeability within the vicinity of the well, revealed by the
time-drawdown and/or time-recovery curves were noted in tests on wells PP25, PP27,
PP30, and PP31 (Parker, 1970).
Estimation of the storage coefficient for the aquifer system is obtained from only
one pumping test on well PP17. A short pumping test of 2 hours and a long pumping test
of 132 hours were conducted on the well. The data were analysed using the Theis (1935)
type curve technique and the Cooper and Jacob (1946) straight line method. The results
show ranges of storage coefficient between 0.9 and 2.5 % for the short test and between
0.01 and 0.93 % for the long test. Unexpectedly, the values obtained from long test are
167
lower than those obtained from short test. However, Parker (1970) suggested that for the
purposes of estimating the total volume of water stored within the aquifer, a storage
coefficient value of between 1 and 10 % might be appropriate.
Well S.Thick Yield Ddown SC Transmissivity Permeability Number (m) (m3/h) (m) (m%) (m%) (m/h) PP15 34.23 137 1.2 114.1 76.0833 2.2227 PP17 33.88 179 0.27 662.9 319.542 9.4316 PP20 30.24 17 23.0 0.7 1.7917 0.0593 PP23 24.96 35 5.33 6.5 3.3125 0.1327 PP25 34.55 42 19.27 2.2 2.8125 0.0814 PP27 35.8 141 7.36 19.1 17.0417 0.4760 PP28 * 14.85 218 0.52 419.2 403.75 27.189
PP30 * 29.96 166 0.44 377.2 289.5833 9.6657 PP31 32.75 97 0.55 176.3 179.167 5.4708 Note: * one hour test
Table (5.11) Results of pumping tests in the Rijam Aquifer System in Jafr Basin
(Parker, 1970)
5.7 LOWER AJLUN GROUP (Al-6) AQUIFER SYSTEM
Ignoring the A4 aquifer system in Upper Zerqa basin, in most of the study area
the Lower Ajlun Group (AI-6) provides the lower confining unit of the extensive B2/A7
aquifer system. Within the group the Na'ur Formation contains limestone beds which
have aquifer potential. The Formation has been penetrated by a number of wells, all of
which have low yields. The wells tested, have experienced continuous drawdown in water
levels without stabilising (Parker, 1970). The low permeability is believed to reflect low
recharge and hence poorly developed karstification.
However, in the Jafr Basin, particularly in the east and southeast, the Group
becomes more sandy and hence modified into an aquifer, since the sand provides
sufficient hydrogeological conditions for movement of groundwater through otherwise
virtually impermeable marls and shales. In general, the AI-6 aquifer is a multi-layered
aquifer, and comprises semi-pervious shaly and marly layers separating discrete aquifer
beds which, in the west, consist of limestones, and in the south and southeast of sandy
limestones and sandstones. Locally, the shaly and marly layers may contain large
168
quantities of sand and silt. The distribution of shale is not unifonn, and in some areas
sand may be the dominant. It is expected that the effectiveness of these layers as a
confining unit may be impaired in these sand rich areas.
All the previous studies, except that of Howard Humphreys (1986), considered the
Lower Ajlun Group in the Western Highlands as an aquiclude, and only as a conduit
allowing flow to pass from the B2/ A 7 aquifer downwards into the underlying Kurnub
aquifer, in the Jafr Basin. Thus, infonnation about the hydraulic characteristics of the
group is very limited. However, Howard Humphreys (1986) carried out aquifer tests at
two sites, but they managed to obtain data from only one of the pumping tests (well
PHT5A); the other test (well PHT11A) was tenninated after five minutes of pumping due
to the excessive drawdown.
Data from test well PHT5A, analysed using the Theis and Jacob methods, gave an
average transmissivity of about 13 m2/h which corresponds to an average penneability of
about 0.05 mIh. Because of the lack of pumping tests in the Al-6 aquifer, it has not been
possible to define the areal distribution of the hydraulic parameters. However, given the
wide variation in the lithology of the aquifer, penneability will have a large range,
between 0.0004-0.0833 mIh (Howard Humphreys, 1986). In the western part of the area,
where the carbonates dominate, the penneability will mainly depend on fracturing. In the
south and east, in the sandy facies, penneability is likely to be both primary and
secondary, and on a large scale will depend on the percentage of marls and shales in the
sequence.
The variable lithological nature of the Al-6 and frequent interbedding of the
aquifer layers with low penneability clays, shales, and silts, suggest that groundwater is
stored under variety of conditions ranging from phreatic to semiconfined and confined.
The average storage coefficient estimated from the pump tests was found to be about
0.005. Where the aquifer is phreatic, specific yields may range between 0.1 and 10 %
depending on the silts and clay content.
169
CHAPTER SIX
RECHARGE
6.1 INTRODUCTION
Arid conditions are said to exist in a region when the potential
evapotranspiration exceeds the rainfall for most of the year. The difference between
annual potential evapotranspiration and annual rainfall can be used to define degree of
aridity: it ranges between less than one metre in semi-arid zones to more than 2 m in
arid zones. Long periods of aridity change the face of the land drastically, which in tum
has a different hydrological response to the atmospheric inputs.
Owing to the climate condition prevailing in a semi-arid to arid area like Jordan,
with surface water limited to a few perennial streams maintained by spring discharges
along the Western Highlands, the increase in water demands to meet the ever-increasing
agricultural and domestic needs, solely depends on groundwater. Before any large scale
exploitation of groundwater reserves takes place it is essential to estimate the amount of
natural recharge from rainfall to set a safe limit on exploitation. The recharge/discharge
relation or the net recharge and its areal distribution is also an important factor in
groundwater movement, and hence in any aquifer system evaluation. Recharge to an
aquifer system is primarily from rainfall, applied irrigation water, and from surface
water bodies. Discharge from the aquifer, excluding pumpage, is mainly to the rivers, by
spring discharges, seepage along deep wadis and canyons, and as outflow to other
adjacent aquifer systems.
6.2 RECHARGE MECHANISMS
Given the climatic condition, the geology, the topography, and the vegetation
cover, natural replenishment of groundwater in Jordan can take place by four
mechanisms:
1. Direct infiltration of rainfall via the soil and unsaturated zone.
2. Indirect infiltration of surface runoff via permeable wadi beds or drainage
systems.
3. Lateral boundary flow
4. Water transfer.
Direct recharge is that amount of rainfall which reaches the water table after
runoff, evaporation and soil moisture deficit have been accounted for. It is likely to
occur in the outcrop area of the high rainfall zone in the Western Highlands. Away from
the Western Highlands, the aquifer is buried under thick low permeability sediments
which prevent replenishment from rainfall or runoff.
Indirect recharge may occur wherever runoff concentrates sufficient water to
exceed the evaporation during the period required for infiltration to take place, once soil
moisture deficit has been exceeded at the locality. Thus indirect recharge may occur in
low rainfall areas as in the eastern and southern part of Jordan if a particular storm, or
water transported into the area by the drainage system are sufficient in intensity and
amount to cause runoff.
Recharge may also occur by lateral flow of groundwater from outside the study,
from aquifer to aquifer, or from one part of the area to another. The major source of
transferred water is the discharge of basalt aquifer waters into the B2/ A 7 aquifer system
in the Wadi Dhuleil - Mafraq area (MacDonald, 1965 and Parker, 1970). Water also
may be transferred laterally within the same aquifer; from the Western Highlands
eastward to replenish the confined part of the aquifer systems in the Central Plateau, or
by vertical leakage into different aquifer systems.
In Jordan, as there are only few surface water bodies, their contribution to the
groundwater budget is extremely small: as a result they are excluded from the
recharge/discharge estimation. However, their contribution to the groundwater flow
system is discussed in more detail in groundwater modelling (Chapter 8).
Irrigated agricultural land usually recharges aquifers by return flow, but as the
irrigated areas in the study area are extremely limited, and occur mostly in confined
areas, and because ofthe regional nature of this study, the effect of irrigation return flow
is ignored in the groundwater budget. But it is considered later for the purposes of
groundwater modelling (Chapter 8).
171
6.3 RECHARGE ESTIMATION
Recharge is controlled by both daily and annual climatic variation and, thus, is
highly variable both temporally and spatially. The magnitude of the natural recharge
component is generally the largest uncertainty in water balance calculations,
particularly in arid regions due to the unpredictable nature of the climate.
There is a wide range of methods of calculating the recharge component. These
methods are based on the water balances, groundwater level fluctuation, baseflow
fluctuation, and chemical and isotopic composition of the groundwater. Should an
attempt be made to estimate the volume of recharge, the methods used should be as
independent as possible of the other methods being used in order to avoid situations in
which an error or wrong assumption would carry through the analysis. In addition the
methods should be applicable throughout the study area so that if there is a bias error, at
least the relative differences between the different hydrological areas would be apparent.
Because recharge is a non-linear process, it is not possible to use the average
value of each controlling factor to derive an average recharge. Recharge should be
estimated separately for each homogeneous zone.
In this investigation, the areal subdivision of the study area (Figure 6.1) and the.
five year period (1980-1985) of meteorological records used in runoff estimations
(Chapter 3) were also used for recharge estimations. The estimates of recharge for each
sub-catchment were obtained by determining the outcrop area of the principal aquifer
system. Outcrops are less than 40 % of the area of the B2/ A 7 aquifer system, but
hydrologically they are significant because of the relatively rapid rates of recharge
(Figure 6.1).
6.4 DIRECT RECHARGE
A number of techniques have been devised for estimating the groundwater
recharge. Different distinctive techniques were used in this study, in order to obtain a
spectrum of values which could then be evaluated on their relative merits.
172
N
W-\rE s
o 50km
100 LEGEND
I;';'/~I ////
outcrop of 82/A7
m basalt . . .. --- faull
/" basin boundary
~ basin subdivision
4 subdivision number
", .. -... - .. wadi , playa
1000
I , I ,
I /
900~--~~ __ ~ __ ~~~~ ____ ~ __ ~~~~ ____ ~ ____ ~ __ -J 200 300
Figure (6.1) The outcrop area of the B2/A7 aquifer system
6.4.1 SOIL-WATER BALANCE
The amount of rain water that percolates through the unsaturated zone to the
water table can be estimated on the basis of climate, soil type, topography, land use, and
consumptive water use by crops and vegetation. These elements have been discussed in
previous chapters.
Recharge was estimated on a daily basis for the different sub-catchments in the
study area by applying a continuity equation that computes daily values of deep
percolation of water below the effective root zone for each sub-catchment. These daily
estimates were used to estimate long-term recharge for the current land use condition in
the study area. The conceptual model of the soil-water balance used in the calculations
is shown in Figure (6.2). The daily water budget is expressed as:
P = R + E + ~SM + I ........................................................................... (6.1)
where P = the rainfall
R= the runoff
E = the evaporation from wet soil and plants
~SM = the change of soil moisture content
I = the infiltration beyond the root zone.
The following data are required for each sub-catchment to estimate recharge:
daily rainfall, daily runoff, daily evaporation, soil-water holding capacity, soil types,
topography, and landuse classification.
Daily meteorological data have been discussed in chapter three. The average
runoff coefficients obtained by using the eN method were adopted here for the
estimation of recharge. Soil information and landuse classifications were obtained from
previous studies and transposed to the study area subdivisions. Hunting Technical
Services (1954) conducted a very comprehensive range classification survey in Jordan.
The classification was simplified by grouping surface types as permeable, semI
permeable, or impermeable to groundwater recharge. Permeable areas, such as wadis
alluvial and river beds, were of negligible extent although such areas play an important
role in indirect recharge calculations. Impermeable areas such as large areas of thin soil
overlying impervious marl, or steep slopes were considered insignificant in terms of
174
recharge but very significant in tenns of local runoff conditions. The semi-penneable
areas, covering the majority of the study area, were classified into different types based
on the soil types and the underlying rocks and the vegetation cover (Table 6.2). Each
catchment has a different soil type/landuse combination. Often, this combination is
found to vary locally within the same catchment, and in these cases the areas were
divided by the ratio of each component comprising the area and the average value was
used. In general the surface condition in Jordan in the wet season changes from bare
soil to a light grass and crop cover.
SURFACE RUNOFF
SUBSURFACE RUNOFF
I RAINFALL
.0-
.0- EVAPORATION
.0- 71
.0- 71 I INTERCEPTION STORAGE
.0-
.0-MOISTURE INCIDENT TO GROUND SURF ACE
.0-
.0- EVAPOTRANSPIRATION
.0- 71
.0- 71 SOIL MOISTURE STORAGE IN ROOT ZONE
I DEEP PERCOLATION
.0-
.0-
I GROUNDWATER
. RECHARGE
Figure (6.2) Schematic diagram showing the conceptual model of the soil-water
balance method.
175
ACTUAL AND POTENTIAL EVAPOTRANSPIRATION
Water loss from a catchment area does not always proceed at the potential rate,
since this is dependent on a continuous water supply. When the vegetation is unable to
abstract water from the soil, the actual evaporation (Et ) becomes less than potential.
Thus the relationship between Et and PET depends upon the soil moisture content.
The true relation between actual and potential evaporation will vary with rooting
characteristics, soil texture, and plant physiology, as well as with the rate of
evapotranspiration itself and the climatological conditions. A popular compromise
between the above factors has been the use of the so-called "root constant" (Penman,
1949). Evapotranspiration is assumed to occur at the potential rate until the SMD
exceeds the root constant; then evapotranspiration proceeds at a slower rate.
The availability of soil moisture for plant growth over a range from field
capacity to permanent wilting point has been treated by a wide range of techniques.
Veihmeyer and Hendrickson (1955) suggested that in some cases evapotranspiration
may proceed at the potential rate until soil moisture approaches the permanent wilting
point. While Thomthwaite and Mather (1955) assume the ratio of actual to potential
evapotranspiration is a linear function of the ratio of available soil moisture to the
available water capacity. Thomthwaite and Mather (1957), Dune and Leopold (1978),
and Mather (1981) suggest a model and provide tables and graphs which allow
calculation of SMD. Palmer (1965) used the root constant in the form of soil moisture
capacity parameters. Palmer's model uses an analogy of the linear approach of
Thomthwaite and Mather (1955) to estimate evapotranspiration from the lower layer.
Another approach is to assume simply that evapotranspiration losses from the lower
layer are equal to some percentage ( often of the order of 10%) of the potential losses
(for example Howard and Lloyd, 1979; Rushton and Ward, 1979; and Calder et aI.,
1983).
When the soil is saturated, it will hold no more water. After rainfall ceases,
saturated soil relinquishes water and becomes unsaturated until it can just hold a certain
amount against the forces of gravity; it is said to be at ' field capacity' (FC). In this
condition, the evapotranspiration occurs at the maximum possible rate, in other words
the actual and potential evapotranspiration are equal (Thomthwaite and Mather, 1955).
176
In Jordan's semi-arid climate, where the groundwater table is far below the reach of
evaporation, the amount of actual evaporation wi11largely be determined by the amount
of rainfall. If there is no rain to replenish the water supply, the soil moisture gradually
becomes depleted by the demands of the vegetation to produce a soil moisture deficit
(SMD) which is the amount of water required to restore the soil to field capacity. As
SMD increases, Et becomes increasingly less than PET. The values of SMD and Et vary
with soil type and vegetation. Often it is assumed that if rainfall occurs, Et=PET up to
the point when the rainfall volume is used up: in this situation Et occurs at the potential
rate even though a non-zero soil moisture deficit may exist.
In early winter, before the establishment of a vegetation cover, the amount of
evaporation will depend on both the amount of rainfall and the moisture conditions of
the soil at the end of the dry season. During this period, the PET exceeds the rainfall, so
the Et was estimated as equal to rainfall + 10% (PET - rainfall) provided that there was
already some water stored in the soil during the current water year to support this (Lloyd
et aI., 1966).
In the later part of the wet season, and in the subsequent dry season, evaporation
becomes dependent on the moisture stored in the soil and will fall progressively below
the potential rates as the water within the root zone or the upper layer of a bare soil is
depleted. Penman and Scholfield (1964) suggested from laboratory experiments that the
evaporation rates from bare soil with a dry layer may be only 10% that of PET after the
first 25 mm have evaporated. When vegetation is present, the lowest soil layers in which
roots are actively growing can be considered analogous to an exposed soil surface as far
as water transfer into it from deeper layers is concerned. Field studies in Pakistan
(Ahmad, 1962) and Libya (Allemmoz and Plove, 1980), however, have shown that
evaporation from bare soils falls almost to zero towards the end of the dry season, so
this 10% rate must decrease further with time.
SOIL MOISTURE DEFICIT
The relative changes of Et with increasing SMD have been the subject of a
considerable amount of studies. Penman (1950) introduced the concept of a ' root
constant' (Re) that defines the amount of soil moisture (mm depth) that can be extracted
from a soil without difficulty by a given vegetation. It is assumed that Et = ET for a
177
particular type of vegetation until the SMD reaches the appropriate root constant plus 25
mm approximately, which is added to allow for extraction from the soil immediately
below the root zone. Thereafter, Et becomes less than PET as moisture is extracted with
greater difficulty. As the SMD increases further, the vegetation wilts and Et becomes
very small or negligible.
To evaluate SMD and Et over a catchment area, the proportion of the different
types of vegetation covering the catchment must be known. This entails a land-use
survey and a classification of the vegetation for allocation of RCs before a water budget
calculation may be carried out. MacDonald (1965), depending on direct observation and
by comparison with other similar environments adopted the following effective rooting
depths for the three main vegetation types in Jordan:
Grass
Cultivated areas
Trees
300 mm
400 mm
1000 mm
Table (6.1 ) shows the potential SMD values calculated at selected stations in
the study area. Potential SMD is the soil moisture deficit that would result if the PET
was always fulfilled. It is the aggregate of the rainfall minus PET considered as a deficit,
and assumed to apply to the riparian lands at or above FC. The values of SMD in Table
(6.1) are only for comparison and to demonstrate the high value of SMD and the low
value of Et with respect to PET, in a semi-arid country like Jordan. The data used are
mean monthly values which do not fulfil the conditions of PET and SMD which assume
that there is abundant water available for evaporation and the lands are at or above FC.
However, the results may be meaningful for the high rainfall period (December-March)
when the rainfall will often exceed the PET for most of the stations. Bearing in mind
that, the high rainfall period coincides with the period of low temperature and thus low
PET.
The soil moisture conditions at the end of the dry season (October-November)
for representative soil types in Jordan were estimated directly from field measurements
by MacDonald (1965). Examination of many soils shows that for most ofthem, only .
178
Amman Airport """'."".""'>''''''''''''''' ,".",:.,>,:::::::::;> ... ,:.::::/ '.""':""'" ":':""",.,: ~tllrh,.,<>h "":':':':.<'::< Mnth Rainfall PET Rain- Poten Rainfall PET Rain- Poten
Min Ave Max PET SMD Min Ave Max PET SMD Oct 0 8 35 109 -101 101 0 6 34 102 -96 96 Nov 1 25 133 67 -42 143 0 17 62 49 -32 128 Dec 1 47 175 52 -5 148 0 26 86 33 -7 135 Jan 18 66 235 55 11 137 4 39 166 26 13 122 Feb 7 45 276 72 -27 164 4 28 75 48 -20 142 Mar 11 61 156 111 -50 214 5 38 108 84 -46 188 Apr 0 19 46 146 -127 341 0 8 63 103 -95 283 May 0 1 7 193 -192 533 0 2 15 132 -130 413 "<>':<,::::/ ""·,,,,''',·Z .. iflln''· ",":></'''' ".,., ,."" .. ,:,,,,,,,,,, , .. :. """""''' .. ''':'><;'' ,,'::::;., ......... ,1)1-;;'1 h"··,'"',,,·,·,, .,<> Mnth Rainfall PET Rain- Poten Rainfall PET Rain- Poten
Min Ave Max PET SMD Min Ave Max PET SMD Oct 0 5 40 105 -100 100 0 5 35 150 -145 145 Nov 0 17 100 63 -46 146 0 14 43 105 -91 236 Dec 3 40 195 48 -8 154 0 31 127 60 -29 265 Jan 0 55 261 49 6 148 0 39 157 62 -23 288 Feb 2 54 119 65 -11 159 0 30 112 70 -40 328 Mar 0 40 113 113 -73 232 0 27 85 109 -82 410 Apr 0 14 262 128 -114 346 0 13 80 165 -152 562 May 0 2 13 180 -178 524 0 1 6 225 -224 786
I::,:,,':> ... >", " ... "'.'", .. , """""" f)lItrllnll.><><><·".'::"::::::> . ";>:.:: .. «" ".:,,',,"'.: . ::'.": .. ':'.':::'>, .. ,. ····,····',:>-Hasa ">::::,,,,:::.,}<" .", .. : . .,".:': .. ' .. ':.' ... ::,.:.,:
Mnth Rainfall PET Rain- Poten Rainfall PET Rain- Poten Min Ave Max PET SMD Min Ave Max PET SMD
Oct 0 3 20 150 -147 147 0 3 21 Nov 0 6 26 105 -99 246 0 5 17 47 -42 42 Dec 0 18 81 60 -43 289 0 9 45 32 -23 65 Jan 0 24 113 62 -38 327 0 12 73 33 -21 86 Feb 0 18 104 84 -66 393 0 10 42 60 -50 136 Mar 0 16 52 124 -108 501 0 8 39 95 -87 223 Apr 0 7 65 210 -203 704 0 2 12 108 -106 329 May 0 1 7 240 -240 944 0 0.3 6 . ':,,: ·':"':·::"·'i"··'''.''::>:'·:/i.:,' .. .. ,."., .. .... r afila'·::· ,:,::>,<":,:,,, ,,:::,;::. .' .... ,',',':':::.:':' :::":::,:::",,,.,,,,:,,,,.,, '." Shaubak:;..:;:' ",,,. " ...
Mnth Rainfall PET Rain- Poten Rainfall PET Rain- Poten
Min Ave Max PET SMD Min Ave Max PET SMD
Oct 0 10 11 102 -92 92 Nov 0 10 19 49 -39 39 0 23 77 49 -26 118 Dec 0 18 35 33 -15 54 0 50 135 33 17 101 Jan 0 63 189 58 5 49 0 31 84 26 5 96 Feb 0 49 160 64 -15 64 0 56 104 48 8 88 Mar 0 60 121 101 -41 105 0 53 79 84 -31 119 Apr 0 48 70 119 -71 176 0 12 22 103 -91 210 May 0 6 9
Table (6.1) Calculation of Soil Moisture Deficit (mm) at selected stations.
179
50% of the soil moisture content at Fe can be taken up by, and evaporated from, plants.
The rest has been considered as hygroscopic and therefore non-available. Thus, the
amount of water which is considered to be freely available to plants or for direct
evaporation in a soil at Fe is the product of 50% Fe x RC. Evaporation is considered to
occur at potential rates until this amount has been removed from the soil, and it has been
termed the drainage factor (e). e for agricultural purposes is defined as the available
amount of water in (mm) which is within the influence of the root depth of a plant. Then
evaporation proceeds at a much slower rate until the rate falls to such a low value that an
almost constant soil water deficit is reached. The amount of water required to bring the
moisture content of such a dry soil back to field capacity over the whole profile depth is
termed the final deficit value(D). For recharge purposes e is. therefore defined as the
limiting amount of water necessary in the root zone before which groundwater can
occur. Before Fe conditions occur, however, the SMD from the summer months must
be satisfied. Under this definition any water passing below the root zone is considered to
be recharge and is potential groundwater. Soil moisture properties for various soil
covers in east Jordan are presented in Table (6.2). The field capacities (Table 6.3) were
measured by determining the residual moisture content of undisturbed soils at different
depth intervals, two days after the soils had been saturated by heavy surface application
of water at the surface and 40 cm below the surface, and compared with a set of control
values measured before the application of water (MacDonald et aI., 1965).
An example of variation of the field capacity with depth is shown in figure (6.3).
Although the example taking for the basalt soil outside the study area, it gives an idea
about the behaviour of the soil moisture at different conditions. Figure (6.4 ) shows the
seasonal variation of the soil moisture content of a 50 cm thick surface layer measured
at three locations.
RECHARGE CALCULATION AND RESULTS
For each year of the record the amount of rainfall in each storm was balanced
against the other parameters in equation (6.1) to calculate recharge. The actual
evapotranspiration for the period October to November was estimated as P+ 1 O%(PET
P), then and during the height of the wet season, when the water is assumed to be freely
180
Soil Type C D Effective Soil Depth (cm) Limestone Cultivated 65 101 lOO
Uncultivated. 52 56 40 Marl-Shale Cultivated 96 107 lOO
Uncultivated. 73 lOl 60 Chalky-Marl Cultivated 74 lOO lOO
Uncultivated. 68 73 40 Chert Cultivated 87 115 lOO
Uncultivated. 59 74 60 Basalt Cultivated 86 127 120
Uncultivated. 57 85 80 Samra Loess 97 151 80 Uneiza Plateau 62 87 60 Rabba Plateau 91 122 120 Hamat Plateau 59 88 80 Source. MacDonald et al. (1965)
Table (6.2) C and D values (mm) for various soil types in Jordan.
Soil Type Field Capacities (%) at Depth (cm) 10 20 40 80 120
Limestone Cultivated 23.5 23.5 22.5 22.0 22.0
Uncultivated. 28.0 24.0 24.0 Marl-Shale Cultivated 35.0 35.0 27.5 27.5
Uncultivated. 35.0 34.0 32.0 Chalky-Marl Cultivated 27.5 24.0 24.0 24.0
Uncultivated. 32.0 32.0 32.0 Chert Cultivated 31.5 31.5 28.0 25.0
Uncultivated. 27.5 27.5 27.5 Basalt Cultivated 28.0 28.0 28.0 22.5 19.0
Uncultivated. 25.0 25.0 25.0 19.0 Samra Loess 35.0 35.0 31.0 30.0 Uneiza Plateau 30.0 30.0 30.0 Rhabba Plateau 31.0 31.0 31.0 25.0 21.0 Hamat Plateau 22.0 22.0 19.0 15.5 Source: MacDonald (1965)
Table (6.3) Field capacities (%) values by weight for various soil types in Jordan.
available for evapotranspiration, the actual evaporation proceeds at the same rate as the
potential (Et = PET) until the amount of the drainage factor (C) has been removed from
the soil, i.e. the accumulated I exceeds C. Then toward the beginning of the dry season
the evaporation from the soil occurs at lower rates substantially less than the potential
i.e. only 10% ofthe potential until a constant soil-water deficit is reached.
181
20
60 e .. ~ 0-
~ 80
100
120
o
c -c
~ ~ A. Q. ~ :- e :: -:
-~ ~ " ~ ~ ~ e e e e c
" e e E : 0 u.s. 2 000 ... ~ ...
10 20
MOISTURE CONTeNT: To
KEY
o ~ SurraCi application,
30
_1(- - -x- Applications at 4:cm &,1_ G.t.
Figure (6.3) Field capacity determination for basalt soil (after Lloyd et al.,1966)
E .§. 0 :E (/)
140
120
100
80 Dhuleil
60
40
20
----------~-~-------------------------------
---- - . - - - - - - - - - - - - - - - - - -,- - - - -- -- .... - - - -.... '- - - - - - --Shaubak • .... _ .. .. - , ,
.... -- - - -O+-~--~--+-~---r--+---r--+--~~r--+--~~ Dec Jan Feb Mar Apr May Jun Jul Aug Sep Oct. Nov
N.B. Data are from the WAJ files
Figure (6.4) Soil moisture content (SMC) variation with time at selected sites for depth range between 0-50 cm.
The estimates of recharge in each of the sub-catchments calculated by solving
Equation (6.1) for I. Recharge occurs when the accumulated I becomes positive, i.e.
the amount of rainfall must exceeds the sum of R, Et , and D. A typical example of the
soil moisture balance calculation is shown in Appendix (D 1). The estimated recharge
and the relation between rainfall and recharge for each basin in the study area during the
record period are shown in Table (6.4).
The mean annual direct recharge to the Amman/Wadi Sir aquifer system in the
study area during the five year period ranges between 5.75-173.55 MCMla with an
average of 50.5 MCMla or 8 % of the average annual rainfall. It varies between the
basins and from year to year within the same basin. This reflects the variation of the
recharge surfaces and the irregular storm patterns. It is worth mentioning that the data in
Table (6.4) are only for part of the basin were recharge had occurred and are not a water
balance for the whole sub-catchments. It is only the western highlands of each basin
which has direct recharge. Elsewhere, the groundwater replenishment depends on
indirect recharge and on water transfer from the other parts of the system.
In general, however, calculation shows that direct recharge occurs only in areas
which receive more than 250 mm of annual rainfall. Above this value the amount of
recharge depends both on the total rainfall and its distribution. The amount of rainfall
183
however, must exceed the evapotranspiration and the balance of water enters the soil
profile.
The relationships between rainfall, potential evapotranspiration, and soil
moisture deficit before recharge can occur is very critical. It is not necessary for large
storms to produce recharge. Isolated large storms and storms at the beginning and end of
the season are often consumed by high evaporation and large soil moisture deficits and
thus give little or no recharge. As such, high rainfall years are not always high recharge
years or may even produce no recharge. But for average years having large storm
incidences during mid season, normally in January and February, the recharge increases
by increasing the rainfall.
The analysis of the relation between recharge and rainfall timing and spatial
distributions, shows that the most likely time for recharge to occur is during the winter
months when potential evaporation is at a minimum, dry periods are more rare, and the
soil moisture deficit is usually small. The most beneficial conditions are during
consecutive storms.
BASIN 1980/81 1981182 1982/83 1983/84 1984/85 AVE. UPPER Rainfall 133.4 112.8 171.6 98.7 116.8 126.7 ZERQA Recharge 0.623 5.1 48.3 0.0 9.01 12.61
% 1 5 28 0.0 8 8 WADI Rainfall 179.7 135.5 250.5 118.8 214.7 179.8 WALA Recharge 4.79 12.23 60.95 5.75 7.08 18.16
% 3 9 24 5 3 9 WADI Rainfall 192.46 139.74 186.88 98.04 139.46 151.32 MUnB Recharge 6.62 13.8 49.2 0.0 5.31 15.00
% 3 10 26 0.0 4 9 WADI Rainfall 20.6 14.0 22.7 10.9 16.5 16.9 HAS A Recharge 0.0 0.0 2.0 0.0 0.0 0.4
% 0.0 0.0 9 0.0 0.0 2 Rainfall 77.5 74 92.25 22.5 53.25 64
JAFR Recharge 0.0 8.8 13.1 0.0 0.0 4.38 % 0.0 12 14 0.0 0.0 5
Rainfall 603.7 476.04 723.93 348.94 540.71 538.7 TOTAL Recharge 12.03 39.93 173.55 5.75 21.4 50.5
% 2 8 24 2 4 8
Table (6.4) Results of direct recharge calculation to the B2/A7 aquifer system
(in MCM) by using soil moisture balance method.
184
Among the other factors which affect the amount of recharge are the topography
and the geology of the sub-catchments. Steep slopes were classified as impervious for
recharge purposes and consequently when present in the area reduce the overall direct
recharge percentage. The western highlands of the Wadi Hasa basin for example, due to
their relatively steep topography, receive less recharge than would be expected from the
amount of rainfall (only 2% of the mean annual rainfall, compared with 9% and 8% in
the Wadi Mujib and Upper Zerqa sub-basins respectively).
The recharge also varies between different recharge surfaces. Uncultivated
limestone and chert soils are the most receptive to recharge in the study area, bearing in
mind that these soils overlay the principal aquifer system. The uncultivated surfaces
absorb more recharge than the cultivated soils due to the shallower rooting depths.
DISCUSSION
This study has been based on the available climatic data in the area for the period
1980-1985. The reliability, density, and distribution of the climatological stations
through the study area are believed to represent satisfactorily the climatic conditions.
This gives advantage to this study over the previous studies which have been based on
less reliable climatic data.
Penman's method for the calculation of potential evaporation is believed to be
the most reliable indirect method. Its suitability for the study area has been tested by
MacDonald et al. (1965).
The mean monthly values of evaporation were adjusted to daily values, which is
the basic period for which calculations of evaporation were made. As storm periods are
normally short and frequently of less than one day it is probable that the daily balance
data are more realistic. Furthermore, most water balance models assume that Et for a
period is equal to the PET whenever P~PET. This assumption is usually made
regardless of whether one is performing a daily, weekly, or monthly water balance.
However, rainfall and evapotranspiration often are distributed within a certain period in
such a way that both periods of deficiency and surplus can occur. These add another
advantage of this study, since all the previous studies used 10 days or monthly values in
their calculations.
185
Morton (1983) reviewed the limitations of the assumption of the threshold
concept, that runoff and recharge does not occur until soil moisture capacity is filled,
found that when the same value of soil moisture value capacity is used, estimates of
recharge will tend to decease as the time step of the computation increases. For example,
Rushton and Ward (1979) found that monthly water balances lead to recharge values
which are up to 25% less than those from daily water balances.
The estimation of actual evaporation from calculated potential evaporation over
the period December to March when recharge is most likely to occur presents some
difficulties, evaporation considered to take place at potential rate. Because of incomplete
vegetation cover and the drying of soils between rainstorms, this may produce an
overestimate of actual evapotranspiration over some periods, and consequently may
reduce recharge estimates a little. The accuracy of actual evaporation rates before and
after this period will depend on whether the assumptions about reducing the actual
evaporation under dry soil conditions are correct, and whether the rooting depth
observation at which this occurs are appropriate. The assumption about reducing the
actual evaporation is accurate for average years, but for some years which receive high
rainfall earlier at the beginning of the wet season when the actual evaporation is
assumed to depend on the amount of rainfall, this leads to an overestimate of the actual
evaporation, and thus an underestimate of the recharge even if the soil moisture deficit is
overcome earlier. The true relationship between the actual and potential
evapotranspiration will vary with rooting characteristics. Evapotranspiration is assumed
to occur at the potential rate until the soil moisture deficit exceeds the root constant;
then evapotranspiration proceeds at a slower rate. Sensitivity analyses performed by
Howard and Lloyd (1979) and Rushton and Ward (1979) suggest that reductions in the
root constant in water balance models may lead to greater reduction in estimates of
evapotranspiration. The rooting depth estimation by MacDonald et al.(1965) was
overestimated, which would also lead to an overestimation of evapotranspiration and
consequently to an underestimate of recharge. Hence it is probable that calculated
recharge is likely to be lower than the true value.
The calculations show that the most significant soil coefficient is the soil
moisture deficit. During the period when drainage occurs the soil moisture is generally
at field capacity irrespective of the rooting depth. Consequently the drainage factor loses
186
its significance except when isolated heavy rainfall periods occur: at such times the
amount of recharge is normally small in comparison with the annual total. The problems
relating to the evaluation ofthe rooting depth are therefore minimised.
The field measurements of field capacity and bulk density were carried out by
MacDonald (1965) using standard field methods. It is important that deficit conditions
should be measured but it is unlikely that the values will vary significantly from year to
year.
The average annual groundwater recharge over the whole study area was 8% of
the average rainfall for the 5 year period. This is consistent with 8.2% estimated by
Lloyd et al. (1966) for northeastern Jordan and 7.4% estimated by Ionides and Blake
(1939) for the northern half of Jordan and 8-10% estimates by Burdon and Quennel
(1959) for the Yarmouk River catchment.
The fair agreement of the regional values with the few hydrological estimates
known for the area indicates that the estimates are of the right order. However Penman
accepts an accuracy of ±1O% for his estimate of potential evaporation as satisfactory;
with the assumptions involved in applying this method of estimating recharge an
accuracy of ±15% might be expected for any single year's estimate.
6.4.2 WATER BUDGET
Groundwater recharge by rainfall occurs when excess rainfall is greater than the
potential evapotranspiration and when the soil moisture storage capacity is full. The
excess rainfall is the amount of rainfall available after surface runoff has been
subtracted. The surface runoff is subtracted even though it may partly become indirect
recharge downgradient (see indirect recharge section below).
The budget equation used in the previous method was again applied, but this
time using the mean monthly values. The monthly rainfall exceeding monthly
evapotranspiration and surface runoff was accumulated and added to the total soil
moisture storage capacity at the end of the season to give the total annual recharge for
the groundwater system.
The estimated groundwater recharge using this method are listed in Table (6.5).
The results are found to be consistent with the results from the previous method. The
differences in using this "water budget" method over the soil moisture balance method
187
are: (i) the "water budget" method deals with mean monthly values without taking into
consideration or including the effects of single events, such as the effect of short tenn
variation in soil moisture deficit or the presence of long dry periods on groundwater
recharge; (ii) potential evapotranspiration values are used; and (iii) during consecutive
storm events and when the soil moisture storage capacity is full, and according to the
previous method, any access rainfall after surface runoff, assumed to be groundwater
recharge, while the total soil moisture storage capacity at the end of the season were
considered in this method.
Hillel (1971) demonstrated that the field capacity is a time dependent parameter,
and the amount of moisture retained for a few days is much more than that retained for
a long period. Thus using variable soil moisture contents with time in recharge
calculations will produce more realistic results than using the total average value at the
end of the season. Season 198011981 for example, received high rainfall, but at the
beginning of the season when the actual evaporation was assumed to be solely
dependent on the amount of rainfall: thus most of the rainfall in that period was assumed
to be consumed by evaporation, and consequently only 12.03 MCM were estimated as
direct recharge by using the soil moisture balance method. This corresponds with 119.45
MCM estimated for the same period by the water budget method.
BASIN 1980/81 1981182 1982/83 1983/84 1984/85 AVE. UPPER Rainfall 133.4 112.8 171.6 98.7 116.8 126.7 ZERQA Recharge 22.47 3.75 33.53 0.384 13.34 14.70
% 17 3 20 0.4 11 10 WADI Rainfall 155.7 108.6 211.6 92.5 155.8 144.8 WALA Recharge 31.3 9.0 37.7 2.1 11.4 18.3
% 20 8 18 2 7 11 WADI Rainfall 192.46 139.74 186.88 98.04 139.46 151.32 MUnB Recharge 48.68 14.32 43.86 0.0 14.04 24.18
% 25 10 24 10 14 WADI Rainfall 54.7 32.3 52.85 30.25 30.55 40.13 RASA Recharge 8.25 0.0 1.0 0.0 0.0 1.85
% 15 2.0 3.0 Rainfall 77.5 74 92.25 22.5 53.25 64
JAFR Recharge 8.75 0.0 20.0 0.0 0.0 5.75 % 11 22 7
Rainfall 613.76 467.44 715.18 341.99 495.86 526.9 TOTAL Recharge 119.45 27.07 136.09 2.48 38.78 64.8
% 20 6 19 1 8 11
Table (6.5) Results of direct recharge calculation (in MCM) for the B2/A7 aquifer system by using the water budget method.
188
45
40 ...... C 35 "E 30 Q) ·u Ii: 25
8 20 Q)
15 e> ra .c 10 0 Q)
a:: 5
0 0 100 200 300 400 500 600 700 800
Rainfall (nm)
Figure (6.5) Relation between rainfall and recharge coefficent ('Yo)
40
35 • ...... ~ ~ 30 "E Q) 25 ·u Ii: 8 20 u Q) 15 e> ra
10 .c 0 Q)
a:: 5
0 0 50 100 150 200 250
Recharge (nm)
Figure (6.6) Relation between recharge and recharge coefficient ('Yo)
Using the water budget method, recharge was found to be most sensitive to
amount of total rainfall. Estimates of the recharge coefficient, the recharge derived from
a given area as a percentage of the annual rainfall, increases from 0 % for annual rainfall
below 200 rom to a maximum of 37 % for annual rainfall of more than 700 rom (Figure
6.5). A better correlation exists between the amount of recharge and the recharge
coefficient, its polynomial relation indicating the change in the rate of change of
recharge as a function of recharge coefficient (Figure 6.6).
189
6.5 INDIRECT RECHARGE
As discussed earlier, significant direct recharge occurs only in the Western
Highlands where the annual rainfall exceeds 200 mm, the threshold value for recharge to
occur. Based on the rainfall record, the majority of the study area to the east of the
Western Highlands into the eastern and south-eastern deserts, lies in an annual rainfall
zone of less than 250 mm, thus groundwater recharge in these areas depends mainly on
indirect recharge through the infiltration of runoff accumulating in the wadi beds during
rainstorms, and the water transfer into the system from other groundwater systems.
Indirect recharge also may occur through man made structures such as dams and flood
control systems. These need not be considered in this study.
6.5.1 RECHARGE THROUGH WADI BEDS
It is believed that the direct recharge estimates are low insofar as they do not
take full account of local concentration of runoff water. Positive evidence that indirect
recharge takes place is available from environmental isotope studies (Parker, 1970).
Groundwater with high tritium concentrations (60-150 TV) were found in several wells
which tap the B21 A 7 aquifer system in areas where the climatic conditions are such that
the majority of the water is unlikely to have originated from direct recharge.
Infiltration of wadi runoff into the aquifer can only occur along the wadi beds
and slopes where the aquifer layers are exposed or at shallow depth. Recharge will only
take place providing storms are of sufficient intensity and amount to generate sufficient
runoff and overcome the soil moisture deficits in the wadi courses. Runoff estimates
show that only long duration rainstorms of intensity which exceed the initial abstraction,
discussed in Chapter 3, are capable of generating runoff. Initial abstraction is found to
range between 8 and 20 mmlday, depending on the sub-catchment hydrological
characteristics.
This mechanism could also account for the part of the recharge which results
from short duration, high intensity storms in desert areas where the conditions necessary
for direct recharge are rarely fulfilled.
In order to estimate the quantities of potential infiltration through the wadi beds,
the total effective rainfall over the sub-catchments must be estimated and converted to
190
an equivalent depth of runoff in the wadi courses. A part of this runoff flows out of the
sub-catchments and another is used to satisfy the soil moisture deficits of the alluvial
deposits in the wadi courses. The rest is assumed to infiltrate and thus become aquifer
recharge. Evaporation of the concentrated runoff during the storm events is usually
negligible because of the short times involved.
The volumes of runoff, as estimated by using the curve number technique
discussed earlier, represent the amount of effective rainfall or the excess rainfall
available for runoff rather than the total amount of runoff reaching the catchment outlet.
The amount of runoff which leaves the catchments must be accurately estimated and
subtracted from the total surface runoff in the catchments to obtain the amount of runoff
available for infiltration through the wadi beds. But due to the complexity of the
drainage systems and the scarcity of the gauging stations, the values of runoff measured
at the gauging stations at the mouth of the main wadis throughout the study area are
doubtful. Consequently, it is believed that using these values for the purposes of indirect
recharge calculations would lead to unreliable results. The average values for Upper
Zerqa and Wadi Mujib basins might be used with caution to give approximate values of
indirect recharge in these basins, as these data show some kind of correlations between
the estimated volumes of runoff in the catchments and those measured at the gauging
stations.
Indirect recharge can be also estimated from the difference between total and
direct recharge. The total recharge can be taken as the total natural losses from the basin
through the spring discharge and baseflow on the major wadis. However, the baseflow
records are of inadequate length and reliability, and furthermore the whole method does
not account for the proportion of the recharge which discharges to the Jordan Valley as
subsurface flow. Therefore, this method has not been used.
An attempt was made to estimate the effect of groundwater recharge through the
wadi beds by using Darcy's law:
K 1= - x H x B x L ............................................................................. (6.2)
M
where I = infiltration (m3/day)
191
K = permeability (m/day)
M = wadi bed thickness (m)
H = head of water in the wadi (m)
B = wadi bed width (m)
L = wadi bed length (m).
The area of aquifer forms outcrop and subsurface layer are estimated from the
geological map by assuming a wadi width of 15 m all over the study area. The
permeability of the alluvial deposits along the wadis was taken as 1 m/day for all the
study area. And thickness of the wadi bed was assumed to be 1 m.
For estimating the head in the wadi beds caused by runoff, the Manning equation
for calculation of the average velocity of water through channels was used:
v = .!..R2/3Sl/2 ...................................................................................... (6.3) n
where V = the average velocity in (m/sec).
R = the hydraulic radius, or the ratio of the cross-sectional area of
flow in (m2) to the wetted perimeter in (m).
S = the slope of the water surface.
n = the Manning roughness coefficient.
The velocity of flow is dependent upon the amount of friction between the water
and the stream channel. The U.S. Geological Survey has published a series of
photographs of rivers for which the value of the Manning roughness coefficient has been
computed (Barns, 1967). In the study area, the Manning roughness coefficient is given
a value of 0.045 for the mountain streams with rocky beds.
The cross section of the wadi beds where infiltration occur were assumed to be
rectangular. Then, the hydraulic radius is given as:
BxH R = ......................................................................................... (6.4)
B+2H
192
and thus
v = ;(::2:) 2/3 /' ...................................................................... (6.5)
The flow in the wadi bed (Q) with a cross-sectional area (B x H) and average
velocity (V ) can be expressed mathematically as:
Q = V x B x H ..................................................................................... (6.6)
and thus
V = Q ........................................................................................... (6.7) BxH
substituting Eq. (6.7) in Eq. (6.5) will give new equation which relates the flow in the
wadi beds to the geometry of the channel bed and the pressure head:
B; H = ;(::2:) 2/3 /' ••••••..••••••••••••.•••••••••••••••••••••••••••••••••••••••••• (6.8)
The slope is the drop in elevation over the length of measurements: it is
estimated from topographic and geological maps to range between 0.01 and 0.005.
By assuming the flow in the wadi bed is equal to the amount of runoff estimated
by using the curve number method (Chapter 3), and since the wadi bed width and the
slope of the water surface are known, the water head can be found by solving Eq. (6.8)
for H.
The amount of indirect recharge can be estimated for each sub-catchment by
solving Equation (6.2). The calculations indicate that the total annual indirect recharge
to the B2/A7 aquifer system ranges between 2.9 and 37.8 MCM with an average of
about 14 MCM (Table 6.6). It varies between the basins and from year to year within the
same basin according to the hydrological characteristics and the annual rainfall amounts
193
and pattern in each basin. The Rasa and Jafr basins, despite their relatively low annual
rainfall, were found to receive the highest indirect recharge in the study area. This is
explained by the presence of longer wadis in the basins.
Although the amount of indirect recharge obtained by using the above method
was found to be reasonable and agreed well with the total recharge and the water
balance for each sub catchment, the results are dependent on the validity of the
assumptions of the permeability and thickness of the wadi beds. If the permeability and
thickness of the alluvial deposits of the wadi beds depart from the values used in the
calculations, then the results will be in error by a factor proportional to the error in the
permeability/wadi bed thickness.
Basin 1980/81 1981/82 1982/83 1983/84 1984/85 Average
Upper Zerqa 1.5 0.6 1.4 1.2 1.4 1.2 Wadi Wala 3.2 0.8 2.3 1.0 3.7 2.2 Wadi Mujib 4.4 0.5 2.5 1.6 2.6 2.3 Wadi Hasa 4.6 0.1 2.9 2.4 1.2 2.2 Jafr 24.1 0.9 1.3 0.3 3.5 6.0 Total 37.8 2.9 10.4 6.5 12.4 14.0
Table (6.6) Results of indirect recharge calculation (in MCM).
6.5.2 LATERAL BOUNDARY FLOW
All natural recharge originates as rainfall, but the routes by which water enters
the aquifer system vary considerably within the study area. Recharge also occurs by
lateral flow within the aquifer system across the study area boundaries from the adjacent
high recharge mounds.
Part of the B2/ A 7 outcrop which occurs beyond the boundary of the study area
receives recharge that moves eastward into the study area. To estimate the quantity of
recharge, the assumption is made that the recharge rates in the recharge mound areas are
equal to the lateral flow in the aquifer system away from those areas. The recharge rate
(Q) can therefore be estimated by Darcy's equation:
Q = -KA( ~~) .......... ~ ................. : ....................................................... (6.9)
where K is hydraulic conductivity
A is cross sectional area of flow
194
dh . h dr 1· . dl IS Y au IC gradIent near the recharge mounds.
The hydraulic conductivity was set as 1m/day for all the areas, while the cross
sectional area of flow and the hydraulic gradient were estimated from the piezometric
map. The estimated recharge over those parts of the outcrop that are assumed to
contribute to the regional system are shown in Table (6.7).
Recharge Mound Basin Recharge (MCM/a)
Amman Upper Zerqa 1.278
Wadi Wala 7.300 Rabba Wadi Muijb 5.000 Mazar Wadi Mujib 4.563
Wadi Hasa 3.000
Tafila Wadi Hasa 6.083
Shaubak Jafr 2.5
TOTAL 29.724
Table (6.7) Recharge estimation from lateral boundary flow.
6.5.3 WATER TRANSFER
It has been shown that recharge to the main aquifer system occurs in the Western
Highlands. To the east, the aquifer is buried under a thick pile of impervious sediments
which prevent replenishment from rainfall or runoff. Hydrological evidence, however,
suggests that the confined part of the aquifer system is recharged by lateral flow from
the Western Highlands. This evidence is corroborated by the westerly decrease in
groundwaters salinity, and the presence of small quantities of tritiated water. However,
as the flow from the west is already within the aquifer system and does not contribute to
the amount of total recharge, it will be discussed in other parts of this study.
Within the study area, the only contribution to the total recharge of the B2/ A 7
aquifer system is the upward leakage of groundwater from the A4 aquifer system
through the intermediate marls of the AS/6 Formation. This phenomenon will discussed
in detail in groundwater flow (Chapter 7) and groundwater modelling (Chapter 8).
The only known source of transferred water into the B2/ A 7 aquifer system is the
basalt aquifer to the north of the study area (Figure 6.1). The basalts directly overlie the
B2/ A 7 aquifer system, and when they are in contact, there is hydraulic continuity
between the two aquifer systems. The piezometeric surface map and hydrochemical
195
studies show that water flows from the basalt aquifer into the B21 A 7 aquifer system in
the Mafraq-Wadi Dhuleil area (Parker, 1970). Raikes and Partners (1962), MacDonald
et al. (1965a), and Parker (1970) assumed substantial amounts of groundwater transfer
from the basalts to the B2/A7 aquifer system in the Upper Zerqa Valley.
The water transfer occurs in the study area at the northern boundary, in the
outflow area of the Upper Zerqa catchment, where it joins the water flowing from the
B2/A7 aquifer system to flow out the system along the lower part of the Zerqa River. Its
contribution to the total recharge of the B21 A 7 aquifer system in the study area is
discussed in more detail in groundwater modelling (Chapter 8).
6.6 TOTAL RECHARGE
The total groundwater recharge is the sum of direct and indirect recharges to
the aquifer system. It can be estimated from the groundwater balance, analytically by
using the groundwater flow equations, or by analysing the response of groundwater to
recharge events. On the assumption that the total outflow from the aquifer system due
to natural losses reflects the replenishment, total recharge can be calculated from the
water balance of the basins; this will be discussed below, in the water balance section.
Direct estimation of the total recharge can be achieved using the groundwater
flow method: the flow across a plane, as calculated by Darcy's law, is assumed to be
equal to the net recharge up gradient from that plane. This calculation is made by
analytical or numerical techniques ( see groundwater modelling, Chapter 8). These
methods assume that the flow system is in equilibrium and that the aquifer properties
are estimated correctly.
6.6.1 GROUNDWATER BALANCE
Total recharge to the groundwater system, as discussed in the previous sections,
includes direct and indirect recharge, and surface and subsurface inflow from adjacent
areas. The soil-water balance method for calculating direct recharge was found to be
reliable under the average conditions, and its results were adopted during this study. It
estimates direct recharge to be 8% of the total annual rainfall, compared with 11 %
estimated from the water budget method. Indirect recharge is an important component of
196
the total recharge, particularly in the eastern and south-eastern areas, and it is likely to
be at least of the same order of magnitude as the direct recharge.
For a steady-state balance of the B2/A7 aquifer system, the following items have
to be considered:
1- The total recharge to the system as:
a) direct recharge (Qd)
b) indirect recharge (Qi)
c) lateral boundary inflow (QnJ
d) vertical leakage (Qv+)
e) water transfer (QJ
2- The total discharge from the system as:
a) spring discharge (Qs)
b) subsurface groundwater outflow (Qt.)
c) vertical leakage (Qv-)
d) hidden discharge and unmeasured seepage (Qh)
Hence:
Some of the terms in this equation, such as the amount of inflow and outflow due to
the vertical leakage and subsurface flow are difficult to measure. However, it must be
stressed that this is a preliminary approach to understand the quantitative relationships
between recharge and discharge. Detailed water budgets will be considered in
groundwater modelling (Chapter 8).
The only source of vertical leakage into the system is about 2.16 MCM/a
which flows from the A4 aquifer system upward through the AS/6 aquitard into the
B2/ A 7 aquifer system in the Upper Zerqa Basin.
The amount of groundwater loss from the system due to subsurface outflow
and vertical leakage is more significant. A subsurface flow term (F s) is introduced to
include the total outflows due to the vertical leakage and subsurface outflow. Fs is
estimated as the net difference between the total recharge and the measured discharge.
197
The results of the water balance calculations for the B2/A7 aquifer system in the study
area are shown in Table (6.8).
Flow (MCM/a) Upper Zerqa Wadi Wala Wadi Mujih Wadi Hasa Jafr Qd 12.61 18.16 15.00 0.40 4.38
Inflow Qi 1.21 2.20 2.32 2.23 6.00
Qv+ 2.16 Qm 1.28 7.30 9.56 9.08 2.50 Qt 3.8
Total 21.06 27.66 26.88 11.71 12.88 Outflow Qs 12.67 15.09 5.89 1.49 1.30
Fs 8.39 12.57 20.99 10.22 11.58 Total 21.06 27.66 26.88 11.71 12.88
Table (6.8) Groundwater balance of the Amman/Wadi Sir aquifer system
6.6.2 THE RESPONSE OF GROUNDWATER TO THE TOTAL RECHARGE
The effect of total recharge on groundwater appears in terms of fluctuating water
levels in boreholes, spring flows, and chemical and isotopic composition of
groundwater.
Groundwater levels, baseflow, and other natural discharges of groundwater will
remain in stable dynamic equilibrium as long as no artificial recharge or discharge is
imposed on the system (Theis, 1940). Stable dynamic equilibrium does not mean that
the system will remain static, but rather that the system will be in a constant state of
flux, with changes in the rate of recharge being compensated for by changes in the
groundwater levels and changes in natural discharge. When artificial discharge is
imposed on the system, the approximately stable dynamic equilibrium is disrupted.
Changes must occur in the groundwater levels and/or natural rate of discharge to
compensate for the artificially imposed discharge.
6.6.2.1 WATER LEVEL FLUCTUATIONS
In most hydrological studies, fluctuations in groundwater heads are used either
to calculate recharge directly (Freeze & Cherry, 1979) or to provide a calibration control
for groundwater resource modelling (Rushton & Redshaw, 1979). In such studies the
head fluctuations are predominantly a function of the recharge, transmissivity, and
storage.
198
Groundwater responds to recharge by rising water levels and increasing spring
flows during and after the rainy season. It is expressed by the equation:
Recharge = Drainage + I1W ......................................................................... (6.11)
where Drainage is the volume of water loss as spring discharges and seepage along
wadis and rivers ( the baseflow); and 11 W is the change in water storage which is a
function of the change in water table levels.
Observation well hydro graphs particularly in the Western Highlands indicate
that the fluctuations in water levels are clearly related to rainfall and thus indicative of
recharge to the aquifer. Water levels begin to rise at the onset of the wet season (October
and November) and reach their peak in March and April. While in the east where the
aquifers are confined by a thick cover of marls, the hydro graphs do not show the same
response. An example of water level fluctuations as a function of recharge is shown in
Figure (6.7). It also shows the influence of changes in transmissivities and the depths to
static water level on the magnitude of water level response to recharge.
To estimate recharge from the change in water levels, the relationship between
storage and water level must be defined. This require delineation for the area
represented by particular borehole hydro graphs. And water level fluctuations not related
to recharge must be recognised (e.g. those resulted from change in barometric pressure
and change in abstractions).
Water level fluctuations are an absolute indication of groundwater recharge. But
the sparcity of data for the groundwater system remote from abstractions precludes this
approach for calculating recharge on more than a local scale.
6.6.2.2 SPRING DISCHARGES
It is difficult to determine the actual groundwater discharge to rivers, because
most of the spring flows are abstracted upstream for agricultural and domestic uses, and
the hidden discharge and seepage along the rivers are difficult to quantify. Furthermore
the methods used in baseflow separation from the total runoff are not consistent. Hence
any analysis of groundwater recharge based on stream flow measurement is tenuous and
subj ect to large errors.
199
E .s ~ c: ' (ij c::
:g; .s
80
70
60
50
40
30
20
10
0
123
~ 123.5 ~
124
19
:g; 19.5
.s Qi 20 > ~ ... 2 ro 20.5 ~
21
6.5
7 :g; .s 7.5 Qi > ~ 8 ... 2 ro ~ 8.5
9
0
0 N D
Lajun No.4 8=840 rrasl
T=728 m2/d
Lajun No.1
8=673 rrasl
T=45 rr?1d
Lajun No.9
8 =686 rrasl
T=23 rr?1d
N D J
Monthly rainfal for Karak station in the year 1985/86
J F M A M J J A s
F M A M J J A S Months
Figure(6.7) Groundwater level fluctuations due to rainfall in the year 1985/86
200
An attempt was made to use measured spring discharges for recharge estimation.
Unfortunately springs with full discharge records are difficult to find. For the period of
which the recharge were calculated (1980-1985), only 1985 was found to have an
acceptable record for some springs. Therefore, analyses of spring flow data for different
periods were used in recharge estimations. Ras el Ain spring in Upper Zerqa catchment
has a continuous record for three years (1987/1988-1989/1990) and therefore is used to
demonstrate the methodology of estimation recharge from the spring hydro graphs.
The catchment areas of many large springs cover tens of square kilometres and
do not necessarily coincide with surface water drainage boundaries. Any estimation of
recharge thus represents only the amount of recharge on the spring catchment rather than
the recharge to the entire system. Therefore, emphasis is' placed upon the relation
between rainfall and recharge derived from spring flow analysis which could be applied
for the entire system providing the other hydrological characteristics do not change
dramatically.
Spring discharge is a response to the recharge undergone by the aquifer primarily
from rainfall. Some springs respond rapidly and clearly to rainfall, while others behave
with more inertia and less correspondence to the amount of rainfall. This difference in
the functioning of carbonate aquifers results from the existence of quickflow due to the
circulation of water through a network of drains formed during the karstification
process, and baseflow due to the circulation of the water through the carbonate matrix
and small fissures.
Ras el Ain spring is a fault spring type situated in the stream bed of Wadi
Abdoun in the western part of Central Amman (Figure 6.8). The spring discharges from
the B2/ A 7 aquifer system, which is separated from the deep A4 aquifer system by the
thick marls of A5/6 aquitard (Figure 6.9). It is believed that the two aquifer systems are
hydraulically interconnected in places where the system is highly affected by tectonics.
The average annual rainfall calculated for 48 years by using the Theissen polygon
method was found to be 500 mm. The infiltration rate for the whole Upper Zerqa
catchment as discussed earlier ranges between 1-28 % with an average of 8 % of the
annual rainfall. Consequently the annual infiltration volume in the catchment area of the
Ras el Ain spring will be in the order of 2,728,000 m3/a for the 68.2 km2 0fthe
201
~JUBEIHA 235
• 30
J/.~ ./~/\
'-...... _/ \.
AMMAN • ..--. ."-
~,,'- . \ / • " . 31
\ I 0 > \ . ( \ ! ... ._ . ..i . I /
• 37
155
150
145 o 2 3Km. ~' ~~S;;;;;iiiii~! ~~'
. . .
Figure (6.8) Location map of the Ras el Ain spring showing the Wadi Abdoun sub-catchment.
catchment. The average discharge rate of the Ras el Ain spring is 0.115 m3/s which
corresponds to 3,612,817 m3/a, 53 mm, or 10.6 % of the annual rainfall. With the above
values, the hydrologic balance shows water surplus of 884817 m3, which is equal to a
discharge of 0.028 m3/s, which may be captured from neighbouring catchments.
HYDRO GRAPH ANALYSIS
For karstic water resources evaluation, the interaction between surface and
subsurface parameters should be examined. A number of methodologies have been
proposed in order to obtain an approach for understanding the hydrodynamic features of
the karstic systems according with their discharge data (Schoeller, 1967; Drogue, 1972;
Galvov, 1972; Mangin, 1970, 1975, 1981, &1984). These methods, usually applied to
surface water hydrology systems, are also a valuable tool to study the flow system of the
karstic aquifers.
The most frequently used method to analyse karstic spring flows in dry season is
recession hydro graph analysis, primarily based on Maillet's simple exponential equation
(Todd, 1959), which also can be derived from linear reservoir theory (Meier, 1980).
The discharge-time curve of Ras el Ain spring for the period 1987-1990,
shows extreme variations in flow between different years as well as within the same
year (Figure 6.10). The correlation between the rainfall and the discharge from the
spring is very significant.
It is well observed that the peak discharge is mostly created by a sequence of
rainfall events but not any individual rainfall event. So rainfall sequences which
correspond to the peak discharge period should be analysed. Plotting the sum of
rainfall onto the area from the beginning of the wet period, versus the maximum flows
recorded in the hydro graph (Figure 6.10) from the first response till the
maximum peak value recorded for that year , shows the threshold rainfall amount
required for flow to occur (Figure 6.11). This threshold value of rainfall, about 70 mm
(Figure 6.11) is assumed to be that infiltration necessary to saturate the soil. It is clear
that if the rainfall amount recorded is less than the threshold value, then no peak will
occur.
203
........ Iii til g II) "C
~ <
E S ~ c:: 'iii a::
1200
Ras el Ain spring
1000
800 A3
600 A3
A3
400
200
O L-____________________________________________________ ~
West.
Figure (6.9) Geological cross section of Ras el Ain spring
250 ,..--------=---------:=----------- - ---,- 0.4
150
100
50
0.35
0.3
0.25
0.2
0.15
0.1
0.05
ODFA JAODFA JAODFA JACD 1987 1988 1989 1990
Figure (6.10) Average monthly values of Ras el Aln spring discharges and rainfall in Wadi Abdoun Basin
204
Eas
Subtracting the threshold value from the accumulated infiltration, calculated as
8 % of the rainfall rate, shows the possible infiltration rate into the system (Figure
6.12). It is clear that the effective recharge to the system occurs after a reasonable
period since the rainfall period elapsed, which agrees with the first response indicated
by the discharge hydro graph shown in Figure (6.10); the delay differs from year to
year depending upon the rainfall intensities.
600
- 500 E .s ~
400
c .~
300 '0 Q) -ro :; 200 E ;j 0 0 « 100
0 0 0.05 0.1 0.15 0.2 0.25
Peak flow (m3/sec)
0.3 0.35
Figure(6.11) Peak flow versus accumulated rainfall
0.4
RECESSION HYDRO GRAPH ANALYSIS
To express the flows in a river fed only by groundwater during a recession
period, Maillet proposed in 1905 the formula (Todd, 1959):
Qt = Qo* exp[-a (t - to )] ........................................................ (6.12)
where Qo is the initial discharge rate at time = to and a is the aquifer discharge
coefficient in units of the inverse of time. This formula can also be derived by the
integration of the first order linear reservoir model due to the initial condition
Qt = Qo at t =to (Meier, 1980). Taking the first logarithms of both side of Eq.6.12
yield for a , which is the slope of a log Qt versus t plot:
205
50,-____________________________________ --.
45
E 40
.s 35 c: ,g 30 I!! ~ 25 .5 "0 20 ~ "5 15 E 10
~ 5
O+-____ ~--~~~~----~----~----~--~ o 30 60 90 120 150 180 210
lime since rainfall ellapsed (days)
Figure (6.12) Accumulated infiltration calculated after subtracting the threshold value
In(Qol Qt) 0.= V-to) ............................................................................... (6.13)
Recession hydrograph analysis was carried out with monthly average March
December flows at Ras el Ain spring for the year 1988. The complete recession curve
plotted on semi-logarithmic paper is shown in Figure (6.13). It begins with an initial
discharge Qo of 0.377 m3 Is and ends with a discharge of 0.118 m3 Is after t =270 days.
After ti =180 days (the begining of exhaustion time) it shows another segment with a
discharge rate of 0.140 m3/s. The extension of the best fitting line to the ordinate gives
the initial exhaustion discharge QEo =0.195 m3 Is. The exhaustion of the reserves of the
karst aquifer system, therefore, can be expressed by the formula:
Q Et = Q Eo* exp [-a. - (t - to) ]. ......................................... {6.14)
and the same for the exhaustion discharge coefficient a. -: .
_ In(QEo I QEt) a. = V _ to) ....................................................................... {6.15)
206
'0 Q)
.!!! Qo M g Q) ~ E!' <11 q, ..r::: U rn is
QEO ~
I- OC ~~Q, - _A
tilQB -
to t, 0.1
o 30 60 90 120 150 180 210 240 270
Tirre (days)
Figure (6.13) Analysis of recession curve of Ras el Ain spring
Subtracting month by month starting from t =0 of the values ofEq. (6.14) from those
of recession curve of Eq.6.l2 gives a new curve (Figure 6.14) described by the
function:
qt = Qt - QEt ................................................................................... {6.16)
Thus the spring discharge Qt at any time of recession can be expressed by the
formula:
Qt = QEt + qt ................................................................................... {6.17)
QEt, the exhaustion discharge rate, is the discharge rate in dry seasons when
infiltration ceases, and corresponds to Maillet's (1905) classic model. qt represents
the system emptying during infiltration. It is approached to nil at the end of the peak
at t =180 days (Figure 6.14) which is an optimum representation of the infiltration
versus time curve shown in (Figure 6.12).
207
c:> Q) If) -C')
S -c-
0.18
0.16
0.14
0.12
0.1
0.08
0.06
0.04
0.02
0 0 30 60 90
time (days)
120 150 180
Figure(6.14) Graph of qt versus time for Ras el Ain spring
However Mangin (1970, 1975) proposed a new function to represent the
system emptying during infiltration:
I-tV qt = qo-- .................................................................................. (6.18) 1 + Et
where qo = Qo - QEo
h . . h 1 11 = t e mfiltration rate coefficIent were ti = -11
E = the outflow heterogeneity coefficient.
Mangin (1970) stated that" 11 determines the outflow rate i.e. it is inversely
proportional to the duration of infiltration for the same qo (infiltration discharge rate
at t =0); and E determines the concavity of the infiltration curve, i.e. the higher its
absolute value, the more the initially rapid infiltration slows down".
Calculation of qt for all the 180 days between t = 0 and ti for the Ras el Ain
spring data by applying Eq. (6.18) shows an excellent fit with the values obtained
from Eq. (6.16).
Integration of the exhaustion function, Eq. (6.14) between t =0 and t =t gives
the volume stored in the system when exhaustion begins:
208
t=t
V = JQEt dt ................................................................................... (6.19) t=O
which corresponds to the potential rather than to a regulating reserve or storage.
Integrating Eq. (6.16) or Eq. (6.18) between t =0 and t =ti gives the infiltration
volume derived from precipitation causing the flood peak:
t=ti
Vp = Jqt dt ................................................................................... (6.20) t=O
Data derived from the general recession curve model are shown in Table (6.9).
Year 1988 1989 1990
Flainfall (rnrnJa) 545 419 444
Flecharge(m) 2591600 1904144 2040544
Q (m'/a) 6815478 5652788 4622000
Qo (m'/sec) 0.377 0.281 0.252
Qeo (m' /sec) 0.195 0.132 0.110
qo (m'/sec) 0.182 0.149 0.142
t(day) 270 210 270
ti (day) 180 150 150
a. 0.005503 0.006313 0.007167
a. - 0.00186 0.001226 0.001641
V (m') 9058065 9153130 5791590
Vp (m') 1144058 793552 737443
VPm (m') 1038922 748556 661651
... Vp and Vpm' denote the results from the mtegratlOn of Eq. (6.16) and Eq. (6.18) respectively.
Table (6.9) The hydrological parameters of the recession curve model for Ras el
Ain spring.
209
HYDROCHEMICAL ANALYSIS
Fluctuation of discharge and water chemistry were studied together during the
hydrological cycles of the Ras el Ain spring system. Figure(6.15) shows the evolution
of the chemistry during the hydrological cycles for the period 1987-1989.
The chemistry of the water flowing out ofRas el Ain spring is characterised by
small electrical conductivity (EC) ranges (540-620 uS/cm) with an average of 582
uS/cm. The EC as well as the major ions generally decreases with high discharges.
The delay between the concentrations and discharge peaks varies with chemical
parameter as well as between different hydrological cycles, and ranges between 60-
120 days.
The fluctuation in water chemistry reflects periods of recharge and draining.
However, some elements (Ca, Mg, and N03) show an increase in concentrations at the
beginning of winter, indicating a possible storage of those ions in the unsaturated zone
during summer and restitution as a recharge takes place. This phenomenon is also
proved by the Ca-Mg and Na-K deficiency graphs (Figure 6.15).
A dilution calculation for the different chemical parameters indicates 12-28 %
fresh water componepts, which is equivalent to 1.08-2.52 MCM of annual recharge
for the year 198711988. This ratio varies from year to year depending on the recharge
amount. The relationships between rainfall, discharge and chemical hydro graphs
indicate that the high rainfall in 1987/1988 gives higher discharges with lower EC,
while the lower rainfall in 1988-1989 gives lower discharges and higher EC.
The expulsion or drainage of the old water stored in the aquifer during summer
IS observed from the increase in the chemical concentration. However, the little
changes in water chemistry after a long period of low discharge are not significant
enough to explain the arrival of water from other aquifer systems, such as the A4.
Although the water chemistry of the A4 aquifer system is characterised by having
relatively higher Mg than the B2/ A 7 aquifer system (Salameh and Khdier, 1985), the
decrease in CalMg ratio is not indicative (Figure 6.15), and most likely to be caused
by the slow seepage from other less permeable geological formations.
210
4.5 -r----------------, --Ca
~ ----teO S 4
5
~ !'l 3.5
\ \..,
\ \ , \ , 'pI
5 <..>
ONDJFMAMJJASONDJFMAMJJA 1987 1988 1989
2,----------------------------, •.••••• N __ CI
~E _ 1.5
~... ..-. ::- ."'.
", ...... p.'oo.. ............ • ............ - ....... ,
0.5 +--i_+_+_~_+_+_+_i_+_+_+_i_+_+_+_+_++_+-+_l
ONDJFMAMJJASONDJFMAMJJA 1987 1988 1989
____ 504
~ __ Il103
S ,g 0.5
I " '" '"' , \" , ... I v ',,_ .#.-'" I " \ I \,... I '--" ./"
O~-++-~~~~~-++-~~~~~~
ONDJFMAMJJASONDJFMAMJJA 1987 1988 1989
0.4,-___________________ ...,. 650
0.3
_Row _8:
! S 0.2
~ u.. 0.1
ONDJFMAMJJASONDJFMAMJJA 1987 1988 1989
T1rre (11"Ilrths)
Figure (6.15) Chemical analysis of Ras Ain Spring
Mg
0.5 ~~~-++-r-+~~~+-i-++-I-+~~_l
ONDJ FMAMJ JA SONDJ FMAMJ JA 1987 1988 1989
5.5
5 Ca/Mg -- ---------------------
E' g 4.5 .§. c 4 ~ ~ 3.5 c Q) <J 3 c 0 0 2.5
2 ON D J F M AM J J A SON D J F MA M J J A
1987 1988 1989
2
~ (Ca+Mg)-(I£03+S04)
CIl
S 1.5 c:
i !'l c:
8 0.5
ONDJFMAMJJASONDJFMAMJ JA 1987 1988 1989
E' CI-(na+k)
I c:
~ 0.5
c: CIl <.> c: 0
<..>
0
ONDJFMAMJ J A SON D J F MA M J J A 1987 1988 1989
T1rre (rronths)
ENVIRONMENTAL ISOTOPES ANALYSIS
After a dry period, the rainfall events stimulate up to a 3-fold increase of
discharge. This effect is well correlated with the evolution of chemistry. The arrival of
the main fresh water component after the maximum discharge is marked by the
increase of oxygen-18, tritium, and deuterium concentration. The ratio between fresh
and old waters can be calculated from the concentration of these isotopes in the spring
water and precipitation in the area. Figure(6.16) shows the evolution of the
environmental isotopes of the Ras el Ain spring during the period 1987-1989.
However the hydro graph of the isotopes is more complicated and the arrival
of different peaks for the same element within the hydrological cycle -due to the
arrival of different types of water- needs more detailed study and interpretation. And
there needs to be more research and investigation on isotopical analysis for every
single event of rainfall together with its corresponding response in the discharge
hydro graph. For example, at the beginning of winter, the early rainfall events infiltrate
through the gravel and the alluvium of the upper system causing the first impulse of
discharge of low EC and environmental isotope concentration. The second part of the
rainfall enters the main aquifer system after a time lag, and pushes out water with
similar chemical and isotopical concentration as in low discharge periods. The
successive rainfall events lead to mixing with the main water body of the reserve
activating the dilution process. This is marked by low EC and higher isotope
concentration. This process lasts till the end of the hydrological cycle depending on
the amount of the fresh water component.
CONCLUSION OF THE METHOD
According to the hydrological balance of the Ras el Ain spring basin, there is
surplus water, even after modifying the different hydrological parameters. To correct
the hydrological balance there are four possibilities:
I-the precipitation is greater than that has been used,
2-the discharge at Ras e1 Ain spring is overestimated,
3-the estimate of the infiltration rate value is too low,
4-the hydrological basin is larger than has been thought.
212
0.4
0.3
~ '" ;::;.
0.2 §.
~ u.
0.1
0
o N 0 J F M A M J J A SON 0 J F M A M J J A 1987 1988 1989
Time (months)
Figure (6.16) Environmental isotopes analysis of Ras el Ain spring
213
The precipitation in the area cannot be increased, because it is already the
highest in the area. Nevertheless, modifying the estimates of precipitation and
discharge rate would have a minimal impact on the balance.
Eq. (6.20) indicates that 1144058 m3 of the water discharged at Ras el Ain
spring is derived from precipitation causing the flood peak, only during the depletion
period, an equal amount or less is assumed to be already discharged before depletion
started. This means that, in the most likely cases a total of about 2288116 m3 of water
would be derived from precipitation for the year 1988. This estimation is very close to
the value obtained by assuming 8 % of infiltration. Although there is a possibility of
increasing the infiltration rate for the alluvial deposits, these deposits are very local
and their water contribution to the system is minimum.
Furthermore, hydrochemical studies and dilution calculations indicate 1.08-
2.52 MCM of fresh water component to be derived from precipitation. This wide
range in estimation is due to the difficulties in interpreting the hydrochemical
evolution. Precipitation does not occur in a measurable amount at the spring, but it
activates stepwisely different water bodies in the aquifer with different time lags.
In conclusion, the only way to correct the hydrological balance is to increase
the assumed size of the hydrological basin. This will only provide a further
undergroundwater flow to the system.
From the recession hydro graph analysis, the discharge coefficient (a) is rather
high and ranges between 0.005503-0.007167. The higher a, the more rapid the
decrease in discharge. Comparisons between a and the total annual discharge for the
period 1988-1990 reflects the dependence of a on the hydraulics of the system. It
suggests that the porosity in the upper system is relatively high. This can be explained
either by the contribution of the highly porous alluvial deposit waters or the high
intensity of karstification in the upper system.
The exhaustion coefficient a- is small and below the average for the
Mediterranean karst system. It suggests the slow depletion of a huge amount of water
stored in the aquifer system. Eq. (6.19) indicates about 9 MCM of groundwater
214
reserves for the Ras el Ain spring system, which can provide a baseflow for the spring
for long time in the ~likely condition of no recharge.
Recession hydro graph analysis has successfully explained the Ras el
Ain spring karstic aquifer system. The spring receives water from two different
systems; the upper system is relatively small and is locally limited to the lower part of
Wadi Abdoun catchment area. The lower system extends over a large area and collects
water from larger catchment area or from leakage from other geological formations.
APPLICATION OF THE METHOD
The original attempt to analyse spring flow data for the purpose of estimating
infiltration rate into the main aquifer system was constrained by the lack of long term
discharge records and the inadequacy of the discharge measurements.
Only a few springs were found to have an acceptable record. These are shown
in Table (6.10). Long term mean monthly discharge data were used to analyse the
flow system of these springs and Eq. (6.19) and (6.20) were applied to estimate the
potential storage (V) and the amount of recharge (Vp).
Except for Sukhna spring, comparison between the amount of recharge
produced from the peak flow analysis and the annual discharge, given the amount of
the annual rainfall in the spring sub-catchments (Table 6.10), indicates that even by
assuming plausible high recharge coefficient, the high discharge rate cannot be
maintained only by the recharge that the surface water catchment area provides.
Therefore groundwater flow from beyond the watershed to the springs must be
assumed. For example a large part of the groundwater flowing from the Mazar
recharge mound must be diverted to Ain Sarah by the drainage system of the Karak
fault line. This scenario could be proved by the observed head drop across the
southern flank ofthe Karak graben and the high discharge rate of Ain Sarah spring.
The steep segment in the recession hydro graph indicates local discharge or
water loss through the permeable karstic surficial aquifer, while the flattened segment
indicates regional baseflow component. Only for the Sukhna spring, the hydro graph
shows that the exhaustion discharge proceeds at a higher rate. However Sukhna spring
215
discharge from the limited reserv~ wadi fill and Hummar aquifer systems. Nimra
spring also discharges from the Hummar aquifer system.
Spring Year Rainfall(mm) Q (mO/a) a. a." Vp (m3) V (m')
Ras el Ain 1988 545 6815478 0.005503 0.00186 1144058 9058065
1989 419 5652788 0.006313 0.001226 793552 9153130
1990 444 4622000 0.007167 0.001641 737443 5791590
Zerqa Ave. 158 7168264 366534 12372943
Sukhna 1987 383760 - -1988 895723 0.001969 0.002751 41006 1884071
1985 183 863784 0.002523 0.001737 17286 1591827
Nimra 1986 1009728 0.003812 0.000353 86187 5384692
1987 897840 0.006404 0.001118 119215 1314049
1988 1116000 0.001627 0.00053 23743 5384692
W.Haidan Ave. 253 15037920 0.001883 0.00076 773376 54591899
1981 245 438120 0.006643 0.000675 5718 1535397
1982 193 358488 - 336552
Lajun 1983 303 355968 0.002068 0.001065 20937 973074
1984 162 305064 0.002571 0.000256 19618 3043097
1985 201 424800 0.003563 0.00095 61652 1245776
Sarah 1985 178 3816000 0.001066 0.000646 465853 16724654
Table (6.10) Results of recession bydrograpb analysis for some springs.
SPRING CATCHMENTS AND RECHARGE COEFFICIENT
The previous discussion and the unusual variation in spring discharges inspire
the idea of analysing the spring discharges and the necessary surface catchment area
needed to provide enough recharge to maintain that discharge. Certainly the spring
discharge depends on the amount of rainfall over the catchment that percolates
through the soil zone to reach the groundwater system. To estimate the amount of
recharge as a percentage of the total rainfall (recharge coefficient, REC), the surface
catchment area has to be defined. It has been shown that it is unnecessary for the
216
catchment area of the spring to coincide with the surface catchment area. Generally
springs receive recharge water from beyond the limit of the surface catchment from
the regional recharge mound area. Although springs with low discharge rate, or those
which issue in the high rainfall zones, might receive recharge water locally from the
surface catchment area of the spring. Therefore, the catchment area of the spring
should be defined as the effective surface catchment which contributes to the spring
recharge area. In a steady state condition, a spring discharge originally generated from
the amount of rainfall over the effective surface catchment area of the spring:
Q = Px RECx A ........................................................................... (6.21)
where Q is the annual spring discharge (m3)
P is the annual rainfall in (m)
REC is the recharge coefficient (fraction)
A is the surface area of the catchment in (m)
The independent variables in the above equation Q and P can be measured
easily with certain accuracy. The REC and A are unknown, but the product of REC
and A can be related to the known variables by rearranging Eq. (6.21):
Q C = - = RECx A ......................................................................... (6.22)
P
C A=
REC or
C REC = A ...................................... (6.23)
The coefficient C is a characteristic factor for the spring which relates its annual
discharge with the annual rainfall as a function of recharge coefficient and the
catchment area. Springs with high C values must have high recharge coefficients or
large surface catchment area. Thus for estimating any of REC or A, one of them
must be known. A reasonable range for the recharge coefficient might be nil and 37%
217
(Figure 6.5). Consequently the possible catchment area can be estimated from Eq.
(6.23) by assuming a recharge coefficient in this range.
The relations between the recharge coefficient and the possible catchment area
for a spring, or a group of springs which are believed to discharge from the same
catchment are illustrated in Figure( 6.17). The relations are distinctive for each spring
or group of spring leading for division of springs in the study area into three types.
The first type represents the local springs with small catchment areas, that do not
exceed the measured surface catchment area even for the lowest possible recharge
coefficient. Most of the low rate discharge springs in the Central Plateau belong to this
type. They are characterised by high response to rainfall, with sudden decline in
discharge rate during the dry season: presumably these systems have small storage.
The second type, referred to here as the intermediate springs, are those springs
in which the catchment area ranges between the local surface catchment of the spring
and the regional recharge mound. Ras el Ain spring, for example, lies in a local
catchment area of 68.2 km2 in the Western Highlands close to the Amman regional
recharge mound. For assumed possible recharge coefficients between 5-15%, the
required catchment area necessary to feed the spring is found to range between 48-145
km2• As this spring lies in high rainfall area, it is expected to receive enough recharge
to maintain its high discharge rate locally. But for recharge coefficient of 8% in the
Upper Zerqa basin as recommended from the recharge estimations in the previous
sections, the spring needs a catchment area of about 90.2 km2 assuming that the
discharge rate and the rainfall measurements are of the right order. This means that
76% of the water discharge from the Ras el Ain spring originates from the local
catchment area of the spring and 24% are derived from the regional groundwater
system. This is consistent with the results obtained from the recession hydro graph
analysis. The discharge rate shows a rapid response to the rainfall over the catchment
which indicates the effect of local recharge, while the discharge rate during the dry
season when the recharge ceases does not change dramatically, and the high storage'
of the spring indicates the relation of this spring partly to the regional groundwater
system.
218
35
30 25
~ 20 i:r 15 w ~
10
5
0 0
35 30 25
~ 20 i:r 15 w ~
10 5 0
0
35 30
25 ~ 20 i:r 15 ::1
10
5 0
0
35
30 25
~ ~ 20 C)
15 w ~
10 5
0 0
35 30 25
~ 20 i:r 15 t ::1 10
5
0 0
35 30 25
~ 20 t 0 15 w 0:
10
5 0
0
Rasel Ain
500 1000 1500 Area
Zerqa
2000 4000 6000 8000 10000 Area
Upper Zerqa Basin
2000 4000 6000 8000 10000 12000 Araa
Wadi Haldan
3000 6000 9000 12000 15000
Araa
50
100
W. MuJib (west)
100 150 200 250 300
Araa
Lajun
200 300 400
Araa
Figure(S.17) Recharge versus spring catchment areas (in Km2
).
35 30 25
~ 20 i:r 15 w ~
10 5 0
35
30 25
~ 20 i:r 15 w ~
10 5 0
35
30 25
~ 20 U 15 w ~
10
5
t
0
0
o
Aln Sarah
.... 1000 2000 3000 4000 5000
Area
Hasa(west)
50 100 150 200 250 300 350 400 450 500 Araa
Hasa(east)
200 400 600 800 1000 1200
Area
35~ __________________________ 1
30 25
iIt 20 id 15 0: 10
Ja fr(northw a st)
5
o~~~~~~~--~--~~ o 100 200 300 400 500 600 700
Area
Ja trIa a at)
100 200 300 400 500 600 700
Are.
35 30 Jafr(southw est) 25 _
~ 20 0 15 I\.. w 0:
10 5 0
0 200 400 600 800
Araa
The third type is the regional springs which are located in the discharge area of
the groundwater system at distance from recharge mounds, but as these springs
discharge the regional groundwater system, they are characterised by high discharge
rates, high storage and smaller variation in the discharge rate between wet and dry
seasons, even though located mostly in a low rainfall zone. Zerqa, Wadi Haidan, and
Ain Sarah springs located in the lower part of Upper Zerqa, Wadi Wala, and Wadi
Karak basins, respectively, belong to this type. The local catchment areas of these
springs are 156, 350, and 15 km2 for Zerqa, Wadi Haidan, and Ain Sarah springs
respectively. The high discharge of these springs needs a much larger catchment area
than the local surface catchment areas provide, even when considering high recharge
coefficients. For a reasonable recharge coefficient of 8%, the catchment areas required
to maintain these spring's discharges are 567, 898, and 254 km2 for Zerqa,
Wadi Haidan, and Sarah springs respectively. By extending the catchment areas to
include the other catchments between the springs and the regional recharge mound,
catchment areas of 551, 840, and 305 km2 for Zerqa, Wadi Haidan, and Ain Sarah
springs, respectively, are obtained, suggesting that this approach may be reasonable.
The area of 305 km2 for Ain Sarah spring is higher than required, but this may be
explained by the high recharge coefficient, since 8% is probably too high for the steep
topography of the Ain Sarah spring area.
It is worth mentioning that the spring discharges used are long term mean
monthly values, reflecting the original condition rather than the present. The present
day spring discharges have been reduced dramatically due to the heavy abstraction
from the aquifer systems.
Although the effective catchment areas were found to range between the
surface catchment area of the spring and the regional catchment area, which might
extends to include the regional recharge mound area, the exact area has to be
delineated in order to estimate the recharge coefficient. There are many techniques
which could provide an estimation of the catchment area such as groundwater tracers,
220
groundwater chemistry, groundwater modelling, water level fluctuation, or recession
hydro graph analysis. All these methods depend on the groundwater velocity and the
time required for the recharge water to reach the discharge points.
It has been shown earlier, that the effective recharge to the system occurs after
a reasonable period since the rainfall period elapsed, which agreed with the first
response of the discharge hydro graph. This period is a function of the groundwater
velocity and the area of the catchment: the distance between the recharge and
discharge points. Under normal conditions, the recharge water infiltrates down the soil
profile to the groundwater table, then moves within the groundwater system toward
the discharge point. The reception of the recharge front at the discharge point can be
detected from the water level fluctuation in an observation borehole or from the
analysis of duration discharge curve of the spring. The rate of propagation of the
recharge front is a function of the distance from the recharge mound to the spring or
the borehole and the time difference between the recharge event and its expression in
spring or borehole hydro graphs.
It is concluded that recharge coefficients can be estimated from the spring
hydro graph analysis provided the springs record is of sufficient length and adequacy.
This method is characterised by: (i) it is simple in that it does not include parameters
which are difficult to quantify such as the actual evapotranspiration and the soil
moisture content; (ii) it does not employ the different hydrological parameters to
proposed recharge to the groundwater system and it estimates the recharge amount
when recharge occurs; and (iii) the recession hydro graph analyses differentiate
between the local recharge amount to the spring storage and the recharge to the
regional groundwater system, local recharge amounts usually show at the peak flow of
the spring hydro graph, which then can be estimated.
Although the limited amount of spring discharge data and lack of time
restricted the application of such methods during this study, this could be the subject
of future research.
221
6.7 RECHARGE MOUNDS
Most of the recharge enters the aquifer systems in the structurally high outcrop
area, in the high rainfall zone of the Western Highlands. Groundwater flows from the
recharge areas to replenish the aquifer systems. Part of the flow is diverted to the
intervening river valleys.
Several extensive recharge mounds have been developed along the Western
Highlands (Figure 6.18). Groundwater hydraulic gradients are solely controlled by
the recharge mounds which are assumed to be developed from the modem recharge.
However, well hydro graphs in the confined areas close to the outcrop do not always
show evidence of seasonal water level fluctuations as would be expected from the
recharge. Burdon (1977) has postulated a variety of mechanisms other than modem
recharge to account for the groundwater gradients found in the arid basins. Some of
the suggestions, though possible, are unlikely to be significant. However, the
hypothesis that the existing gradients can be attributed to the creation of recharge
mounds in the Pluvial Pleistocene periods and subsequent long-term head decay under
distant groundwater discharge might explain the evolution of groundwater flow
mechanisms in the study area.
AMMAN MOUND
This mound is located in the high rainfall zone west of Amman. The water
flows from this mound westward (giving rise to the springs of Wadi Sir), north
eastwards down the Amman-Zerqa syncline (to discharge to upper Wadi Zerqa
Valley), eastwards into Azraq Basin area, or southwards to contribute to the baseflow
of Wadi Haidan.
RABBAMOUND
This small mound is developed at Rabba. It forms a groundwater divide, from
which the water flows eastwards to Wadi Mujib and westwards to feed the springs
along the foothills between Wadi Mujib and Wadi Karak.
222
100
1000
't7
'" CD
o
IU .... ::I )( IU U.
"' c: CD ~
=? "' ";:::
Y <t
... .... ...
N
W-\rE s
o 50km ;
LEGEND
~ basin boundary
r'\S outcrop of 82/A7 aquifer system
-_ .. fault
...... flexure
0 recharge mound
=C> flow direction .. water transfer from the basalt aquifer system
" ........... wadi
• city; town
I 900L-__ ~ ________ ~~~ __ ~ ________ ~~~ ________________ __
300 200
Figure (6; 18) Distribution of recharge to the B2/A7 aquifer system;
MAZARMOUND
This mound is developed in the mountainous area between Karak: and the
Wadi Hasa. The western side of the Mazar mound marks the unsaturated limit of the
B2/ A 7 aquifer system. Water discharges southwards to the Wadi Hasa, north
eastwards to the Wadi Mujib, and northwards to feed Ain Sarah in Wadi Karak.
TAFILA MOUND
The crest of this mound is a few kilometres south of Tafila on the mountains
between the Wadi Hasa and the Dana fault. Groundwater flows northwards to
discharge in the Wadi Hasa and eastwards to the Wadi Hasa Basin. The Dana fault is a
hydraulic barrier which prevents the groundwater flow south-eastwards to the Jafr
Basin.
SHAUBAK-RAS EN NAQB MOUND
This elongated recharge mound is developed on the continuous range of the
Western Highlands between Shaubak and Ras en Naqb. Part of the recharge flows
westwards and gives rises to springs along the escarpment and the rest flows
eastwards to the Jafr Basin.
6.8 RECHARGE TO THE RIJAM (B4) AQUIFER SYSTEM
The Rijam aquifer system is located in an area where the mean annual rainfall
IS less than 50mm and the evaporation is high. It is therefore considered that
significant direct recharge to the aquifer is unlikely. It is possible for indirect recharge
to take place from transmission losses from floods in wadis which cross the Rijam
Formation and drain towards the Jafr playa. Such floods have very low frequency,
occurring once or twice a year and originating mainly from precipitation in the
highlands on the western side of the Jafr Basin. Indirect recharge calculations (Table
6.6) indicate about 6 MCMla as mean annual indirect recharge for the whole Jafr
Basin. However, occasional flash floods also occur in wadis draining other part of the
basin. Within the outcrop area of the aquifer, runoff calculations using the curve
224
number method during the period 1980-1985 show that only during the season
1980/1981 there was enough water to produce flood flow.
Throughput calculations using flownet analysis carried out by different authors
indicate ranges of recharge between 3-10.7 MCMla (Abujamieh, 1967; Parker, 1970;
and AHG, 1977). Environmental isotope analyses carried out by Abujamieh (1967)
showed low tritium values and carbon-14 apparent ages of25,000 years. The analyses
of Howard Humphreys (1986) did not detect any tritiated water but carbon-14 dating
showed the groundwaters to be less than 500 years old. It appears that the Rijam
waters are a mixture of old and new water.
The long term monitoring of water levels does not show any measurable
change in water level due to recharge for the period 1962-1967, before and in the early
stages of groundwater development when the abstraction was less than 1 MCMla,
which suggests that if recharge occurs, it will be at considerable distance from the
observation well, and that recharge pulses have levelled out to give minimal
fluctuations before reaching the observation wells. While the water levels fluctuations
after development shows that for average abstractions of approximately 1.5 MCMla
water levels have shown a consistent decline at rates between 0.1-0.36 mia, indicating
that abstraction exceeds replenishment, despite the fact that approximately 25% of the
total abstraction returns back to the aquifer as irrigation return flows.
It is believed that the Rijam aquifer system in the Jafr Basin receIves
intermittent indirect recharge with an average of 1 MCM/a (Howard Humphreys,
1986). The higher estimates reported by some different authors demonstrate that using
flownet analysis can be misleading. This method assumes that the rate of flow
represents the rate of recharge. This can only be correct where a source of continuous
replenishment is available.
However, recharge to Rijam aquifer occurs sporadically in the form of pulses
of runoff infiltration into the aquifer which helps to maintain a hydraulic gradient that
has been established for a long time. Furthermore, the method is sensitive to
transmissivity and hydraulic gradients, which are not accurately known for the entire
aquifer.
225
6.9 RECHARGE TO THE HUMMAR (A4) AQUIFER SYSTEM
The Hummar Fonnation outcrops in a narrow zone on the northwestern flank:
of the Amman-Zerqa syncline (Figure 4.3). Although much of the outcrop lies in a
fairly high rainfall zone, the area which recharges the aquifer is small, only about 20
km2•
Direct recharge has been calculated using the soil-moisture balance method for
the period 1980-1985. The results indicate annual direct recharge to the aquifer range
between 0.11-2.9 MCM with mean annual value of about 0.83 MCM. The throughput
of the aquifer at the outlet of the area north ofZerqa is estimated at 5 MCM/a.
It is believed that the indirect recharge is the most important constituent of the
recharge. Figure (6.19) shows the density of the drainage pattern on the outcrop, and
this could provide the conditions for indirect recharge. Parker (1970) suggested that
water may collect in numerous small pools and then enter the aquifer through joints
and solution channels. Calculations indicate that the mean annual indirect recharge for
the A4 aquifer system in the Amman-Zerqa area is about 2.3 MeM.
Natural vertical leakage from the overlying or underlying strata is not possible,
as the water level of the Hummar aquifer is higher than the others in the area.
However, in the Zerqa over flow area, west of Zerqa, the Hummar Fonnation is
affected by folding and faulting, which result on outcropping the Fonnation, in some
areas it just underlying the alluvial of the Zerqa River (Figure 6.20). In other areas the
B2/ A 7 and A4 aquifer system meet each other. These in turn allow for indirect
recharge to the Hummar aquifer system either by direct water transfers from the
B2/ A 7 aquifer system, or by infiltration from the Zerqa River via the wadi fill of sand
and gravel: this will be discussed in detail in section on Aquifer Interrelationships
(Chapter 7).
6.10 RECHARGE TO THE LOWER AJLUN GROUP (A1-6)
The thick limestone of the upper part of Na'ur Fonnation (A1I2) interbedded
with the thick sequences of marls provide aquifer potential in the northern part of the
study area. However; the limestone beds are not continuous, and lithology changes to
226
230 2315 240 2415
o 150Kme 170 ,=1======::::::11
les
Rlnr 160
no 235 ~40 15
After Parker (1970)
Figure (6.19) outcrop of Hummar Formation showing local drainage pattern.
900
800 ~ ~
700 ~
... !! 600 k. ~.,&s GI E .5 500 I-GI "0
~ 400 r-< 1511 -
300 t- u 200
Figure (6.20) Geological cross-section in Hummar Formation to the NW of Zerqa.
marly facies and decreases in thickness are well observed southward from the
Wadi Mujib. In the south east the whole group (Al-6) is replaced by the sandy facies
of the Fassu'a Formation, the latter considered as a potential aquifer system in the Jafr
area.
In the northern part of the study area, where the limestone beds are thick and
believed to form a potential aquifer system (A 112), the flow dynamics are
insufficiently well known to establish the flow pattern. Due to low permeability and
poor recharge, the aquifer does not seem to have economical value. However, a small
number of springs and wells provide water for local domestic supply. In general the
limestone of Na'ur Formation is exposed in small areas, of steep topography, in low
rainfall zones. Thus natural direct recharge is expected to be very small, and most of it
will be rejected as spring flows. There may be some vertical leakage of water via fault
conduits from the overlying aquifers, but it considered that the total recharge to the
aquifer is small.
In the southeastern part of the study area, the recharge mechanism for the Al-6
aquifer must be very similar to that of the B2/ A 7 aquifer in the area, although the
small extent of outcrop would indicate little total recharge. The potential for recharge
is limited to the Western Highlands where the aquifer outcrops: to the east the Lower
Ajlun occurs under a great thickness of sediments and therefore cannot receive
recharge from rainfall. Thus recharge must derived from lateral flow, which, given the
resistance to and retardation of flow by the relative impermeability of the group in the
west, is unlikely to reach the aquifer in the east. However downward leakage of
groundwater from the overlying B2/ A 7 aquifer may be a source of recharge.
6.11 CONCLUSION
Jordan is semi-arid to arid climate has high potential evaporation rates which
can exceed mean annual rainfall by more than one order of magnitude. Thus, if the
rainfall rate were constant, then there would be little or no recharge. However
groundwater reserves exist because the rainfall distribution is far from uniform. Net
228
recharge occurs during prolonged periods of rainfall higher than the average
intensities.
All natural recharge originates as rainfall, but the routes by which water enters
the aquifer system vary considerably within the study area. Recharge occurs by direct
infiltration of rainfall in outcrops areas, indirect recharge through the transmission
losses of the flood flow via the wadi beds, vertical leakage through the underlying and
the overlying strata, water transfer from adjacent aquifer systems, or by lateral
boundary flow from outside the study area.
The unpredictability of rainfall events, the large variety in soil, the geology,
the topography, the landuse, and the combined difficulties in estimating
evapotranspiration, in Jordan presents great difficulties for accurate estimation of
recharge. A number of methods have been used, and the results of each method were
then evaluated according to their merits. The empirical methods used were found to be
corroborated by the results obtained from the analysis of the groundwater response to
recharge (such as fluctuation in water levels and spring discharge rates).
Calculations show that direct recharge is only dominant in the Western
Highlands. The recharge appears to occur to a series of mounds with groundwater
flow moving partly toward the intervening river valleys and partly toward the east.
Recharge generally decreases from the northwest to the southeast.
In the Western Highlands and Central Plateau, long term, average estimates of
recharge for the present landuse conditions indicated about 100 MCMla, most of
which was probably discharged as spring flows. The direct recharge and the lateral
boundary flow occupy the major part of the total recharge in the Western Highlands.
Direct recharge calculations suggest that 8% of the total rainfall percolates downward
to recharge the aquifer systems. While in the eastern and southern parts, indirect
recharge and lateral boundary flow constitute the majority of the total recharge. The
main precipitation occurs in the higher part of the wadi catchments with the major
recharge to the groundwater occurring through wadi bed transmission losses during
flood runoff.
229
Generally direct recharge does not occur when annual precipitation is less than
200-250 mm. However, due to the presence of penneable materials which have low
field capacities in the wadi alluvial fans, localised direct recharge can occur after large
intensive stonns even though the annual rainfall is less than 200 mm.
The estimates of recharge do not attempt to distinguish the recharge to the
water table and the recharge to the deep aquifers. The estimates of recharge presented
herein include water that locally recharges groundwater only to be discharged nearby.
Recharge estimates were found to be most sensitive to the amount of total
precipitation and its temporal distribution, and to soil type; uncultivated limestone and
chert soils allow for deeper percolation.
230
CHAPTER SEVEN
GROUNDWATER FLOW
7.1 GENERAL
It has been shown that the Amman-Wadi Sir (B2/A7) aquifer system is the
most extensive and continuous aquifer system in the area. It is hydraulically
interconnected with the Hummar Aquifer in Amman-Zerqa area. The Hummar
Aquifer wedges out in the south. The marls of the lower Ajlun Group (Al-6), which
form an aquiclude system in the north and west become more permeable to the south
and south-east, where they become hydraulically connected with the overlying B2/ A 7
aquifer system. However, the low permeability and the regional extension of the Al-6
aquitard is the main reason why the B2/ A 7 is delineated as regional aquifer system.
The Group separates between the regional B2/ A 7 and deep sandstone aquifer systems.
Figure (7.1) shows a generalised hydrological section through the study area.
An understanding of groundwater flow in the principal aquifer system IS
needed to evaluate the potential of the resource for use. Recharge, discharge, and flow
are important factors relating to water-yielding capability, spatial and temporal
variations of the water quality, and the response of the system to development.
The major source of groundwater in the aquifer system in the study area
includes recharge in and near the outcrops in the north, north-western and western
edge of the Highlands. Most of the recharge enters the aquifer in the structurally high
outcrop areas (recharge mound), then flows downward toward the lower areas of
discharge. Groundwater in the western Highlands is generally under unconfined
conditions, whilst in the east the aquifer become confined by the thick marls of the B3
Formation.
The total recharge to the Amman-Wadi Sir system within the study area is
estimated to be 94.3 MCMla, of which approximately 36.5 MCMla may be accounted
for as spring discharge. This leaves some 57.8 MCMla which discharges as sub
surface flow. A proportion of the sub-surface flow is transferred by various conduits
to deeper aquifers. Some of the water which discharges at the springs may be
intercepted along its flow path between the recharge and discharge points.
-It) '" w
-It) ~
, , , I I
VI CIl e> ro
.s::: u VI '6 Ol c:: o§. VI
\ ~\ ~. -;,\ ~.
~\ ~\ i:I\. ~\
"'
Jordan River Valley and the Dead Sea
\ , \
1/1 E IV .... 1/1 >. 1/1 Lo
~ ::s C" C'CI
C'CI r:::: o Cl IV Lo
IV .r:::: .... -o . .9l ;;:: o Lo C.
C'CI o .-Cl o o IV Cl o Lo
'0 >.
.r:::: '0 IV .~
C'CI Lo IV r:::: IV Cl -~ r--: -
Under normal conditions, recharge to the water table occurs mainly through
short-term events during heavy winter rains, whereas groundwater discharges occurs
almost continuously as steady downward leakage to underlying rock units, lateral flow
through boundaries out of the study area, and discharge in outcrop areas as spring flow
and baseflow.
Long-term water level records from observation wells in unconfined aquifers
show seasonal trends that correlate with general patterns of groundwater recharge and
discharge. In the long dry period of summer and autumn there is no recharge.
However, discharge from the aquifer system continues causing the water table to
decline. Recharge occurs in winter when heavy rains occur. This period of recharge
generally replenishes groundwater storage causing the water table to rise. Extreme
conditions of drought or greatly above normal precipitation cause departure from this
generalised pattern.
7.2 FLOW MECHANISMS
The flow dynamic in the B2/ A 7 aquifer system is very complex. It is the net
result of interaction of various factors, which are mainly related to the
hydrogeological framework of the system. The flow pattern is found to be strongly
influenced by the following features:
1. The recharge mounds.
2. The regional dip of aquifer strata.
3. The geological structure in the area (faults & flexures).
4. The hydraulic characteristics of the aquifer systems.
5. The wadis existing in the area.
6. The Jordan Rift Valley, being the final base level for all flows.
2. The existence of two major aquifer systems; the B2/A7 and the sandstone
aquifer system, separated in the vertical direction by the low permeability
A1-6 aquitard.
The regional groundwater movement is dominated principally by the recharge
mounds along the Western Highlands and the regional dip of strata toward the east
and north-east. High hydraulic heads along the Western Highlands suggest the major
recharge areas for the regional aquifer system, from which, the groundwater flowed
233
east and north-eastwards with the regional slope of the base of the aquifer beds. This
original pattern has been changed slightly as a result of the tectonic movements and
the subsequent creation of new drainage system and deep wadis which cut down to the
saturated zone, and hence induced groundwater flow westwards by spring discharge
and baseflow. Therefore, a regional groundwater divide separating the eastern and
western flows, and a new hydrodynamic equilibrium were developed, which resulted
the present pattern of groundwater flow.
The above mentioned features suggest a rather intricate flow system has
developed. Within the highland block, groundwater flow is directed to the east until it
enters the zone of influence of the draining wadis: here steep hydraulic gradients
develop, and hence the flow turns towards the centre of the wadis. The groundwater
appears as baseflow in the wadis. However, the wadis probably do not fully intercept
the regional north-easterly flow. The degree of interception depends, among other
factors, on the depth of the wadi, and how far it penetrates through the saturated zone
ofthe aquifer system.
7.2.1 FLOW DISTRIBUTION
The B2/ A 7 is an extremely heterogeneous aquifer unit that transmits water
most readily through the joints and fractures that commonly constitute a considerable
percentage of an individual bed. These beds are separated by the less transmissive
marls and marly limestone, in which the fractures are more or less vertical. Lateral
groundwater movement in the marl and marly limestone interbeds is probably
negligible when compared with the volume of water that moves laterally through the
limestone beds. This is because lateral movement of water in the limestone beds is
controlled by fracture and joint systems, whereas lateral movement of water in the
interbed zones is controlled by primary features. Vertical movement of groundwater
between the limestone beds is much less per unit area than lateral movement but is
large over the entire aquifer area. Vertical movement of groundwater varies because of
the structure of the individual bed and the hydraulic characteristics of the interbeds.
Except for B 1, most interbeds within the aquifer are of very limited extent. This gives
rise to a further complication: head changes occur within the vertical section of the
234
system and changes of water level of up to two metres have been recorded while
drilling (parker, 1970 and Howard Humphreys, 1986).
Because the rocks that make up the B2/ A 7 aquifer system vary greatly in
penneability, the system resembles a group of layers composed of alternating zones of
low and high permeability. Vertical flow between permeable zones probably occurs
through sink holes and fractures. However, the amount of vertical flow is probably
small compared with the amount of horizontal flow. The zones of high permeability
generally are at or near unconformities and are generally parallel to bedding planes.
Thus the apparent gradient, as indicated by the change in water level elevation
between two wells, may reflect a composite of the regional gradient and head changes
in the sections penetrated by the wells. Such effects weigh heavily in calculations
involving low gradients and high transmissivities.
7.2.2 GROUNDWATER STRATIFICATION
The previous paragraphs describing the vertical distribution of flow give rise
to the hypothesis of groundwater layering or stratification through the B2/ A 7 aquifer
system. The groundwater in the B2/ A 7 aquifer system can be divided into three
general zones: (1) the upper shallow high permeable zone, which contains fresh
groundwater, (2) the middle zone with intermediate permeability, which contains
mixed fresh and old groundwater, (3) the lower zone with lower permeability, which
contains old higher salinity groundwater. This hypothesis is supported by indirect
evidence from hydrochemical and environmental isotope analysis.
The quality of water in the aquifer system near the outcrop reflects recharge
conditions. Water in these areas of the aquifer system contains small concentrations of
dissolved solids and chloride and is commonly a calcium carbonate type. The calcium
and bicarbonate ions probably are derived chiefly by dissolution of calcium carbonate
in the outcrop area by carbon-dioxide-charged meteoric water that constitutes the
recharge. The areal distribution of the total dissolved solids indicates a general trend
of increasing the salinity eastwards and north-eastwards from the outcrop area with
the direction of regional groundwater flow. The salinity substantially increases as flow
conditions change from unconfined to confined. Water salinity also increases by
increasing well yield, the higher pumping rate presumably drawing water from a
235
greater depth and distance, the depth probably being more significant. Furthermore,
chemical analysis of groundwater discharges from the springs along the Western
Highlands indicates low salinity values (Salameh and Khdier, 1984). This can be
explained by the fact that the springs rapidly drain large areas of the karstic outcrop,
which suggests that groundwater contributions to the springs are chiefly from the
upper zone. The implication is then that stratification of water occurs through the
B2/A7 groundwater system. For detailed water quality analysis, one can refer to the
. previous studies such as Parker (1970) Agrar und Hydrotechnik (1977), Howard
Humphreys Ltd (1986), and GBR (1987).
Further indications of groundwater stratification can also be revealed by
environmental isotope analysis. Recharge occurs in the Western Highlands, thus water
with high tritium values is expected in the recharge mound areas. The tritium levels in
the precipitation are around 100 TV, however, interestingly, Lloyd (1980) found the
samples from the B21 A 7 aquifer system in the Western Highlands to be non-tritiated,
with the highest tritium concentrations located in the eastern part of the area and along
the wadi courses away from the recharge mound. Lloyd however concludes the
importance of indirect recharge as transmission losses from floods into the aquifer
system in the eastern region. Given the fact that the tritium samples are pumped
samples taken from wells penetrating into the deeper zones of the aquifer, integrated
samples are obtained resulting in reduced tritium levels. The non-tritiated water in the
recharge mound areas, then implies that the large volumes of direct recharge of high
tritium levels can only move through the upper permeable section of the aquifer over
the top of underlying older waters. The hypothesis is supported by other evidence.
Groundwater to the east of the recharge mounds and in the valleys is found at
shallower depths than that in the recharge mound areas; therefore, tritium samples
taken from wells in the eastern parts probably represent the upper zone of the system.
Tritium levels decline with increasing well yield, for example, tritium reducing from 7
to 0 TV with the yield increasing from 80 to 120 m3/h. These are consistent with a
stratified groundwater condition.
236
7.2.3 CONCEPTUAL FLOW MODEL
The previous discussion reflects a concept of a multiple flow system of
different areal and subsurface extent, applying to regional groundwater flow in the
Central Plateau of Jordan. The flow systems, which are driven by the hydraulic head
at the water table (Figure 7.1), range from local to intermediate to regional in scope.
These conceptual flow systems are illustrated in Figure (7.2) as generalised
(predevelopment) flow paths along an east-west hydrogeological section in the
Central Plateau.
The local and intermediate groundwater flow system are closely related to the
water table and local topography and surface drainage features. Regional flow is
controlled mainly by elevation differences between major regional topographic
features and by the regional framework of the aquifers and the confining units.
Relatively shallow, local flow systems, mostly less than few kilometres in
length, dominate the hill slopes of the recharge mound areas and account for a
substantial part of the overall volume of groundwater recharge and discharge in the
study area, much of this flow does not pass through the groundwater bodies, it is
dominant in the uppermost part of the aquifer and discharge as spring flow and
seepage along the main wadis and cliffs. The intermediate flow system dominates the
unconfined part of the aquifer system, it is considered to be the main mechanism of
the lateral flow of groundwater to lower areas of discharge. However, this flow is
mainly deeper, slower, and occurs over much larger distances than the local flow. The
groundwater transfers from the recharge mound into the confined part of the aquifer
system by the intermediate and regional flow systems. The latter is the deepest and
slowest flow. Part of the intermediate and regional flow systems flow downward
through the Al-6 aquitard into the deep sandstone aquifer system. Thus, the
groundwater flow in the aquifer is partly controlled by the vertical leakage coefficient
of the underlying Al-6 aquitard.
Therefore, it must be considered that, beneath the main flow in the B2/ A 7
aquifer system from the recharge mounds in the west to the east, a second flow
component exists in the deep sandstone from east to west, appearing partly in springs
at the rim of the rift or flowing undetected into the Dead Sea. The sandstone aquifer
system contributes a considerable amount to the baseflow in the main wadis. The
237
'o a. III ::> ;0 ,," ~
~ CD '< III ::> 0.
~ ro o ro III 0. en ro III
East West Recharge
11111111
local flow
intermediate flow
regional flow
B2/A7 aquifer
~ =A1<; ",<Om = _ ~
Kurnub-Disi aquifer
--:;: ~ : ~ -
Figure (7.2) Conceptual model of groundwater flow in the regional aquifer systems
groundwater body in the sandstone, which has relatively high hydraulic conductivity,
must necessarily be replenished by downward leakage from the B2/ A 7 through the
intermediate Al-6 aquitard. Indeed, going eastwards in the carbonate aquifer and
westwards in the deep sandstone the groundwater age increases, which can be
explained by the longer distance from regions which precipitation occurs and the very
low flowing rate of groundwater.
7.2.4 GEOLOGICAL STRUCTURES AND GROUNDWATER MOVEMENT
The large fault and lineament system that have developed in the many bedrock
units of the study area during geological time are important features in the analysis of
the hydrogeologic system. Both faults and lineaments appear to provide paths for
increased water movement, both horizontally through the aquifer and vertically
through the confining beds. These factors also may act as barriers to flow normal to
the direction of the fault or lineament.
The present relief in the area is almost entirely the result of the tectonic
movements which took place during late Tertiary to Recent times. It was started by
the formation of the Jordan Rift Valley and the Dead Sea Graben, which caused the
gentle eastern to north-eastern dip of strata. followed by a sequence of movements
which resulted in faulting and tilting of the block to the east of the graben. Its effect is
more severely along the Western Highlands parallel to the Jordan Rift Valley, whilst
in the eastern part, its effect is of minor importance.
The late Tertiary pre-faulting relief was a sub-horizontal peneplain, drained by wide
shallow wadis, a relief which is still present in the east and south-east where the B3
Formation is exposed.
The difference in elevation between the rift valley and the highlands enhanced
a backwards incising erosion, thus creating the main east-west trending wadis such as
the Zerqa, Mujib, and Hasa wadis, which are cut deeply into the otherwise almost
undisturbed peneplain of the highland.
FaUlting and tilting of blocks changed the pattern, with new drainage channels
following the new structural pattern such as the straight north-south course of the
Wadi Nukheila, Wadi Yubbs, and Wadi Sultani.
239
The structural history of Jordan is also reflected in the sediments, and geologic
structure is one of the important factors that controls porosity and permeability in
carbonate and sedimentary rocks. Movement along structural zones creates porosity
and increases permeability by fracturing; porosity and permeability may be modified
at a later time by chemical processes that occur in the aquifer as water moves through
the fractures. Thus, the faulting system triggered the processes responsible for the
considerable contrast between the hydraulic characteristics of the B2/ A 7 and the other
lithological units. The faulting increased hydraulic gradients in the fault zone, which
enhanced the percolation of meteoric water from the land surface and increased the
velocity of shallow groundwater flow. A dynamic regime of shallow groundwater
flow evolved that promoted dissolution. Dissolution along fractures and bedding
planes formed joints and solution channels that became principal conduits of regional
groundwater flow in the B2/ A 7 aquifer system.
Movement along major faults and lineaments may affect the porosity and
permeability of rock over large area and through a long span of geologic time.
Structural adjustments between large blocks of geologic materials may have modified
the primary porosity. Structural adjustments also may result in a decrease in porosity
and permeability and hence modify the flow system so that materials in the water
precipitate in the rock pores.
Faults which displace the strata vertically so that impermeable rock is
juxtaposed against permeable strata, will impeding groundwater flow in directions
normal to the faults.
From the up gradient parts of the outcropping recharge area, groundwater
generally flows downdip in north-easterly direction. The barrier faults typically block
the north-eastward flow of groundwater and divert it southwards and northwards,
along flow paths aligned with the fault zone. In some places, a secondary network of
transverse faults obstruct the major south-east trending flow paths, imposing internal
boundaries that further divert or compartmentalise the flow system. As a result, local
patterns of groundwater flow can be extremely complex, making predictions about
future response to prolonged drought or additional pumping difficult to determine.
The following sections will discuss the effects of barrier faults in the groundwater
movements in the study area.
240
Many of the structural features in the study area are associated with the
present-day physiographic features that affect the deep and shallow groundwater and
surface water flow systems. Structural movement has a major effect on deposition of
clastic sediments such as those in the lower part of Upper Amman-Zerqa Basin which
forms a shallow aquifer system with hydraulic continuity with the underlying B2/ A 7
aquifer system. The presence of these alluvial deposits on the top of the B2/ A 7 aquifer
system plays a major role in modifying the permeability within the B2/A7 aquifer
system.
7.3 REGIONAL GROUNDWATER FLOW
The preceding sections pertain to the effects of recharge and discharge to or
from the main aquifer systems, as well as the groundwater flow mechanisms. The
most important controls on hydraulic head in the B2/ A 7 aquifer system are the slope
in the base of the aquifer system, topographic relief, and location of springs and
streams.
The mam features of the regional flow system are reflected in the pre
development potentiometric surface (Figure 7.3). The surface portrays a general
easterly and north-easterly component of lateral flow through the aquifer system. The
potentiometric surface map is based on the earliest recorded heads in the aquifer,
depict the approximate steady-state conditions of head and flow prior to the beginning
of large scale groundwater abstraction. However, water levels in the Amman-Zerqa
area might reflect the effects of minor groundwater development. Head changes
caused by pumping are discussed in section (7.5).
The potentiometric surface of the confined part of the aquifer system depicts a
broad, regional pattern of uniform hydraulic gradient. The character of the
potentiometric surface provides another insight for the conceptual model of flow.
7.3.1 UPPER ZERQA BASIN
Two main aquifer systems have been delineated in the Upper Zerqa Basin; the
unconfined B2/ A 7 and the confined A4 aquifer systems, separated by the confining
A5/6 Formation. The two aquifer systems are found to be hydraulically
241
10
so
40
20
000
.... '"
I I
.. ,.oJaf, I rrr·-7IO- ---" ,
I
" ,
I ePHTt ePNTS
: ePHTtt
I ,
s
o 10 20"'"
W """" f
LEGEND
o city, town
"L wadi, river
weU
fault
ftexure
fault, inferred from groundwater modeling
",/ equipotentiallina 850 m
_ 0 !low Una number 5
,/ saturation IIm~
dry area
\
\ .....
\ '20L---~2~OO~-------2~20--------~2~4~O------~2±60~------~2~80~------~3~OO~------~3~20~------~~~o----J
Figure (7.3) The potentiometric surface map of the B2/A7 aquifer system
in the study area.
interconnected. An upper shallow water table zone was found in the alluvium deposits
along the course of the Zerqa River; it is hydraulically interconnected with the
underlying B2/ A 7 system.
The regional groundwater flow movement in the B2/ A 7 is strongly influenced
by the recharge/discharge areas, topography, and the geological structure in the area.
Recharge occurs in the south-western part ofthe area, at the Amman recharge mound.
Part of the water flows westwards and gives rise to the springs of Wadi Sir. Of the
remainder, some of the water flows north-eastwards down the Amman-Zerqa syncline
to recharge the upper aquifer system in the area, and the rest flows into the desert
regions to the east, north-east, and south-east of Amman. However, the Qihati fault,
which has a maximum displacement of about 300 m, places the B2/ A 7 against the
impermeable Muwaqqar Formation. This structure is believed to form a groundwater
barrier which separates water discharging to the Upper Zerqa basin from water
flowing to the Azraq basin. The main natural discharge area for the two main aquifers
is at the Zerqa River outlet near Sukhnah, where the mean annual groundwater
discharge under natural conditions was estimated to be about 12.7 MeM.
The flow patterns of the aquifer are highly controlled by the geological
structure and the topography of the area, in particular, the Amman-Zerqa flexure and
the Wadi Seil-Zerqa River (Figure 7.4). In the south-west and south-east of the
Amman Zerqa Syncline, a watershed is formed by an anticlinal structure following the
Amman Zerqa flexure. The dip of the formation strata in these areas is 20-40%
forming a clear boundary to the water bearing formation, and the potential water
bearing formations are uplifted above the saturation zone, eroded in some places, or
being drained by deep wadis.
Along the syncline, from its recharge to its discharge, the groundwater flow
quite naturally follows the main direction of the Wadi Seil-Zerqa River (Figure 7.4).
In the south-western part, in the Amman area, the water flows towards the north-east
until Awajan. Thereafter, where the influence of the Wadi Zerqa as a drainage system
increases, the groundwater flow direction turns north with the direction of the wadi
until it reaches the Wadi Dhuleil Valley in the far north ofthe area, where the flow out
243
180 I N
\ W-\>-E 170 s
160
.../" basin boundary
-600- equipotential line 600 m
150 "'fTTTTTTT n Flexure
__ --<: wadi
o 10km I
2~u 250 260 J
140 I 240 270 280
Figure (7.4) The potentiometric surface map of the Amman-Wadi Sir aquifer system in Amman-Zerqa area.
I f
of the Amman Zerqa groundwater basin joint the flow from the east to be directed
westwards with the main direction ofthe Zerqa River.
The regional hydraulic gradients in the aquifer system were steeper in Amman area
(2%) than in the Zerqa Area (1.5%). In the Ruseifa area, where the B2/A7 system is
highly affected by the overlying alluvium deposits, the hydraulic gradient is only
about 0.5 %.
The irregularities in water levels and hydraulic gradients lead to the conclusion
that the aquifer system consists of number of sub-basins with limited connection to
each other through the overlying alluvium deposits. From the hydrogeological section
along the Zerqa River (Figure 7.5), three main sub-basins, ordered stepwise along the
river, can be delineated: the Amman sub-basin upstream of Ruseifa, Ruseifa sub-basin
upstream of Zerqa, and the Zerqa sub-basin upstream of Sukhna. The groundwater to
each sub-basin would then be recharged by infiltration from the Zerqa River, lateral
movement of groundwater between the sub-basins through the alluvium deposits, or
locally by infiltration of rainfall. This hypothesis is supported by the hydrogeological
section, the variation of spring discharge with water level of the Zerqa River, the
different in water level fluctuation patterns between the basins, and by the
groundwater flow model. However, this again indicates that the groundwater level
essentially follows the ground surface and the groundwater flow through the Amman
Zerqa basin is rather small.
The natural pattern of groundwater movement, the hydraulic gradient, and the
rate of recharge/discharge have been significantly altered by water development. The
long-term water level variations show a 5-20 m water level decline due to abstraction.
HUMMAR AQUIFER SYSTEM (A4)
The piezometric surface map of the A4 aquifer system is shown in Figure
(7.6). The Amman -Zerqa flexure which extends north-east of Amman-Ruseifa-Zerqa,
is believed to form a hydraulic barrier east of which the A4 aquifer becomes less
saturated.
The direction of groundwater flow within the basin is from the outcropping
recharge areas towards the south-east until the syncline is reached, then the flow turns
north-east parallel to the syncline, gradually turning north downstream of Zerqa.
245
'Q)
(jj E .5 Q) "0 :::l ~ ~
900 I c: III
c:
a: J!!
:::l a c:
Q)
"Qj
:g « Cl
!II
Q)
III
:::l
« "Qj u ~
0::
III Q)
!II "iii en
III c.. 0::
800
_ top of A7 on right hillside
" 700
500
400 l- - - - piezometric level in B2/A7
-r--- piezometric level in A4
300
modified from VBB (1977)"
Figure (7.5) Hydrogeological profile along the Zerqa River
Q) Cl "0 "C CD III e-Q) N
~ a 'E "Qj Q)
:; > .c 0
III a III c: 'i5 e- .c x ~
Q) :::l N en
- top of A7 on left
hillside
-topofA70n right hillside
boUom of A7 on left
hillside
_ boltom of A7 on right hillside
180,r----,----------,----------,----------,----------,-------____ ~--~
N
\ W-</-E 170 s
160
~ basin boundary
-600- equipotential line 600 m
Flexure 150 .....-nTTTT T ...
~~.--::, wadi
o 10km I
140 I , '"
230 240 250 260 270 280
Figure (7.6) The poten'tiometric surface map of the Hummar aquifer system in Amman-Zerqa area.
The annual variations in the piezometric level seem to be very small, which is
natural considering the long distance between the recharge areas and the existing
wells. Substantial decline in the piezometric level has taken place during the early
stages of development, 50-100 m in Amman and around 50 m in Zerqa. This can only
be the result of considerable overdraft.
7.3.2 WADI W ALA BASIN
The B2/ A 7 aquifer system is unconfined and is underlain by the impermeable
A5/6 Formation. Only in isolated areas especially in the eastern part is the aquifer
confined by the impermeable Muwaqqar Formation (B3). However the degree of
confinement is small with the piezometric level a few metres above the base of the B3.
The major source of groundwater in the B2/A7 in the Wadi Wala basin
includes recharge in and near the outcrop areas in the north-western and western edge
of the Western Highlands.
The regional pattern of groundwater movement is dominated principally by the
recharge mound, geological structure, and the Wadi Haidan drainage system. Most of
the groundwater flows southwards from the recharge mound to Jiza then south
westwards along the main drainage system of the Wadi Haidan, where the major
outflows from the aquifer occur as spring discharge in the lower reaches of the Wadi.
There is a direct hydraulic connection between the baseflow in the Wadi Haidan and
the B2/ A 7 aquifer in the area· between elevations of 250 and 450 mas!. In the lower
part of the basin, between Wadi Haidan and Wadi Mujib, where both wadis cut deep
into the saturated zone, the groundwater splits between the two drainage systems. A
local groundwater divide is developed between the two flows. This can be deduced
from the steep hydraulic gradients and the configuration of the 400 m equipotential
line (Figure 7.3). However a small part of the southwards flow from the recharge
mound flows to the east into the Azraq groundwater basin.
The geology of the area is very complicated and hence the groundwater flow
system is very complex. It seems it is highly structurally controlled. Apart from the
Siwaqa fault line, which is considered a barrier fault separating the Wadi Wala and
Wadi Muijb groundwater systems, very little is known about the structure in the basin.
Attention was given when flow modelling to attempt to understand situations not
248
explained by the hydrogeological framework. The model inferred the extension of the
Qihati fault system into the area and the occurrence of a fault system striking NW-SE
crossing the basin through the Qastal area (Figure 7.3); this fault has not been
observed before. Both fault systems are thought to be barriers which impede the flow
from the recharge mound toward the east, and hence increase the magnitude of the
south-westwards flow component. This can be verified by the number of dry
boreholes which have been drilled in the north-eastern part of the basin, and by the
high spring discharges in Wadi Haidan.
A small potentiometric high is present in the Khan Zabeeb area to the north of
the Siwaqa fault line, where the north-easterly flow is thought to be impeded by the
barrier Siwaqa fault. Apparently, the potentiometric high is maintained as a result of
the structural high at this location. A groundwater divide has developed in this area,
separating the eastern flow into the Azraq groundwater basin from the western flow
into Wadi Haidan-Wadi Mujib drainage system. However, the Siwaqa fault line
probably does not completely block the north-eastern flow, so part of the flow
probably crosses the fault line into the Khan Zabeeb area.
Hydraulic gradients are generally steep close to the recharge mounds and tend
to become much milder downward then steep again in the discharge areas. This
pattern may reflects an interaction of several factors, including flow divergence,
discharge from springs, and permeability changes.
7.3.3 WADI MUJIB BASIN
The B2/ A 7 is the most important aquifer system in the Wadi Mujib. It is under
unconfined conditions in the western part, while in the east it is confined in some
localities by the overlying B3. The impervious marls and shales of the AS/6 form the
lower confining unit. Although, the A4 and A2 may have a groundwater potential, it is
very limited, so the whole Lower Ajlun Group (Al-6) is considered as one unit, acting
as an aquitard separating the B2/ A 7 from the deep sandstone aquifer system. The
latter is not considered in detailed in this study.
As indicated by the groundwater flow model (Chapter 8), the permeability of
the Al-6 aquitard is rather higher than thought before and hence it is considered to be
leaky. The steady st~te simulation suggests that in order to reconcile the composite
249
relationships between the recharge, transmissivity distribution, and discharge, it was
essential to increase the vertical permeability of the Al-6 and consequently the
quantity of downward leakage from the B2/ A 7 into the deep sandstone aquifer system.
The lateral flow pattern in the B2/ A 7 (Figure 7.3) is strongly influenced by the
Mazar recharge mound and the existing fault systems in the area. The Tertiary-Recent
regional tectonic events developed a group of north-south normal faults which control
the aquifer system geometry. Typically numerous horst and graben structures are
present.
The groundwater flows radially away from the Mazar recharge mound, mostly
in a north-east direction. The flow is intercepted by erosional channels which lead to
the rift valley, and is reduced by downward leakage into the lower sandstone aquifer
system.
A highly permeable drainage line provided by the Karak-Wadi el Fiha fault
intersects the north-easterly flows from the recharge mound and diverts the flow
south-westwards and north-eastwards. The latter flow gives rise to the Ain Sarah
spring in the Wadi Karak with a measured discharge of about 5 MCMla. The
influence of this drainage system on the groundwater flow can be deducted from the
equipotential contour map (Figure 7.3) and verified by the unusually high discharge of
Ain Sarah spring. As has been discussed before, the Ain Sarah Spring needs a much
larger catchment area than the surface catchment area provides. This leads to the
conclusion that a substantial ~ount of the groundwater discharge at Ain Sarah spring
must be diverted from the Mazar recharge mound by the drainage system of Karak
Wadi el Fiha fault line. This interference reduces the north-eastwards groundwater
flow from the recharge mound.
A similar situation is presented by the Wadi Yubbs and Sultani-Qatrana
grabens, where the highly permeable drainage system diverts the flow into the
downward deep canyon which leads to the Wadi Mujib. The 700 m equipotential line
indicates the strong receding tendency of the hydraulic head in the wadi area (Figure
7.3). However, within the Wadi Mujib Basin, apart from the eastward flow, the Wadi
Mujib acts as the base level for the B2/A7 aquifer system.
None of the fault systems described as causative factors of creating drainage
system maintain their draining characteristics along their entire extensions. In some
250
areas these faults act as barriers, impeding the water flow or at least reduce the
saturated thickness ofthe aquifer significantly. This pattern can be easily noticed from
the water head variation along these fault systems.
Further evidence ofthe nature ofthese structures also can be obtained from the
variations of the well yield-drawdown relationships for wells drilled along and around
the fault lines. Well LA19 drilled in the north-western part of the Karak-Wadi e1 Fiha
fault line, in aregion where the fault was thought to be acting as a drainage system,
shows a yield of 41 m3/h with 10.4 m drawdown, while to the south-east along the
fault line, in the Wadi Batra, wells PP49, LAI6, LA 17, and LA18 have very low
yields and high drawdown. Well LA15 drilled to the east of the fault line was found to
be dry. The effect of the fault in reducing the saturated thickness is demonstrated in
Figure (7.7). This phenomenon extends further to the south-east till Wadi Gheith
where the yield of well S78 is 12m3/h with a drawdown of 18.7 m. Further to the
south-east, in the Wadi Abiad, the Karak-Wadi el Fiha fault line regains its
characteristics as a drainage system. Most of the wells drilled in Wadi Abiad are
highly productive. A similar pattern of specific capacity variations were found along
Wadi Yubbs and Wadi Sultana fault systems. Wells LA9, LA13, and LA14 drilled
along the northern extension of the fault line which extends from Wadi Yubbs to
Wadi Nukheila were found to have very low yields. And also the same for wells
LAlO, LAl1, and LA12 which were drilled along the northern extension of the Wadi
Sultana fault line.
Unsaturated areas of the aquifer system are also present along anticlinal and
horst structures caused by the adjacent fault systems in many localities in Wadi Mujib.
The most extensive were found in Jebel Rueifa, Jebel Mutaramil, Jebel Saqrat, and
two areas which were found in Wadi Patra between wells LA18 and LA15 (Figure
7.3).
In the north-eastern part of the Wadi Mujib, the Wadi Tuwal fault line reduces
the eastern flow significantly, well AP18 drilled in the western side of the fault line
was found to have a yield of 18 m3/h and drawdown of 29 m, while well AP 17 in the
eastern side yields only 8m3/h for a 32 m drawdown. However, the natural water flow
251
Q) c ::;
w 3 z CO u. CO cr CO ~ rn
AP15
c CI)
.D ~
(!) c :::J
"iij' -J
iIi
co .r:. u:: Qj
'6C1) COc ~"-,-J .)t{.-
~:; COCO ~u.
;: (/)
0 0 0 0 N to
"0
"' "OJ .c "~ S
"' .e >-.c
0 0 ~
'iii is , .D ::J c: .... ::J ~
c: CI) ro
OJ .0 ........ ::J
::E '0 ro ;:: c: Q.)
;;: 0 .... e. ro 0
Cl 0 0 Q.)
Cl 0 .... '0 >-:t: -"-: I"--Q.) .... ::J Cl U.
is inferred to continue to the east and north-east, with low hydraulic gradients of about
1 % compared with 5% in the western area.
The Siwaqa fault line running west-east at the northern boundary of the Wadi
Muijb watersheds is a barrier which blocks the flow passage into Wadi Wala.
Hydraulic head differences of up to 50 m were observed around the fault line.
The preceding descriptions of the groundwater flow in the Wadi Mujib Basin,
concludes that the regional potentiometric contours indicate that under typical,
isotropic conditions, most of the groundwater should flow eastwards. However many
of the faults are barrier faults, which impede or block the eastern flow of groundwater,
and given the presence of the Wadi Karak and Wadi Mujib drainage systems, most of
the water is diverted north-eastwards. The fracture network, as well as the associated
joint cavities and solution channels that are subparallel to the barrier faults, impart an
anisotropic pattern of hydraulic conductivity and a dominant west-east component of
transmissivity. Although the west-east gradients are comparatively small, the
transmissivity tensors aligned with the fault zone are great enough to move large
amount of groundwater from the recharge areas to the north-east.
7.3.4 WADI HASA BASIN
The groundwater flow pattern of the B2/ A 7 aquifer system in the Hasa Basin
is mainly affected by the Tafila recharge mound, the Wadi Hasa drainage system, and
the Salwan fault system (Figure 7.3). The piezometric surface elevation in the Western
Highland, as high as 1200 masl, decreases north, north-east, and eastwards to less than
750 masl in Wadi Hasa and in the eastern part of the basin. The hydraulic gradient is
very steep in the recharge mound areas in the Western Highlands, getting milder as
the flow continues eastwards.
The original eastward flow from the recharge mound is bounded by the Wadi
Hasa drainage system in the north and the Salwan fault line in the south. The Wadi
Hasa deep canyon cuts deep into the saturated section of the aquifer system causing
divergence of the flow north and north-eastwards from the recharge mound into the
wadi, where part of the groundwater flow component appears as spring discharge and
baseflow along the Wadi Hasa. The Salwan fault system is an extensive groundwater
253
barrier separating the Rasa groundwater basin from the Jafr Basin. To the north of the
fault line, the flow is generally eastwards parallel to the fault line.
The Karak-wadi el Fiha fault line intersects the basin in the eastern part.
Although the vertical displacement along the fault line is about half the thickness of
the aquifer system, given the lithological nature of the B21 A 7, it acts locally in some
areas, particularly in the north-eastern part, as a barrier. This view is proposed by the
configuration of the 750 m equipotential line, and has been verified by the steady state
calibration of the groundwater flow model.
Locally, the groundwater flow model suggests that some lineaments are
groundwater barriers, such as the north-west trending fault line which extends from
the south of Jurf Darawish north-westwards to the south of Tafila, and the north
south fault line at the eastern boundary of Qa Jinz (Figure 7.3).
The volcanic eruptions in the south-western part of the basin, along the
Salwan fault line, cause a slight, very local, disturbance for groundwater flow.
7.3.5 JAFR BASIN
7.3.5.1 INTRODUCTION
Five aquifers have been recognised in the argillaceous, arenaceous, andlor the
carbonate rocks of the Cambrian to Paleogene age, such as the Disi, Kurnub, Lower
Ajlun (AI-6), Amman-Wadi Sir (B2/A7) and the Rijam (B4). The deep sandstone
aquifer systems of the Kurnub and Disi groups are not considered in detailed in this
study. The B2/A7 is the main aquifer system in the area, and is hydraulically
interconnected with the underlying AI-6 system especially in the eastern part. The B4
aquifer system with limited potentiality has been under development in the centre of
the Jafr Basin since 1964.
7.3.5.2 AMMAN-WADI SIR AQUIFER SYSTEM (B2/A7)
In the southern part of the Jafr Basin, the B2/A7 and AI-6 are both thin and
unsaturated. The B21 A 7 aquifer is unconfined in the south, where the water table
occurs in the A7. Further north, the water table occurs in the B2. In the central and
northern areas, the :821 A 7 is confined by the overlying thick impervious argillaceous
254
unit of the Muwaqqar Fonnation. The aquifer occurs at shallow depth (70-100 m) in
the southern areas, at a greater depth (100-250) in the central areas, and up to 800 m in
the Jafr trough in the north-west.
The B2/ A 7 sequence can be considered as a single hydraulic system with
hydraulic continuity between the B2 and the A7. However, recent drilling in the area,
at borehole number (JTl) shows that locally, the marl of the Bl Fonnation can
develop as a confining layer causing a head difference between the two fonnations of
up to 35 m (JICA, 1990).
The extent and the hydrology of the B2/ A 7 sequence are markedly controlled
by the geological structures in the area, although, it is considered regionally
continuous. The regional direction of groundwater flow is from west to east, from the
Shaubak-Ras Naqb recharge mounds in the Western Highlands through the Arja
Uweina flexure into the Ma'an-Shidiya-Jafr areas and continues to the east and north
east (Figure 7.3). The piezometric groundwater surface elevation ranges between 1500
mas I in the Western Highlands to 1100 masl in the west of Arja-Uweina flexure. The
water levels drop to 900 masl immediately east of the flexure and to 800 masl at
Ma'an (Figure 7.3). This drop in piezometric elevation indicates that the Arja-Uweina
flexure acts as a hydraulic barrier to flow. Groundwater flow modelling however,
suggests that the flexure is not entirely impenneable but allows some groundwater to
pass across, particularly in the northern and southern ends.
Salwan fault intersects the northern part of the Jafr Basin from west to east. It
is an extensive but not a continuous structure which is frequently cut by a series of
north-south trending discrete fault systems. The maximum vertical displacement is
estimated to be about 200 m. Recent drilling in the area confinned the perfonnance of
the Salwan fault as a hydraulic barrier, impeding the groundwater flow across into the
Upper Hasa Basin. JICA (1990) reported groundwater levels in borehole JT2, drilled
in the north-western part of the Jafr Basin to the south of the fault line, of 794.5 masl,
and 790.5 masl for borehole J03 in the north-eastern part to the north of the fault line.
This gives a difference in water levels around the fault line of more than 150 m.
In the central areas, piezometric elevations range between 770-800 masl, but to
the east towards the Karak-Wadi el Fiha fault zone they drop gradually from 770 to
255
740 masi. The piezometry in the eastern areas suggests that the Karak-Wadi el Fiha
fault line does not impede the eastwards flow component.
SHIDIYA AREA
The piezometric elevations in Shidiya area indicate that the hydraulic
conditions are very uniform and the groundwater movement is very slow, since the
piezometric elevations are very similar over a relatively large area, particularly in the
southern part, where the piezometry remains unchanged (on average 789 masl) for a
distance of 15 Ian northwards of the southern limit of saturation (Figure 7.3). In the
northern part towards the Salwan fault line, the water levels are slightly higher but the
hydraulic gradient is still very small. This situation is difficult to explain; four, not
necessarily mutually exclusive, explanation are considered: the permeability is
unusually very high, the aquifer thickness is large, the aquifer receives quantities of
recharge, or the aquifer is structurally controlled.
At first sight, the equipotential map of the central Jafr area suggests that the
permeability is very low in the Western Highlands, very high in the central Jafr areas,
and intermediate in the east. This correlates with the distribution of the hydraulic
gradients which are very steep in the Western Highlands (1-20%), very gentle in the
central areas (0.03%), and increasing eastwards to 0.1% (Figure 7.11).
There is no plausible reason for larger than expected permeability. The lateral
changes in lithology from· ~arbonate dominant in the west to sandy facies in the
central and eastern areas, exc~ude any possibility of permeability increase due to
bedding plane, solution channels, or karstification. One would expect decrease in
permeability due to the depth of burial of the aquifer system in the central Jafr area.
The permeability values obtained from pumping test analysis (on average 1-8 mid) are
not unusually high.
Furthermore, groundwater flow modelling failed to reproduce the water level
distribution, by using plausibly high values of permeability ( up to 7 times the
measured values), without introducing some sort of impermeable structure,
intersecting the central Jafr area from west to east, or by increasing the saturated
thickness of the aquifer system, which is not proved by drilling. This could be
256
w
1500
1000
500
Arja·Uweina Flexure
Figure(7.8) Hydrogeological profile in Jafr Basin
- - .s.. _ hydraulic head of B2/A 7 aquifer
..... t.... hydraulic head of A 1-6 aquifer Karak·Wadi el Fiha Fault line
"U J: -i IC
"U :r -i 01
"U J: -i ~
m
7///////A///////////j n' ~ ~,. 1111////////////"////////////
J 7 7)/~;111111~~/11 //////////////////////// I 1111111 1III11 ~~~//////////////////~// ~/IIIIII ., I
Kurnub-Dis.
"U J: -i ~
N
E
possible by including part of the underlying Al-6 aquifer system, but the groundwater
level data (Figure 7.8) indicate a difference in water levels between the two aquifers
and the water bearing formations among the Al-6 are believed to occur in A4 and A2,
in the middle and lower part of the Al-6. In addition to that, the effects of saturated
thickness sounds a possible reason rather than to explain the situation in the Jafr areas,
the occurrence of very gentle hydraulic gradients occur in the southern part, close to
the saturated limit of the aquifer system, where the saturated thickness is less.
Analytically, the same results can be obtained by only increasing the permeability of
the Al-6 aquifer system and hence the increase in downward leakage from the B2/A7
into the underlying Al-6 aquifer system. This can explain the drop in water levels in
the central Jafr areas rather than the uniformity of the groundwater heads or at least
the configuration of 790 m equipotential line.
At this stage it is possible to conclude that the eastwards increase in
permeability in the B2/A7 and Al-6 aquifer systems gives rise to the drop in water
levels. So the general picture is therefore that the steep hydraulic gradients in the west
governed by the low permeability, the Atja-Uweina flexure, and the movement of
groundwater with the dip of the aquifer beds, giving way to an area of very gentle
gradients governed by the increase and uniformity of the B2/A7 and Al-6
permeabilities and the horizontal movement of groundwater parallel to the aquifer
beds. Further east, the intermediate permeability and groundwater flow against the dip
of the aquifer beds result in slightly increased head gradients (Figure 7.8).
It has been discussed earlier that the central Jafr area lies in a very low rainfall
zone, and as the B2/ A 7 aquifer system is confined, any source of local recharge
should be very limited. Groundwater replenishment depends mainly in groundwater
lateral flow from the Western Highlands. However, the water level distribution map
indicates a local southward hydraulic gradient in the northern part of the central Jafr
area (see the 790 m equipotential line, Figure 7.3). This area is located to the south of
the eastern end of the Salwan fault line and coincides with the saturation limits of the
B4 aquifer system. This leads to the possibility that there is source of replenishment in
that area, in the form of water transfer from the overlying B4 aquifer system or by
lateral flow from the Wadi Hasa groundwater basin. Downward leakage from the B4
258
aquifer system via the thick marls of the B3, though not proved, is possible, as
discussed in the following sections. Groundwater leakage from the storage of the thick
B3 marl will be very slow, and has been taking place for a long time. However, small
leakage would be expected to be uniform all over the area. The only option left is that
a proportion of the easterly flow in the Upper Hasa basin crosses the eastern end of
the Salwan fault line into the central Jafr area. The continuity of the 790 m
equipotential line from the northern part of the central Jafr area into the Hasa Basin,
and the hydrochemistry support this view. The electrical conductivity for the Jafr
groundwater increases from about 700 f.lS/cm in the Western Highlands to about 1500
JlS/cm in the western part of the central Jafr areas, then drops to about 1300 f.lS/cm in
the central Jafr area: the electrical conductivity in the Upper Hasa Basin is about 1100
f.lS/cm.
The groundwater replenishment in the northern part of the central Jafr area
from any source are bound to be very small and without economical value. However,
this must be considered if the groundwater flow in the central Jafr area are to be
understood.
7.3.5.3 LOWER AJLUN GROUP AQUIFER SYSTEM (Al-6)
The equipotential map and limit of saturation of the Lower Ajlun aquifer are
shown in Figure (7.9). The regional groundwater flows are confined by the three
major fault systems; the Arja-Uweina flexures, Salwan fault and Karak-Wadi el Fiha
fault. These faults act as impervious barriers, since displacement exceeds the total
thickness of the water bearing zones within the Al-6 group.
In the Western Highlands, the groundwater flows from south-west to north
east across the Arja-Uweina flexure. In the area north of the Salwan fault and west of
the Karak-Wadi el Fiha fault, the groundwater flows from north-west to south-east.
The easterly direction of groundwater movement in the Lower Ajlun aquifer
system would suggest that groundwater is discharged into the underlying Kurnub
sandstone aquifer, since the water table elevation in the AI-6 is higher than that in the
Kurnub.
259
"" ..... ....
~ .... .... .... ... ..........
• Tafila
~ ....
o ... I 20km - -"" ' I -'" .... ... .... --- ..... --. -,
,1 ...
1000 u~ak
, , \
..... ,\ .... , \
...... , " '\ ,,',
\ I ... ,
~ A1-6 Outcrop
fault
r,.,,-n1 flexure
\ " , , \ , '\ -850- equipotential line 850 m
, '" A~<:) , \ ~
A"~ '\<::s ~
Sq/wq <o~ - "\ .: ~qlJlt . , ,~ (,n. ~,-f.
,~" ...., ~ \,\l" ......... ' ,~
. \ "l.... .... , ~ \ \, ~
\ \ ~~. \~\~ \\~.
\ ~ ;(\ • Jafr .~
\~ \\~
\\(,\l \ \ \
Limit of saturation A 1-6 \\
" '\ \ \ , \
\ \ \\ \\
\
900~,-----~~---------------~------~~,~,~J 200
, "
300
Figure (7.9) The potentiometric surface map of the A1-6 aquifer system in Jafr Basin.
7.3.5.4 RIJAM AQUIFER SYSTEM (B4)
The pre-development equipotential map of the Rijam aquifer system is shown
in Figure (7.10). The groundwater flows from the north-west to the south-east
following direction of the main wadi courses, where aquifer is thought to receive
recharge by infiltration of flood runoff. The hydraulic gradient is relatively flat in the
eastern part of the saturated zone (0.1 %), while it becomes steeper upstream in the
west. The flattening of the gradient occurs where the aquifer thickness, and therefore
transmissivity, is greatest.
Groundwater flows eastwards into an area where the Rijam aquifer wedges
out, and as there is no evidence of discharging groundwater, such as springs, seepages
or marshlands, it is thought that the groundwater discharges to the underlying
Muwaqqar Formation (B3). It should be noted that the amount of this downward
leakage is very small, since the total annual recharge to the B4 aquifer system does
not exceed one million cubic metres.
The Rijam aquifer has been exploited mainly for irrigation since 1964. Prior to
1967, abstractions were just over 1 MCMla and at present abstraction is
approximately 2 MCM/a. Long-term water level data (Table 7.1) indicate the water
level to be declining at rates between 0.1-0.36 mlyr.
Well No. Period (year) Decline (m) Decline rate (mly) 112 1966-1983 2.80 0.l65 117 1962-1982 3.90 0.195 J23 1967-1983 1.30 0.080 J26 1968-1981 1.70 0.l30 110 1966-1983 6.l0 0.360 After NRA (J 985)
Table (7.1) Long term groundwater level fluctuations ofthe Rijam (B4) aquifer system.
7.4 HYDRAULIC GRADIENTS
Hydraulic gradients generally are steeper close to the recharge area and tend
to become milder downflow. This pattern may reflect an interaction of several factors,
including flow divergence, discharge from springs and permeability changes. The
major outflows from the aquifer system are flowing springs and discharge to pumping
wells. Most ofthe large springs that discharge regional flow within the study area are
261
980
960
~
I I \ , I , I \
I
, ... )
o co co
·PP29
.........
~, . CQ \ ,
I I
I I I I \
.• PP23
.PP28
\ • S54 \ /g.S52 , // co
'--
260
'0c6> ,
280
N
W-<rE s
o 5km I
• well
-850- equipotential line 850 m
limit of saturation 84
Figure (7.10) The potentio~etric surface map of the Rijam aquifer system in Jafr Basin.
located in the lower reaches of the wadis. Spring discharge and water level fluctuation
are highly affected by the rainfall in the study area, as well as by borehole abstraction.
The maximum hydraulic gradient (dh I dl) along a flow path indicates the
direction of groundwater movement, whereas the velocity is dependent on the
hydraulic gradient and the hydraulic properties of an aquifer specifically, hydraulic
conductivity and porosity. The average effective linear velocity of groundwater ( v) in
a homogeneous and isotropic porous medium is given by Darcy's law, which is
expressed by the following equation:
-K(dhl dl) v = ~ ................................................................................ (7.1)
where v = average effective linear velocity (mid)
K = hydraulic conductivity (mid)
dh I dl = hydraulic gradient
~ = effective porosity
The groundwater level map (Figure 7.3) shows the configuration of the
contours representing the prevailing equipotentials, and groundwater movement is
downgradient, from the recharge mound areas to the discharge areas and
approximately normal to the contours if the aquifer system is assumed to be isotropic.
Where hydraulic conductivity, saturated thickness, and effective porosity are
constant, changes in hydraulic gradient indicate relative changes in velocity. To gain
an insight into the rate of groundwater movement eastwards and north-eastwards from
the recharge mounds in the Western Highlands, estimates of velocities and transit
times were calculated for estimated pre-development hydraulic gradients along
selected flow lines (Figure 7.11; see Figure 7.3 for locations of flowlines). The results
are listed in Table (7.2).
The estimates are further based on the assumptions that the pre-development
hydraulic gradients were unaffected by abstraction, the permeability distributed
according to the results obtained in Chapter 4, the porosity is constant through the
study area and equal to 0.25, and that vertical flow is negligible.
263
800 Flow Line No.1
750
700
iii 650 ;;; ca .§. <>
::>'" 600 ...
G> .a 'C '" ~ 550 ~
c: « 500 => <>-
450
400
0 5000 10000 15000 20000 25000 30000 35000 40000
Distance from Am m an recharge mound (m)
800
750 Flow Line No.2
700
iii 650 ;;; <> co ::>'" g
600 .a .. '" ." 3 :l 550 g E =>
oC <>-500
450
400
0 5000 10000 15000 20000 25000 30000 35000 40000 45000
Distance from Am m an rech.rge mound (m)
800
750 Flow Line No.3
700
iii 650 ;;; ca .§. 600 <>
::>'" ... ~ ...... :dI:sdEa_
G> 550 .a 'C '" ~ 500 3
0 Ci: c:
450 a. 400
350
0 10000 20000 30000 40000 50000 60000
Distance from Am m an recharge mound (m)
800
750 Flow Line No.4
700 ::0 ... <T - 650 <T
iii ... ca ;;; .§. 600 ()
::>'" ... G> 550 .a 'C '" ::I 3 ~ 500 0
~ § 450 <>-
400
350
0 5000 10000 15000 20000
Distance from Rabba recharge mound (m)
Figure (7.11) Estimated predevelopment hydraulic gradients along selected flow lines from the recharge mounds to the discharge areas
264
950 Flow Line No.5
850
0; 750 "" :s: ;;; ... '" A: :s: .5. <> ,... VI ::s- .s. 650 ... ;(: ~ 41 .a i'T 5' "C ... .." \.C
~ 3 5' !!! e,
550 0 ... r- VI
c .." ... <> :::> ... c· ::s-o.. !i- :::> ...
VI .a 450 ~ "0 ... ... S· VI
\.C
350
0 5000 10000 15000 20000 25000 30000 35000 40000 45000
Distance from M azar recharge mound 1m)
950
900 Flow Line No.6
850 :s:
(/)
~ -< c ... Cl~ 0"
0; 800 0" (/) .0
VI ~Il EO 'e
'" ;;; @ ... 0
.5. 750 ... ... .0 C <> ... :::> ~ ... :::> ::s- O" 41
... ... ... .." 0..
700 .a :::> ... EO "C ... ~
$ ... !i-
650 3 ;1 ~ 0 « c ... :::>
600 0.. .." ... 550 !i-
r-5'
500 ... 0 10000 20000 30000 40000 50000 60000 70000 80000 90000
Distance from M azar recharge mound 1m)
1200
Flow Line No.7 1100
0; 1000 ;;;
'" <>
.5. ::s-
900 .a CD ...
"C 3 ::I 0
= 800 c :::> « 0..
700
600
0 2000 4000 6000 8000 10000 12000 14000 16000
Distance from Taflla recharge mound (m)
1500
1400 Flow Line No.8
1300
'" 1200 0; c
_.0"
'" 5 ;g .5. 1100 ;;; i ~fr 41 <>
"C 1000 ::s- (/)'" !!( =;. 0
::I ::s-c = \.C EO !: s 900 ... ... « 3 5' 0
0 ... E:
800 c ::!! :::> ... 0.. ~ 700 ;;;
600
0 10000 20000 30000 40000 50000 60000 70000 80000 90000 1 E+05 1 E+05 1 E+05
Distance from Shaubak recharge mound(m)
Fig ure (7.11) Continued.
265
Flow Recharge Discharge area Distance Velocity times line No. mound (km) (mlyear) (year) 1 Amman Zerqa River Overflow 38 104 365 2 Amman subsurface outflow into Azraq Basin. 42 6 7000 3 Amman W.Haidan springs 52 5.3 9811 4 Rabba W.Mujib springs 19 15 1267 5 Mazar W.Mujib springs 42 6.5 6462 6 Mazar subsurface outflow into Azraq Basin 89 3 29667 7 Tafila W. Hasa springs 14.5 64 227 8 Shaubak subsurface outflow to the east of Jafr Basin 120 2.1 57143
Table (7.2) Estimated groundwater velocities and transient times along flow lines from the recharge mounds to discharge areas.
Groundwater moves very slowly through the B2/ A 7 aquifer system in most
areas. Calculations indicate velocities generally in the range of 2.1-104 mla. So
groundwater would require 7000 - 57000 years to move across the study area.
Velocities are least in the deep structural basins. Therefore, much of the regional
groundwater flow systems are virtually stagnant except in terms of geological times.
The presence of saline water in the deep part of the aquifer system indicates
incomplete flushing of connate water from the aquifer system. Even the present
distributions of dissolved solids and chloride appear to be related largely to
environment of deposition with limited modification by postdepositional flow.
Incomplete flushing of formation fluids seems probable under conditions
imposed the depositional and postdepositonal history. The nearshore depositional
environment probably included original formation water of fresh to moderately saline
composition. The fluctuating continental-to-marine conditions would have permitted
some introduction of fresh meteoric water into the sediments during and soon after
their deposition.
Furthermore, the small permeability of the B3 have effectively reduced
flushing of the aquifer system by impeding the infiltration of precipitation into the
underlying B2/ A 7.
The estimates are based on the average regional distribution of hydraulic
conductivity and constant porosity. However, the groundwater moves along fractures
and joints, where the horizontal hydraulic conductivity and vertical leakage are
expected to be high. Therefore, the actual groundwater velocities in the B2/ A 7 aquifer
system might depart radically from this general picture.
266
7.5 WATER LEVEL FLUCTUATIONS
Groundwater monitoring has been carried out by W AJ since the early sixties at
the observation wells shown in Figure (7.12).
The potentiometric surface fluctuates continuously as a result of natural
variations in recharge and discharge, and because of external effects such as
barometric pressure. The most obvious pattern of natural water level change is the
seasonal cycle of fluctuation. The potentiometric surface is higher in winter and
spring than in summer and autumn. Water level fluctuation trends in the study area are
shown in Figure (7.13). Fluctuations caused by groundwater recharge range between 0
and 5 m. The range is high in the Western Highlands where recharge occurs, and
decreases eastwards. Observation well hydro graphs in some areas in the south-eastern
part of the study area, where the aquifer system is confined by the thick marls of the
B3, do not show any significant water level variations. The potentiometric surface
also fluctuates through long-term cycles during which water levels show recovering or
declining trends caused by climatic cycles of several years duration. Water level
records from Jarba well (S65), for example, indicate a water level decline of about 5
m during the period 1967-1969 when the total abstraction from AIja well field does
not exceed 2.5 MeM. It is believed that only small proportion of the head decline can
be attributed to abstraction. However, Parker (1970) reported recharge values in the
region in rainfall years 1963/64-1964/65 as high as nearly seven times above the
average, followed by drought over the period 1965-1968 with very small recharge
amounts. Therefore, it seems probable that the decline in water level at S65 is mainly
in response to changes in annual recharge pattern.
Fluctuation of the potentiometric surface can be caused by pumping of wells.
Short-term, rapid fluctuation of groundwater levels are caused chiefly by pumping,
and the amplitude of the fluctuations decreases with distance from the pumped wells.
Such fluctuations usually have a duration of few hours or days. Seasonal fluctuations
are characterised by several months of rising groundwater levels followed by
declining levels. In recent years, as the abstraction increased, decline in water levels is
observed across almost all the study area. The greatest decline are in the confined part
of the aquifer.
267
150
100
50
000
950
Ruseifa 8 Abdoun 20 • RC 29
I - RC13 A~man
0: .as
ER4 e • Braik 1
• Faliq 1
o Tafila
• . ..IT1 • S121
o 5haubak I JT2
5108. 565 • o Petra I S118
• PPS4 I PP65
o Ma'an
.5100
• S83
• AB3
• Hasa 15
• Hasa 11
• S53
• JT4
eJT3
o Jafr
PH05
• observation well
o city, town
~ river, wadi
N
W-</-E s
o 40 km
900~~~------________ L-____________ ~ ______________ -L ________ ~
200 250 300 350
Figure (7.12) Location map of the observation wells in the study area
'iii '" .c .5. .... ;::
'iii '" .c .5. .... ;::
'iii '" .c .5. .... ;::
iii z E :r ==
20
25
30 __ Aw sjan Observation IJ\I9tt
35
40
45
lf72 1m 1~ 1m 1~1~1~1~1~1~1~1~1~1~1~1~1~1~1~1~ MonthlYear
40 -o-Zerqa Obs.
45 -+ - ReflnaryNo.9
50 __ Hoshemah No.8
55
80
8/87 10187 12187 2188 4188 8188 8/88 10188 12188 2189 4/89 8189 7/89 9/89 11/89 1/90 3/90 5/9D 7/90 9/90 11/90 1/91 MonthlYear
9/85 12185 3186 6186 9/86 1~ 3187 6187 9/87 12187 3188 6/88 9/88 12188 3189 6189 9/89 12/89 3190 6190 9190 12/90 MonthlYear
:~1-: =': ":' . ~-~":3 •
I
9/85 1/86 5/86 9/86 1/87 5/87 8/87 12187 4188 8/88 12/88 4189 8/89 12189 4190 8/90 12/90 MonthlYear
!~j~J 9/85 1/86 5186 9/86 1187 5187 8/87 12187 4/88 8/88 12188 4189 8/89 12/89 4190 8/90 12/90
MonthlYear
160 r---------------------------~--------------------------------------------_, 165
j 170
~ 175
== 180
__ Oastal No. 7
-0-Oastal No. 6
185+---+---+---+---+---+---+---+---+---r---r-~~~--~--~--~--~--~ __ ~ __ ~
10/84 2185 6185 9/85 1/86 5/86 9/86 1/87 5187 9/87 1/88 5188 9/88 1/89 5189 9/89 1/90 5190 9/90 1/91 MonthlYear
4/88 6/88 8/86 10/86 12186 2169 4/89 6/89 8/89 10/89 12189 2190 4190 6/90 8190 10/90 12190 MonthlYear
Figure (7.13) ObselVation well hydrographs in the study area.
269
.. '" .., .s. ....I
~
154 156 158 160 162 164 166 168
__ Braik No.1
-<>- Ohabe No 70
1/88 3/88 5/88 7/88 9/88 11188 1/89 3/89 5/89 7/89 9/89 11/89 1190 3/90 5/90 MonthlYear
7/90 9/90 11/90 1191
_
~ 197 •• •• 198
; 199 ,. 200
__ Um Rasas No.9 .. ~~~f == '" 196 I " • • • • • • == 201 +---~ __ ~ ____ ~ __ +-__ ~ __ ~ __ -+I ___ ~ __ ~I __ -+ __ ~ ____ ~ __ +-__ ~ __ -+ __ ~ __ ~;:~
12/87 2188 4/88 6/88 8/88 10/88 12/88 2189 4/89 6/89 8/89 10/89 12/89 2190 MonthlYear
4/90 6/90 8/90 10/90 12/90
1170 1171 1172 1173 1174 1175 1176 12176 12177 12178 12179 12180 12181 12182 12183 12184 12185 12186 12187 12188 12189 12190 MonthlYear
9185 12185 3186 6186 9186 12186 3187 6187 9187 MonthlYaar
1::j~ -o-SW2 I .. 102
~ :~: -:-F811q NO.4
~ 108 •••• ~"'¥~"'''''''''_'''''_ 110 . • ••• ~ 112 +---+-__ +-__ ~ __ ~ __ ~ __ ~ __ +-__ +-__ +-__ +-__ +-__ +-__ +-__ +, __ -+ __ -+ __ -+ __ -+, __ ~,~~~~~
9185 12185 3186 6186 9186 12186 3187 6187 9187 12187 M~~~hr$::r 9188 12188 3189 6189 9189 12189 3190 6190 9190 12190
34
36
j 38
~ 40
~ 42
44
V .I, /
V __ UdruhSl18
.. A !\---- ~ ,,-~
\. I -.. 6/82 12/82 6183 12/83 6184 12/84 6185 12/85 6/86 12/86 6187 12/87 6188 12/88 6189 12/89 6190 12/90
MonthlYear
85r-------------------------------------------------------,
t ... ·
!:~ 4 CO ~.~~ 12/68 6169 12/69 6170 12170
~ .. ~.: 12171 6172 12172
.. ~. • JarbaS65
. ~.---t- r r
6/73 12173 6174 12174 6175 12175 6/76 12176 6171
MonthlYear
Figure(7.13) Continued.
270
7.6 AQUIFER INTERRELATION
Fracture systems and lineaments transverse the entire area and act either as
conduits or barriers to groundwater flow, depending on their location and orientation.
Vertical leakage from the aquifers is restricted by the low permeability of intervening
marls and shales. However, interaquifer leakage appears to occur through and along
some of the major lineaments and fractures, and it may be a major consideration in
determining the quality of water produced from wells.
Vertical leakage is probably greatest where the aquifer system is highly
affected by tectonic. In most parts of the study area, the small permeability of aquitard
units restricts interchange of water between aquifers. Because of the large area of
contact, however, cumulative leakage is an important element of the regional aquifer
system.
Recognition of the vertical flow component is important to improve
understanding of the water quality distribution, hydraulic head distribution, and
overall budget.
7.6.1 LEAKAGE BETWEEN B2/A7 AND A4 AQUIFER SYSTEMS
The Amman-Wadi Sir (B2/A7) and the Hummar (A4) aquifer systems in the
Amman-Zerqa area are separated by the thick very low permeability A5/6 Formation.
Therefore, significant interrelation between the two aquifers is unlikely unless there
are fractures that convey water between them. Under the pre-pumping conditions the
piezometric level of the Hummar aquifer was usually higher than the water table of
the overlying Amman-Wadi Sir aquifer, so that leakage was most likely to be upwards
direction; this was demonstrated by the steady-state calibration of the model
(Chapter 8). Due to the extensive water extraction from the lower aquifer and the
subsequent decline in the piezometric level, the possible natural leakage has now
reversed, since in some parts the piezometric level is lower than the base of the
overlying A5/6 Formation.
Field observation has not directly revealed any interrelation between the
aquifers, except for direct contacts between them via the alluvial deposits in the north
west of the Zerqa, where the A5/6 is missing.
271
At the Zerqa overflow west of Zerqa the geological formations are abruptly
interrupted thus giving a sharp end to the Zerqa basin. The formations are lifted 120 m
in total over a short distance by folding and several faults. This causes the A4 to reach
the ground surface in some places along the Zerqa River and the two aquifers meet
each other in one location. The A5/6 Formation, acting as an aquiclude between the
two aquifers, forces the water of the upper aquifer to flow into the A4 Formation and
the overlying wadi fill of sand and gravel (Figure 7.5).
The Sukhna spring discharges from the A4; if its flow measurements are
correct, the large amount of water discharging from the A4 cannot be explained by the
annual recharge. The most likely source of water to the spring is the transition of
water from the upper aquifer at the Zerqa overflow directly or via the sand and gravel
in the wadi bottom to the A4. Chemical analysis, however, shows that the water
emerging from the Sukhna spring is not typical water from the A4 aquifer system
(VBB, 1976).
7.6.2 LEAKAGE BETWEEN B2/A7 AND Al-6
In the south and south-east, as the A 1-6 is gradually modified into an aquifer
system, the difference in water level between the Al-6 and the overlying B2/ A7
aquifer gradually decreases suggesting an increase in the degree of interrelation
between the two aquifers south-east wards.
The Lower Ajlun aquifer is multi-layered, and comprises semi-pervious shale
and marly layers separating discrete aquiferous beds which in the west, consist of
limestones, and in the east and south of sandy limestones and sandstones. In the
south-eastern areas it has been possible to delineate an uppermost arenaceous layer
(20-50 m thick) in direct hydraulic continuity with the overlying Wadi Sir aquifer
system. This layer is separated from the main aquifer by clays and silty sands. The
underlying main aquifer is arenaceous but generally impure and therefore poorly
productive.
Piezometric levels in the Lower Ajlun Group aquifer system were invariably
lower than in the Amman-Wadi Sir aquifer system. The difference in the piezometric
heads between the two aquifers vary considerably, from 44 m in the west to about 7 m
272
in the east. However, Parker (1970) reports a 171 m difference in head between the
two aquifer at borehole S 1.
The variable differences in piezometric heads between the' two aquifers
suggest that there is little hydraulic continuity in the west and north-west but
significant hydraulic connection in the south and east. However, the regional
hydrological environment tends to suggest that slow downward movement of
groundwater from the B2/ A 7 into the AI-6 may be taking place, particularly in areas
where clay-shale separating layers are absent and the two aquifers lithologically
merge into each other. This can be proved by the difference in piezometric heads, and
by the fact that the limit of saturation of the Lower Ajlun is located beyond that of the
Amman-Wadi Sir in areas where natural recharge believed not to be occurring;
therefore, the source of recharge is most probably the Amman-Wadi Sir. It is not
possible to quantify the rate of replenishment of the AI-6 by leakage because the lack
of sufficient data; it is thought, however, that it is small, very slow and has been
taking place for long time (tens of thousands years).
7.6.3 LEAKAGE BETWEEN B2/A7 AND K-D AQUIFER SYSTEMS
Recharge by downward leakage from the B2/ A 7 to the sandstone aquifer
system through the overlying AI-6 occurs over most of the study area. The rate and
distribution of this leakage are a function of head gradient between the two adjacent
rock units and the vertical hydraulic conductivity ofthe intervening materials.
Head differences between the regional water table and the potentiometric
surface of sandstone aquifer system indicates a downward hydraulic gradient and thus
a downward potential for flow. Head differences are greatest in the Mazar (~ 1000 m) ,
and Tafila recharge mounds (~ 1100 m), and gradually decrease eastwards to less than
200 m. Assuming a plausible range (for example, 4.17 x 10-7_ 2.08 X 10-
6 m/h) for
vertical hydraulic conductivity of marl (the principal rock type of the AI-6 confining
system), an average head difference between the two aquifer systems of about 500 m,
and an average thickness of 250 m for the AI-6 aquitard, the downward leakage rate
per unit area of the aquifer system is very small, perhaps ranging between 1825-9125
m3/a/km2•
273
CHAPTER 8
GROUNDWATER MODELLING
8.1 INTRODUCTION
Computer model are tools that can be used effectively to help understand the large
and complex groundwater systems within simulated area. Numerical simulation, by a
digital computer, was used in this study to evaluate conceptual models of the hydrologic
system. The regional groundwater flow in the Western Highlands and Central Plateau of
Jordan was investigated by simulating both steady state pre-development conditions and
transient conditions since significant pumping began.
However rarely are computer models used to simulate groundwater flow over
areas as large as 22,350 km2 as was done in this study. Therefore , it must be stressed
that the computer simulation discussed in this report is conceptual in nature. Only broad
concepts and large scale features can be inferred from the results of the model study.
However, the regional groundwater flow model was supplemented by four sub-regional
models, so local flow systems and water problems can be studied in more detailed. In fact
the objective in presenting a detailed analysis of groundwater flow is to examine the
hydraulic inter-connection between the different aquifer systems through the carbonate
rocks, and how the regional geologic features and the lateral changes in lithofacies affect
the direction of flow and the water level.
8.2 MODEL DEVELOPMENT
The code used to simulate the groundwater flow system of the different
subregional study areas and the entire Western Highlands and Central Plateau of Jordan
was the US Geological Survey's modular, three-dimensional, finite-difference
groundwater flow model (McDonald and Harbaugh,1984) with the processing software
PM (Wen-Hasing and Kinzelbach, 1991). The code uses a modular programming
structure comprising a main program and various subroutines to simulate aspects of the
aquifer system.
The general equation for steady-state three-dimensional flow based on Darcy's
law and the principle of continuity (Darcy, 1856; De Wiest, 1965) takes the form of:
where:
~(K ~) + ~(K ~) + ~(K ~) = 0 ......................... (8.1) a x xx a x a y YY a y a z zz a z
x,y,z = Cartesian co-ordinate corresponding to the major axis of hydraulic
conductivity (L).
K = hydraulic conductivity (LIT)
h = hydraulic head ( L)
This equation is one of the most basic partial differential equations. It is called
Laplace's equation. The solution of the equation is a function hex, y, z) that describes the
value of hydraulic head (h) at any point in a three-dimensional flow field. The reader is
referred to Freeze and Cherry (1979) and Rushton and Redshaw (1979) for a detailed
discussion of the derivation of the equation.
In many seepage and groundwater problems, the variation of the groundwater
potential with time (t) is of considerable significance. In considering time-dependent
problems, it is important to consider separately the two alternative mechanisms for
confined and unconfined aquifers.
For a confined aquifer, a fall in the groundwater potential (piezometric head)
results in a reduction in pressure. The volume of water released per unit volume of
aquifer due to a unit decrease in head is termed the specific storage coefficient, 8s. De
Wiest (1965) indicated that both the compressibility of the water and the change in pore
volume due to vertical compression of the aquifer contribute to the specific storage (8s)
and Eq. (8.1) becomes:
a (K a h) a (K a h) a (K a h) = a x xx a x + a y YY a y + a z zz a z s ~ ................ (8.2) Sat
275
For a horizontal confined aquifer of thickness, b, integrating Eq. (8.2) over the aquifer
thickness gives:
ah S - .................................................... (8.3) Cat
where Txx and Tyy are the principal components of the transmissivity tensor defined by
T = Kb, and Sc is the confined storage coefficient defined by Sc = Ssb.
Eq.(8.3) is the equation for transient flow through a saturated anisotropic porous
medium. It must be noted that the above equation holds only for an element within a
saturated aquifer.
With an unconfined aquifer, two mechanisms apply; due to the compressibility of
the aquifer and the water, the specific storage coefficient applies to all elements within
the saturated portion of the aquifer. In addition, the fall of the free water surface (water
table) leads to a dewatering ofthe aquifer. A unit fall in the free surface position results in
a release of water from storage equal to the specific yield (Sy) per unit plane area. The
form of the flow equation for transient flow through an unconfined aquifer, therefore
expressed as:
a h a h s -+ S - .................................... (8.4) Cat Yat
The release of water due to specific yield occurs at the water table, unlike the
effect of the specific storage which is distributed uniformly throughout the saturated
volume of the aquifer. In fact the release due to the specific yield may not be
instantaneous, as indicated by the concept of delayed yield (Boulton, 1963). In practice
the confined storage coefficient is very much smaller than the specific yield; hence the
differential equation for an unconfined aquifer is usually written as:
276
~(T ~) + ~(T ~) = a x xx a x a y YY a y ah s - .................................................... {8.5)
Y at
For a confined aquifer, the saturated thickness remains constant, but for an
unconfined aquifer the saturated thickness is a function of groundwater potential.
However, in many practical situations the saturated thickness and therefore the
transmissivity are assumed to remain constant, either because the change in saturated
thickness is small or alternatively because the manner in which the permeability changes
with depth is unknown.
Usually and in many groundwater situations, water enters the aquifer from above
or below depending on the hydrogeological setting of the aquifer system. Inflow to the
aquifer system (W) arises from recharge which may result from precipitation or perhaps
the presence of a stream, causes an increase in groundwater potential. This is equivalent
to an additional recharge W to the upper surface. The units of W are volumetric flow per
unit volume of sources or sinks of water (liT). In addition, water may flow between
layers and in the computer program, this flux is identified as leakage, L, which is
calculated from the vertical hydraulic conductivities and the head differences between the
layers. Detailed discussions on evaluating L and the average vertical hydraulic
conductivity used, are given in sections (8.7) and (8.8.2.1.2). Finally a fall in water table
(decreasing groundwater potential) results from water released from storage. Assuming
that this flow is immediately distributed throughout the full depth of the aquifer, then the
basic partial differential equation for an anisotropic, heterogeneous porous medium with a
constant water density, as used in the computer program, becomes:
~(T ~) + ~(T ~) - W -L = a x xx a x a y YY a y s~ ................................. {8.6)
a t
The program described herein distinguishes between layers in which storage
coefficient values ,S, remain constant throughout the simulation, and those which the
storage coefficient may convert from confined value to water table value, or vice-versa,
277
as the water level in a cell falls below or rises above the top of the cell. For a confined
layer, the storage coefficient values ,S, are given by the specific storage of the cell
material multiplied by the layer thickness of the cell. For an unconfined layer they are
equal to the specific yield of the material in the cell. The incorporation of layer thickness
into the confined storage term maintains the flexibility of the program to represent layers
of varying thickness, and to implement either the direct three-dimensional or quasi three
dimensional conceptualisations of vertical discretisation. The specific yield is closely
related to the porosity of the rock but not exactly equivalent. The computer program only
uses kinematic porosity, $ , in transport and pathline calculations to calculate the average
linear velocity of groundwater, v = q / $ , where q is the flow rate.
The partial differential equation for groundwater flow can be easily approximated
by finite - difference equations, which are sets of algebraic expressions that are solved
simultaneously by using, in this model, the strongly implicit procedure ( SIP ). The
solution of this algorithm involves setting up a three dimensional grid system in which
each model cell within the grid exhibits specific hydrologic properties that best
approximate the true physical setting of the area, within which any property, hydraulic
head, or flow rate associated with that cell is applied uniformly over the extent of the cell.
For a detailed discussion of the solution technique used in the model, the reader,
however, is referred to McDonald and Harbaugh (1984).
Thus the code allows for spatial variations of aquifer properties, hydraulic heads,
and flow rates, and temporal variations of hydraulic heads and flow rates. For simulations
that do not include changes in head with respect to time ( steady state), the right side of
the equation goes to zero and estimates of storage are not needed. This was the case for
simulations used to conceptualise the groundwater flow in the area.
The model requires values for hydraulic properties, boundary conditions, sources
and sinks, and initial hydraulic head distributions. The following sections which describe
the initial and simulated hydraulic properties, contain discussions on the formulation and
method of calculating these hydraulic properties. Primary results from the model consist
of head distributions and volumetric water budgets.
278
8.3 GENERAL ASSUMPTIONS AND LIMITATIONS
In addition to the assumptions and limitations in groundwater flow modelling,
reported by Rushton and Redshaw (1979), McDonald and Harbaugh (1984), and others,
several simplifying and necessary assumptions are made during the conceptualisation and
simulation of groundwater flow in the area.
Given the nature of the carbonate aquifer system in the study area, the
groundwater not only flows through the porosity of the carbonate rocks but also flows
through fractures and solution openings. It was assumed that the fractures and solution
opening through which water flows could be represented as a porous medium and that
Darcy's Law was applicable from a regional perspective. These assumptions may be
reasonable because the model grid spacing is large.
The model simulations assumed steady-state conditions prior to development in
which estimates of current recharge (1980-85) is assumed equal to estimates of current
discharge (assuming no groundwater abstractions). During the pluvial Pleistocene periods
(10,000-20,000 years before present), climate in the area was significantly wetter than
today, and numerous lakes and rivers were present as indicated by the occurrence of
extensive fluviatile deposits in many places throughout the study area, even in areas
where the present annual rainfall does not exceed 50 mm. It is possible that both
groundwater levels and spring discharge are not in equilibrium, because of the long
distances between areas of recharge and discharge, and both are still declining since the
climate has become drier.
Burdon (1977) discussed evidence of water table declines in the Arabian Basin.
He concluded that the existing gradients can be attributed to the creation of recharge
mounds in the pluvial Pleistocene periods and subsequent long-tenn head decay under
distant gradual discharge. The author in 1984 recorded a calctic veins and hot spring
deposits (travertine) 20-50 m above and up gradient from the present water table of Wadi
Afra hot springs in Wadi Rasa Basin. These deposits are associated with other features
indicative of paleo-groundwater discharge. Furthennore, the well hydro graphs in the
confined areas of the B2/A7 aquifer system close to outcrop do not always show evidence
279
of seasonal water levels fluctuation as would be expected with such recharge. Lloyd
(1980), based on environmental isotope data, demonstrated that mixed recharge and fossil
gradient decay conditions probably exist in the area with much of the old water possibly
having been recharged some 10,000-20,000 years ago. Thus, the evidence suggests that
change in water levels have occurred slowly, and the water table must have declined at
least 50 m over the past 10,000-20,000 years as the Dead Sea shrank to the present level
with high salinity.
Whether or not current groundwater discharge is in equilibrium with the current
recharge is unknown. The time that a change in recharge might be reflected in heads and
in a change in discharge is also unknown. However, the change in water level decline is
slow through geologic time, therefore the assumption of steady state might still be valid
for the length of simulation time. If the steady state assumption is not valid, then the
results of the conceptual model described in this thesis could be affected.
It has been discussed before that recharge occurs in or near the mountain ranges
along the Western Highlands, and groundwater flows downdip with the regional direction
of the hydraulic gradients to discharge at the land surface by springs, or infiltrate
downward to replenish the deep sandstone aquifer system from which eventually it
discharges at the land surface by springs. Therefore, for groundwater simulations, it is
assumed that the areal distribution of discharge is known as well as amount of discharge
from springs and seepages.
The nodes are assumed to be in the centre of the model cells and the permeability
in each model cell was assumed to be homogeneous and isotropic. The model can
simulate heterogeneity caused by differences in hydraulic properties of the rocks by
varying the permeability values of each model cell. It can also simulate heterogeneity
between the vertical and horizontal direction, but the permeability in each model cell is
assumed to be isotropic. The model simulates anisotropy in the horizontal direction
except that the anisotropy is one uniform value that is used throughout a model layer.
Different layers can have different values for anisotropy, but anisotropy cannot vary
within a layer.
280
The steep dip of the aquifer system into the Hasa and Jafr basins along their
western flanks violates an assumption of layer horizontality in the model. The steepness
also make it impossible to properly represent rapidly changing features of the aquifer
system, such as the altitude of the top, bottom, or hydraulic head, without more detailed
data and finer discretization. Consequently, model results along the western margin could
be in doubt. However, in the Jafr Basin, emphasis was given only to the area east of the
Arja-Uweina Flexure which acts as impervious barrier for groundwater flow. This barrier
marks the end of the steepness in the dip of the aquifer system, where eastward from the
flexure the dip becomes milder. In other areas, given the regional nature of this study, the
dip in the aquifer system is believed to be acceptable for the relevant grid intervals.
Geological structures that could result either in preferred directions of! or produce
a barrier to groundwater flow, are not uniform throughout the modelled area, thus each
model cell was assumed isotropic. Geological structure designated as a conduit or a
barrier to groundwater flow are simulated by changing the permeability of the cells
representing that structure. Linear features that could be barriers to flow as indicated by
zones of steep gradient of the piezometric surface, were given very low permeability.
Although, these cells have been given reasonably small grid intervals, they could not be
adequately simulated in the model.
The model also incorporates the assumption that an impervious layer underlies the
lower most aquifer layer. The model also assumes that leakage through the confining
beds occurs simultaneously as heads in the aquifer changed and that no water is released
from the storage ofthe confining beds.
8.4 APPROACH
Commonly, model layers are separated on the basis of permeability contrast.
However, due to the complexity of the geologic structures, the lack of data, and to ensure
that the regional model would properly represent the aquifer system, the layering scheme
for the model was linked to the hydrogeological framework described' in previous
chapters. The regional hydrogeological framework of the carbonate aquifer system
defines a layered sequence of 3 aquifers and 3 confining units, underlain by the deep
281
sandstone aquifer system (Figure 8.1). The first upper layer was used to simulate the B4
aquifer system, the second to simulate the B3 aquiclude, the third to simulate the B2/ A 7
aquifer system, the fourth to simulate the A5/6 aquiclude, the fifth to simulate the A4
aquifer, the sixth to simulate both the A1I2 aquifer and the A3 aquiclude and the seventh
to simulate the deep sandstone aquifer system. Although, the Al-2 contains some water,
it is modelled as part of the sixth layer which is considered as an aquitard, since the water
bearing horizons within the Al-2 Formation are interbedded within thick sequences of
impermeable marls and all the Formation is separated from, and thus has no effect on, the
overlying aquifer sequence by the thick impermeable marls of the A3 Formation.
However, one layer or more are missed in some areas and (or) emerged into one layer
somewhere else through the study area.
Figure (8.1) Conceptualisation of the regional groundwater flow model.
Simulated flow within each layer is strictly horizontal and is perpendicular to
grid-cell faces. Simulated flow between layers is strictly vertical between vertically
adjacent cells. The quasi-three-dimensional approach described by Trescott (1975) is
282
hence being used; it is assumed that the vertical flow between aquifers are controlled by
the hydraulic characteristics of the confining unit and there is no horizontal flow within
confining units.
The approach used in this study was to simulate groundwater flow at both a
regional and subregional scale. In addition to the regional model, six are ally smaller, or
subregional (Figure 8.2), models were developed for Amman-Zerqa area, Amman-Zerqa
and Wadi Wala, Wadi Wala, Upper Wadi Mujib and Wadi Rasa basins, Wadi Mujib
(Wadi Wala and Upper Wadi Mujib basins) and Wadi Rasa basins, and Jafr Basin.
In the regional model, the extensions of, and the lateral changes and modification
in, the different aquifers and confining units are represented by changes in
hydrogeological characteristics within a layer. The A4 Formation for example, is an
aquifer in Amman-Zerqa area, underlain by the A3 and overlain by the AS/6, aquicludes.
To the south in Wadi Mujib Basin, the A4 aquifer wedges out and the whole Lower Ajlun
Group (AI-6) is considered to be an aquitard. In the model, the A4 is modelled as a
continuous unit with different hydrogeological properties, highly permeable in areas
where it acts as an aquifer, and having the same properties as the AI-6 elsewhere. And
the same for the AI-6 unites): an aquiclude in the north to an aquitard in Wadi Mujib
and Wadi Rasa basins and then to an aquifer (Fassu'a Formation) in the Jafr Basin. This
approach is found more realistic, since the lateral changes in the various aquifer and
confining units occur gradually.
The model simulate the regional flow system by approximating potentiometric
surface defined by the head measurements from wells, and the flow to streams provided
by spring discharges. The model was designed to simulate both steady-state and transient
conditions.
Prior to pumping long-term average head conditions prevailed in the system, and
the system was in an assumed state of equilibrium. Therefore a steady-state simulation
was used for pre-pumping conditions. The early head data and the modified
potentiometric surface map of Parker (1970) were used to calibrate the model under
steady-state conditions. After significant pumping began and at present, heads have
declined and are declining in the aquifer system. Its worth mentioning here that the
283
groundwater developments and significant pumping from the different aquifers began at
different times in various basins. Significant pumping from the B2/ A 7 and A4 aquifers
started as early as 1970 in Amman-Zerqa area, as well as in the Jafr area from the B4
aquifer. Elsewhere, significant pumping began later, in 1985. Thus, the modelling
approach was to simulate the transient conditions only for the subregional models.
Aquifer characteristics derived from the initial calibration of the steady-state
model were used as initial conditions for the transient model. The additional stresses
provided by simulated pumping in the transient model allowed refinement of the aquifer
characteristics which were then used to recalibrate the steady-state model.
To ensure continuity of the simulated aquifer properties between subregional
models, the Amman-Zerqa and Wadi Wala subregional models overlapped the Amman
Zerqa and Wadi Wala subregional model, and the same for the Wadi Mujib and the Wadi
Rasa subregional models which overlapped the Upper Wadi Mujib and the Wadi Rasa
subregional model. Modelling was independent from subregion to subregion, but
differences in calibrated values in overlapping areas were resolved by mutual agreement.
To provide an overview of the flow system in the study area, and to ensure the
compatibility of the different subregional models, a regional model was constructed using
the data from the subregional models. The regional model has a coarser mesh than the
subregional models.
The model was calibrated using measured heads and estimated flows to define the
areal distribution of hydraulic characteristics. Sensitivity analysis was performed to
assess the effect of ranges of hydraulic characteristics on model behaviour and, thus, to
determine the reasonableness ofthe calibration.
Groundwater development from the Al-6 aquifer system is negligible and no
drawdowns have been recorded throughout the aquifer: therefore, transient conditions for
the Al-6 aquifer system were not simulated.
The regional simulation is discussed herein, the subregional-based models and
results are described in regional basis.
284
8.5 MODEL GRID AND LAYERS
Analysis of a complex nonhomogeneous aquifer system is accomplished by
subdividing the aquifers into large number of rectangular cells, which constitute a finite
difference grid. The horizontal grid consists of number of rows and columns depending
on the model area and the cell dimensions. Vertically, the modelled aquifer system is
divided into number of layers, each layer represents an aquifer or a confining unit.
A grid (Figure 8.2) was superimposed on a map of the aquifer system to facilitate
discretization of data and enable finite-difference computations. The cell dimensions and
the orientation of the grid system were determined according to the following criteria:
1. The north-south orientation of the grid system parallels the prevailing direction
of the Western Highlands and adjacent valleys and, thus, the regional
geological structures. It coincides approximately with the principal regional
direction of hydraulic conductivity and groundwater flow.
2. The grid cell size was variable and the largest as possible that could be used
while still maintaining the topographic and structural expression of the study
area.
3. The width-to-Iength ratio selected avoids stability problems during solution of
the finite-difference equations. When the cell dimensions differ by more than a
factor of 2, the model may become unstable, creating erroneous solutions or
convergence problems (Remson and others, 1971).
4. A smaller grid si~e was used in areas of special interest, such as where the cell
represents certain geological structure or topographic feature.
The distribution of active cells within each layer is defined by the extent of the
saturation limits of the geohydrologic unites) within the model area. The limits of the
model area are related to the extent of the occurrences of the Ajlun Group in the study
area. The limits coincide with the Western Highlands in the west, and extends a few cells
beyond the eastern boundary of the study area where the aquifer system continues farther
in the east. The isolated volcanic eruptions in many places throughout the study area are
too small to be considered as internal limits for a regional groundwater flow model. Since
these eruptions do not have a significant influence in groundwater flow mechanisms.
285
160
140
100
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240
\\ \ \
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LEGEND
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260
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280
\ \ \ \
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300
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320
Figure (8.2) Regional and subregional model areas.
Amman-Zerqa
Wadi Wala
Amman-Zerqa and Wadi Wala
Jafr Basin
Regional
Wadi MY.iiJ21lliin
Figure (8.3) Conceptualisation of the subregional groundwater flow models.
287
The various conceptual models used to simulate the groundwater flow
system ofthe subregional and entire study area are following (Figure 8.3):
1. Amman-Zerqa subregion, occupying the area between Wadi Dhuleil-Zerqa
River in the north and Amman-Zerqa flexure in the south (Figure 8.2). In this area the
groundwater flow system was modelled in three layers, two aquifers (the B2/A7 and
A4) and a confining unit (A5/6). Horizontally, the grid system consists of 41 rows and
59 columns with variable cell size ranging between 416.67-833.33 m (Figure 8.4).
2. Wadi Wala subregion, occupying the surface catchment area of the Wadi Wala
Basin from the Amman recharge mound in the north to the Swaqa fault line in the south.
The flow in this region was simulated by four layers (Figure 8.3, Wadi Wala), from top to
bottom; B3 confining unit, B2/A7 aquifer system, AI-6 aquitard, and the deep sandstone
aquifer system. The model was discretisized horizontally into 49 rows and 52 columns
with a variable node spacing ranging between 1250-2500 m (Figure 8.5).
3. Amman-Zerqa and Wadi Wala subregion, this model including the Amman
Zerqa and the Wadi Wala groundwater basins, between Wadi Dhuleil-Zerqa River in the
north and the Swaqa fault line in the south (Figure 8.2). The flow in this subregion was
simulated by 6 layers (Figure 8.3, Amman-Zerqa and Wadi Wala) , the top layer
representing the Muwaqqar confining unit (where it is present), the second layer
representing the B2/A7 aquifer, the third layer representing the A5/6 aquiclude (aquitard
in the Wadi Wala) , the fourth layer representing the A4 aquifer (aquitard in the Wadi
Wala), the fifth layer representing the A3 and A1I2 aquiclude (aquitard in the Wadi
Wala), and the sixth layer representing the deep sandstone aquifer system. The model was
discretisized horizontally into 72 rows and 52 columns with a variable node spacing
ranging between 625-2500 m (Figure 8.6).
4. Upper Wadi Mujib and Wadi Hasa subregion. This area lies between the Swaqa
fault line in the north and Salwan fault line in the south (Figure 8.2). Vertically, the
flow system was simulated by 4 layers, similar to the Wadi Wala; the top layer
represents the Muwaqqar confining unit (where it is present), the second layer represents
the B2/A7 aquifer, the aquifer of primary concern. The third layer, the AI-6 aquitard,
allows hydraulic interchange between the B2/ A 7 and the deep sandstone aquifer system
288
(the fourth layer). Horizontally, the model consists of 70 rows and 64 columns. The
model cell size ranges between 1250-2500 m (Figure 8.7).
5. Wadi Mujib and Wadi Hasa subregion. This represents the whole surface
catchment area of the Wadi Mujib and Wadi Hasa basins, from the Amman recharge
mound in the north to Salwan fault line in the south (Figure 8.2). It has the same vertical
subdivisions as in the Wadi Wala subregion (Figure 8.3, Wadi Wala), The horizontal grid
consists of 114 rows and 64 columns with variable node spacing range between 1250-
5000 m (Figure 8.8).
6. Jafr subregion. The area between Salwan fault line and the southern boundary
of the study area (Figure 8.2) was simulated in 5 layers representing 4 aquifers and one
confining unit (Figure 8.3, Jafr Basin). The horizontal grid consists of 53 rows and 71
columns with cell size ranging between 625-2500 m (Figure 8.9).
7. The regional model. The entire study area was modelled in 4 layers (Figure 8.3,
Regional), two main aquifer systems and two confining units. The aquifer systems are the
B21 A 7 and the deep sandstone regional aquifer systems, separated vertically by the Lower
Ajlun group (Al-6), which has variable hydrogeological characteristics (changing from
aquiclude to an aquifer) throughout the study area. Simulated vertical flow between the
two main aquifer systems was controlled by the intervening Al-6 unit in which vertical
leakage could be varied areally. In most of the study area, where the B2/A7 aquifer is
overlain by the Muwaqqar confining unit (B3), it is confined: elsewhere, along the
Western Highlands, the aquifer is under water table conditions. The Basement Complex
forms the lower boundary of the regional aquifer system groundwater flow model. The
horizontal finite-difference grid for the regional model consisted of 99 rows and 36
columns with a variable node spacing ranging between 1250- 5000 m (Figure 8.10).
In the regional groundwater flow model, the A4 aquifer in Amman-Zerqa area is
not simulated: it is included in the Al-6 aquiclude. This omission of the A4 aquifer in the
simulation creates no significant hydrologic errors, because the Al-6 aquiclude generally
is about 6-8 times as thick as the A4 aquifer. Therefore, the vertical leakage into the deep
sandstone aquifer system, is dominantly affected by the hydrological characteristics of the
Al-6 aquiclude not the A4 aquifer.
289
.~ ~
P ,
E3 E3 E3 1<> 1(0)) III
; ~ . i .. ·,·
fiJI ellB ,Ii
GliB
Wadi Dhuleil ,. IS. J'6. ;s J .. ~~ . ~_ . s.
1'.
H He' ,. 1(0)) III
, .
..
Figure (8.4) Finite-difference grid for the Amman-Zerqa subregional model.
25 30 35 40 45 50
Figure (8.5) Finite-difference grid for Wadi Wala subregional model.
35 40 45 50
Figure (8 .6) F nee grid for the Amman-Wadi Wala subregional model
Figure (8.7) Finite-difference grid for the Upper Wadi Mujib and Hasa subregional model.
Figure (8.8) Finite-difference grid for the Wadi Mujib and
Wadi Hasa subregional model.
- " " J6 ,. .. ,.
" " .. .. .. .. .. 6~~1 . eM' . ",::",
" ~ II ...
82/J!.:7 aquifer Syste,~~
I. 11'
J6
,.
" .. "
"
'. ~ . >
s s S ' .. '"
Figure (8.9) Finite-difference grid for the Jafr subregional model.
Figure (8.10) Finite-difference grid for the regional model.
8.6 BOUNDARY CONDITIONS
Boundary conditions are one of the most important inputs to a simulation model.
The boundary conditions were selected to best represent the groundwater flow conditions
near the boundary of the aquifers. The boundary conditions between the different
subregional models were chosen to coincide as closely as possible with assumed no-flow
boundaries or with the groundwater divides. Because the . boundary specifications
represent observed or inferred conditions at the limits of the aquifer system, the simulated
conditions for the interior parts of the flow system are reasonably free of boundary error.
This assessment assumes that the model results are used in conjunction with other sources
of information and are tempered with the understanding that the model is a learning tool
for regional application, rather than a management tool with local application.
Several different types of boundary conditions for the regional and subregional
models, as shown in figures (8.4) through (8.10), can be assigned. They are as follows :
1. Constant head boundary. In general, the model boundaries of the carbonate
rocks were extended to the mountain ranges in the west, where the aquifers crop out in a
zone where annual rainfall exceeds 400 mm. The western model boundaries along the
Western Highlands were designated as constant head boundaries. A break in this line of
constant head cells coincides with areas where the groundwater flows westward to
discharge at the surface as spring discharges, such as in Wadi Haidan, Wadi Mujib, Wadi
Karak, and Wadi Rasa. In these locations, general head boundaries were used to permit
simulation of discharge of groundwater to the west. The lines between the recharge
mounds and the discharge areas which coincide with the western saturated limits of the
B2/ A 7 aquifer system were designated as flow line. The same constant head boundary
were given to cells representing the north-eastern comer of the model boundary, where
the groundwater transfer between the basalt and the B2/ A 7 aquifer systems occurs.
The assumption of a constant head water table for simulation purposes can be
justified by calculation of the amount of water table change represented by the downward
leakage rate. A downward leakage rate range between 1825-9125 m3/a/km2, as estimated
previously, would represent a negligible water table decline, far less than the usual
groundwater recharge rate from precipitation over most of the Western Highlands.
297
Although constant head boundaries have been used, generally they are not
desirable. However, the location of constant head boundaries is here associated with
recharge areas. In each case, the use of constant heads to simulate groundwater recharge
to the aquifer systems can be compared with the amount of estimated recharge and the
groundwater discharge from the aquifer systems.
The aquifer system extends eastward beyond the eastern limit of the study area,
beneath the Muwaqqar Formation confining unit. So the eastern boundary is considered
as an equipotential line with fixed head for the preliminary stage of the model calibration,
just to allow the model to calculate the outflows of the system along the eastern
boundary, which then will be designated as specific flow boundary, in the final stage of
steady state calibration and transient conditions, with the same previously simulated flow.
2. Specified flow is given, as discussed above, to the nodes along the eastern
boundary, where groundwater flows out to further east from the study area, and for the
cells representing the spring discharge in the western highlands of the Jafr Basin, where
the steep topography precludes simulation of the spring discharges by using any of the
boundary conditions. In these nodes, the observed spring discharges were used as
specified flows.
3. General head boundari~s are given to the spring discharge points throughout
the model area. By using this type of boundary, the model can simulate the effect of
regional decline in water levels on spring discharges. General head boundaries are also
used in the upper layer in four locations, where perennial streams cross the area such as
Wadi Dhuleil-Zerqa River, Wadi Haidan, Wadi Mujib and Wadi Hasa. The Wadi
Dhuleil-Zerqa River, considered as an outflow for the main aquifer system, approximates
the northern boundary of the model area. Part of the water which flows in the Zerqa River
is assumed to infiltrate downward to recharge the underlying B2/A7 aquifer. The major
wadis ofHaidan, Mujib and Hasa, are eroded canyons which have cut the saturated zones
of the aquifer system.
The general head boundary allows flow to occur either to or from the model cell
depending on whether the head in the model cell adjacent to the boundary is less than or
greater than the specified boundary head (McDonald and Harbaugh, 1984). The
298
simulation of flow across the general head boundaries was computed by multiplying the
head difference across the boundary with a conductance term. The head difference was
determined by comparing a specific boundary head to that in the adjacent model cell. The
conductance term is the hydraulic conductivity times the cross-sectional area of the
boundary through which flow is simulated divided by the length of the flow path. The
conductance terms were adjusted during the model calibration.
4. No-flow boundary. The southern and parts of the western limits of the modelled
area were approximated by the west-east, southwest-northeast, and south-north flow lines.
These flow lines are chosen as closely as possible and parallel to the saturated limit of the
aquifer system. The nodes along these flow lines were simulated as no-flow boundaries.
No-flow boundaries are also used to simulate the boundaries between the different
subregional models, since these boundaries were chosen to coincide with regional
structural barriers or flow lines. Additionally, the major fault and flexure systems in the
area, which act as a barrier for groundwater flow, are considered to be internal no-flow
boundaries.
6. Stream boundaries were assigned to cells where the Zerqa River traverses the
outcrop areas of the B2/ A 7 and A4 aquifer systems in the Amman-Zerqa area. Other
streams occur below the saturation zone of the main aquifer systems, thus, they were
simulated as general head boundaries.
8.7 INPUT DATA
The model requires values of the different hydrogeological properties of the
aquifer system. These properties are spatially distributed, one value per cell in each model
layer. For each cell, values of altitude of the top and bottom, hydraulic head, hydraulic
conductivities (lateral and vertical), conductance for stream bed and general head
boundaries, recharge, and discharge were supplied. These values are assumed to be
constant everywhere in the cell and are the average values. For transient simulations, in
which conditions are time-dependent, changes in groundwater storage can occur, and the
model also requires storage coefficient data.
299
The altitude of the top and bottom of the layer as well as the layer thicknesses are
derived from the structure contour maps, well drilling and other information explained in
the previous chapters. Transmissivity distribution depends on the saturated thickness and
the hydraulic conductivity of the layer. In the unconfined part of the aquifer, the saturated
thickness is the difference in altitude of the hydraulic head and bottom of the layer, while
in the confined part the saturated thickness is equal to the thickness of the layer.
Hydraulic head data were obtained from the measured water heads in the wells and water
level contour maps in the study area.
The aquifer properties varied during model calibration were hydraulic
conductivity, leakance, conductance, and recharge. Initial estimates of hydraulic
conductivity for the aquifer system were derived from aquifer test data and specific
capacity data. The regional distribution of the hydraulic conductivity discussed in chapter
five was used to obtain a first indication of the likely hydraulic conductivity distribution.
Lateral hydraulic conductivity values for the confining units are assigned values for the
purposes of estimating the vertical hydraulic conductivity, as explained later. They are
assigned to be a certain fraction of the regional average values for the overlying aquifers.
The fraction used was 0.01 because the hydraulic conductivity of marls and shales, the
major components of the confining units, is typically several orders of magnitude less
than that of the limestones which are the principal rock type of the aquifer system. These
fractions were applied uniformly within each layer. Departures from the procedure are
invoked for areas where the lower confining unit becomes an aquifer as in Jafr area,
where the B21 A 7 aquifer directly overlies the A 1-6 aquifer. Then the lateral hydraulic .
values of the A 1-6 aquifer were increased.
Vertical conductance (leakance), L, is a property that controls the rate of vertical
flow between layers. It is calculated by one of two methods (Figure 8.11). In the first, the
vertical hydraulic conductivity has to be input and the model calculates L according to the
relation (McDonald and Harbaugh, 1984):
300
where
1 dj{ dYz' L = kf v, I + k fv22 ............................ ..... .... ................... .............................. (8.7)
d 1 = thickness of the upper aquifer
d2 = thickness of the lower aquifer
kfv, I = vertical hydraulic conductivity of the upper aquifer
kfv, 2 = vertical hydraulic conductivity of the lower aquifer
This relation controls vertical flow between two adjacent geohydrological units where an
intervening layer does not exist, such as the case between the B2!A7 and A1-6 aquifers
in the J afr area.
In the second method, the L must be applied directly according to the relation
(McDonald and Harbaugh, 1984):
where
d l /
~=_1_22_ + L kfv, I
dYz' + ~ ...... ................................................................ (8.8)
k fv,2 kfv, c
dc = thickness of the confining layer
kfv, c = vertical hydraulic conductivity ofthe confining layer
This relation defines vertical flow through a well-defined confining unit. It was applied to
control flow between B4 and B2!A7 aquifers in the Jafr area, where the aquifers are
separated by the B3 confining unit, between B2! A 7 and A4 aquifers in Amman-Zerqa
area, where the aquifers are separated by the A5!6 confining unit, and between the B2! A 7
and the deep sandstone regional aquifer systems, where the aquifer systems are separated
by the A1-6 aquiclude(aquitard).
Vertical hydraulic conductivity values for each layer were assumed to be a certain
fraction of that layer' s lateral hydraulic conductivity. Ratios of vertical to lateral
301
d I
C
without semiconfining unit
/~
, , , /
with semiconfining unit
ktv,1
Figure (8.11) Diagrams for calculation of verticalleakance
hydraulic conductivity were assumed about 0.001 for unites) considered as an aquiclude
and about 0.01 for unites) considered as an aquifer. It should be noted that these estimates
are initial values which are applied to the model and were open to modification during the
model calibration to improve simulation within limits permitted by geological evidence.
The conductance term for general head boundaries and streambeds is the
hydraulic conductivity of the boundary or the streambed times the cross-sectional area of
the boundary through which flow is simulated divided by the length of the flow path.
Owing to the lack of any values for this term, the conductance value were entirely based
on the model simulations.
Recharge was assumed to occur in the western highlands except in wadis, where
indirect recharge is expected to occur. Initial recharge estimates were based on the
previously discussed recharge calculations. It is applied as a percentage of the estimated
annual precipitation. The percentage increases from 0 % for the annual precipitation zone
below 200 mm to a maximum of 20 % for annual precipitation zones of more than 600
mm, but excludes precipitation that falls on the deep valley floors.
Natural groundwater discharges were derived from spring flow measurements:
however, emphasis was given to discharge quantities as an easily measurable or
estimatable item of the groundwater balance. Nevertheless, uncertainties remain in the
overall balance. The amount of subsurface discharge and recharge and the quantity of
water lost as hidden springs and seepages are still a matter of rough estimate. The visible
discharge in the area appears as spring and Wadi base flows. Spring distribution,
partiCUlarly in Wadi Mujib Basin shows that not only the B2/A7 contributes to the entire
spring flow, some springs discharge from the AI-6 series, which means that the AI-6
series cannot be considered as a perfect aquiclude. Groundwater abstractions from wells
are used for transient simulations, and were obtained from water consumption data
throughout the study area.
The model also requires information about the type of layer under simulation. For
the top layer, confined/unconfined conditions were chosen, in which the transmissivity of
the layer varies, it being calculated from the saturated thickness and the hydraulic
conductivity. The storage coefficient may alternate between the confined and unconfined
303
value. The other layers in the model were assumed to be under confined conditions.
Initial values of storage coefficient, used in transient simulations, were based on estimates
previously discussed in chapter five. Where confined conditions prevail, storage
coefficient was set equal to 0.0001. To represent unconfined conditions, storage
coefficient was set equal to 0.1 (specific yield).
8.8 MODEL SIMULATIONS
8.8.1 STRATEGY
The basic goal for calibration was to obtain a model that could simulate actual
hydrologic conditions within acceptable limits of error. The calibration criteria used to
calibrate the model were the calculated water balance and the groundwater head. The data
used as calibration standards were based on field observations, as well as on the
previously discussed conceptual models of the groundwater flow systems. The results of
calibration were used to re-evaluate and improve the conceptual models of the system, as
well as to compile a water budget for the groundwater regime. Calibration was largely a
trial-and-error process in which the input data were modified in response to shortcomings
in the model, as determined by the importance of the difference between the simulated
conditions and the observed or (inferred) conditions. Numerous sets of input data were
required to simulate the aquifer system. The accuracy of those data determined the
reliability of the simulated conditions. In tum, the accuracy of the input data was strongly
influenced by the availability and validity of the control data, which for the most part
consisted of field observations made during previous hydrogeological investigations.
The overall strategy of the model calibration was to delineate a plausible set of
boundary conditions, initiate simulation using preliminary estimates of recharge,
hydraulic conductivity, leakance, streambed conductance, and storage coefficient, and
refine the original parameter estimates by trial-and-error simulation until the model's
output data satisfied the calibration criteria. Throughout the process of model
development, an attempt was made to achieve a physically meaningful characterisation
of the flow system. Therefore it was essential to keep the ensuing adjustments within the
limits of sound hydrologic judgement and geological principles.
304
The model was calibrated for both steady state and transient conditions. The
steady state model was developed first, and provided the initial conditions for the
transient simulations. The transient model, a functional extension of the steady state
model, represents the stresses of pumpage and considers the effect of times and
groundwater storage. Although the steady state and transient versions of the model
contain much of the same program logic and input data, the individual versions depict
different sets of conditions, thus requiring adjustment of different data sets during the
calibration process. The hydraulic conductivity, leakance, riverbed conductance and
recharge data were calibrated during steady state runs, the storage coefficient data were
calibrated during transient runs. Because both the boundary conditions and the pumpage
were considered known components, neither of these data sets was modified as part of the
calibration process. The validity of the transient results proved to be very dependent on
the distribution of recharge and leakage, which were functions of the steady state
calibration. Consequently, during the later stages of calibration, steady state runs were
alternated with transient runs, and adjustments were made to the appropriate data sets in
the appropriate model so that the calibration of each model appeared to improve from
each change.
8.8.2 STEADY STATE CALIBRATION
The steady state model was calibrated to simulate conditions in the groundwater
flow system prior to the beginning of significant pumpage, when the aquifer system was
for the most part in its natural, predeveloped state. In this state, recharge was
approximately equal to discharge, water levels were essentially stable, and there were no
significant changes of groundwater in storage. Groundwater development from the
different aquifer systems in the different groundwater areas began at different times,
between 1960 and 1985, but groundwater level declines in some of these areas started
more recently because the early stages of pumping involved small groundwater
development. Verification data for the preliminary predevelopment model consists of
water level measurements given in the earliest reports available (Parker, 1970 and
305
Mudallal, 1973). Prior to 1970, the small amount of abstraction is assumed to have taken
place without significant changes in groundwater levels or storage.
In the preliminary steady state model, recharge was introduced by constant flux,
in the up dip areas of the respective aquifer outcrops, as a percentage of precipitation. This
might result in the injection of too much recharge, since the model will introduce the
recharge irrespective of whether the other hydrological parameters will allow for recharge
to occur or not. But it is reasonable from the perspective that the model simulates the
long-term average recharge.
During a simulation, recharge was held constant and the hydraulic conductivity
and leakance values were adjusted until simulated water levels in the model layers
approximated the general water levels trends. Changes to estimates of these aquifer
parameters were, in general minor. Significant changes that deserve discussion include:
hydraulic conductivity in the aquifer outcrop areas that control the amount of recharge
that enters the aquifer; the aquifer hydraulic conductivity along the lineament features
which control the groundwater flow distribution; and the vertical hydraulic conductivity
of the confining units that effects the amount of vertical groundwater flow and thus the
amount of recharge to the deep aquifer systems. In the preliminary steady state model,
hydraulic conductivities in the outcrop areas of the model aquifers, and the vertical
hydraulic conductivity, were set to high values. This allowed the aquifers to accept as
much recharge from the outcrop as they could transport downdip under the hydraulic
conductivities used in the model; that is, the aquifer never became starved for recharge at
the outcrop. But the amount of recharge is dependent, among the other previously
discussed parameters, on the amount of rainfall, and the ability of the aquifer materials to
transmit water. Though groundwater rejection from the aquifer system due to access of
recharge is uncommon, to avoid an overestimation of recharge, the maximum allowed
annual recharge was limited to 20% of the annual precipitation and balanced against the
amount of natural groundwater discharges. This configuration of parameters provided a
good match to the available predevelopment water levels data.
During the transient simulation, however, it was discovered that in some of the
areas subject to heavy pumping it was necessary to restrict the availability of recharge
306
from the outcrop to match the pattern and distribution of observed water level declines.
Varying the storage coefficient within a reasonable range of values could not produce the
desired declines. Restricting the quantity of recharge that could be captured by the
aquifers was possible by reducing either the hydraulic conductivity in the outcrop areas or
the vertical hydraulic conductivity for the confining units, thereby reducing the ability of
the aquifer to move water downdip from the outcrop areas. The adjustments of the aquifer
parameters according to the two possibilities were made bearing in mind that under
steady state conditions, enough water has to be moved vertically to replenish the deep
sandstone aquifer system, in which, in tum will maintained a base flow of about 55
MCMla discharge from the sandstone aquifer system along the western escarpment into
the Jordan Rift Valley.
To verify that the final selection of parameters made during transient simulation
continued to represent an acceptable calibration for steady state predevelopment
conditions, the model was rerun as a steady state simulation without pumpage, and the
results were analysed using the same technique used to evaluate the preliminary
predevelopment model.
Water levels were a major control on the distribution of simulated hydraulic
conductivity. The direction of flow reflects the presence of major internal barriers. The
effect of these barriers on groundwater movement has been discussed earlier in chapter
seven. Model cells along these barrier were set to a very low hydraulic conductivity.
Once the simulated heads approximated the estimated heads, model calibration
was focused toward simulating the observed discharge from springs and the estimated
water budget. Simulated spring discharges were calibrated by first adjusting the
conductance value of the head dependent function used to simulate springs in the model
and then by adjusting the hydraulic conductivity and leakance values in the vicinity of the
simulated springs. In a few instances, the altitudes of the springs were adjusted,
particularly in areas of large relief where estimated of land surface altitude as well as the
estimated head values were averaged for the cell.
Changing the conductance values of the springs had little effect on simulated
discharge rates but tended to increase or decrease the simulated head in the model cell.
307
The conductance values of simulated springs were then set high enough such that the
heads in the model cells were only slightly greater than the altitude of the spring. The
discharge to the springs was then calibrated to by adjusting the hydraulic conductivity
and leakance values in the vicinity of the spring. Comparison between observed and
calibrated discharge rates for each spring or spring areas is discussed below.
The model is accepted as being calibrated when the difference between simulated
and observed hydraulic heads is acceptable and simulated discharge from springs
approximated the known discharges. The agreement between observed water levels and
calculated water levels, for selected control points, under steady state conditions is shown
in Table (8.1) and Figure (8.12). However, the disagreement between the simulated and
observed water levels could be attributed either to the unreliability of the data, or that
observed water levels refers to the water levels of the observation well itself while the
calculated water levels refers to the water levels of the centre of the block which contains
the well. Therefore, higher differences between calculated and observed water levels are
obvious where the hydraulic gradient is steepest, particularly around the internal
groundwater barriers and in the deep incised wadis. The largest differences between the
simulated and observed hydraulic heads occur along the western boundary of the Wadi
Hasa and Jafr groundwater areas, where the model cannot reproduce the hydrogeological
complexities of the steep, faulted western highlands.
However, the calibration of the model for the predevelopment conditions was
judged to be adequate. The distribution of the final calibrated hydrologic parameters for
the aquifers and confining units is discussed in the following sections.
8.8.2.1 SIMULATION RESULTS
This section discusses the results of the model simulations, in particular the
direction and magnitude of groundwater flow in the different subregions, and the
magnitude and distribution of hydraulic properties determined by model simulations. The
groundwater budget for the different subregional model areas will discussed later.
308
Observation well Observed (masI) Simulated (masl) Error(%) Refmary No.1 488.3 490 0.35 Zerqa Obs. 490.7 495 0.88 Awajan Obs. 574 575 0.17 RC29 688.5 685 0.51 Q7 570 570 0.0 W14 450 450 0.0 SW2 665.7 665 0.11 LA2 711 710 0.14 Qatrana No. 10 (S124) 679 685 0.88 nCA1 526.7 525 0.32 nCA2 582.7 582 0.12 nCA3 664.1 665 0.14 JICA4 756.4 760 0.47 AB3 759 760 0.13 Rasa No. 11 791 790 0.13 PR05 789.1 790 0.13 JT3 876 875 0.11 JT4 785 785 0.0 S53 916.3 915 0.14 S65 1193 1180 1.1
S70 578 580 0.35 S83 740 740 0.0 S100 1463 1300 12.5 S108 1309 1250 4.72 S118 1271 1225 3.82 S121 988.5 990 0.15
Table (8.1) Observed and simulated water levels for selected observation wells.
8.8.2.1.1 FLOW SUBREGIONS
AMMAN-ZERQA AREA .
The steady state simulation in the Amman-Zerqa area model shows that the
groundwater flow through the aquifer is small and mainly restricted to a narrow strip
along the Zerqa River. Assuming the width of this strip ranges between 1000 and 2000 m,
average hydraulic conductivity of about 1 mIh, average saturated thickness of 50 m, and
average hydraulic gradient of 0.01, the flow of groundwater ranges between 4.6 and 9.1
MCMla. This agrees with the model prediction of the outflow of groundwater from the
downstream end ofthe basin of about 7.2 MCMla.
309
160
140
120
100
80
60
40
20
000
980
960
940
'. --..... "".';"~"" .'.
200 220
Zerqa River '- "".r-'=+-~:..--,--:---....,
240
\ ~ \ \ '>" ,
\ , \ ' \ \
\ \
"
\ \' ,
..... .... ......
• Jat'
260
10 20 :\O'cm
LEGEND
• crty, town
~~:. wadl,river
--- fault
- fautt. Inferred from groundwater modehng
/,00"'" eqUtpotentialltne 800 m
model boundary
~ dry.rea
'>Go
, , , \ '
" \ \' \ \
\ \ \ ,
\ , \
\ \
\ \
280 300 320
Figure (8.12) Simulated steady state water levels for the B2/A7 aquifer system
Groundwater levels were found to be very sensitive to the recharge distribution,
the hydraulic conductivity, and the leakance of the AS/6 confining unit. Adjustment of
these parameters within a range which provides a simulated recharge amount
approximating the natural discharge of about 12.7 MCMla, and more importantly,
maintain a continuous groundwater flow through the Amman-Zerqa groundwater basin
was difficult. As discussed in the previous chapter, the aquifer system in the Amman
Zerqa area consists of number of sub-basins with limited connection to each other.
Decreasing the recharge and increasing the hydraulic conductivity lowers the simulated
water levels, and the groundwater appears as isolated areas with dry nodes at the rim of
each area. To allow the model to simulate the limited connection between these areas
through the alluvial deposits without affecting the groundwater budget, a great effort was
spent in adjusting the relation between hydraulic conductivity and recharge in order to
obtain a reasonable distribution of simulated water levels.
The leakance of the AS/6 confining unit controls the upward leakage into the
B2/ A 7 aquifer system from the deeper A4 aquifer system. The simulated piezometric
levels for the A4 aquifer were found to be too high with small hydraulic gradients toward
the west and northwest of Amman. However, the steady state simulation indicates that the
flow through the A4 would be much larger than the potential recharge capacity for the
applied hydraulic conductivity. Even with a general reduction in hydraulic conductivity
by 20%, the piezometric levels in Amman area were still too high. This cannot be
explained by the recharge /discharge relation and the hydraulic conductivity distribution
assuming that the map presented by Parker (1970) was relatively close to the original
conditions. Therefore, the assumption was to made to increase the leakance value of the
AS/6 confining unit and thus allow a leakage of around 2.16 MCMla to the upper B2/ A 7
aquifer under steady state conditions.
The fixed head cells at the northeastern comer of the model area indicate that,
under steady state conditions, 3.8 MCMla were simulated to enter the groundwater
system as subsurface inflow from the adjacent basalt aquifer system.
311
The simulated flow through the A4 aquifer is approximately 9.86 MCMla.
However the increased abstraction may transform part of the A4 aquifer outflow areas,
particularly downstream of Amman, into recharge areas.
The direct contact between the two main aquifer systems at the downstream end
of the area, was not simulated because of the large grid spacing used in the model. It is
thought that this direct contact is not of significance for the groundwater flow system.
WADI WALA BASIN
The general features of groundwater flow system in the Wadi Wala Basin are the
Amman recharge mound in the north, the impermeable Swaqa fault line in the south, and
the Haidan springs in the lower reach of Wadi Wala in the western part. Under steady
state conditions groundwater flows mainly southwards then south-westwards from the
recharge mound, where the groundwater flows out via the Haidan springs. The dry nodes
at the north-eastern comer of the basin indicate that the south-eastward groundwater flow
component from the recharge mound is rather small. The Salwan fault line is a
groundwater barrier which separates the groundwater of the Wadi Wala Basin from that
ofthe Wadi Mujib Basin.
The criteria for calibrating the steady state model were that the groundwater level
configuration, the Haidan spring discharges of about 15 MCMla, and the dry spots as
revealed by drilling in the north-eastern part of the basin were to be represented. It was
impossible for the model, for any combination of data input, to reproduce the steady state
conditions according to these criteria, without proposing a northwest-southeast trending
barrier features which might be a fault line as shown in the water level contour map
(Figure 8.12). Nodes along this feature were given a very low hydraulic conductivity.
Projection of this lineament on the geological map shows that it coincide with
southeastern sudden termination of the B3 Formation: furthermore, without previous
knowledge, Howard Humphreys (1977) reported that the NRA (without reference), based
on geophysical investigations, recorded the existence of a northwest-southeast lineament
passing the Qastal pool. This coincides with the modelled proposed lineament.
Accordingly the model simulates the steady state conditions with difference between
312
simulated and observed water levels less than 5 m, and 17 MCM/a for the Haidan spring
discharges.
WADI MUJIB AND WADI HASA BASINS
The hydraulic situation, particularly in Wadi Mujib is very complex, being highly
controlled by the structures in the area. Besides the hydraulic heads and outflow from the
aquifer system which occurs as Wadi baseflow and spring flow, and thus calibrated in
accordance with measured values, the previous discussions in the chapter on groundwater
flow ( Chapter 7), provide the general criteria for the overall steady state simulation of the
aquifer systems in this area. The calibration of the model appeared very sensitive to
variation of the hydraulic parameters.
The calibration indicated that several major fault lines had an influence on
groundwater flow. For faults considered as a groundwater barrier, the flow across them
was reduced substantially by decreasing the hydraulic conductivity. Sometimes, in the
vicinity of the fault systems, a higher flow to the deeper sandstone aquifer had to be
assumed, in this case the vertical hydraulic conductivity of the AI-6 had to be increased
up to 0.00000625 mIh
The highly permeable drainage lines provided by the Karak-Wadi el Fiha, Wadi
Yubbs, and Sultani-Qatrana fault lines were given higher hydraulic conductivity, either to
obtain good agreement between calibrated and measured hydraulic heads or to maintain
the observed spring flows such as the case in Sarah spring in the Wadi Karak Graben. The
simulated flow from the Sarah spring was estimated to be 5.3 MCM/a, which is only 6%
higher than the observed flow of 5 MCM/a.
Considering the overall groundwater balance of the Wadi Mujib and Wadi Hasa
basins, to reconcile the composite relationships between the hydraulic heads, hydraulic
conductivity, recharge and discharge, it was essential to increase the vertical hydraulic
conductivity of the AI-6 and consequently the vertical flow into the deep sandstone
aquifer system. The average calibrated vertical hydraulic conductivity of the AI-6 was
found to be about 0.0000021 mJh which resulted in vertical leakage into the deep
sandstone aquifer of about 39.32 MCM/a. The amount of vertical leakage into the deep
313
sandstone aquifer system is assumed to discharge from the system as spring discharges.
The simulated outflow from the sandstone aquifer under steady state conditions, for an
average calibrated hydraulic conductivity of about 0.08 mIh, was about 65 MCMla
against observed discharge from springs issuing from the sandstone aquifer system of
about 30 MCMla. The difference between observed and calculated discharge however is
assumed to discharge as hidden outflow from the system into the Dead Sea.
Generally, there is a good agreement between simulated and observed water
levels. Large differences between the simulated and observed water levels were observed
in some areas with steep topographic relief and high hydraulic gradients. But for the most
of the modelled area, the difference was in the range of about 14 m, which is believed to
be a good approximation for an area of such hydraulic head differences of up to several
hundreds of metres. Furthennore, it has to be considered that the model results are the
mean water levels values for the cell areas while the observed are spot measurements of
water levels of the wells.
JAFRBASIN
The groundwater of the B2/A7 and the A1-6 aquifer systems in the Jafr area are
bound between groundwater barriers of the Arja-Uweina flexure in the west and the
Salwan fault line in the north. The Karak-el Fiha fault line, however, intersecting the
eastern part of the area, is proved to be penneable, allowing the groundwater to outflow
further to the east. The southern boundary is defined by the limit of saturation of the
aquifer systems and thus has been modelled as a no flow boundary.
Steady state simulation shows the Arja-Uweina flexure is not a continuous
groundwater barrier but allows some groundwater to pass to the east. A simulated flow of
about 2.95 MCMla across the Arja-Uweina flexure was derived from the final calibration.
This amount was assumed to represent the rate of lateral recharge to the aquifer system,
transferred from the Western Highlands recharge area. It is the same for the Salwan fault,
which is frequently cut by a series of north-south trending discrete fault systems, which
allow a small amount of groundwater to pass from the Rasa Basin to the Shidiya area in
the J afr Basin.
314
Due to the steep topography and thus the high hydraulic gradient in the western
part of the Jafr Basin, the model fails to produce the predevelopment steady state
conditions with a adequate agreement between the simulated and observed data.
Emphasis was given only for the area to the east of the Arja-Uweina flexure. But to
simulate the effect of the constant head boundary on the groundwater levels in the central
part of the Jafr Basin, the model was tested by reducing the water levels of the constant
head boundary by 100 m: there are no significant effect on the simulated water levels in
the central part of the area. However, in the western area, as indicated by the high
hydraulic gradient, low hydraulic conductivities were assigned to both the B2/A7 and AI-
6 aquifer systems. The observed 1.3 MCMla of spring discharge was adopted here as a
constant discharge for the steady state and transient simulations.
The groundwater flow system in the Shidiya area was described early in chapter
seven. The model simulation reflects the geology and the structure of the area, where the
hydraulic conductivity, leakance, and the thickness of the aquifer systems, increase
eastwards. The water levels for the B21 A 7 aquifer system in the Shidiya area found to be
controlled by the composite relationships between the high, uniform hydraulic
conductivity of the B2/A7 and AI-6 aquifer systems (which results in the observed
uniform, very small hydraulic gradient), the vertical groundwater flow into the AI-6
aquifer (which is obvious from the significant drop in water levels), and the calibrated
hydraulic conductivity of the Salwan fault line, since it was essential to allow a small
amount of water to pass cross the fault line into the Shidiya area.
The simulated flow to the system was found to be in the range of about 12.4
MCMla, including 6 MCMla which is believed to enter the aquifer system as indirect
recharge through the Wadi beds on the outcrop areas ofthe B2/A7 aquifer system.
However, it is worth mentioning that due to the lack of data from the AI-6 aquifer
system, the steady state calibration was only based on a few observed hydraulic heads in
the latter aquifer system.
In steady state calibration of the Rijam aquifer system (B4), although, the vertical
hydraulic conductivity of the lower confining layer (B3 Formation) was set to an
extremely low value to reduce the downward leakage and thus to reduce the flow in the
315
B4 aquifer, the model indicates unreasonably high amount of groundwater flows out from
the system into the east.
8.8.2.1.2 SIMULATED HYDRAULIC PROPERTIES
HYDRAULIC CONDUCTIVITY
Initial estimates of permeability for the model layers were based on geology,
structure, water level configuration and pumping test analysis. The estimates were made
to provide a starting point in the calibration process as permeability values were changed
during model calibration.
The final calibrated distribution of hydraulic conductivity in the models is shown in
Figure (8.13). Differences between the initial input and the final, calibrated hydraulic
conductivity are substantial. Many of the most significant differences result from
calibration improvements that were identified in one or more of the subregional models
and later incorporated in the data base of the regional model. The calibrated hydraulic
conductivity values are closely correlated with the depositional environment and the main
structural features in the study area.
Calibrated hydraulic conductivity values range from about 0.0021 mIh to about
1.9 mIh. The hydraulic conductivity pattern in the B2/ A 7 aquifer system reflects the
characteristics of the sedimentary sequence that deepens from west to east and grades
from mostly marine deposits in the north and west to sandy facies toward the southern
limits of the model area.
In Amman-Zerqa area the largest hydraulic conductivity values are in the central
areas of the syncline, along the Zerqa River, where the influences of alluvial deposition
are predominant and the accumulated thickness of sediments is the greatest. Outside the
Amman-Zerqa syncline the hydraulic conductivity values are moderate, but are generally
low in the eastern parts.
Generally the hydraulic conductivity varies from low in the far west where the
aquifer materials are much thinner to higher values in the central areas along the western
highlands of the Central Plateau. The largest hydraulic conductivity values in this area
result in part from the degree of karstification and in part from the effect of tectonics.
316
1160
1140
1120
1100
1080
1060
1040
1020
1000
980
960
• •• • •• • •• • •• • .~~~~! ~~ ••••• ;;; ..... i~ii:i ii:i:ii:iiii~iii: ==== = ••••• ••• • ••• •
• • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • I I I I I I I •• •••• • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • •
940 ••••
L-~ ____ ~ ____ ~--~~--~----~----~ 300 200 220 240 260 280
metres
0 10000 20ii 30000
V.di
Fault
1-¥h~1c c:onducIMly (nVh)
o 0 .000 to 0.001
• 0 .001 to 0.005 • 0.005 to 0 .010 • 0 .010 to 0.050 • 0.050 to 0.100
0.100 to 0.250
• 0.250 10 0.500 • 0 .500 to 0 .750 • 0 .750 10 1.000 • 1.000 10 2.000
Figure (8.13) Areal distribution of calibrated hydraulic conductivity of the B2IA7 aquifer system.
Further to the east, the hydraulic conductivity values are low. This can be attributed to
increase in percentage of marls and shale, and to the influence of depth of burial and,
therefore to the decrease in karstification.
In the Jafr area the hydraulic conductivity pattern reflects the depositional
environment, since the carbonate of the Ajlun Group grades easterly, in a downdip
direction, into sandy facies. Thus it becomes more uniform in pattern and intermediate in
magnitude. The B2/A7 and the Al-6 merge together to form one more-or-Iess continuous
conduit for groundwater flow.
Locally, in some areas, the reduced thickness and hydraulic conductivity combine
to cause reduced transmissivity, with a coincident steepening of hydraulic gradient, as
shown on the potentiometric surface map of the B2/A7 aquifer system (Figure 8.12).
During the calibration process, the hydraulic conductivity of the A4 aquifer
system was reduced substantially to obtain realistic simulated recharge and flow. The
final calibrated hydraulic conductivity ranges between 0.0021-0.33 mIh and averages
about 0.15 mIh. Higher values were found along the central area and downstream of
Amman-Zerqa syncline.
The calibrated hydraulic conductivity for the Rijam aquifer system in the Jafr area
ranges between less than 0.021 mIh in the western part and at the rim of the aquifer to
more than 1.7 mIh in the central areas decreasing eastwards to less than 0.083 mIh in the
east. The variations in the hydraulic conductivity reflect the karstic nature of the Rijam
limestone.
The hydraulic conductivity of the regional Al-6 aquitard was found to be in the
range of 0.0021 mIh. In the Jafr area, where the Al-6 is modified into an aquifer system,
the hydraulic conductivity ranges from 0.00042 mlh in the western part to 0.083 mIh in
the central area of the Jafr Basin and less than 0.0.0042 mIh in the east.
The deep sandstone aquifer system in the regional groundwater flow model was
given a uniform hydraulic conductivity value of 0.083 mIh which is found adequate to
simulate the general distribution of groundwater levels and flow.
318
LEAKANCE
The vertical resistance to groundwater flow was simulated in the model with a
leakance term: leakance is defined as the vertical conductivity divided by the length of
the path ( Lohman, 1972). No attempt was made initially to distinguish leakance value
according to hydrogeologic conditions because of the uncertainties in estimating the
vertical conductivity and the length of the flow path. Vertical conductivity was changed
during the model calibration.
Leakance data are provided to the model of the carbonate aquifer system to
simulate conditions of vertical groundwater flow across each of the B3, AS/6 and Al-6,
simulated confining units or aquitards. Although good estimates of confining unit
thicknesses were available from the delineation of the hydrogeological frame work,
virtually no quantitative information was available for the vertical hydraulic conductivity
of the confining materials within the study area. Consequently, the initial input of
leakance was based on the mapped confining unit thicknesses and an assumed vertical
hydraulic conductivity of 2.1 x 10-7 mIh, which was also suggested by a list of clay
hydraulic conductivities compiled by Bredehoeft and Hanshaw (1968). Although the
initial estimates of leakance were appropriate at the beginning of model simulation, they
had to be calibrated through trial-and-error simulation to provide a reasonably accurate
distribution of vertical leakage, water levels, and head gradients in the model.
The results of calibration indicate that the vertical hydraulic conductivity values
generally are smaller for model confining layer B3, intermediate for the AS/6, and higher
for the entire Al-6 aquitard.
The vertical hydraulic conductivity of the B3 confining unit is about 4.2 x 10-8
mIh in the Jafr area where the thick sediments of the confining unit are essentially
impermeable owing to their high marl, shale, and clay content. The AS/6 confining unit in
Amman-Zerqa area, lying between the B2/A7 and A4 aquifer systems, is considered
leaky. To achieve a good agreement between the simulated and initial water level
distribution for both the A4 and B2/ A 7 aquifer systems, it was essential to allow a
vertical leakage of about 2.16 MCMla from the A4 to the B2/A7 via the AS/6 confining
unit by increasing the vertical hydraulic conductivity values of the AS/6 confining unit
319
up to 10 times the initial input. The final calibrated vertical hydraulic conductivity values
of the A5/6 range between 4.2 x 10-7 and 2.1 x 10-6 mIh. The largest values of vertical
hydraulic conductivity are along the Amman-Zerqa syncline, where it is affected by the
geological structure along the syncline, and in the shallow downdip area to the west of
Zerqa.
The regional groundwater flow model simulation shows that, among other aquifer
parameters, the regional vertical leakage through the confining layers is the most
important factor affecting the regional groundwater flow system. For example, vertical
hydraulic conductivity of the AI-6 aquitard controls the amount of water flowing
downward to the deep sandstone aquifer system, and thus is the main control on water
level configuration and groundwater budget for both the B2/ A 7 and the deep sandstone
aquifer systems.
The deep sandstone aquifer receives recharge mainly from the infiltrated water via
the AI-6 aquitard, of which about 55 MCMla discharges as spring flow along the western
boundary to the Jordan Rift Valley (WAJ, 1985). To maintain this discharge and the
water levels of the sandstone aquifer, it was essential to increase the vertical hydraulic
conductivity for the AI-6 aquitard. The model distribution of the vertical hydraulic
conductivity of the AI-6 aquitard shows that only two values (1.3 x 10-6 and 1.7 x 10-5
m/h) are assign~d to the nodes. The larger values are assigned to the areas where the AI-6
is affected by tectonic features, along the major fault lines, and to the areas where the Al-
6 contains a large fraction of permeable sandstone and limestone.
In areas where confining units are absent, vertical flow rates are highest, and the
vertical hydraulic conductivity of the aquifers controls the vertical flow. An example of
this type of situation occurs between the B2/A7 and the AI-6 aquifer systems in the Jafr
Basin, where vertical flows are as much as 3.5 MCMla.
The vertical flow rate normally was small, however, the area of aquifer was very
large, therefore, the amount of water moving vertically was very large and also very
important in maintaining the regional flow system.
320
STREAMBED CONDUCTANCE
The only interaction between groundwater regime and surface water occurs along
the Zerqa River in Amman-Zerqa groundwater basin. The river deposits of sand and
gravel mostly overlie the B2/ A 7 aquifer system, and to a lesser extent the A4 aquifer
system (where it outcrops between Zerqa and Sukhna areas). Under steady state
conditions, the water table altitude in each grid block was simulated as lower than the
altitude of the stream, therefore the stream was acting as recharge points for the
groundwater flow system.
MODFLOW employs Darcy's law in the vertical direction to simulate leakage
through the reach of a streambed (Mac Donald and Harbaugh, 1984). The relation
between the simulated leakage and the associated field conditions for each grid block of
the model can be expressed as:
where
QRIV = KLW(HRIV - HAQ)/ M ................................................................. (8.9)
QRIV = leakage through streambed (m3/h)
K = vertical hydraulic conductivity of the streambed (m/h)
L = length of s1!eambed (m)
W = width of streambed (m)
HRIV = head on the river side ofthe streambed (m)
HAQ= head on the aquifer side of the streambed (m)
M = thickness ofthe streambed (m)
The effects of K, L, W, and M were considered in combination and incorporated into
a single conductance parameter ( CRIV), where
CRIV = (KLW)/ M ...................................................................................... (8.10)
321
The difference between the head in the stream and the head in the aquifer
determines the hydraulic gradient across the streambed thickness and the direction of
flow between a given stream reach and the adjacent aquifer. A positive head gradient (the
case in this study) provides leakage into the aquifer from the stream; a negative gradient
results in aquifer discharge to the stream or base flow. The head in the aquifer is the head
simulated by the model for the aquifer included in a particular grid block. This simulated
aquifer head can vary during a model run, depending on the net effect of recharge into,
and discharge out of, that grid block from one time step to another. Water head in the
Zerqa River is maintained by the spring discharge, effluent, and flood flow, is variable
and ranges between less than one metre in winter to almost nil in summer: however water
flow through the year through the alluvium of the river course and for most of the year
the Zerqa River is not more than sand river. For modelling purposes the head in the river
is taking as the altitude ofthe land surface ofthe river course.
A uniform value of CRIV (m2/d) was applied for all stream grid blocks, and
calibrated through an iterative, trial-and-error process to reflect the net effect of stream
bed geometry and permeability on stream- aquifer leakage. The calibrated streambed
conductance was found to be very small, about 60 m2/d. This value thought to be
underestimated since it is based on large grid block size of minimum width of about 417
m· while the actual streambed width does not exceed 50 m. Therefore, the actual
streambed conductance should be as high as 10 times the calibrated value.
Results of the simulation indicate that under steady state conditions, the Zerqa
River recharges the B2/A7 aquifer system by about 2 MCMla, and the A4 aquifer system
by only 0.1 MCMla. For an average hydraulic head difference between the river and the
B2/ A 7 aquifer system of about 10m, the calculated streambed conductance will be about
548 m2/d. Assuming an average streambed thickness of about 1 m, streambed length in
contact with the B2/A7 of about 10,000 m and an average streambed width of about 50
m, the conductance data suggest that the vertical hydraulic conductivity value of the
simulated streambed average about 0.00005 mIh, which is thought to be reasonable for
the streambed materials of silt, sand, and gravel.
322
RECHARGE
Recharge to the model was simulated as a constant flux to the highest active cell
in each vertical column. Recharge was not simulated in model cells that corresponded to
valleys because either these cells are below the saturated zone of the aquifer or much of
that recharge does not infiltrate into the deep part of the aquifer system and discharge as
evapotranspiration or issues as a small springs.
The rates of input recharge were estimates derived from a gIven area as a
percentage of the estimated annual rainfall. The percentage increases linearly from 0% for
the annual rainfall zone below 200 mm to a maximum of 20% for annual rainfall zones of
more than 500 mm and were refined later by trial-and-error simulation.
Initially, the percentage was obtained by trial- and-error calculations, such that
estimates of recharge were balanced against estimates of groundwater discharge from
natural losses in different areas of the model. Consequently, the percentage of recharge
applied to each rainfall zone varies considerably between hydrographic areas.
The simulated total recharge for the B2/ A 7 aquifer system over the actively
simulated outcrop area in the Western Highlands averages about 120.76 MCMla of which
about 41.02 MCMla enters the aquifer system as lateral flow. The A4 aquifer system is
recharged along narrow strip of outcrop at rates averaging about 6.4 MCMla. The Al-6
aquifer system receives the least amount of recharge, averaging about IMCMla across
the outcrop area in the western highlands of the Jafr Basin. The major recharge to the
AI-6, averaging about 3.5 MCMla, occUrs as downward leakage from the B2/A7 aquifer
system. The simulated recharge to the B4 aquifer system in the low rainfall zone of the
Jafr Basin is estimated to be about 2 MCMla.
Simulated recharge to the deep sandstone aquifer system, as downward leakage
from the B2/A7 aquifer system, averages about 50.42 MCMla, which is less than 42 % of
the simulated total recharge to the regional B2/ A 7 aquifer system. The difference between
the two is equal to groundwater discharge from the B2/ A 7 aquifer system to relatively
shallow surface drainages and the downgradient subsurface outflow to the east. As
previously explained, the deep flow regime is the predominantly confined, less dynamic
part of the flow system, including all of the subcrop area and part of the outcrop area that
323
discharges to the surface along the Jordan Rift Valley and via subsurface seeps into the
Dead Sea.
From the groundwater used for irrigation, a substantial amount infiltrates
downward back to the groundwater system as irrigation return flow. The amount of return
flow is estimated by the VBB (1977), Howard Humphrey (1986), and JICA (1987-90) to
range between 20-30%. For the purpose of the model simulation, the recharge from
irrigation has been taken into account by reducing the reported abstraction by 25%.
Although the recharge and abstraction are not distributed in exactly the same manner, it is
believed that this approximation seems permissible for the model resolution of such a
regional study.
8.8.3 TRANSIENT STATE CALIBRATION
The transient model was calibrated to simulate the response of the aquifer system
to the withdrawal of groundwater through industrial, irrigation, and public supply wells.
Groundwater abstractions from the aquifer system and consequently the transient
simulations began at different times in different areas. Heavy abstraction started as early
as 1970 in the Amman-Zerqa area and in the Jafr Basin, while in the Wadi Mujib, year
1980 was considered for the beginning of transient simulations. Between about 1970 and
1990, the calibration period for the transient model, water levels in some areas have
declined by more than 20 m. The transient model was calibrated primarily against
hydrographs drawdown from water level measurements made since the early 1970's on
the premise that if the model could be calibrated to replicate long-term patterns of head
change, then it would inherently simulate the important changes in the distribution of
flow.
Transient simulations of the groundwater flow system in the different subregions
require an accurate description of the distribution of groundwater abstraction, both are ally
and in time. Inventories of groundwater use were sufficiently detailed to allow abstraction
from each aquifers in the subregions to be determined, but not to determine both the
spatial and temporal distributions of abstraction that are need for simulations.
Publications and files containing water-use data rarely specify the abstraction rate and
324
pumping periods for each well. Data concerning well abstraction mostly consists of the
total abstraction from a group of wells. This problem is significant, particularly in areas
where a large number of wells pump water into a main reservoir.
However, for the purposes of transient simulation, the total water abstractions
from each wellfield were approximately equally distributed between the wells over the
year. The total pumpage simulated in the transient models for each aquifer and time
period is shown in table (8.2). Figure (8.14) shows the major areas of groundwater
abstraction.
Groundwater development from the Rijam aquifer system, mainly for
irrigation, started in 1964. Prior to 1967 abstractions were just above 1 MCMla, and
reaches a maximum of about 2 MCMla in 1975. The present abstraction from 5 boreholes
around the town ofEI Jafr is about 1.14 MCMla. The average abstraction for the 20 years
modelled period was about 1.5 MCMla.
For modelling purposes, groundwater abstraction for irrigation in the Jafr area and
in parts of the Amman-Zerqa and Wadi Wala areas, were reduced by 25% to account for
irrigation return flow.
Table (8.2) Abstraction used in simulations, by aquifer, area, and time period.
325
150
IQastal1
IErneibal
Figure (8.14) The major areas of groundwater abstractions.
Simulation of the transient behaviour of the aquifer system required also
discretization of time and time dependent stresses, such as groundwater abstractions.
Because groundwater development commenced at different times with variable rates,
abstractions were discretized into pumping periods, whereas time was discretized with
time steps. Abstractions were changed abruptly at the start of each pumping period and
were held constant for the duration of the pumping period. For the subregional models,
various pumping periods of various duration were used and each pumping period was
divided into a number of equal time steps, each of 3 months duration. The total
simulation of 20 years was completed with 80 time steps. The number of time steps and
duration was constrained by the computer capacity: however, it is believed that the
number oftime steps and duration were reasonable for such a regional study.
Table (8.3) Stress periods and time steps used in the simulations (days).
327
Pumping of groundwater, particularly for municipal use, has locally caused water
level declines in most parts of the study area. Significant declines in water levels were
observed in Amman-Zerqa area as early as 1973: elsewhere, the water level decline
started after the onset of heavy pumping after 1984. Regionally the declines in water
levels are less significant as the flow system equilibrated in response to increased
groundwater abstractions. The observed well hydro graphs discussed in chapter seven
were used to compare with the simulated declines.
The preliminary steady state models provided the starting point for the simulation
of transient conditions. In the process of calibration to reproduce the observed well
hydro graphs, several changes were made in the hydrological parameters of the
preliminary model. Rerunning the steady state model with the final parameters calibrated
under transient conditions produced results similar to the initial predevelopment
conditions, and these parameters were accepted as the best calibration to represent both
steady state and transient conditions.
In addition to the hydrologic parameters necessary for steady state simulation,
calibration of the transient simulation model required the areal distribution of the value of
storage coefficient for each of the aquifers in the system. Only storage in aquifers was
considered in this model. Although storage in the thick marls of the confining units that
separate the aquifers is per~aps significant, no direct or indirect evidence is available to
quantify flow derived from confining unit storage. The assumption that all of the storage
in the system is from aquifer material may, therefore, result in some overestimation of the
aquifer storage coefficients. The distribution of storage coefficient values estimated in the
process of transient calibration will be discussed in a following section.
Comparisons of observed and simulated hydro graphs for the calibrated models are
shown in figure (8.15). General patterns of drawdown were approximated fairly well, but
in a few areas the simulated drawdown exceeded that suggested by observed data. This
may be due to poor matches for these areas in the steady state simulation, that is, the
simulated steady state heads that were used as starting heads for the transient simulation
were lower than the measured steady state heads. Because groundwater development
328
predated water level measurements in most of the areas, some of the measured heads used
for the steady state calibration may reflect prior stresses in those areas. Therefore, the
simulated steady state heads may have been closer to the true predevelopment heads than
the heads defined by the observed measurements. Thus the excessive drawdown
simulated by the transient model may in fact be the correct amount.
The simulated condition in the B2/ A 7 aquifer system relates most directly to the
proximity of pumping. The simulation indicates that hydraulic head declines were
affected significantly by pumping in localised places, in the centre of the wellfields. In
the other areas head changes in response to pumping were negligible.
Water table declines in the B2/ A 7 aquifer system have been substantial only in
Amman-Zerqa area where the abstraction was high. Apart from node No. 8-37 in the
Amman-Zerqa model which shows the greatest drawdown of 24 m, the rest of the area
experienced regional decline in water level ranging between 1-13 m. The largest cone of
depression have developed in up dip areas of the Amman-Zerqa syncline: this pattern
reflects the effect of storage coefficient values being generally larger in shallow downdip
areas, and the fact that captured base flow in the downdip areas more effectively
compensates for abstraction. However the modified cone of depression has diverted some
of the water toward pumping centres that would otherwise have discharged as base flow.
Field evidence indicates significant reduction in spring flows have occurred by
commencing the heavy abstraction from the basin.
In the Wadi Wala-Wadi Mujib-Wad Hasa basins, water level declines started at
the beginning of heavy abstraction from the different wellfields in the basin, after 1985.
The simulated decline in water levels varies between the wellfields, depending on the
abstraction rates and the aquifer properties: it is about 0.5 min Qastal, 3 m in Lajun, 11 m
in Qatrana, 5 m in Wadi Abiad, and about 2.5 m in the northern part of the Wadi Hasa.
In the Jafr Basin, the decline in water levels started as early as 1973, since the
beginning of the groundwater development from the aquifer system. Successive decreases
in the water levels have been monitored in borehole S121 since 1973. The simulated
drawdown for the model period is about 9 m which is consistent with the observed
drawdown of about 8.3 m for the period between 1973 and 1988.
329
(j) Cl .0
E. ..J
~
(j) Cl .0
E. ..J
~
15 20
25 30 35
40 45
70 71
170
175
180
185
Awajan observation well
72 73 74 75 76 77 78 79 80 81 82 83 84 85 86 87 88 89 90
Year
Qastal No.7
80 81 82 83 84 85 86 87 88 89 90 Year
160
(j) 170
\~ Cl .0 Arainba (ER4) E. 180 ..J
~ 190
200 80 81 82 83 84 85 86 87 88 89 90
Year
95
(j) 100 Cl .0
E. 105 ..J
~ 110
115 80 81 82 83 84 85 86 87 88 89 90
Year
--Observed -Sirrulated
Figure (8.15) Observed and simulated drawdowns in observation wells.
330
20r----------------------------------------------.
Ii) Cl .c 5 25
Wala No.14
~+-----~----~----~----~----~----_r----_r----~ 80
118
........ 119 I/) Cl 120 .c E 121 ....... ...i 122 s:
123 124
80
81 83 84 85 Year
87
Lajun No.2
----- --A....r
-81 82 83 84 85 86 87 88
Year
88 89 91
89 90
9Or----------------------------------------------.
Ii) 95 Cl .c 5100 ...i
s: 105
Qatrana No. 512
110+-~_r_+~--+_~_r_+~--r_+_~_r_+~--+_~_r_+_4
nn~N~mn~~ro~~~84MOO~MOO90
Year
55r-----------------------------------------------~
Ii) 60 Cl .c .s 65 s: 70
Hasa No. 5121
72 73 74 75 76 76 77 78 79 80 81 82 83 84 85 86 87 88 89 90
Year --- observed -sirruiated
Figure (8.15) Continued.
331
The average abstraction from the B4 aquifer system of about 1.125 MCMla, after
the irrigation return flow has been accounted for, caused a water level decline ranging
between 2 and 7 m during the model period which corresponds to water level decline
rates ranging between 0.11 and 0.40 mla.
In the A4 aquifer system water levels have declined steadily since pumping began
in the 1960s. In some areas, the water levels have declined by more than 150 m and thus
the distribution of groundwater flow has changed significantly.
For the Al-6 aquifer system in the Jafr area, transient simulations were not made
because no long-term change in water heads or flow has been documented.
The accuracy of the transient simulation is limited by the accuracy of the
estimated aquifer characteristics and abstraction rates. This model has been tested against
a period of stresses during which abstraction was low. However, inflows and outflows
that were not accounted for in the simulation may be occurring, therefore, greater future
stresses may cause unanticipated responses by the aquifer. If any of these effects becomes
evident in the future, the model should be recalibrated to include them. Nevertheless, the
overall results of the model indicate that most of the flow adjustments in response to
pumping during the model period were adequately simulated with respect to the effect on
the regional flow regime.
8.8.3.1 STORAGE COEFFICIENT
Storage coefficient, as used herein, is taken to mean either the storage coefficient
of a confined aquifer or the specific yield of an unconfined aquifer. Prior to construction
of the computer model, the storage coefficient had been estimated from pumping tests for
fewer than 20 sites in the aquifer system. Most of the individual storage coefficient data
were provided through use of the Theis (1935) type-curve or the Cooper-Jacob (Cooper
and Jacob, 1946) straight-line approximation for nonleaky confined aquifers and are
listed in various reports. Using data from previous pumping tests, storage coefficient
values were obtained by using the method of Remson and Lang (1955) as modified by
Ramsahoye and Lang (1961) (Chapter five).
332
Owing to the limited number of the values of storage coefficient with which to
estimate the model input data, average values were applied uniformly throughout each
model layer at the beginning of the calibration process. The initial input values were 0.1
and 0.0001 for the unconfined and confined parts of the aquifer systems, respectively.
The model input values of storage coefficient were adjusted by trial and error
during the latter stages of calibration, with the purpose of simulating measured
hydro graph trends (Figure 8.15). The degree of adjustment was dictated by the apparent
improvement in the calibration of the model and the sensitivity of the model output to
changes in the input data. Although the simulated and observed hydro graph data differ in
some cases, most of the discrepancies cannot be corrected through adjustment of the
storage coefficient data alone. The model is not very sensitive to changes within one
order of magnitude of the original input values of storage coefficient, rather than to local
changes in the storage coefficient values within the wellfield. The calibrated values of
storage coefficient for the unconfined part of the B2/ A 7 aquifer system, range from
0.015-0.07 and average about 0.03, and between 0.00005-0.001 and average about
0.0001 for the confined part. The calibrated storage coefficient values average about
0.025 for the B4 aquifer system, 0.0005 and 0.05 for the confined and unconfined parts of
the A4 aquifer system respectively.
Although reducing the storage coefficients to values below those initially thought
improved the simulated hydro graphs in some areas, it is believed that small values of
storage coefficient, however, are not compatible with those calculated from aquifer test
data, nor are they considered appropriate to represent regional aspect of the aquifer
system hydrogeology. It is believed that the largest hydro graph mismatches evident in
figure (8.15) are due to errors inherited from the pumpage data and the relatively coarse
model grid, rather than to input values of storage coefficient.
Nodes with largest values of storage coefficient are found in the western part of
the study area, along the Western Highlands, where the aquifer is unconfined. To the east,
the aquifer gradually becomes confined, and therefore the storage coefficient is smaller.
In the southern desert, the groundwater regime is for the most part confined: however, in
the shallow, up dip areas of the outcrops, even in the case of the Al-6 aquifer system, the
333
aquifers may be unconfined. It is believed that the largest values of storage coefficient are
up dip in the shallowest parts of the aquifer system, where they reflect semiconfined
conditions. Moderate storage coefficients are expected in the central and downdip areas
where intermixing of sands with marine silts and clay occur.
Generally, the small values of storage coefficient in the eastern parts of the
modelled area inhibit the capacity of the subcropping parts of the aquifer system to
release stored water. The larger storage coefficient values associated with the updip,
unconfined and semiconfined aquifers make them more efficient in terms of water supply
and development, compared with the completely confined downdip aquifer.
8.8.4 REGIONAL GROUNDWATER BUDGET
The regional groundwater budget determined by the model simulations is
discussed here in conjunction with the estimates of recharge to the water table and
discharge from the aquifer system due to natural losses. The simulated water budget for
the aquifer systems under steady state conditions is shown in Figure (8.16). The water
budget shows the relative importance of vertical and lateral aspects of inflow and
outflow. The relation between the simulated regional flow water budget and the water
budget from estimates of recharge to the water table and surface water and groundwater
interaction for each subregional model areas is shown in Table (8.4).
Analysis of the cal~brated model indicates that before development (pre.1970),
about 120.67 MCMla of water flowed through the B2/A7 aquifer system. More than
79.74 MCMla ofthis amount represented recharge from the aquifer outcrop areas, and the
. remainder, less than 41.02 MCM/a represented water flowing from vertically and laterally
adjacent aquifers and streams. In the steady state condition, nearly 44.95 MCM/a of the
water flowing through the aquifer system discharge to the regional drains in the major
river valleys, about 49.62 MCMla discharged to the deep sandstone aquifer system, about
3.5 MCMla discharged into the A1-6 aquifer system in the Jafr Basin, and about 22.69
MCMla left the study area laterally to the east. The substantial amount of the water flows
in the A1-6 aquifer system discharged into the east (about 2.7 MCMla) and the remaining
334
0.8 MCMla infiltrated downward increasing the total vertical leakage into the deep
sandstone aquifer system to 50.42 MCMla.
Because the deep sandstone aquifer system outcrops only in very limited areas,
downward leakage from the B2/ A 7 aquifer system is the sole source of recharge to this
aquifer. However, the simulated water inflow from the eastern boundary was about 65.8
MCMla. This gives total flow in the deep sandstone aquifer system in the study area of
about 116.22 MCMla, which has to flow out of the system along the western boundary as
spring discharges and subsurface outflow into the Dead Sea. According to the W AJ
spring discharges data (1985), approximately 55 MCMla of water was discharged from
~~~
Spring
Subsurface
the Dead Sea
Al-6 aquitard
'''''' discharges (44.95)
Lateral outflow to
the east (22.69)
~~~~
Lateral outflow to
the east (2.7)
~ ~~~~
Lateral inflow from
the east (65.8)
¢::l¢::l¢::l¢::l¢::l
Figure(S.16) Simulated predevelopment water budget for the regional aquifer
system (MCMla).
335
the sandstone aquifer as base flow. This leaves about 61.22 MCMla to discharge into the
Dead Sea as subsurface outflow. The total discharge from the sandstone aquifer system
was combined with the total discharge from the B2/ A 7 aquifer (without the total
downward leakage into the deep sandstone aquifer system) to account for the total
groundwater discharge of about 186.56 MCMla, which is assumed to be equal to the total
areal recharge for all the aquifer systems under steady state conditions.
The regional water budget indicates that the system rapidly approaches a steady
state condition because the amount of water removed from storage was negligible
compared with the amount of water withdrawn. In 1990 the average pumping rate from
the regional aquifer was about 80.5 MCMla. This pumping rate was balanced by an
increase in total recharge, a decrease in discharge to river valleys, a decrease in storage,
and a decrease in downward leakage and water flowing out of the system outside the
study area. The head declines however, are localised in the pumping centres. Only in the
Amman-Zerqa area does the water balance analysis sho~ a regional decrease in storage as
a result of heavy abstraction in the area.
The amount of water stored in the B2/ A 7 aquifer system in the Amman-Zerqa
area is calculated to be about 780 MCMla. The accumulated abstraction during the
modelled period mounted to 549 MCM against 407 MCM of recharge, leaving a deficit in
the groundwater balance of about 142 MCM (18 % of the storage). Of this value, as the
water balance indicates, 54 % was derived from decreased natural discharge from the
aquifer, about 10 % was derived from increased recharge, and about 32 % was being
supplied from the aquifer storage. The remaining 4 % includes all other factors, including
interchange with adjacent aquifers and lateral flow moving out of the study area.
The base flow from the aquifer system is simulated to have decreased from about
18.6 MCMla prior to 1970 to about 13.0 MCMla in 1990, which represents a reduction in
the discharge of base flow totalling about 5.6 MCMla. This decrease in base flow
indicates a reduction in hydraulic gradients between the regional aquifers and major
springs, owing to water level decline caused by pumpage. The simulated reduction of
336
Outflow
Outflow
Outflow
12.39 Outflow 1.30
7.59 3.50 12.39
* These values are spring discharge measurements.
Table (8.4) Simulated steady-state groundwater flow budget.
337
base flow appears to be reasonable, however, considering the observed decreases in base
flow and head decline in wells adjacent to the major springs.
The simulated loss of water from storage in the Anunan-Zerqa area was 54 MeM.
A change in storage in the model, as well as in the aquifer, results directly from head
change. Assuming the system underlies about 350 km2 and a storage coefficient of 0.03,
the change in storage during the transient period indicates that during the transient period
heads decline an average of about 5 m over the entire area. Although the hydro graphs
(Figure 8.15) suggest that within the major pumping centre declines average between
1-13 m, the average decline over the entire system was probably about 5 m. In other
words, the estimated loss of groundwater in storage appears to be consistent with the
observed conditions during the transient period.
The steady state budget of the Hummar (A4) aquifer system in the Amman-Zerqa
area indicates a total recharge of about 6.4 MCMla, of which about 2.16 MCMla was
simulated to leak upward into the B21 A 7 aquifer system. The simulated spring discharges
from the system were about 7.7 MCM/a, which indicates a water deficit of about 3.5
MCM/a. However, the model provides that amount from the lateral flow from the western
part of the model area: it is believed that a substantial part of the 7.2 MCMla Zerqa River
overflow from the B2/A7 aquifer system in the northern part of the Amman-Zerqa area,
where the A4 aquifer crops ~ut, infiltrates downward into the A4 aquifer system.
The amount of drainable water in the B21 A 7 aquifer system is a function of
specific yield or storage coefficient and the thickness of the aquifer. The estimated
specific yield for the B21 A 7 aquifer system is 0.03. Estimates of drainable water stored in
the B2/A7 aquifer system are about 780, 9900, 17100, and 14400 MCM in the Amman
Zerqa area, the Wadi Wala Basin, the Wadi Mujib and Wadi Hasa basins, and the Jafr
Basin respectively.
8.8.5 SENSITIVITY ANALYSIS
In order to assess the importance of the variation and uncertainty associated with
the definition of parameters used in the model, a sensitivity analysis was conducted on
338
those parameters that were routinely adjusted during the calibration process. The
sensitivity analysis identified the parameters that were most important in controlling
groundwater flow and assessed the reliability of the model by demonstrating the effect of
a given range of uncertainty in a hydraulic factor on the simulated heads and flow in the
groundwater flow system.
The sensitivity of the regional steady state model was made on recharge, hydraulic
conductivity, and confining unit leakance. Locally, the model was also tested with respect
to changes in the location of the no-flow boundaries, extension of groundwater barriers,
and to variation in the altitude of constant head boundaries. A single data set, storage
coefficient, was tested in the transient model. Sensitivity was measured by varying the
model input parameters through factors (0.1, 0.5, 2.0, and 10) both greater and less than
the calibrated value of each parameter and observing the resultant change in simulated
hydraulic head and flow. Each parameter was tested independently of the others. The
results of the sensitivity analysis are given in figures (8.17), (8.18), and (8.19), which
show the effects of variation of the model inputs as changes in the mean absolute
residuals for the B2/ A 7 aquifer system ( the mean absolute values of the difference
between simulated water levels from the sensitivity and steady state calibration
simulations).
8.8.5.1 HYDRAULIC CONDUCTIVITY
Simulated heads of the B2/ A 7 aquifer system were tested for sensitivity to
hydraulic conductivity values that were 0.1, 0.5, 2, and 10 times as large as the values
used in the steady state calibration. With reduced hydraulic conductivity of the B2/ A 7
aquifer, simulated heads were up to 100 m higher than the calibrated heads. Increasing
the hydraulic conductivity resulted in a decrease in simulated water levels of about 25 m
only (Figure 8.17). This leads to the conclusion that the water level is more sensitive to
decrease than increase in hydraulic conductivity.
In contrast, when the hydraulic conductivity was varied uniformly for all the
aquifer system units, the relationship between water levels and hydraulic conductivity
was found to be linear, with the water levels increasing with increasing hydraulic
339
conductivity. This effect must be arise through the combined effects of the hydraulic
conductivity changes for the different layers. However, independent sensitivity analysis
shows that the water levels of the B2/ A 7 aquifer system were insensitive to changes in
hydraulic conductivity of the Al-6 aquitard ( ± 2 m) and the deep sandstone aquifer
system( ± 0.5 m). Therefore, the only explanation of the combined effects of uniform
150r------------------------------------------------.
co -5 100 ·iii ~ Q) -::J (5 en ~ 50 -------- --------------~-------------------------------co c:: co Q)
~
O~----------~--------~~~---------+----------~ 0.1 0.5 Calibrated value 2
Multiple of hydraulic conducti\Aty for the B2IA7 aquifer
Figure (8.17) Sensitivity ofthe B2/A7 aquifer system to changes in hydraulic conductivity.
10
changes in all aquifer units on water levels is that the decrease in water levels is a
consequence of decrease in recharge as a result of decreasing hydraulic conductivity.
The sensitivity analysis shows that the model is more affected by changes in
updip hydraulic conductivity values, in the western highland outcrop areas, than by
changes in the downdip hydraulic conductivity values.
In the Jafr area, the water levels of the B2/A7 aquifer system is sensitive to
changes in the Al-6 hydraulic conductivity at higher vertical hydraulic conductivity, were
the Al-6 aquifer itself becomes sensitive to hydraulic conductivity.
The piezometric levels of the A4 aquifer system in the Amman-Zerqa area are
more sensitive to changes in vertical hydraulic conductivity of the A5/6 confining unit
than the hydraulic conductivity and least sensitive to recharge.
340
The deep sandstone aquifer system is relatively insensitive for any change in
hydraulic parameters, water levels vary by around ± 12 m, ± 5 m, and ± 1 m against
changes in hydraulic conductivity, Al-6 vertical hydraulic conductivity, and B2/A7
aquifer system recharge respectively.
8.8.5.2 VERTICAL HYDRAULIC CONDUCTIVITY
Simulations were made to evaluate the sensitivity of the aquifer system to
confining unit leakance. In the simulations the vertical hydraulic conductivity of the Al-6
confining unitlaquitard was changed. A decrease in the Al-6 vertical hydraulic
conductivity generally increased simulated water levels in the overlying B2/ A 7 aquifer
system by up to 200 m with a mean residual of 15 m (Figure 8.18), and lowered simulated
heads in the underlying sandstone aquifer relative to the simulated heads from the
calibrated model. Similarly, increasing the vertical hydraulic conductivity lowered heads
in the B2/ A 7 and increased them in the sandstone aquifer system. The simulations show
that increases in the vertical hydraulic conductivity by 10 times decrease the water levels
by up to 100 m causing desaturation in the aquifer system in many areas, the calculated
mean absolute residual being 45 m. The model system is insensitive to changes in the
vertical hydraulic conductivity of the B2/ A 7 aquifer system: increasing the vertical
hydraulic conductivity by 10 times decreases the water levels by only 2 m.
This general trend ~as only provoked in the Amman-Zerqa area, where the water
levels of the A4 aquifer systems are higher than those for the overlying B2/A7 aquifer
system, and therefore increasing the vertical hydraulic conductivity of the intervening
A5/6 confining unit increase the water levels in the B2/ A 7 aquifer system and decrease
them in the A4 aquifer system.
The water levels of the sandstone aquifer are less sensitive to changes in the Al-6
vertical hydraulic conductivity (± 5 m) than the B2/A7 aquifer system.
Sensitivity analyses were also made to simulate the effects of changes in the
vertical hydraulic conductivity on the recharge to the B2/ A 7 aquifer system. The
simulation shows that decreasing the vertical hydraulic conductivity by a factor of 0.1
341
increases the recharge very slightly (1 x 104 m3/d), while increasing the vertical hydraulic
conductivity by a factor of 10 decreases the recharge to almost half the calibrated value as
many cells went dry.
The model was particularly sensitive to confining unit leakance for several major
reasons. First, due to the uncertainty associated with the estimation of the vertical
hydraulic conductivity, the range of values tested (0.1-10), was considerably larger than
the uncertainty associated with the other better known aquifer parameters. Second, the
leakance affects heads in two aquifers. Although changes in the hydraulic conductivity
values in an aquifer have some effect on adjacent aquifers, the effect was significantly
smaller than the effect of leakance changes for the range of values tested in this
sensitivity analysis. Third, the vertical flows were generally greater than horizontal flows,
and changes in vertical hydraulic conductivity, a control on vertical flow, had a greater
effect on simulated heads than did changes in hydraulic conductivity, a control on
horizontal flow.
-,S 'iii ::l "0 'iii ~ Q) -::l "0 II) ..c ro c: ro Q)
::E
2i
0.1 0.5 Calibrated value 2
Multiple vertical hydraulic conductivity for the A 1-6 aquitard
Figure (8.18) Sensitivity of the B2/A7 aquifer system to changes in vertical hydraulic conductivity of the A1·6
aquitard.
342
10
8.8.5.3 RECHARGE
The model was tested for sensitivity to changes in the recharge values that were
adjusted during calibration. The distribution of simulated hydraulic head changes
associated with a change in value of recharge is significant. Figure (8.19) shows that the
hydraulic head is more sensitive to increases in recharge than decreases.
g (ij ::J
"C ·iii ~ 2 ::J (5 CI) .D C1l c C1l Q)
:2
250
200
150
100
50
0 0.1 0.5 Calibrated value 2
Multiple recharge
Figure (8.19) Sensitivity ofthe B2/A7 aquifer system to changes in recharge.
10
Comparison between the input and calculated values of recharge shows that the
calculated value is insensitive to increases in recharge input while it decreases
approximately in the same ratio by decreasing the recharge input. Magnitudes of
hydraulic head change increase to the west where a decrease in the recharge caused
de saturation for large portion of the aquifer system along the western highland outcrop
areas.
This distribution reflects a sensitive balance between the small hydraulic
conductivity along the Western Highlands and recharge rates at outcrop area. The
magnitude of the computed hydraulic conductivity and leakance values were dependent
on the amount of recharge used in the simulations, because steady state condition were
343
assumed in the simulations. Increasing recharge results in a corresponding increase in
discharge and a proportional increase in hydraulic conductivity and leakance values.
8.8.5.4 STORAGE COEFFICIENT
The estimates of storage coefficient (S) are only approximations, thus during
model calibration, storage coefficient was increased by a factor of 2 and decreased by a
factor of 2 to evaluate the effect of storage coefficient on computed drawdown values.
Resulting differences in simulated hydraulic head declines were not especially
significant (Table 8.5).This can be attributed to the overall small values of hydraulic head
decline, since the dynamic state started only recently after the start of heavy abstraction.
Use of either the modified or original (calibrated) value of storage coefficient results in
simulated hydraulic head declines that are comparable to inferred regional decline. The
calibration value however, results in slightly better agreement in overall head decline
distribution.
Although the Amman-Zerqa area has experienced the greatest abstraction and
consequently the maximum drawdown, the simulated hydraulic head declines were found
to be insensitive to changes in storage coefficient. The areas that showed greater
sensitivity to the changes in storage coefficient typically had abstractions that were
continuing to increase till the end of the simulation. In contrast, the abstraction in the
Amman-Zerqa area remai~ed relatively constant during the period 1976 to 1990;
therefore the system has reached a steady state condition for both values of storage
coefficient (calibration and modified). Another reasons which might affect the sensitivity
of the hydraulic head decline to changes in storage coefficient includes:
a- the storage coefficient values used are the maximum,
b- the declines in water levels are tempered by induced recharge provided by the
upward vertical leakage from the A4 aquifer system, or downward leakage from
the Zerqa River,
c- the hydraulic conductivity used in the simulation is extremely high, which means
that the groundwater flow is great and thus quicker tapping of the aquifer during
344
and shortly after pump mg. Generally reducing the storage coefficient and
increasing the hydraulic conductivity has about the same effect.
Table (8.5) shows that in the Amman-Zerqa area, the water level declines in the
A4 confined aquifer system are very sensitive to changes in storage coefficient.
Observation well ·0.5 X 8 8 (calibrated value) 2x 8 Wala No. 14 (B2/A7) 4 3 2.5 8W7 (B2/A7) 6.9 5 3 8124 (B2/A7) 19 13 10 8121 (B2/A7) 10.2 8 6.7 9-36 (A4) 120 85 45 47-12 (A4) 28 20 13
Table (8.5) Maximum drawdown (m) for selected observation wells due to different
storage coefficients at the end of transient simulations.
8.8.5.5 SUMMARY AND DISCUSSION
The steady state model of predevelopment conditions is most sensitive to
increases in the rates of recharge in the outcrop area. After recharge the model is most
sensitive to decrease in hydraulic conductivity, especially in the outcrop area, and is more
affected by changes in updip hydraulic conductivity values than by changes in downdip
hydraulic conductivity values.
The sensitivity analysis shows that model calibration was least affected by large
changes applied to all aquifer system layers compared with the changes in individual
layer hydraulic parameters. The model is insensitive to the variation in hydraulic
conductivity of the Al-6 and sandstone units, and less sensitive to the changes in the
vertical hydraulic conductivity of all the aquifer system layers than to changes in the
vertical hydraulic conductivity applied to the Al-6 aquitard.
Simulated discharge to streams is most sensitive to hydraulic conductivity and
much less sensitive to other parameters including the hydraulic conductance of the
dependent flow boundaries used to simulate the spring discharges.
345
The model is relatively insensitive to changes in the location of downdip no-flow
boundaries and to moderate changes ( ± 10m) in the altitude of constant head boundaries.
But it is very sensitive to the extension and the hydraulic conductivity of the internal
groundwater barriers. The transient model is less sensitive to increases in storage
coefficient than to decreases.
The sensitivity of the model to the different hydraulic parameters is generally less
where the hydraulic conductivity is small, which it is in most downdip parts of the aquifer
system.
The results of the sensitivity analysis show that the calibrated values of the model
input are, for the most part, consistent and within the range of reasonable possibilities.
The simulated response to departures from the calibrated input suggests that the capacity
of the model to simulate field conditions deteriorates as the departure increases.
Although a sensitivity analysis such as that performed in this study provides a
general idea of the model sensitivity to changes in model inputs (aquifer properties), it
cannot demonstrate the effects of interaction between these properties, nor can it show the
relative differences in degree of head change among different areas of the model.
However, a comprehensive analysis that would demonstrate model sensitivity to areal
variation of aquifer properties or their interaction would be impractical for a model of this
size and complexity.
8.9 MODEL RELIABILITY
The reliability of model results depends on the accuracy of the model data used to
describe the hydraulic characteristics, the distribution of data used for calibration, and the
degree to which the model design represents the physical system. Because model
reliability is not easily quantified, it is included in the previous descriptive discussion of
the model sensitivity and the possible sources of error in estimating the hydraulic
characteristics. Model results are also dependent on the general assumptions and
limitations discussed previously in section (8.3).
346
The discussion of the model sensitivity analysis evaluates the model's sensitivity
to both hydraulic characteristics and model design. Several potential sources of error may
affect the model reliability. These sources are either measurement errors such as in layer
thicknesses, water table elevations, and discharge, or estimation errors such as in the
estimates of aquifer hydraulic properties and recharge.
The magnitude of errors associated with water level measurements which is
generally on the order of tenth of a metre, is usually insignificant regionally. The
estimation of land surface altitudes may be a major source of error in the conversion of
measured depth to water head. Although the error from altitude estimates was highly
location dependent, errors of several metres probably were common. This could affect the
simulated hydraulic properties of the aquifer system by at most an order of magnitude and
the model results might still be reasonable.
Most of the regional hydraulic characteristics used in the model were estimated
from point data which, because of poor areal distribution, may not represent the regional
hydrogeologic system. For example, most wells were drilled in the most productive parts
of an aquifer; therefore, point data represent this bias. Another significant source of error
might be the vertical hydraulic conductivity values for the confining units for which no
field data are available.
Although the recharge values used in the model were bound by reasonable limits
depending on the amount .of annual rainfall, the uncertainties in the distribution and
amount of recharge remains the most significant source of error in the model's
simulations. Hydraulic conductivity values computed for the model layers are in part
dependent on the amount and distribution of recharge used in the model, particularly for
model cells that correspond to the Western Highlands. The assumption that most of the
recharge occurs in the Western Highlands, which consist of carbonate rocks, is probably
reasonable because little surface water flows to the nearby valleys. However, in areas that
consist of low permeability rocks, much of the water flows into the nearby valleys where
recharge occurs mostly in the adjacent alluvial fans.
The errors in the estimates of recharge are unknown but could be well in excess of
100 %. If recharge is increased in the model by 100 %, a similar distribution of head
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could be simulated by proportional increases in lateral and vertical hydraulic conductivity
values. Because the model assumes steady state conditions, the simulated discharge
would also increase by the same percentage. A different distribution of hydraulic
conductivity near springs would be simulated to force the groundwater flow to equal the
estimated discharge from the springs. Furthermore, because of the large size of the model
cells in the regional model, in some areas the water recharging the aquifer system
discharges to sinks within the same cell receiving the recharge and cannot be simulated
by the model.
In the transient simulation, the storage coefficient, because it was derived from
aquifer test data, had an estimated range of uncertainty similar to that for hydraulic
conductivity. The lack oflong-term regional hydraulic head declines in the aquifer system
preclude accurate simulation of storage coefficient. Abstractions are a time-dependent
variable; therefore, the uncertainty is also time dependent. The greatest uncertainty was
for the earlier abstractions. The uncertainty of abstraction was high for individual wells,
low for the total abstraction for all wells for a specific time period.
The effect of model design on model reliability is also discussed in terms of the
scale dependence of the results. Martin and Leahy (1983) discussed the impact of areal
discretization scale on model results. The methodology of interfacing regional and
subregional models was discussed by Martin (1987). The larger size mesh uses areally
averaged values for hydrau.lic characteristics, and thus the larger regional model blocks
average local variations in potentiometric surface and in recharge or discharge.
Comparisons of heads and flows simulated at regional and subregional scales by a model
having equivalent hydraulic parameters has demonstrated the effect of discretization scale
on model results. The potentiometric surface simulated by the regional model showed the
same general configuration as the surface simulated by the subregional model. The
regional simulated surface lacks the resolution or detail of the subregional simulated
surface, because heads were averaged over a larger cell area in the regional model. Thus,
local flow features tend to be lost in the regional model. An increase in model resolution
provided by the finer mesh of the subregional models provided a more accurate
calibration than did the regional model.
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Estimates of the uncertainty in the initial estimates of hydraulic characteristics are
subjective. These can only be based on hydrologic experience, judgement, and knowledge
of the method used to develop the hydrologic data base. The data used in this study are
believed to be acceptable and within the range allowed by the climatic and geologic
conditions prevailing in the study area. However the model results, particularly the
transient simulations, should be considered conceptual.
8.10 DISCUSSION
Groundwater flow in the study area was conceptualised as relatively shallow
flow primarily through the carbonate sediments of the Mountain Ranges and the Central
Plateau superimposed over deeper flow through primarily sandstone sediments. A three
dimensional groundwater flow model was used to simulate this concept of groundwater
flow in the area. The area was subdivided into 5 subregions, where each subregion was
modelled separately and then compiled in one regional model. Six model layers were
used to simulate relatively shallow and deep flow. The upper five layers were used to
simulate the flow in the carbonate aquifer system. The lowest model layer was used to
simulate the concept of deep flow in the sandstone aquifer system.
Regional flow in the carbonate aquifer system is controlled by the
hydrogeological properties of the aquifer materials, the topography within their outcrop
areas, and the geological structures. Topography within the outcrop area determines the
locations of major natural recharge and discharge areas, whereas the transmissivity
primarily determines the total amount of flow through the aquifer, the storage coefficient
determines the amount of water stored in the aquifers, and the geological structures affect
the rate and direction of the groundwater flow.
The model was calibrated using the idea of developing the simplest model that
could account for the principal features of flow in the aquifer system. A steady state
model of the principal aquifer system in the study area was constructed to simulate the
average of long-term, equilibrium conditions that are inferred to have existed prior to
about 1970. The steady state model was calibrated principally against water levels and
natural discharge data based on observations dating from about 1970 to about 1990.
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Predevelopment potentiometric surfaces show the general direction of
groundwater flow within the aquifers and indicate the areas of recharge and discharge.
These maps and the details of aquifer geometry previously described guided the
delineation of the boundaries of the study area.
The steady state model explicitly depicts a state of hydraulic equilibrium. Solution
of the steady state flow equation requires that recharge and discharge and that the
boundary conditions and stresses do not change with time. The steady state condition is a
relatively simple idea, but one that, owing to the complexity of the hydrogeological
systems, may never exist in the real system. The condition has been approximated for
simulation purposes; thus, the historic observations on which the model is based represent
the average of actual conditions over along period of time. If the hydrological system
undergoes uniform and cyclic changes (such as seasonal fluctuation in precipitation and
evapotranspiration), then the average of a resulting hydrologic response (such as the
decline and recovery of hydraulic head) can define a steady state condition for modelling
purposes.
The allocation of recharge between local and regional flow systems is scale
dependent and continuous flow component. The discretization of the finite-difference grid
used in the digital model determined the scale of the features which were represented in
the model.
Because recharge es.timation as discussed in a previous chapter, is not definite,
recharge in the digital model was simulated by constant flux and specified head nodes in
the outcrops of the modelled aquifers; thus a better understanding of recharge to the
regional aquifer system was one of the benefits derived from the digital model
calibration. Hydrographs of water table wells generally show annual water level
variations of less than 10m without any discernible long-term trend. This observation
and the variable base flow in the outcrop areas, as previously discussed, does not indicate
that recharge from precipitation in the Western Highlands provides all the recharge that
the aquifer can accept and that much of the total precipitation is rejected by the aquifers
and is diverted to the surface runoff, rather than the amount of water surplus available for
infiltration after evapotranspiration and surface runoff were accounted for. However the
350
situation might be different very locally in some areas in the Western Highlands. In these
locations, the amount of recharge accepted by the aquifer is more dependent on the ability
of the aquifer to accept infiltration in the outcrop area and the rate at which the aquifer is
able to move the infiltrated water downgradient out of the outcrop area than on the gross
amount of precipitation available.
Although initial estimates of hydraulic conductivity and verticalleakance values
used in the model were derived from aquifer tests and specific capacity data, the
hydraulic conductivity was allowed to change during calibration. The calibrated hydraulic
conductivities were therefore dependent on water levels, distribution and amount of
recharge, and the amount and distribution of spring discharge but were independent of the
geology. Increasing recharge in the simulations resulted in a corresponding increase in
discharge and a proportional increase in the computed hydraulic conductivity and vertical
leakance values. Thus the magnitude of the simulated hydraulic conductivity and vertical
leakance values include an uncertainty equal to the uncertainty of the estimated recharge.
Groundwater flow to the deeper sandstone aquifer system via the Al-6 aquitard
occurs all over the area. The rate of vertical flow varies between areas depending on the
vertical hydraulic conductivity as well as on the hydrogeological setting of the aquifer
system. It accounts for about 41 % of the total recharge to the B2/A7 aquifer system.
There is a general upward leakage from the lower A4 aquifer to the overlying
B2/ A 7 aquifer. Only small .amount of upward leakage occurs between the two aquifers
because of the very low vertical hydraulic conductivity of the AS/6 confining unit that
separate them.
As the groundwater development in the A4 aquifer affects the piezometric level,
patterns of leakage are disturbed, and the general trend of upward leakage from the lower
to the upper aquifers is reversed and downward leakage between these units occurs.
Although the A4 aquifer system in Amman-Zerqa area, the Al-6 aquifer system in
the Jafr Basin, and the regional deep sandstone aquifer system, represent aquifer systems
deeper than the regional B2/ A 7 aquifer system, and have less complete data sets, the
calibration appears better for these aquifer systems. This probably because the lateral
hydraulic gradients are generally less in the deep aquifers than in the B2/ A 7 aquifer
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system, owing to the fact that the deep aquifers have limited outcrop area and interact to a
much lesser extent with the shallow groundwater flow regime, which is thought to be
highly affected by topography and geological structures.
The effect of the subdivision of the groundwater flow system in the Amman
Zerqa area into sub-basins, with limited connection between them, might weigh heavily
in the future development associated with head decline.
The transition from predevelopment conditions to the 1990 pumping conditions
was accomplished mainly by an increase in recharge to, and a decrease in discharge from,
the regional flow system and by a small decrease in groundwater storage.
Although the assumptions associated with the groundwater flow model of the
regional aquifer systems are probably valid for parts of the study area, the validity of each
assumption is not known for the entire area. Therefore, the results of the simulations
should be considered as conceptual and interpreted with caution.
Admittedly, different methods could be used to synthesise the sparse data, and
different approaches could be used in the simulation of groundwater flow in the carbonate
aquifer system in the study area that may produce different results. However, the overall
trends in the simulation of groundwater flow in the study area would be similar, at least in
a conceptual sense.
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CHAPTER 9
SUMMARY AND DISCUSSION
This study has described the hydrogeological framework and associated
groundwater flow system of the carbonate aquifer systems of the Western Highlands
and Central Plateau of Jordan.
The geology of the study area IS complex. Rocks range in age from
Precambrian to Recent, and the history of the area includes many episodes of
sedimentation, volcanic activity, and tectonic deformation.
The beds dip gently toward the east and northeast. Each formation is
overlapped eastward by the next younger formation, and their eroded edges are
exposed in an updip to downdip succession of older to younger zones. Updip to
downdip variation in lithology within the formations was caused by the succession of
depositional environments. Some formations also exhibit significant lateral
lithological variation. In the Upper Cretaceous-Cainozoic section, there is a general
lateral transition from marine deposits (mainly carbonates) in the north and west to
continental deposits (sandy facies) in the south and southeast.
The National Water Master Plan of Jordan (1977) divides the groundwater
systems in the country into three major aquifer systems or complexes: the deep
sandstone aquifer complex, the Upper Cretaceous carbonate aquifer, and the shallow
aquifer complex.
The Upper Cretaceous-Cainozoic strata, the Belqa and Ajlun groups, comprise
a regional aquifer system of three aquifers and three confining units. The aquifers
from top to bottom are: the Rijam Aquifer System (B4), the Amman-Wadi Sir Aquifer
System (B2/A7), and the Hummar Aquifer System (A4). The B2/A7 is the most
extensive and continuous aquifer system in Jordan. It is the main source of water in
the country. The other aquifers have economical importance only in limited areas,
such as the A4 aquifer system in Amman-Zerqa area, and the B4 aquifer system in the
Jafr area.
The aquifers are separated vertically by three confining units which are
continuous over large areas and affect regional patterns of groundwater circulation.
The confining units from top to bottom are; the Muwaqqar Formation (B3), the
Shue'ib Formation (AS/6), and the Fuheis Formation (A3). The B3 confining unit
confines the downdip parts of the B2/ A 7 aquifer system in most of the eastern parts of
the Central Plateau and in the Southern Desert of Jordan, where it separates the B2/ A 7
and the B4 aquifer systems. Along the Western Highlands however, the B2/ A 7 is
under water table conditions. The AS/6 is a continuous formation that comprises the
lower confining unit of the B2/ A 7 aquifer system and separates it from the A4 aquifer
system in Amman-Zerqa area. The A4 aquifer system is underlain by the A3 confining
unit, which together with the Na'ur Formation (Al/2) separate the Upper Cretaceous
Cainozoic carbonate aquifer system from the deep sandstone aquifer system.
Because of the regional scope of this study and the need to generalise from
site-specific data, the aquifers include some confining strata such as the B 1 Formation,
and the confining units contain some strata permeable enough to supply small
amounts of water to few wells in limited areas such as the thick limestone strata of the
A1I2 Formation.
The aquifer units are mostly composed of limestone, silicified limestone,
sandy limestone, dolomitic limestone, dolomite, chert and sandstone. The confining
units are composed of marl, shale and chalk. In the south and southeast, the limestones
of the aquifer systems and the chalks, marls, and shales of the confining units which
are relatively continuous and hydraulically tight over most of the northern part of the
study area and along the Western Highlands, grade southeastward into the
comparatively permeable sandy facies of the Fassu'a Formation. Accordingly, the
effectiveness of the confining units diminishes and the whole lower part of the Ajlun
Group (Al-6) is modified into a single aquifer system. Therefore, over most of the
Jafr Basin to the east of Arja-Uweina Flexure, the Ajlun Group is connected
hydraulically to the overlying B2/ A 7 aquifer system.
The deep sandstone aquifer system is also an extensive and continuous aquifer
system in the study area. But it is only exploited in its outcropping areas, in the
southern desert of Jordan. Elsewhere, the aquifer is deeply buried and contains saline
water. However, it should be noted that changing economic conditions and increasing
demands for water will almost certainly ensure that the sources of groundwater
presently regarded as too expensive for development will in the future be utilised.
These aquifer and confining-unit divisions are based on regional contrasts in
hydraulic conductivity that determine the relative capacity of the different rock units
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to transmit groundwater. The hydraulic conductivity of the strata was inferred largely
from aquifer tests and specific capacity data and an inherent relation between the
stratigraphy and hydraulic conductivity. A general relation between the stratigraphy
and the hydraulic conductivity exists because the stratigraphy reflects the spatial
distribution of the individual rock units, and each rock unit resulted from a unique
combination of depositional, tectonic, and diagenetic conditions. These same
conditions control the distribution of hydraulic conductivity.
As calculated from pumping test analysis, hydraulic conductivity in the B2/ A 7
aquifer system ranges from less than 0.00005 to more than 45 mIh, with an average of
about 1.4 mIh in the B2/A7. Hydraulic conductivity of the B2/A7 aquifer system was
also calculated from the specific capacity data. The empirical relation between
transmissivity (m2/h) and specific capacity (m3/h1m) is: T = 1.0566(Q I s) 1.0655. Values
ofthe hydraulic conductivity obtained from the latter method -ranging between 0.0002
and 36 mIh with an average of about 1.5 mIh- were found to be consistent with the
values obtained from the pumping test analysis. The hydraulic conductivity ranged
between 0.002 and 39 mIh with an average of about 0.4 mIh in the A4, between 0.06
and 27 mIh with an average of about 6 mIh in the B4, and between 0.0004 and 0.08
with an average of about 0.05 mIh in the Al-6 aquifer systems.
These initial estimates of hydraulic conductivity, which are based on pumping
test analysis, specific capacity data, and hydrogeological data, were refined through
model calibration. The groundwater flow model suggests hydraulic conductivity
values ranging between 0.02 and 1.7 mIh for the B4, between 0.002 and 1.9 mIh for
the B2/A7, between 0.002 and 0.33 mIh for the A4, and between 0.0004 and 0.08 mIh
for the AI-6, aquifer systems.
The hydraulic conductivity of the aquifer systems mainly results from fracture
and joint cavities, solution channels, and fabric-selective forms of porosity caused by
the dissolution of the relatively unstable carbonate constituents. Within the fault
zones, however, the hydraulic conductivity of carbonate strata has increased over
times as the result of large-scaling normal faulting, coupled with associated fracturing
and subsequent dissolution. The faulting vertically displaced the terrain, which
increased hydraulic gradients and helped initiate a dynamic regime of shallow
groundwater flow. In addition to forming new porosity (within the fractures), the
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fracturing increased the hydraulic conductivity by interconnecting voids that, before
the faulting, had been isolated. The dissolution of soluble calcareous constituents
formed moldic and other forms of fabric-selective porosity (Choquette and Pray,
1970) that increased hydraulic conductivity locally. Dissolution along fractures and
bedding planes formed joint cavities and solution channels that eventually became the
principal conduits of regional groundwater flow. The increases in hydraulic
conductivity were greatest in shallow parts of the fault zone because fractures
typically close with increasing depth below land surface and dissolution is most active
near the interval of water table fluctuation (LeGrand and Stringfield, 1971).
A dynamic regime of shallow freshwater circulation probably has existed in
the fault zones areas since Pleistocene time after the formation ofthe fault systems and
has exposed the relatively permeable strata to meteoric conditions. The concentration
of the high angle faults and associated fractures facilitated the percolation of meteoric
water and extended the depth of freshwater diagenesis. The partial pressure of
dissolved carbon dioxide, derived from the atmosphere and soil, increased the
solubility of calcareous constituents.
The areal distribution of hydraulic conductivity generally reflects the
characteristics of the sedimentary sequence that deepens from west to east and grades
from mostly marine deposits in the north and west to sandy facies in the southeast.
However, the hydraulic conductivity distributions depart locally from the general
trend due to tectonic features and karstifications. In the northern parts of the study
area, where the carbonate is dominant, the hydraulic conductivity is of primary and
secondary origins, the latter controlled by the tectonics and the degree of
karstification. Thus a random distribution of hydraulic conductivity is expected, with
local areas of high or low hydraulic conductivity being quite common. In the
southeast, the hydraulic conductivity is more uniform due to the increase in sand
content in the sedimentary sequence (Fassu'a Formation). Furthermore, the increase in
sand content in the Lower Ajlun Group (AI-6), improves the hydrological
characteristics which allow the development of an aquifer system within the group
which is in hydraulic continuity with the overlying B2/ A 7 aquifer system. In general,
the hydraulic conductivity of the A-6 throughout the study area was found to be higher
than thought before, and it is better considered an aquitard which transmits water
downwards into the deep sandstone aquifer system than an aquiclude.
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Vertical hydraulic conductivity of the confining units is the most important
factor affecting the regional groundwater flow system. It controls the amount of
vertical leakage between the aquifers, and thus is the main control on the water level
configuration and groundwater budget. The vertical hydraulic conductivity as
estimated from the groundwater flow model was found to be about 4.2 x 10-8 mIh for
the B3 confining unit, between 4.2 x 10-7 and 2.1 x 10-5 mIh for the A5/6 confining
unit, and between 1.3 x 10-6 and 1.7 x 10-5 mIh for the Al-6 aquitard.
Storage coefficient is important to aquifer system development. In areas where
the system is unconfined, and near outcrop areas, storage coefficient is nearly equal to
specific yield. The specific yield and storage coefficient of the B2/ A 7 aquifer system,
as determined from pumping test analysis, were found to range between 0.002 and
0.13 with an average of about 0.027 and between 0.00001 and 0.03 with an average of
about 0.006 respectively. The groundwater modelling suggests value of storage
coefficient for the unconfined part of the B2/A7 aquifer system ranging from 0.015-
0.07 with an average of about 0.03, and between 0.00005-0.001 with an average of
about 0.0001 for the confined part. The calibrated storage coefficient values average
about 0.025 for the B4 aquifer system, and 0.0005 and 0.05 for the confined and
unconfined parts ofthe A4 aquifer system respectively.
The study area is classified to have variable climatic conditions ranging from a
Mediterranean type in the Western Highlands with rainfall reaching a maximum of
about 650 mm1a to a semi-arid to arid type on most of the Central Plateau and eastern
desert with annual rainfall ranging between 50-200 mm1a.
Most of the precipitation is evaporated from the land surface, is transpired by
vegetation, or moves directly to nearby streams and wadis as overland flow.
Depending on the amount, duration, and intensity of the precipitation, as well as on
the nature of the terrain, soil, and hydraulic gradient, part of the precipitation
infiltrates the land surface; some of the infiltrated water may eventually recharge the
groundwater system.
Estimates of average annual evapotranspiration rate within the study area vary
from a maximum of about 2490 mm1a in the southeastern desert to a minimum of
nearly 1153 mm1a in the Western Highlands. The annual evapotranspiration rate
decreases to the north and west, reflecting regional climatic trends. The linear
357
regression equations of the relationships between potential evapotranspiration (PET)
and Class-A Pan evaporation (Eo) were found to be; PET = 0.5822 (Eo) + 13.023 for
the wet period, PET = 0.5221 (Eo) + 27.342 for the dry period, and PET = 0.5537
(Eo) + 17.046 for the whole year.
Runoff is the second largest element of the water budget in the study area
(after evapotranspiration). It averages about 9.3 % (179 MCMla) of the amount of
precipitation: 4.9 % (94 MCM/a) as surface runoff, and about 4.4 %( 85 MCMla) as
baseflow. The curve number method (CN) was applied during this study to estimate
the volume of surface runoff in the study area. Depending on the curve number values
and the amount of rainfall, the method combines infiltration losses with the initial
abstractions to estimate the runoff volume. The curve number value assigned for an
area indicates the runoff potential of that area. It incorporates the important catchment
properties (such as soil type, landuse/treatment, surface condition, and antecedent
conditions) that affect the runoff volume. Estimated surface runoff coefficients vary
between the different subcatchments, ranging from about 3.15 % in the desert areas to
more than 6.2 % in the west and north. The areal pattern of runoff is similar to that of
precipitation. Runoff generally increases from east to west and in the northern part of
the study area. However, this pattern reflects the changes in climate, physiography,
and geology of the study area. Surface runoff is generally most important where the
terrain is steep, the soil texture is fine, and there is little plant cover.
The amount of recharge to an aquifer is limited by the amount of infiltration,
which in tum is limited by the difference between precipitation and overland flow
after evapotranspiration has been accounted for. Infiltration and overland flow are
inversely related. The ratio of infiltration to overland flow decreases as the rate of
precipitation exceeds the infiltration capacity of the soil. The areal distribution of
recharge to the regional groundwater flow regime of the carbonate aquifer system was
calculated from what is known about precipitation, total runoff, and
evapotranspiration, and analysed by groundwater flow model simulation.
Recharge occurs by direct infiltration of rainfall in outcrop areas, indirect
recharge through the transmission losses of the flood flow via the wadi beds, vertical
leakage through the underlying and the overlying strata, water transfer from adjacent
aquifer systems, or by lateral boundary flow from outside the study area.
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Generally direct recharge does not occur when annual precipitation is less than
200-250 mm. However, due to the presence of permeable materials which have low
field capacities in the wadi alluvial fans, localised direct recharge can occur after large
intensive storms even though the annual rainfall is less than 200 mm.
Long term average estimates of recharge for the present landuse conditions
indicated about 100 MCMla, most of which was probably discharged as spring flows.
The direct recharge and the lateral boundary flow occupy the major part of the total
recharge in the Western Highlands. Direct recharge calculations suggest that 8% of the
total rainfall percolates downward to recharge the aquifer systems. While in the
eastern and southern parts, indirect recharge and lateral boundary flow constitute the
majority of the total recharge.
The effect of recharge on groundwater appears in terms of variation in water
levels, spring flows, and chemical and isotopic composition of groundwater. The
attempt to analyse these variations for the purpose of estimating recharge into the
main aquifer system was constrained by the inadequacy of the data. However,
recession hydro graph analysis for some springs, has successfully explained the
relationship between the recharge rate and spring flow systems. The recession
hydro graph analysis shows that the discharge coefficient (a) for the main springs in
the study area is rather high and ranges between 0.001066-0.007167 with an average
of about 0.003794. While the exhaustion coefficient a- is small and below the average
for the Mediterranean karst system, it ranges between 0.000256-0.002751 with an
average of about 0.00112.
Spring hydro graph analysis was also conducted to estimate the recharge
coefficient from the relationship between the spring discharges and the necessary
surface catchment area needed to provide enough recharge to maintain that discharge.
The method entails establishing the spring coefficient C which relates its annual
discharge with the annual rainfall as a function of recharge coefficient and the
catchment area. Springs with high C values must have high recharge coefficients or
large surface catchment area. The distinctive relation between the recharge coefficient
and the possible catchment area for a spring, or a group of springs that are believed to
discharge from the same catchment leads for division of springs in the study area into
local, intermediate, and regional on the basis of their surface catchment area.
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Recharge also has been estimated from the groundwater flow model
simulations. The simulated total recharge for the B2/ A 7 aquifer system averages about
120.76 MCMla, of which about 42 % (50.42 MCMla) flows downward to recharge the
deep sandstone aquifer system. The simulated recharge to the A4 aquifer system is
about 6.4 MCMla. The AI-6 aquifer system receives the least amount of recharge,
averaging about IMCMla across the outcrop area in the western highlands of the Jafr
Basin. The major recharge to the AI-6, averaging about 3.5 MCMla, occurs as
downward leakage from the B2/ A 7 aquifer system. The simulated recharge to the B4
aquifer system in the low rainfall zone of the Jafr Basin is estimated to be about 2
MCMla.
Much of the groundwater from the carbonate aquifer system is discharged to
the land surface by numerous springs. This water, abstracted for domestic and
irrigation uses, seeps back into the ground, discharges as evapotranspiration, or flows
to wadis and rivers that leave the study area. In addition to the main springs that occur
in the wadis which cut deep into the saturated thickness of the aquifer system, many
small springs occur in the mountains. The locations of these springs are controlled by
permeability variations in the rocks and water levels related to the land-surface
altitude which causes the water to discharge at the surface.
Baseflow is controlled largely by the underlying geology, the degree of stream
entrenchment, and the head relations between groundwater levels and water levels in
the surface water bodies. Shallow headwater streams receive baseflow from locally
occurring, principally unconfined aquifers. The major, more deeply entrenched
streams-such as the lower parts of the Wadi Mujib and Wadi Hasa- receive baseflow
from the deep, principally confined aquifers. Although over the long term the shallow
streams drain off a significant a mount of groundwater, many dry up during extended
periods of little precipitation. Because the major streams tap flow paths deeper in the
regional groundwater flow regime, they are less affected by either droughts or periods
of above average rainfall.
Groundwater flow in the study area was conceptualised as relatively shallow,
intermediate, and regional flows primarily through the carbonate sediments of the
Western Highlands and the Central Plateau superimposed over deeper flow through
primarily sandstone sediments.
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The most important controls on hydraulic head in the B2/ A 7 aquifer system
are the slope on the base of the aquifer system, topographic relief, the location of
springs and streams, and the geological structures. The base of the aquifer system
generally slopes from west to east and northeast, and this is the prevailing direction of
groundwater flow as indicated by the potentiometric contours. The altitude of land
surface decreases between 500-800 m from west to east, and the potentiometric
surface typically is a subdued replica of the associated topography. The strong
influence of springs and streams on the shape of the potentiometric surface indicates
that the distribution of hydraulic head and the direction of groundwater flow largely
are controlled by the areas of the groundwater discharge. The influence of the main
wadis is apparent from the steep hydraulic gradients toward these regional drains. The
potentiometric contours sweep upstream where the wadis draining the Central Plateau
are sustained largely by base flow.
The geological structures and the resulting distributions of transmissivity in
the fault zones make the regional potentiometric surface map a misleading indicator of
the direction of groundwater flow in many areas, particularly in areas where the fault
lines trend north-south in a direction perpendicular to the regional groundwater flow.
The regional potentiometric contours indicate that under typical, isotropic conditions
most of the groundwater should flow northeastward. However, some of the fault
systems are barrier faults, which impede or block the northeastward flow of
groundwater, so that most of the water is diverted north and south wards, and around
both ends of the fault line, towards the closest discharge area.
The philosophy adopted in modelling the carbonate aquifer systems in the
Western Highlands and Central Plateau of Jordan was to develop the simplest model
that could account for the principal features of flow in the aquifer systems. Three
dimensional groundwater flow models were used to simulate the concept of
groundwater flow in the area. The area was subdivided into five subregions that
approximately cover, individually, the Upper Zerqa, Wadi Wala, Wadi Mujib, Wadi
Hasa, and Jafr basins. Each subregion was modelled separately and then compiled in
one regional model. An intermeshing finite-difference grid system was used to
coordinate the entry and calibration of model data. The regional grid has 98 rows and
36 columns with variable node spacing ranging between 1250-5000 m. The
subregional grids are meshed with the regional grid such that the subregional blocks
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fit within one regional block. The number of subregional blocks in one regional block
is varied according to the variable node spacing chosen for the subregional models.
Six model layers were used to simulate the shallow and deep flows. The upper five
layers were used to simulate the flow in the carbonate aquifer systems. The lowest
model layer was used to simulate the concept of deep flow in the sandstone aquifer
system.
Steady state calibration of the models was achieved by adjusting hydraulic
parameters and then comparing simulated heads and flows with those measured or
estimated prior to pumping. Calibration of the models for the transient conditions was
achieved by further adjustment of hydraulic parameters including the storage
coefficient until the computed response of the models for the 1990 pumping
conditions approximated the measured heads. Parameter adjustments made for
calibration were used to resimulate prepumping conditions to obtain the initial
conditions for the transient simulation. This procedure ensured calibration
compatibility between steady state and transient conditions.
Sensitivity analysis was performed on hydraulic parameters and model
assumptions to evaluate the reliability of the model calibration. The steady state
model of predevelopment conditions is found to be sensitive to increases in the rates
of recharge in the outcrop area, to the changes in hydraulic conductivity, to the
changes in the vertical hydraulic conductivity of the confining units, and to the
extension and the hydraulic ·conductivity of the internal groundwater barriers.
The sensitivity analysis shows that model calibration was least affected by
large changes applied to all aquifer system layers compared with the changes in
individual layer hydraulic parameters.
Simulated discharge to streams is most sensitive to hydraulic conductivity and
much less sensitive to other parameters including the hydraulic conductance of the
head dependent flow boundaries used to simulate the spring discharges.
Definition of the flow system was accomplished through examination of the
following results derived from the calibrated model: (1) regional water budget, (2)
potentiometric surfaces, (3) vertical leakage between aquifers, and (4) lateral flow
directions in the aquifers. Analysis of the calibrated model indicates that before
development (pre 1970), about 120.67 MCMla of water flowed through the B2/A7
aquifer system. More than 79.74 MCMla ofthis amount represented recharge from the
362
aquifer outcrop areas, and the remainder, less than 41.02 MCMla represented water
flowing from vertically and laterally adjacent aquifers and streams. In the steady state
condition, nearly 44.95 MCMla of the water flowing through the aquifer system
discharge to the regional drains in the major river valleys, about 49.62 MCMla
discharged to the deep sandstone aquifer system, about 3.5 MCMla discharged into
the AI-6 aquifer system in the Jafr Basin, and about 22.69 MCMla left the study area
laterally to the east. A substantial amount of the water flows in the AI-6 aquifer
system discharged into the east (about 2.7 MCMla) and the remaining 0.8 MCMla
infiltrated downward increasing the total vertical leakage into the deep sandstone
aquifer system to 50.42 MCM/a.
The steady state budget of the Hummar (A4) aquifer system in the Amman
Zerqa area indicates a total recharge of about 6.4 MCMla, of which about 2.16
MCMla was simulated to leak upward into the B2/A7 aquifer system. The simulated
spring discharges from the system were about 7.7 MCMla, which indicates a water
deficit of about 3.5 MCMla. However, the model provides that amount from the
lateral flow from the western part of the model area: it is believed that a substantial
part of the 7.2 MCMla Zerqa River overflow from the B2/A7 aquifer system in the
northern part of the Amman-Zerqa area, where the A4 aquifer crops out, infiltrates
downward into the A4 aquifer system.
In 1990 the average pumping rate from the B2/ A 7 aquifer was about 80.5
MCMla. This pumping rate was balanced by an increase in total recharge, a decrease
in discharge to river valleys, a decrease in storage, and a decrease in downward
leakage and water flowing out of the system outside the study area. The head declines
however, are localised in the pumping centres. Only in the Amman-Zerqa area does
the water balance analysis show a regional decrease in storage as a result of heavy
abstraction in the area.
The amount of water stored in the B2/ A 7 aquifer system in the Amman-Zerqa
area is calculated to be about 780 MCMla. The accumulated abstraction during the
modelled period mounted to 549 MCM against 407 MCM of recharge, leaving a
deficit in the groundwater balance of about 142 MCM (18 % of the storage). Of this
value, as the water balance indicates, 54 % was derived from decreased natural
discharge from the aquifer, about 10 % was derived from increased recharge, and
about 32 % was being supplied from the aquifer storage. The remaining 4 % includes
363
all other factors, including interchange with adjacent aquifers and lateral flow moving
out of the study area.
The base flow from the aquifer system is simulated to have decreased from
about 18.6 MCMla prior to 1970 to about 13.0 MCMla in 1990, which represents a
reduction in the discharge of base flow totalling about 5.6 MCMla. This decrease in
base flow indicates a reduction in hydraulic gradients between the regional aquifers
and major springs, owing to water level decline caused by pumpage. The simulated
reduction of base flow appears to be reasonable, however, considering the observed
decreases in base flow and head decline in wells adjacent to the major springs.
The amount of drainable water in the B2/ A 7 aquifer system is a function of
specific yield or storage coefficient and the thickness of the aquifer. The estimated
specific yield for the B2/A7 aquifer system is 0.03. Estimates of drainable water
stored in the B2/A7 aquifer system are about 780, 9900, 17100, and 14400 MCM in
the Amman-Zerqa area, the Wadi Wala Basin, the Wadi Mujib and Wadi Rasa basins,
and the Jafr Basin respectively.
Groundwater levels in the B2/ A 7 aquifer system mostly vary in response to
short-term fluctuations in recharge and long-term variations in discharge. Most of the
fluctuation in recharge results from cyclic patterns in precipitation, and most of the
variation in discharge results from pumpage trends. Water levels have declined where
and when the rates of recharge and natural discharge have not compensated for
increasing rates of groundwater abstraction.
The results of the groundwater modelling show that the calibrated values of the
model input are, for the most part, consistent and within the range of reasonable
possibilities. The simulated response to departures from the calibrated input suggests
that the capacity of the model to simulate field conditions deteriorates as the departure
mcreases.
364
CHAPTER 10
CONCLUSIONS AND RECOMMENDATIONS
10.1 CONCLUSIONS The fundamental, but largely unresolved, problem in investigating the
hydrogeology of regional groundwater flow system has been establishing relationship
between geology, morphology, climate, and the hydrological response of the aquifer
system.
The geological and morphological (physical) attributes that will be focused
upon are the hydraulic properties of the catchments and the aquifer systems. The
climatic (the atmospheric input) attributes control the amount of surface runoff and
groundwater recharge which are controlled to an important degree by the physical
attributes. The physical attributes are vary in space fixed in time, while the
atmospheric input are vary in space and time.
The focus of the hydrogeological studies is the hydrological response of the
geological framework of the aquifer system as a function of the atmospheric input.
This required investigating the physical characteristics of the aquifer systems and the
catchment areas, and to determine what effect they have on hydrological response. As
such, the hydrogeological study is intended to answer question about lateral flow of
groundwater from recharge to discharge areas, its vertical movement, and the ground
water yielding properties of the aquifer system.
This study describes the hydrogeological framework and associated
groundwater flow system of the Mesozoic aquifer systems of Jordan. The study area is
classified to have variable climatic conditions ranging froin Mediterranean type in the
Western Highlands to a semi-arid to arid type on most of the Central Plateau and
Eastern Desert. The aquifer systems are developed in a thick sequence of Upper
Cretaceous-Cainozoic carbonate rocks. The sequence exhibits vertical and lateral
variation in lithology. Since deposition, however, compression, extension, intrusive and
volcanic episodes, and erosion have greatly modified the distribution and thickness of
the carbonate rocks.
The possible effect of major structures and change in rock type and lithology
on groundwater flow was the subject of this study. The general purpose of this study
was therefore to produce a conceptual evaluation of groundwater flow and to better
define the relationships between recharge, discharge, water level, and aquifer
characteristics.
In order to properly address the objectives of the study the regional approach
was required. Therefore, and because the scarcity and variability of the hydrogeological
factors which are directly affect the groundwater flow system, generalisations from site
specific data were essential in many parts of this study.
The objectives ofthe study was best met by the construction of the regional scale
digital model of the aquifer systems, supplemented by more detailed subregional
models. These models provided a framework for the interpretation and evaluation of the
distributions of observed aquifer characteristics and their relation to present and past
patterns of groundwater flow.
The physical properties of the aquifer system have been refined and the
relationship between the variable atmospheric input and the hydrogeological
framework with certain characteristics has been established. The hydraulic parameters
of the aquifer systems were inferred from aquifer tests, groundwater flow modelling,
and the inherent relation between the stratigraphy and hydraulic parameters. The results
of the model calibration and sensitivity analysis show that the calibrated values of the
model input are, for the most part, consistent and within the range of reasonable
possibilities.
As in many of the hydrogeological studies, investigation might be constrained
by the lack of full continuous records of the different hydrogeological parameters.
However, the study shows that the absence of direct measurements for the different
hydrogeological parameters must not be an obstacle. Indirect methods provide
reasonable estimations for the different parameters required for such regional study.
Because many of these parameters were dependant on each other, the relationships
between them were studied to derive an equations that, in tum, are used to calculate
these parameters. For examples potential evapotranspiration was calculated from the
relationship between Class-A pan evaporation and potential evapotranspiration, the CN
method were employed to estimates the volume of surface runoff, transmissivity was
calculated from the relationship between specific capacity and transmissivity, The
storage coefficient was calculated by comparing the volume of dewatered material in
the cone of depression and the total volume of discharge water, and spring hydro graph
366
analysis and water level fluctuations, with conjunction with the recharge estimations,
were used to determine the recharge coefficient and to explain the effect of recharge on
groundwater flow system.
It should be noted that the values of the regression coefficients produced through
this study, are specific for the units and types of data used in this analysis. The degree of
accuracy of these coefficients are proportional to the reliability of the input data used in
the calculations. Therefore, future investigations were required to recalibrate and
verified the validity of the values of these regression coefficients.
The groundwater flow systems of the carbonate aquifer systems of the Western
Highlands and Central Plateau of Jordan are complex. They reflect the changes in
climate and geology of the study area. The flow within the regional aquifer system, in
general, is controlled by the altitude of major recharge areas, major discharge areas,
and major structural features. Thus topography provides the major control for the
regional aquifer system.
The effects of interaction of the physical attributes and atmospheric input on
groundwater flow system were demonstrated, for example, on the areal distribution of
hydraulic parameters; in the northern parts of the study area, where the rainfall is the
highest and the aquifer systems are made of limestone, the permeability is high. In this
area the permeability of the limestone, as in many carbonate rocks, is provided by
solution enlargement of bedding planes and joints as a result ofkarstifications, which is
in tum related to the quantity of water flowing through the system. While in the south
and southeast, where the rainfall is lower and the limestone tend to be silicified or
sandy, the degree of karstification becomes lower.
Although the aim was to describe the framework hydrogeology of the carbonate
aquifer system in the Western Highlands and Central Plateau of Jordan and the model
simulations were entirely conceptual, this study presents estimate of the direction and
magnitude of flow from recharge to discharge areas and discusses where the results
agree and disagree with the hypotheses and hydrological estimates reported by other
investigators.
The results of this study intend to fully document and demonstrate the different
aquifer parameters and the description of the hydrogeologic framework and associated
367
flow systems of the carbonate aquifers, so that can be used by others to evaluate specific
groundwater management for the principal aquifer systems.
Although this study includes fairly comprehensive details about the regional
carbonate aquifer system in the Western Highlands and Central Plateau of Jordan, it
represents rather a methodology for investigating a framework hydrogeology of a
regional complex aquifer system.
10.2 RECOMMENDATIONS
As in any hydrogeological study, the availability of adequate data records for
the different hydrogeological parameters is essential for better understanding of the
groundwater flow system. Two sets of recommendations are drawn particularly from
this study. The first set concerns the need for more data and information, particularly
in the remote eastern parts of the study area:
1- potential evapotranspirations,
2- continuous records of spring flow and overland flow measurements,
3- land classifications including soil types and properties,
4- storage coefficients of the different aquifer systems,
5- vertical hydraulic conductivities of the confining units,
6- continuous records of water levels,
7- groundwater abstractions by time, location, and aquifer.
The second set of recommendations address the need for more detailed
investigations or studies on particular aspects of the groundwater flow systems:
1- the relationships between recharge and water level fluctuations and spring
discharges,
2- the interrelationships between the different aquifer systems,
3- the storage of the confining units and its effect on the flow system,
4- the groundwater flow system and development of the AI-6 aquifer system,
5- the continuity ofthe flow systems into the remote eastern parts of the country,
6- the effect of future abstractions on groundwater flow and quality,
7 - groundwater modelling with smaller grid blocks and more accurate input data,
which will include calibration of the groundwater flow model after long term
water development.
368
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APPENDICES
Appendix (AI) Well list in the study area
Appendix (BI) Definition ofSCS Hydrologic Soil Groups (HSG)
Appendix (B2) Runoff curve numbers
Appendix (B2.1) Runoff curve number for Urban Areas
Appendix (B2.2) Runoff curve number for cultivated Agricultural Lands
Appendix (B2.3) Runoff curve number for other Agricultural Lands
Appendix (B2.4) Runoff curve numbers for Arid and Semiarid Rangelands
Appendix (B3) Surface Water in Jordan
Appendix (B4) Runoffmeasurments in the study area.
Appendix (B4.1) Runoff measurments for Zerqa River at Sukhna Gauging
Station in MCM.
Appendix (B4.2) Runoffmeasurments for Wadi Wala at Karak Road in MCM.
Appendix (B4.3) Runoffmeasurments for Wadi Wala at weir in MCM.
Appendix (B4.4) Runoffmeasurments for Wadi Swaqa in MCM.
Appendix (B4.5) Runoffmeasurments for Wadi Mujib at Karak Road in MCM.
Appendix (B4.6) mean annual observed flood flow of Hasa River at Tannur in
MCM.
Appendix (B4.7) Observed runoff discharge of Has a River at Ghor Safi in MCM.
Appendix (B4.8) Mean annual observed flood flow of Wadi Jurdhan in MCM.
Appendix (CI) Results of pumpin test analysis in the B2/A7 aquifer system.
Appendix (DI) Soil moisture balance (mm) for West Amman sub-catchment for the water year 1982/1983.
Appendix (AI) Well list in the study area
381
Well No. Coordinate Groun Well Water Water Yield D.dwn. SC S. Th. T K
East North level depth depth level (mJ/h) (m) (m2/h) (m) (m2/h) (m/h)
A61 255.257 171420 70 15.1 4.7 18 5.495799 0.305322
A62 256.273 170.206 527 85 38 489 40 47
A63 256.454 170.163 527 194 38 489 77 17.7 4.3 156 4.998863 0.032044
A64 255.780 170.100 530 108 36 494 70 14.7 4.7 72 5.5 0.076
A65 255.904 170.100 531.1 95 40.6 490.5 0.5 54.4 0.505 0.0093
A66 255.604 170.608 188 2.19 66 97 91.75604 0.945939
A67 241.006 152.556 60
A68 251.49 159.5 40
A69 250.312 169.778 25
A70 242.3 152.8 751 261 55.9 695.1 250 1.69 147 205.1 215.371 1.050078
A71 240.8 152.3 350 5.92 121 III 175.0324 1.576868
A72 242 152.2 250 9 28 110 36.80082 0.334553
A73(PPI09) 247.815 158.842 621 39 23 598 50 0.13 14 0.120176 0.008584
A74 245.27 156.325 95 28 3.3 15 3.770398 0.25136
A75 242.56 154.403 30
A76 253.408 159.485 618 200 44.7 573.3 120 51 155.3 69.71506 0.448906
A77 253.31 159.222 635 90 61.9 573.1 80 66 28.1 91.75604 3.26534
A78(PP208) 253.26 158.95 630 92 47.88 582 III 0.24 463 44 731.2858 16.62013
A79 253.02 159.972 603 119 30 573 90 11 89 13.59924 0.1528
A80 252.8 157.85 642 107 67 150 8 19 40 24.34572 0.608643
A81 241.25 151.6 130 5.25 25 109 32.61487 0.299219
A82 241.075 151.42 30 6 5 25 5.870339 0.234814
A83 250.04 158.75 601 21 15.8 585.2 20 5.2
A84 252.745 159.6 598 155 24 574 80 131
A85 247.5 158.7 613 38.2 23 590 90 8.8 10.2 15.2 12.54799 0.825526
A86 255.9 164.85 60 0.5 120 11 173.4915 15.77196
A87 255 159 170 9.4 18 18 22.98283 1.276824
A88 254.5 159.55 230 2.5 92 16 130.7153 8.169708
A89 253.637 160.018 92 5.85 15.8 5 20.00229 4.000458
A90 258.17 165.715 45 7.1 6 25 7.129035 0.285161
A91 251.11 159.04 621 65 44.8 75 4.7 15.9 20.2 20.13721 0.996891
A92 246.285 156.345 25 30.6 0.8 70 0.833015 0.0119
A93 253 158.42 92 0.6 153 10 224.7498 22.47498
A94 250.855 158.93 594 37 17.5 576.5 15 19.5
A95 252.852 160.6 152 3.3 46 30 62.45671 2.08189
A96 254.887 166.872 569 136 66.6 502.4 167 4.8 34.7 69.4 46.25212 0.666457
A97 255.53 157.82 63 4.6 13 17 16.24865 0.955803
PP118 256.66 160.06 638.3 170 97 541.3 12 14 0.86 73 0.899743 0.012325
A98 252.548 162.142 575 65 13.5 561.5 15 51.5
A99(PP325) 254.835 165.96 573 136 57.1 515.9
AI00 240.986 152.535 13.7 42 17.1825 0.409107
A101 242.535 155.563 127 100 184.295 1.84295
AI02 238.732 150.493 103 100 147.431 1.47431
AI03 243.239 152.324 6.5 35 7.763719 0.221821
AI04 250.986 159.366 621 65 44.8 576.2 165 20.2 243.5789 12.05836
AI05 251.409 159.365 592 50 18.5 573.5 51 31.5 69.71506 2.213177
A 106 252.65 161 35 28
AI07 254.789 166.056 569 136 66.6 502.4 3.1 69.4 3.527414 0.050827
AI08 254.815 166 573 137 57.1 515.9 40 14 79.9 17.58369 0.220071
AI09 238.31 150.563 0.08 0.07164
AllO 246.338 156.338 175 76.2 2.5 98.8 2.804889 0.02839
Alll 246.409 156.197 212 79 0.6 133 0.613099 0.00461
All2 247.676 156.197 696 135 59.7 636.3 2.9 75.3 3.285456 0.043632
382
Well No. Coordinate Groun Well Water Water Yield D.dwn. SC S. Th. T K
East North level depth depth level (m3/h) (m) (m2/h) (m) (m2/h) (m/h)
AI13 252.394 159.577 3.1 3.527414
A114 255.775 164.93 I 1.0566
A115 254.93 159.014 50 68.25951
A116 252.535 162.254 570 70 5 565 0.15 65 0.13997 0.002153
A117 257.606 169.85 539 \06 50 489 1.9 56 2.093739 0.037388
A118 251.338 173.521 474 29 4 470 I 25 1.0566 0.042264
A119 255.211 170.845 543 114 32 511 1 82 1.0566 0.012885
A120 252.817 172.507 486 7.6 4.9 481.1 5.6 2.7 6.623766 2.453246
AI21 (SI6) 254.800 167.000 586.64 116 57.95 528.7 132 15 8.8 40.1 10.72 0.2674
AI22 (Zerqa No.5) 255.750 170.180 160 60 3.8 15.8 120 20.00 0.16669
AI23 (Khaw) 258.17 165.71 170 42 1.5 28 100 36.80082 0.368008
AI24 (W. Rimam) 241.25 151.6 120 5 24 80 31.22667 0.390333
AI25 250.5 171.81 485 40 10 475 30
A126 251.386 159.391 482 22 6.1 475.9 15.9
AI27 253.84 173.65 490 71 21 469 50
AI28 254.185 174.09 499 85 29 470 56
AI29 254.875 174.09 504 70 16.8 487.2 53.2
A130 253.238 172.696 487 40 4 483 36
AI31 250.875 171.87 516 107 41 475 66
AI32 255.8 170.13 528.5 104 39 489.5 65
AI33 252.714 162.044 563 2 561
AI34 252.8 161.783 569 5.8 563.2
A135 252.886 161.9 572 15 6.2 565.8 8.8
A136 252.955 161.188 582 20 3.8 578.2 16.2
A137 257.191 173.364 518 75 19 499 56
A138 252.186 173.647 477 80 8.3 468.7 71.7
A139 251.586 173.382 50 17 33
AI40 252.477 173.7 47 4 43
AI41 255.214 172.64 518 91 28 490 63
AI42 255.129 172.94 500 50 14.3 485.7 35.7
A43 255.071 171.217 518.4 90 29.83 489 42.7 3.3 12.9 62 16.11551 0.259928
AI44 255.371 170.826 525.3 78 34 491.3 44
AI45 255.211 172.324 533 85 25 508 60
AI46 252.817 173.38 484 22 11.9 472.1 10.1
AI47 255.634 172.535 515 32 17 498 15
AI48 251.268 171.972 514 106 40.4 473.6 65.6
AI49 251.127 170.282 72 39 33
AI50 250.5 170.5 499 40 23 476 17
AI51 250.6 171.831 505 112 31.3 473.7 80.7
AI52 250.4 170.6 50 36 14
AI53 506 95 41.5 464.5 53.5
AI54 254.648 171.69 513.3 70 25 488.3 45
AI55 252.817 172.676 482 39 0.6 481.4 38.4
AI56 255.916 172.394 516 85 39.7 476.3 45.3
AI57 256.9 173.521 540 102 23 517 79
AI58 497 115 84 413 31
AI59 257 173.521 543.5 120 43 500.5 77
AI60 252.747 162.394 580 31 12 568 19
AI61 258.099 165.704 589 130 100 489 30
AI62 251.831 163.521 74 17 57
AI63 250.1 159.8 502 40 22.5 479.5 17.5
AI64 250.5 169.5 510 60 31 479 29
AI65 251.409 159.575 595 22 5.8 589.2 16.2
383
384
385
Well No. Coordinate Groun Well Water Water Yield D.dwn. SC S. Th. T K
East North level depth depth level (m31h) (m) (m21h) (m) (m21h) (m/h)
PP81 239.94 125.23 714.55 477 143 572 10 50 0.2 157 0.190177 0.001211
PP82 239.49 129.96 745.66 209 168 578 8 35 0.23 41 0.220715 0.005383
PP85 238.38 108.46 623.19 281 87.6 536 26.5 55.4 0.48 121.4 0.483363 0.003982
PP86 246.28 110.43 699.72 265 90 610 6.2 25 0.25 175 0.241221 0.001378
PP87 242.65 131.7 720 162 152 568 10
PP88 229.7 54.35 968 250 152 816 80
PP89 223 52.6 1060 300 79.9 980 20 57 0.35 151 0.345235 0.002286
PP480 239.95 129.79 738.1 245 152.3 586 79 8.3 9.51 92.7 11.6456 0.125627
GW3 220.94 62.35 985.7 216 179 37
GW4 220.75 62.1 1000 156 94.74 905 170 0.6 283.3 61.6 433.2907 7.033939
GW5 220.94 61.58 980 267 99.81 880 28 22.6 1.2 167.2 1.283152 0.007674
MU1 227.7 94.1 180 500 77.4 103 103 10.34 9.9 422.6 12.15514 0.028763
MU2 227.45 94.6 180 750 66.95 113 683.1
Ql 239.89 129.71 737.05 216 148.7 588 67.3
Q2 240.63 129.65 723.72 245 158 566 81 11.02 7.36 87 8.862759 0.101871
Q3 242 128 730 367 148 582 DRY 219
Q4 239.08 129.34 730 288 156 574 57 1.75 32.6 132 43.27569 0.327846
Q5 239.55 128.76 725 202 149.6 575 56 15 3.73 52.4 4.296021 0.081985
Q6 242.75 131.7 740 363 151.8 588 30 4.5 6.67 211.3 7.980253 0.037767
Q7 238.3 131.4 745 325 175 570 DRY 52 150
Q8 239.65 129.25 730 309 162 564 147
Q9 238.25 129.75 326 161.9 21 50.35 0.4 164.1 0.398021 0.002425
QIO 239.3 127 720 363 160.4 560 DRY 202.6
Ql1 239.25 130.7 205 162.4 63 1.45 43.45 42.6 58.77447 1.379682
Q12 238.95 125.85 720 343 146.6 573 39 61.8 0.6 196.4 0.613099 0.003122
Q13 239.97 130.52 332 158.2 DRY 173.9
Q14 239 131.5 750 200 166.5 583 67 1.2 55.8 33.5 76.7272 2.290364
Q15 239.05 132 745 224 167 578 55 33 1.7 57 1.859747 0.032627
Q16 239.65 132.15 735 201 163.2 572 65 9.75 6.7 37.85 8.018503 0.211849
Q17 239.5 132.54 742 204 165.6 576 75 2.35 31.2 38.4 41.29831 1.075477
Q18 239.2 132 740 204 171.2 569 70 5.6 13.2 32.8 16.51514 0.50351
Q19 239.29 133.29 222 185 DRY 37
Q20 239.29 133.18 203 173 40 30
Q22 238.06 132.65 278 187.9 11 34.7 0.32 90.1 0.313796 0.003483
ABI 247 45 880 161 111.1 769 91 0.25 364 49.9 565.9318 11.34132
AB2 250 45 855 161 113.8 741 100 1.53 65.4 47.3 90.86752 1.921089
AB3 248 46 855 232 95.79 759 61 7.01 8.7 136.2 10.59177 0.077766
AB4 247.38 47.8 850 180 98.48 751 51 8.92 5.72 81.5 6.775106 0.08313
AB5 249.7 49.3 845 180 103.5 742 80 1.17 68.38 76.6 95.28566 1.243938
AB6 248.61 44.77 850 250 110.9 739 139.1
AB7 251.44 47.66 850 180 108.7 741 85 12 7 71.3 8.401611 0.117835
AB8 251.91 46.07 860 180 113.8 746 50 15 3.33 66.2 3.80693 0.057507
AB9 247.1 44 165 117.9 40 23.72 1.69 47.2 1.848093 0.039155
API 249.34 72.4 780 140 95 685 15 45
AP2 253.45 78.1 808 163 135 673 28
AP8 248.89 74.94 785 230 103.3 682 115 27.3 4.2 126.7 4.875091 0.038477
AP9 249.44 76.12 780 223 102.6 677 123 3.79 32.4 120.4 42.99286 0.357084
API0 247.94 77.44 793 226 113.6 679 50 13.2 3.8 112.4 4.381976 0.038986
AP11 247.5 76.1 780 263 103.5 676 60 70.9 0.85 159.5 0.8886 0.005571
API2 249.3 75.1 778 251 101.2 677 111.5 10.28 12.25 149.8 15.25175 0.101814
AP13 248.68 .74.95 778 215 95 683 120 9.3 12.9 120 16.11551 0.134296
AP14 249.5 75.75 765 232 97.55 667 120.5 3.58 33.1 134.5 43.98325 0.327013
AP15 248.95 78.9 781.5 233 98.58 683 114.3 2.94 38.9 134.4 52.23985 0.388689
386
Well No. Coordinate Groun Well Water Water Yield D.dwn. SC S. Th. T K
East North level depth depth level (m31h) (m) (m21h) (m) (m%) (m/h)
API6 249.65 76.51 775 240 100.1 675 126 1.92 65.6 140 91.16364 0.651169
API7 269.2 72.92 858 280 195 663 8 32 0.22 85 0.210505 0.002477
API8 262.82 75.02 826 266 161 665 18 29 0.62 105 0.634898 0.006047
ERI 241.1 121.95 704.97 294 152.7 552 87 1.07 81 141.3 114.1304 0.807717
ER2 242.9 119.2 718.01
ERJ 243.32 118.725 718.97
ER4 242.4 118.41 714.03
SW6 251.8 87.25 743.34 160 79.1 664 130 35.52 3.7 80.9 4.259215 0.052648
SW7 252.95 87.05 727.55 262 63.95 664 146 0.63 232 198.1 350.2177 1.767883
SW8 254.6 86.94 743.62 260 81.95 662 54.2 79.15 0.7 178 0.722541 0.004059
SW9 251.1 87.28 714.21 163 50 664 113
SWIO 250.84 87.84 710.23 237 47.34 663 146 24.3 6 189.7 7.129035 0.037581
SWII 250.2 88.25 712.09 254
SWI2 249.82 84.7 752.27 192 87.2 665 109 22.93 4.76 105 5.570585 0.053053
SWI3 253.05 85.15 767.29 283 108 659 75 2.56 29.3 175 38.62407 0.220709
SWI4 252.11 85.85 745.27 206 84.6 661 89 3.45 25.9 121.4 33.86737 0.278973
SWI5 251.67 85.76 750.94 200 90.9 660 68 2.89 23.5 109.1 30.53398 0.279871
SWI6 249.55 86.35 752.04 186 73.6 678 101 1.03 98.1 112.4 139.9697 1.245282
SWI7 253.2 84.05 768.06 227 102.1 666 70 61.4 1.1 125 1.169538 0.009356
SWI8 252.1 92.15 786.31 400 186.6 600 213.4
WI 224.09 106.65 458.82 77 8.9 450 47 14.55 3.2 68.1 3.648782 0.05358
W2 223 107.25 445.81 200 20 426 DRY 180
W3 224.45 106.75 476.59 167 21.4 455 44 53.3 0.9 145.6 0.9444 0.006486
W4 221.75 107.1 433.35 204 23.4 410 120 46 2.6 180.6 2.924588 0.016194
W5 220.81 107.7 425.94 237 44 382 110 25 4.6 193 5.371296 0.027831
W6 223.38 107.25 455 217 25.68 429 140 15 9.6 191.3 11.76307 0.06149
W7 223.09 107.18 446.68 430 167.2 280 262.9
W8 222.38 107.4 438.3 299
W9 219.38 108.1 350.31 166 10.15 340 134 0.58 231 155.9 348.6095 2.236109
WIO 219 104.35 260.76
WII 219.9 107.53 344.23 200 9.2 335 190.8
WI2 222.24 106.64 4~1.34 225 36 415 38 66.02 0.6 189 0.613099 0.003244
WI3 224.93 106.99 496.94 166 36.8 460 124.5 3.6 35 129.2 46.67831 0.361287
WI4 223.9 107.05 476.54 241 26.22 450 77 73.35 1.1 214.8 1.169538 0.005445
WI5 220.5 108.25 413.96 305 63.63 350 61.6 71.49 0.9 241.4 0.9444 0.003912
WI6 224.68 106.76 475.33 202 21.4 454 108 52.83 2 180.6 2.211353 0.012244
SMI 228.1 79.7 750 315 68.4 682 58 33.7 1.7 246.6 1.859747 0.007542
LAI 234.52 70.84 673.17 224 22 651 90 58.4 1.5 202 1.627556 0.008057
LAI(Observation) 234.52 70.8 673.1 250 22 651 6.5 228
LAIA 233.86 72.19 670.98 223 24.2 646 90 41 2.2 198.8 2.447721 0.012312
LA2 228.48 68.86 831.54 235 121 711 83.7 2.8 29.9 114 39.46738 0.346205
LA2(Observation) 228.4 68.8 831.54 155 121 711 0.9 34
LA3 224.94 67.71 873.21 265 121 752 44
LA4 228.51 67.28 839.69 243 123 717 61 28.8 2.1 120 2.329353 0.019411
LA4(Observation) 228.5 67.2 839.69 224 123 717 0.9 101
LA5 231.4 69.36 781.01 201
LA6 227.4 67.55 782 147
LA7 232 65.93 729.46 197 23.1 706 53 32.7 1.6 173.6 1.743414 0.010043
LA8 225.95 64.75 865.31 246 95.8 771 55 48.1 1.1 155 1.169538 0.007545
LA9 234.59 69.245 686.21 207 8.17 678 68.6 81.7 0.8 199 0.833015 0.004186
LA9(Observation) 234.5 69.24 686.21 201 8.17 678 33 192.8
LA 10 239.22 69.39 816.98 250 151.5 665 10 45 0.04 98.48 0.03423 0.000348
LA 11 239.93 67.96 826.05 227 147.6 678 79.4
387
388
Well No. Coordinate Groun Well Water Water Yield D.dwn. SC S. Th. T K
East North level depth depth level (m3th) (m) (m2th) (m) (m2th) (m/h)
PV37 248.87 107.325 758.38 282 190.6 568 60 26 2.3 91 2.566443 0.028203
PV38 244.14 103 694.38 255 123.7 571 62 1.27 48.8 131 66.51536 0.507751
PV39 247.9 113.66 713.68
PV40 252.65 110.4 746.67 350 199.3 547 60 22 2.7 150 3.044589 0.020297
PV41 252.625 115.39 710.11 275 155.7 554 60 1.95 30.8 125 40.73441 0.325875
PV42 252.52 116.07 703.84 300 149.7 554 75 0.3 250 150 379.2414 2.528276
PV43 259.53 106.49 790.73 297 210.4 580 50 17.75 2.8 87 3.164882 0.036378
PV44 258.575 108.825 774.56 316 197 578 60 46 1.3 119 1.397389 0.011743
PV45 254.315 106.805 779.09 350 32 46.78 0.68 0.700566
PV46 253.103 108.39 762.97 350 196.2 567 55 52 1.06 154 1.124279 0.007301
PV47 254.52 111.13 751.43 266.7 204 485 40 31.8 1.3 63 1.397389 0.022181
PV48 247.69 115.15 704.98 240 156.2 549 91 4.4 20.7 84 26.67332 0.317539
PV49 257.61 111.25 801.71 285 232.3 569 40 28.27 1.41 53 1.523714 0.028749
PV50 239.275 99.415 763.51 350 199 565 70 25.75 151
PV51 237.915 101.615 747.06 350 204.8 542 36 68.2 0.53 145 0.537188 0.003705
PV52 238.115 104.5 695.33 303 139.9 555 50 59.2 0.85 163 0.8886 0.005452
PV53 239.27 102.75 717.01 323 158.8 558 75 0.85 88.24 164 125.0309 0.762383
PV54 240.15 104.43 707.67 322 148.4 559 76 0.6 126.7 174 183.8311 1.056501
PV55 235.22 105.54 695.54 300 210 486 35 42 0.83 90 0.86634 0.009626
PV56 234.785 99.04 745.86 322 183.7 562 61 27.66 1.8 138 1.97653 0.014323
PV57 236.255 98.345 768.04 328 206.1 562 50 41.25 1.2 122 1.283152 0.010518
PV58 232.255 99.08 779.68 323 285 495 35 39 1 38 1.0566 0.027805
PV59 232.12 96.915 755.44 401 218.6 537 19 35.4 0.54 175 0.547994 0.003131
PV60 233.9 111.02 603.6 288 72 532 30 173 0.17 216 0.159939 0.00074
PV61 232.83 111.06 590 255 89.66 500 45 8 5.6 165 6.623766 0.040144
PV62 234.085 109.44 579.73 60
PV63 232.575 109.7 569.36 152 82.8 487 50 44.7 1.1 69 1.169538 0.01695
PV64 231.56 110.3 637.99 250 144.6 493 25 76.4 0.3 105 0.292943 0.00279
PV65 230.95 111.3 603.57 225 108.7 495 15 52.5 0.3 116 0.292943 0.002525
PV66 249.345 116.575 689.58 210 136 554 60 0.12 500 74 793.7124 10.72584
PV67 251.47 114.886 711.29 290 156.9 554 80 0.35 229 133 345.3944 2.596951
PV68 251.69 115.815 702.69 273 148.3 554 60 28.8 2.1 125 2.329353 0.018635
PV69 250.762 115.765 700.59 310 146.2 554 45 62.12 0.7 164 0.722541 0.004406
PV70 251.072 116.64 697.95 225 143.3 555 80 2.27 35 82 46.67831 0.569248
PV71 246.365 114 701.68 323 152.9 549 50 22.1 2.3 170 2.566443 0.015097
PV72 250.34 113.935 742.29 326 192.2 550 60 53.18 1.13 134 1.203554 0.008982
PV73 251.575 113.92 731.64 300 183.2 548 DRY 117
PV74 257.03 114.31 727.99 358 168.6 559 30 92.8 0.33 189 0.324255 0.001716
PV75 250.3 116.54 695.56 220 140.9 555 50 48.9 1 79 1.0566 0.013375
PV76 250.175 115.Q75 710.19 304 155 555 71 0.81 87.6 149 124.0648 0.83265
PV77 252.77 114.677 717.15 280 164 553 32 56 0.57 116 0.580491 0.005004
PV78 245.255 118.265 688.47 267 130.5 558 70 3.6 19.4 136 24.8922 0.183031
PV79 243.95 118.26 717.65 257 158.9 559 62 98
PV80 246.77 116.973 687.13 210 128.8 558 75 19.5 3.8 81 4.381976 0.054098
PV81 243.74 117.065 696.12 241 138.1 558 65 2.8 23.2 103 30.11883 0.292416
PV82 256.6 121.8 768.25 280 211.6 557 72.4 33 2.2 68 2.447721 0.035996
PV83 256.285 120.18 743.47 205 40
PV84 250.16 124.765 730.18 352 162.4 568 33 45.9 0.72 86 0.744558 0.008658
PV85 253.46 124.725 759.28 400 197 562 35 70 0.5 203 0.504851 0.002487
PV86 251.495 119.7 712.91 250 155.9 557 55 38.8 1.42 94 1.535231 0.016332
PV87 252 121.85 721.94 225 164.1 558 50 28.5 1.75 61 1.918084 0.031444
PV88 250.78 122.96 724.11 328 162.6 562 40 62.5 0.64 165 0.656743 0.00398
PV89 253.63 119.52 721.22 300 190.7 531 18 27.77 0.65 109 0.667682 0.006126
389
Well No. Coordinate Groun Well Water Water Yield D.dwn. SC S. Th. T K East North level depth depth level (m% ) (m) (m2/h ) (m) (m2/h) (m/h)
PV90 253.795 117.365 715.96 275 159.3 557 70 60 1.2 116 1.283152 O.oI 1062
PV91 254.575 116.295 713.77 248 160 554 60 31.5 1.9 88 2.093739 0.023792 PV92 249.72 117.53 698.69 245 142.8 556 65 31.95 2 102 2.211353 0.02168
PV93 256.45 116 719.04 275 163 556 65 9 7.2 112 8.657617 0.0773
PV94 247.88 122.45 705.94 345 144.7 561 38 56 0.68 190 0.700566 0.003687
PV95 246.96 120.05 694.84 250 133.9 561 18 23.5 0.77 116 0.799773 0.006895
PV96 248.745 120.585 723.25 300 162.1 561 40 40.8 0.98 138 1.034099 0.007493
PV97 247.92 117.95 690.9 200 139.5 551 60 0.18 333 60 514.7246 8.578744
PV98 247.95 119.66 701.9 342 132.2 570 DRY 210
PV99 247.325 118.92 705.9 297 150.8 555 76 16.4 4.6 148 5.371296 0.036293
PV100 248.265 118.875 695.64. 305 139.6 556 72 18.6 3.9 165 4.50495 0.027303
PV101 261.5 120.16 769.98 365 220.2 550 DRY 145
PV102 239.89 129.71 737.05 300 148.7 588 80 151
PV103 239.46 129.885 744.56 25
PV104 237.74 130.165 737.3 264 162.3 575 50 0.8 62.5 102 86.58063 0.84883
PV105 240.3 133.97 753.75 215 179.1 575 50 0.3 166.7 36 246.2538 6.840384
PV106 238.39 134.33 789.37 305 181.6 608 32 123
PV107 242.88 132.7 736.28 300 159.1 577 DRY 141
PV108 241.245 130 732.23 70
PV109 239.35 129.35 732.21 335 160.6 572 55 7 7.9 174 9.557235 0.054927
PV110 245 119.55 708.46 243 150.8 558 50 46.3 1.08 92 1.146895 0.012466
PV111 244.055 119.92 719.69 235 161.1 559 62 1.18 52.5 74 71.9019 0.971647
PVII2 143.135 120.45 719 400 163.3 556 20 67 OJ 237 0.292943 0.001236
PV113 242.878 121.02 723.25 250 165.4 558 40 23.5 1.7 85 1.859747 0.021879
PVI14 241.955 121.175 716.75 283 159.7 557 65 17.65 3.68 123 4.234688 0.034428
PVI15 237.94 123.285 705.88 350 140.8 565 70 1.4 0.5 209 0.504851 0.002416
PVI16 236.03 123.55 724.77 370 152.5 572 50 82.7 0.6 200 0.613099 0.003065
PVII7 239.1 75 120.11 785.08 270 121.3 664? 90 46 I 149 1.0566 0.007091
PVI18 242.065 123.315 714.61 270 152.2 562 40 2.35 17 118 21.62489 0.183262
PVII9 243.9 123.77 708.88 280 143.5 565 45 75.7 1.75 136 1.918084 0.014104
PV120 244.41 123.835 707.12 330 147 561 45 68 0.66 183 0.678633 0.003708
PV121 243.055 123.655 7J5.58 265 153.3 562 59 0.58 110.7 112 159.2025 1.421451
PVI22 243.89 123.135 714.99 266 151.2 564 55 1.9 28.9 115 38.0625 0.330978
PVI23 240.15 126.98 725.88 245 155.7 570 44 22.3 1.97 89 2.176027 0.02445
PV124 240.935 119.965 709.36 370 152.5 557 50 82.7 0.6 217 0.613099 0.002825
PV125 236.8 129.88 757.49 306 134 623 40 35.6 1.12 172 1.1 92209 0.006931
PV126 241.355 119.795 708.62 248 145.7 563 79 32.3 2.5 102 2.804889 0.027499
PV127 241.55 120.81 718.52 243 155.5 563 50 9.45 5.3 87 6.246353 0.071797
PV128 235.22 130.865 758.36 236 182.4 576 60 0.95 63.2 54 87.61423 1.622486
PV129 238 130.925 737.37 260 161.9 575 45 0.15 300 98 460.5569 4·69956
PV130 234.275 134.5 774.87 250 148.4 626 40 39.9 I 102 1.0566 0.010359
PVI31 230.26 131.88 773.82 308 219.1 555 80 3.25 24.6 89 32.05915 0.360215
PV132 229.86 130.225 769.57 350 219 551 10 131
PV133 234.29 131.14 761.93 385 196 566 32 44 0.73 189 0.755581 0.003998
PVI34 236.15 133.06 766.71 370 198.4 570 20 52.55 0.38 172 0.376851 0.002191
PV135 236.335 133.055 774.03 370 196.4 578 20 52.55 0.38 174 0.376851 0.002166
PV136 237.2 131.645 744.89 285 169.1 576 25 53.29 0.47 116 0.47264 0.004074
PV137 233.805 131.8 755.9 227 189.1 567 70 71.7 0.98 38 1.034099 0.027213
PVI38 234.645 132.04 786.31 300 209.8 577 18 90
PV139 235.34 131.59 769.04 271 195 574 60 0.5 120 76 173.4915 2.282783
PVI40 236.64 . 132.27 760.1 395 191 569 25 89 0.28 204 0.272181 0.001334
PVI41 235.3 124.71 741.19 316 177.2 564 DRY DRY 139
PV142 232.83 125.81 756.28 232 181.5 575 30 42 0.7 50 0.722541 0.014451
390
391
Well No. Coordinate Groun Well Water Water Yield D.dwn. SC S. Th. T K
East North level depth depth level (m3/h ) (m) (m2/h ) (m) (m%) (m/h)
S58 230.33 14.81 1025.8 200 173.1 852.7 11.87 7 1.7 27 1.859747 0.06888
S59 205.16 950.4 1317.8 250 34.5 1283 180 4.8 2.1 89 2.329353 0.026173
S60 204.1 950.5 1330.8 103 46.5 1284 104 0.8 133 57 193.5863 3.39625
S61a 204.63 951.38 1322.4 95 37.8 1285 111 1.9 59.55 57 82.23319 1.442688
S62 222.07 930.2 1188.9 155 140.5 1048 15
S63 214.24 989.69 1198.9 782 26
S65 210.14 977.52 1275.8 207 83 1193 63.5 14.8 4.29 71 4.986477 0.070232
S67 197.19 947.6 1450.1 290 50 1400 171
S68 235.26 945.24 993.9 240 79 914.9 101
S74 226.46 0.64 1070.2 145 81.5 988.7 82 0.4 390 64 609.1019 9.517217
S79 228.48 2.9 1044.6 155 56.8 987.8 155 2 77.5 67 108.8834 1.625125
S80 229.29 1.13 1045.3 159 57.5 987.8 10 39
S86 228.88 1.31 1047 201 59.1 987.9 157 1 131.9 80 191.8808 2.39851
S88 291.9 963.5 860 803 83 777 28
S94 196.23 941.14 1483.3 104 19.8 1464 III 5.2 21.3 84 27.49787 0.327356
S100 196.13 941.72 1494.3 163 30.9 1463 67 3.2 20.67 85 26.63213 0.313319
S101 207.78 977.13 1298.8 159 F 1299 8 80 0.1 12 0.090868 0.007572
S102 211.875 979.44 1218.4 363 37.1 1181 42 32 1.3 77 1.397389 0.018148
SI03 206.73 979.835 1316.2 37 12.2 1304 25
S104 206.55 976.935 1320.1 48 16.8 1303 120 0.3 445.5 32 701.8719 21.9335
SI05 215.53 981.225 1208.9 333 37.2 1172 23
S106 209.27 979.79 1270.5 203 67.9 1203 34
S107 213.435 982.155 1190.8 282 15.6 1175 57 44.2 1.29 58 1.385938 0.023895
S108 204.77 976.63 1358.3 62 49.4 1309 12
S109 214.85 980.93 1172.1 306 F+1 1173 153 46.1 3.32 80 3.794751 0.047434
SilO 209.7 982.34 1259.2 210 74.6 1185 14
Sill 211.65 981.41 1221 213 41.7 1179 1.5 62.5 0.024 13 0.019862 0.001528
S112 214.13 982.01 1184.6 266 14.7 1170 5
SI13 212.69 978.14 1227.8 210 45 1183 17
S115 215.39 970.56 1160.3 200 F+16 1176 19.2 16 1.2 26 1.283152 0.049352
S116 207.76 970.03 1287 170 34.5 1253 10
S117 207.71 977.7 1432.4 62 4 1428 58
S118 206.79 969.76 1302 259 31.3 1271 43 93
S121 228.7 2.08 1046.7 171 58.2 988.5 44.26 0.5 92.2 90 131.0181 1.455757
S136 195.23 947.54 1516.4 105 FO 1516 93.5 19 4.92 105 5.770314 0.054955
S137 199.94 958.24 1462.5 92 FO 1463 29 54.7 0.53 87 0.537188 0.006175
PP8 255 976.8 899 515 103 796 211
PP16 B3 243 969 860 95 DRY
PP19 B3 289.7 966.7 850.5 80 DRY
PP21 266.3 960.5 860.1 50 13 847.1
PP29 B3 261.9 981.06 882.2 70 42.3 839.9 28
PP37 245.4 24.5 812.1 120 22.7 789.4 26 30.2 0.9 97.3 0.9444 0.009706
PP38 252.6 20 832.3 103 31.4 800.9 113.6 0.13 847 59 1391.779 23.58948
PP40 269 5 913.2 191 113.6 799.6 77
PP41 251.21 20.43 830.5 148 29.8 800.7 203 24.8 8.2 118.2 9.944416 0.084132
PP42 250 21 828.3 150 28 800.3 210 0.2 1312 113 2218.549 19.63318
PP43 251.2 25 852.4 185 55.3 797.1 160 13.6 11.8 127 14.65551 0.115398
PP44 248.23 27.14 903.8 212 101 802.8 90 15 6 III 7.129035 0.064226
PP50 230.63 8.14 1002.3 172 43.9 958.4 41 6.6 6.2 57 7.382508 0.129518
PP51 215.32 22.91 1240 162 59 1181 8 84.2 0.095 103 0.086035 0.000835
PP52 218.78 22.24 1185 151 54 1131 56 40 1.4 97 1.512203 0.01559
PP55 223.5 0.3 1116.9 229 128 988.9 87 0.8 114 101 164.2641 1.626377
PP56 200.38 992 1435 66.5 34.1 1401 107 32
392
Well No. Coordinate Groun Well Water Water Yield D.dwn. SC S. Th. T K
East North level depth depth level (m3/h) (m) (m2/h) (m) (m2/h) (m/h)
PP57 200.74 991.31 1420 91 21.5 1399 119 2.1 57.7 70 79.51397 1.135914
PP58 200.46 989.67 1437 123 39.2 1398 113.6 2.1 55.1 84 75.70205 0.901215
PP59 200.4 988.53 1455 120 52.8 1402 81.5 2 40.75 67 54.89106 0.819269
PP60 200.81 989.58 1353 105 43 1310 99 0.2 660 62 1066.927 17.2085
PP61 204.07 983.25 1423.3 232 93.5 1330 70 32.1 2.18 138 2.424019 0.017565
PP62 198.78 980.82 1550 109 47 1503 62
PP63 206.94 973.14 1319.1 72 15.1 1304 58 29 2 24 2.211353 0.09214
PP64 203.5 967 1380 94 42 1338 106 3 35 52 46.67831 0.89766
PP65 211.85 961.39 1205.7 144 88.5 1117 115 3 38.3 136 51.38175 0.377807
PP66 195.95 947.18 1482.7 259 11.5 1471 100 75.2 1.33 109 1.431774 0.013136
PP67 197.7 935.6 1517 102 52 1465 49
PP90 219.33 26.2 1130 80 1050 11.58 89.1 0.13 145 0.120176 0.000829
PP457 248.4 27.06 900.8 230 115.6 785.2 31 22.4 1.4 114.4 1.512203 0.013219
PP449 255.1 14.2 837.4 97 42.03 795.4 37
PH01 221.76 941.17 1110.7 500 DRY
PH02 232.7 920.73 1119.5 500 DRY
PH03 Ku 249.24 909.38 999.8 350 148 851.8 149
Kh 999.8 306 693.8
PH04 A 255.11 935.2 905.2 304 162.2 743 107
PH05 278.29 956.58 860.2 400 71.1 789.1 93
PH06 250.15 923.37 948.2 200 DRY
PH07 293.7 927.61 893 300 DRY
PH08 278.55 919.89 899.9 330 DRY
PH09 AB 259.08 955.94 870.6 402 33.9 836.7 110
A 870.6 81.6 789 98
PH010 266.69 941.41 877.4 120 88.5 788.9 75
PHOll 247.81 951.86 905.5 322 116.2 789.3 89
PH012 311.4 967.45 869.5 200 125 744.5 41
PH013 A 300.98 940.29 871.8 175 136.8 735 38
PH014 250 962.9 889.2 284 100.2 789 114
PH015 233.08 960.11 979.3 365 187.6 791.7 116
PH016 292.01 963.29 856.9 200 96.4 760.5 54
PH017 272.47 949.11 861.2 380 72.3 788.9 69
PH018 233.96 947.71 985.5 310 188.9 796.6 8
PHT5 278.3 956.6 859.8 335 71.5 788.3 93 6.548175 0.7041
PHT9 259.07 955.91 870.6 290 81.8 788.8 100 2.353 0.02353
PHTI1 247.79 951.86 905.2 340 116.3 788.9 78 19.7596 0.2533
PHTI4 250.07 962.91 888.3 295 99.4 788.9 114 11.502 0.101
PHTI5 233.09 960.12 978.6 300 189.1 789.5 106
PHTI6 292.03 963.28 857.1 165 96.8 760.3 53 0.9444 0.01782
PHTI7 272.49 949.12 861 161 72.2 788.8 90
PHTI9 280.2 975.93 849.6 210 67.6 782 107
W22 205.48 987.28 1351 99 41 1310 122 61 2 58 2.211353 0.038127
W23 205.85 986.26 1346 101 45 1301 190 63.3 3.2 56 3.648782 0.065157
W24 207.45 977.95 1300 195 7 1293 94 58.8 1.6 188 1.743414 0.009273
W25 206.02 985.62 1358 162 51 1307 250 43.9 5.7 111 6.749868 0.06081
W26 200.18 987.7 1471 244 73 1398 150 10.6 9.4 171 11.50213 0.067264
W27 200.4 989.86 1437 122 40 1397 200 118 1.7 82 1.859747 0.02268
W28 200.2 992.8 1455 51 1404 25 25 1 1.0566
W32 206.89 968.86 1299 114 34 1265 78 21.1 3.7 80 4.259215 0.05324
W34 235.93 0.86 1080 102 102 978 100 14.5 6.9 8.273787
W35 219.34 26.3 1131 83 83 1048 5 50 0.1 0.090868
W36 220.7 27.9 1100 85 54 1046 140 34.1 4.1 31 4.751512 0.153275
393
* indicate that the well penetrate more than the B21A 7 aquifer system: A the whole Ajlun Group, A 112 the Na 'ur Formation, AB the Belqa and Ajlun Groups, B3 the Muwaqqar Formation, Ku Kurnub Group, Kh Khreim Group, D Disi Group.
394
Appendix (Bt) Definition of SCS Hydrologic Soil Groups (HSG) Source: Soil Conservation services (1986)
Group Description A Soils have low runoff potential and high infiltration rates even when
thoroughly wetted. They consist chiefly of deep, well to excessively drained sands or gravels and have a high rate of transmission (greater than 0.30 in./hr)
B Soils have moderate infiltration rate when thoroughly wetted and cosist chiefly of moderately deep to deep, moderately well to well drained soils with moderately fine to moderately coarse textures. These soils have a moderate rate of water transmission (0.15-0.30 in./hr).
C Soils have low infiltration rates when thoroughly wetted and cosist chiefly of soils with a layer that impedes downward movement of water and soils with moderately fine to fine texture. These soils have a low rate of water transmission (0.05-0.15 in./hr).
D Soils have high runoff potential. They have very low infiltration rates when thoroughly wetted and cosist chiefly of clay soils with a high swelling potential, soils with a permanent high water table, soils with a claypan or clay layer at or near the surface, and shallow soils over nearly imprevious mateaial. These soils have a very low rate of water transmission (0.0-0.05 in./hr)
Remarks Some soils in the list are in group D because of a high water table that creates a drainage problem. Once these soils are effectively drained, they are placed in a different group. For example, Ackerman soil is classified as AID. This indicates that the drained Ackerman soil is in group A and the undrained soil is in group D
395
Appendix (B2) Runoff curve numbers (eN) *Source: Soil Conservation service (1986) * Average runoff condition *Ia= 0.2S
396
Appendix (B2.1) Runoff curve number for Urban Areas
Cover description CNforHSG Cover type and hydrologic condition Impervious A B C D
area (%) a
Fully developed urban areas(veget. established) Open space(Lawn,parks,golf courses,cemetries) Y<> Poor condition(grass cover <50%) 68 79 86 89 Fair condition (grass cover 50%-75%) 49 69 79 84 Good condition (grass cover >75%) 39 61 74 80 Imprevious areas ..••••..••.•.•... · .•••• · ...•. >:i« ·.·.........i..«.-;Z Paved parking lots,roofs,driveaway, etc 98 98 98 98 Street and roads»< : ....... < ........................ ......... . .......•. :.: ......... : ... ...: Paved; curbsand storm sewers 98 98 98 98 Paved; open ditches (including right-of-way) 83 89 92 93 Gravel ( incuding right-of-way) 76 85 89 91 Dirt ( incuding right-of-way) 72 82 87 89 Western desert urban areas .•.••........... ..... } ./>
Natural desert landscaping (pervious area only) C 63 77 85 88 Artificial desert landscaping (impervious) 96 96 96 96 Urban dsert •••••.•••.. > .. . .. .................... :: ....... : ....... Commercial and businness 85 89 92 94 95 Industrial 72 81 88 91 93 Residential districts by average lot size »< 118 acre or less (town houses) 65 77 85 90 92 114 acre 38 61 75 83 87 113 acre 30 57 72 81 86 112 acre 25 54 70 80 85 1 acre 20 51 68 79 84 2 acre 12 46 65 77 82 Developing urban areas Newly graded areas(pervious areas only,no vege) 77 86 91 94 Idle lands(CN determined as in Appendix (B2.3» e
Note: a The average percent impervious area shown was used to develop the composite CN's. Other assumptions are as follows: impervious areas are directly connected to the drainage system, impervious areas have a CN of 98, and pervious areas are considered equivalent to open space in good hydrologic conditions. CN 's for other combinations of conditions may be computed. b CN's shown are equivalent to these of pasture. Composite CN's may be computed for other combinations of open space cover type. C Composite CN's for natural desert landscaping should be computed based on the impervious area percentage (CN=98) and the pervious area CN. The pervious area CN's are assumed equivalent to desert shrub in poor hydrologic condition. d Composite CN's to use for the design of temporary measures during grading and construction should be computed, based on the degree of development (impervious are percentage) and the CN's for the newly graded pervious areas.
397
Appendix (B2.2) Runoff curve number for cultivated Agricultural Lands
Cover description CNforHSG Cover type Treatment a Hydrologic
condition b
A B C
Fallow Bare soil 77 86 91 Crop residue cover (CR) Poor 76 85 90
Good 74 83 88 Row crops Straight row (Sr) Poor 72 81 88
Good 67 78 85 SR+CR Poor 71 80 87
Good 64 75 82 Contoured (C) Poor 70 79 84
Good 65 75 82 C+CR Poor 69 78 83
Good 64 74 81 Contoured & terraced(C &R) Poor 66 74 80
Good 62 71 78 C& T +CR Poor 65 73 79
Good 61 70 77 Small grain SR Poor 65 76 84
Good 63 75 83 SR+CR Poor 64 75 83
Good 60 72 80 C Poor 63 74 82
Good 61 73 81 C +R Poor 62 73 81
Good 60 72 80 C&T Poor 61 72 79
Good 59 70 78 C&T+CR Poor 60 71 78
Good 58 69 77 Close-seeded or SR Poor 66 77 85 broadcast Good 58 72 81 legumes or C Poor 64 75 83 ration meadow Good 55 69 78
C&T Poor 63 73 80 Good 51 67 76
Note: a Crop residue cover applies only if residue is on at least 5% of the surface throughout the year. b Hydrologic condition is based on combination offactors that affect infiltration
and runoff, incuding (a) density and canopy ofvegitative areas, (b) amount of year-round cover, (c) amount of grass or close-seeded legumes in rotations, (d) percent of residue cover on the land surface (goocP-.20%), and (e) degree of surface roughness. Poor,' Factors impair infiltration and tend to increase runoff. Good,' Factors· encourage average and better than average infiltration and tend to decrease runoff.
398
D
94 93 90 91 89 90 85 88 86 87 85 82 81 81 80 88 87 86 84 85 84 84 83 82 81 81 80 89 85 85 83 83 80
Appendix (B2.3) Runoff curve number for other Agricultural Lands
Cover description CNforHSG Cover type Hydrologic A B C D
condition Pasture, grassland, or range-continuous Poor 68 79 86 89 forage for grazing a Fair 49 69 79 84
Good 39 61 74 80 Meadow-continuous grass, protected 30 58 71 78 grazing and generally mowed for hay Brush-brush weed-grass mixture with Poor 48 67 77 83 brush the major element b Fair 35 56 70 77
Good 30c 48 65 73 Woods-grass combinatio (orchard or Poor 57 73 82 86 tree farm) d Fair 43 65 76 82
Good 32 58 72 79 Woods" Poor 45 66 77 83
Fair 36 60 73 79 Good 30 55 70 77
F armsteads-buildings,lanes,drivewayes, 59 74 82 86 and surrounding lots Notes: a Poor <50% ground cover or heavily grazed with no mulch. Fair: 50%-70% ground cover and not heavily grazed. Good:> 75% ground cover and lightly or only occassionally grazed. b Poor <50% ground cover. Fair: 50%-70% ground cover. Good:> 75% ground cover. C Actual curve number is less than 30; use CN=30 for runoff computations. d CN's shown were computed for areas with 50% woods and 50% grass (pasture) cover. Other combinations of conditions may be computed from the CN 's for woods and pasture. e Poor: Forest litter,small trees, and brush are destroyed by heavy grazing or regular burning. Fair: Woods are grazed but not burned, and some forest litter covers the soil. Good: Woods are protected from grazing, and litter and brush adequately cover the soil.
399
Appendix (B2.4) Runoff curve numbers for Arid and Semiarid Rangelands a
Cover description Cover type Hydrologic
condition b
Herbaceous-mixture of grass,weeds, and Poor low growing brush, with the minor Fair element Good Oak-asp en-mountain brush mixture of Poor oak brush,aspen,mountain mahogany, Fair bitter brush,malpe,and other brush Good Pinyon-Juniper-pinyon,juniper,or both; Poor grass undesroy Fair
Good Sagebrush with grass undestroy Poor
Fair Good
Desert shrub-major plants include salt- Poor brush,greasewood,creosotebrush,bursage Fair palo verde,mesquite, and cactus Good Notes: a For range in humid regions, use Table 3B3 b Poor: <30% ground cover (/itter,grass, and brush overstory). Fair:30%-70% ground cover. Good: >70% ground cover
400
CNforHSG A B C D
80 87 93 71 81 89 62 74 85 66 74 79 48 57 63 30 41 48 75 85 89 58 73 80 41 61 71 67 80 85 51 63 70 35 47 55
63 77 85 88 55 72 81 86 49 68 79 84
Appendix (B3) Surface Water in Jordan (Data are compiled from WMP and WAJ)
Group Drainage Area Areas Rainfall e Estimated Streamfow e (lan2
) mm MCM Runoff Baseflow A I Yarmouk River 6790 a 370 2512 182 218
Zerqa River 3530 b 219 773 35.85 31.5 II W. Wala 2030 c 189 383 29.4 21
W. Mujib 4500d 128 576 47.64 20 W. Hasa 2198 92 202 13.8 25.5
B I Small areas 1017 157 4.57 5.1 W.Arab 267 467 124.7 6.48 24.9 W. Ziglab 106 494 52.4 2.2 8.3 W.Jurum 22 429 9.6 0.23 11.5 W. Yabis 124 525 65.1 1.63 6.2 W. Kufrinja 111 542 60.2 1.02 5.8 W.Rajib 85 515 43.8 1.31 3.0 W. Shu'eib 178 398 70.8 1.77 8.0 W. Kafrein 189 397 75 1.35 12 W. Hisban 82 312 18.1 0.34 6.3
II Small areas 871 203 177 1.03 29 W. Zarqa Main 272 302 82.1 2.96 20 W.Karak 190 278 52.8 3.17 15
C I Small areas 1306 24.9 0.19 0.25 W. Feifa 161 206 33.2 1.16 10 W. Khuneizir 183 234 43 1.18 3 W. Dahl 96 197 18.9 0.3 0.04 W. EI-Feidan 280 235 65.8 1.32 4.1 W.E1-Buweirida 518 213 54.3 2.44 0.8 W.Musa 165 168 27.7 0.14 2.6 W.Huwar 229 144 32.9 0.29 0.5
II Small areas 960 19.4 0.29 W. Abu Barqa 136 139 18.9 0.22 .12 W. Rakiya 182 128 23.3 0.09 0.2 W. Yutum 4443 67 296 5.98
D Azraq 11588 90 1043 12.52 15.1 EI Jafr 13427 51 685 22.9 W.Hammad 19271 W. Sirhan 15155 Southern Desert 4153 45 187 3.4
E W.Dhuleil North 1305 171 223.2 7.59 W. Dhuleil South 500 147 73.5 2.5 W.Hammam 340 139 47.26 2.17 W. Siwaqa 520 84 43.44 2.02 W. Qatrana 1300 64 83.2 4.4 W. Sultaneh 680 110 74.8 3.45
Notes: a: Including 4791 km2 of Syrian Territory.
b: including 1305 km2 of Wadi Dhuleil North in Group D. W. Dhuleil South in Group D were excluded as its restricted drainage area.
c: including 340 km2 of Wadi Hammam in Group D.
d: including 520 km2 of W. Siwaqa, 1300 km2 of w. Qatrana, and 680 km
2 of W.
Sultaneh of Group D ..
e: Figures are the mean annual in MeM.
401
Runoff Total (%) 400 7.3 67.35 4.6 50.4 7.7 67.64 8.3 39.3 6.8 9.67 31.38 5.2 10.5 4.2 11.37 2.4 7.83 2.5 6.82 1.7 4.31 3.0 9.77 2.5 13.35 1.8 6.64 1.9 30.03 0.6 22.96 3.6 18.17 6 0.44 0.8 11.16 3.5 4.18 2.8 0.34 1.6 5.42 2 3.24 4.5 2.74 0.5 0.79 0.9 0.29 1.5 0.34 1.2 0.29 0.4 5.98 2 27.62 1.2 22.9 3.3
3.4 1.8 7.59 3.4 2.5 3.4 2.17 4.6 2.02 4.66 4.4 5.3 3.45 4.6
Appendix (B4) Runoff measurments in the study area. Source: WAJ records, Jordan.
402
. I
Appendix (B4.1) Runoff measurments for Zerqa River at Sukhna Gauging Station in MeM.
year Oct Nov Dec Jan Feb Mar Apr May Jun Jul Aug Sep
71/72 BF 0.05 0.06 0.90 1.89 1.60 1.26 1.73 2.10 1.29 0.95 0.70 0.45
FF 4.34 0.86 1.30 1.13 1.17 0.53
72/73 BF 0.35 0.31 0.60 0.57 1.51 4.00 0.41 0.29 0.17 0.09 0.08 0.08
FF 0.15 1.77 0.18 1.34
73/74 BF 0.08 2.58 1.16 2.08 4.41 5.60 3.47 2.09 1.58 1.15 0.67 0.26
FF 0.72 0.22 31.4 6.00 0.80
74/75 BF 0.29 0.48 1.30 1.25 1.47 1.33 0.61 0.47 0.32 0.27 0.26 0.27
FF 0.32 0.17 3.10 1.28
75/76 BF 0.10 0.31 0.54 1.97 2.15 1.70 0.49 0.20 0.0 0.0 0.0 0.0
FF 0.15 2.02
77/78 BF 0.0 0.0 0.41 0.67 0.78 0.85 0.66 0.22 0.10 0.06 0.0 0.0
FF 0.90 2.55 0.23 3.03
78/79 BF 0.0 0.0 0.53 0.58 0.40 0.54 0.15 0.0 0.0 0.0 0.0 0.0
FF 1.42 1.66 0.26
82/83 BF - - - - 2.13 4.14 2.25 1.76 1.69 1.75 1.66 1.61
FF 9.66 4.17
83/84 BF - - - - - - - -- - - - -FF 0.22 0.97 0.91 2.07
Source: WAJ records, Jordan.
403
Annual
12.98
9.33
8.45
3.43
25.13
39.13
8.34
4.87
7.45
2.17
3.75
6.71
2.19
3.34
---4.17
Appendix (B4.2) Runoff measurments for Wadi Wala at Karak Road in MeM.
Year Flow Oct Nov Dec Jan Feb Mar Apr May Jun Jul Aug Sep Ann. 63/64 BF 0.14 0.22 0.59 0.63 0.39 0.37 0.38 0.21 0.16 0.16 0.16 0.16 3.56
FF 35.9 7.22 4.02 0.05 47.21 64/65 BF 0.14 0.22 0.59 0.62 0.38 0.36 0.38 0.21 0.16 0.16 0.16 0.16 3.54
FF 0.30 8.36 1.89 10.54 65/66 BF 0.16 0.16 0.16 0.16 0.13 0.12 0.16 0.11 0.13 0.13 0.11 0.10 1.63
FF 0.68 0.17 3.65 4.50 66/67 BF 1.15 2.70 2.82 1.21 0.03 0.14 0.14 0.13 0.10 0.20 0.10 0.10 8.83
FF 2.30 4.07 6.88 4.60 1.84 19.69 67/68 BF 0.29 0.42 0.34 0.55 0.43 0.34 0.29 0.28 0.28 0.21 0.08 0.10 3.59
FF 0.38 0.45 0.37 0.43 1.81 68/69 BF 0.23 0.96 5.66 4.97 0.36 0.37 0.98 0.32 0.32 0.43 0.44 0.54 15.60
FF 1.52 2.23 1.38 12.0 3.37 20.53 69.70 BF 0.07 0.09 0.10 0.11 0.10 0.10 0.08 0.07 0.06 0.09 0.09 0.08 1.03
FF 0.79 0.48 1.27 70171 BF 0.40 0.41 0.43 0.44 0.40 0.37 0.40 0.30 0.19 0.18 0.19 0.18 3.89
FF 1.53 0.75 45.8 48.07 71172 BF 0.24 0.37 0.41 0.34 0.33 0.25 0.24 0.22 0.28 0.23 0.23 0.18 3.34
FF 0.12 23.0 23.13 72173 BF 0.19 0.18 0.19 0.20 0.15 0.19 0.22 0.25 0.21 0.12 0.08 0.08 2.06
FF 0.00 73174 BF 0.08 0.19 0.27 0.46 0.57 0.60 0.43 0.34 0.26 0.27 0.27 0.23 3.98
FF 1.95 0.71 11.3 3.89 0.6 18.42 74175 BF 0.24 0.29 0.53 0.47 0.45 0.49 0.36 0.28 0.22 0.16 0.13 0.08 3.71
FF 0.01 75176 BF 0.08 0.11 0.09 0.08 0.08 0.14 0.08 0.08 0.08 0.08 0.08 0.08 1.05
FF 0.01 7.05 7.06 76177 BF 0.27 0.26 0.27 0.27 0.24 0.28 0.23 0.20 0.20 0.18 0.12 0.15 2.67
FF 2.37 2.37 77178 BF 0.44 0.64 0.86 0.71 0.55 0.71 0.55 0.26 0.15 0.13 0.13 0.12 5.25
FF 0.03 2.53 4.83 0.77 8.16 78179 BF 0.17 0.19 0.28 0.31 0.27 0.25 0.29 0.23 0.17 0.13 0.17 0.17 2.62
FF 0.29 2.04 2.33 79/80 BF 0.23 0.32 0.45 0.60 0.32 0.33 0.27 0.35 0.24 1.24 1.69 0.17 6.22
FF 15.6 24.6 3.93 13.3 60.69 80/81 BF 0.17 0.17 0.27 0.41 0.56 0.49 0.37 0.30 0.22 0.19 0.16 0.16 3.46
FF 46.9 1.34 1.94 50.16 81182 BF 0.21 0.16 0.16 0.19 0.30 0.48 0.50 0.54 0.39 0.28 0.23 0.20 3.62
FF 1.42 20.9 10.8 33.17 82/83 BF 0.19 0.31 0.35 0.44 0.63 0.66 0.48 0.40 0.34 0.27 0.24 0.23 4.54
FF 1.71 0.12 4.18 3.40 13.0 22.40 83/84 BF 0.19 0.18 0.24 0.27 0.33 0.47 0.26 0.24 0.21 0.19 0.16 0.16 2.89
FF 2.56 2.56 84/85 BF 0.16 0.23 0.27 0.27 0.36 0.62 0.51 0.34 0.26 0.24 0.24 0.21 3.71
FF 1.39 1.94 18.4 39.4 60.07 85/86 BF 0.23 0.25 0.46 0.39 0.66 0.50 0.70 0.08 0.05 0.08 0.05 0.05 3.51
FF 0.43 1.33 0.18 1.93 86/87 BF 0.05 0.14 0.23 0.40 0.34 0.26 0.18 0.19 0.05 0.06 0.05 0.05 2.01
FF 9.60 0.89 2.35 12.85 87/88 BF 0.15 0.13 0.19 0.29 0.27 0.29 0.18 0.05 0.03 0.03 0.03 0.05 1.68
FF 3.37 0.08 6.32 3.91 5.08 19.76 88/89 BF 0.12 0.10 0.18 0.23 0.16 0.16 0.15 0.13 0.08 0.08 0.06 0.06 1.50
FF 9.05 1.52 10.57 Ave. BF 0.23 0.36 0.63 0.58 0.34 0.36 0.34 0.24 0.19 0.21 0.21 0.15 3.84
FF 0.2 1.3 5.2 2.2 1.8 4.7 3.3 0.5 19.1
Source: WAJ records, Jordan.
404
Appendix (B4.3) Runoff measurments for Wadi Wala at weir in MeM.
year Flow Oct Nov Dec Jan Feb Mar Apr May Jun Jul Aug Sep Ann. 71/72 BF 0.24 0.35 0.30 0.31 0.33 0.25 0.24 0.21 0.28 2.54
FF 0.16 23.2 0.03 23.46 72173 BF 0.11 0.20 0.24 0.33 0.30 0.36 0.34 0.31 0.27 0.14 2.62
FF om om 73/74 BF 0.75 0.41 0.19 0.15 0.39 0.21 0.15 0.15 0.10 0.14 0.14 2.69 5.45
FF 2.04 0.82 11.5 3.98 0.60 18.94 74/75 BF
FF 75/76 BF 0.08 0.11 0.09 0.08 0.08 0.20 0.08 0.08 0.08 0.08 0.08 0.08 1.12
FF 7.05 7.05 76177 BF 0.32 0.27 0.58
FF 2.34 2.34 77/78 BF 0.44 0.64 0.75 0.70 0.55 0.67 0.55 0.45 0.42 0.54 0.44 0.28 6.43
FF 0.07 1.69 3.26 0.53 5.54 78/79 BF 0.09
FF 1.45 79/80 BF
FF -80/81 BF 0.24
FF 41.76 81/82 BF 0.54 0.50 0.42 0.29 0.29 0.29 0.29 2.91
FF 3.53 2.13 6.09 82/83 BF 0.54 0.53 0.54 0.58 0.54 0.34 0.29 0.29 0.29 0.15 0.29 0.29 4.68
FF 0.98 3.18 6.03 23.6 33.77 83/84 BF 0.38 0.44 0.46 0.46 0.44 0.51 0.40 0.40 0.39 0.29 0.29 0.29 4.74
FF 1.77 1.77 84/85 BF 0.39 0.51 0.50 0.46 0.41 0.45 0.43 3.18
FF 1.37 3.37 3.01 10.5 1.94 20.15 Ave. BF 0.37 0.40 0.38 0.38 0.38 0.39 0.33 0.29 0.27 0.23 0.26 0.65 2.90
FF 0.20 0.80 3.60 2.20 4.10 3.80 0.40 0.30 15.30 Source: WAJ records, Jordan
Appendix (B4.4) Runoff measurments for Wadi Swaqa in MeM.
year Flow Oct Nov Dec Jan Feb Mar Apr May Jun Jul Aug Sep Ann. 62/63 FF 0.07 3.71 0.04 3.83 63/64 FF 7.27 7.27 64.65 FF 0.13 0.52 0.04 0.69 65166 FF 1.79 0.79 0.24 0.02 1.63 4.45 66/67 FF 1.73 1.50 3.23 67/68 FF 0.17 0.12 0.15 1.55 2.00
68/69 FF 1.35 0.26 1.61
69/70 FF 0.67 0.03 3.30 4.00
70171 FF 0.04 3.29 0.09 3.42
71172 FF 0.05 0.05
72/73 FF 0.04 3.51 0.16 0.09 3.80
73/74 FF 1.23 1.23
74/75 FF 1.75 1.75
75/76 FF 0.00
76/77 FF 0.82 0.82
77/78 FF 0.00
78/79 FF 0.51 0.97 0.84 10.1
79/80 FF 0.84
80/81 FF 0.84 0.85
81182 FF 0.17 35.2 0.37 35.72?
82/83 FF 0.00
83/84 FF 3.01 0.07 1.16 4.24
Ave. FF 0.30 0.2 0.50 0.7 0.1 0.3 2.00 0.2 4.30
Source: WAJ records, Jordan
405
Appendix (B4.S) Runoff measurments for Wadi Mujib at Karak Road in MeM.
year Flow Oct Nov Dec Jan Feb Mar Apr May Jun Ju1 Aug Sep Ann. 64/65 BF 0.14 0.34 0.72 0.54 0.61 0.36 0.27 0.25 0.21 0.13 3.58
FF 107 0.05 0.80 108.18 65166 BF 0.56 0.65 0.64 0.66 0.48 0.39 0.26 0.22 0.21 0.31 4.38
FF 0.1 1.69 1.79 66/67 BF 1.03 1.54 1.44 0.56 0.74 1.22 0.59 0.37 0.36 0.34 0.29 0.29 8.78
FF 7.03 12.0 4.15 1.53 5.18 29.87 67/68 BF 0.46 0.72 0.86 0.92 0.87 082 0.60 0.52 0.27 0.19 0.08 0.05 6.36
FF 6.17 0.46 2.15 0.65 1.65 11.08 68/69 BF 0.21 0.41 0.67 0.64 0.48 0.63 0.67 0.11 3.83
FF 0.20 1.37 1.05 0.22 0.02 13.1 15.95 69170 BF 0.16 0.29 0.46 0.55 0.37 0.73 0.35 0.04 0.Q3 0.Q3 0.03 0.Q3 3.04
FF 1.19 7.48 0.03 8.70 70171 BF 0.19 0.27 0.60 0.92 0.47 0.46 1.25 1.20 0.13 0.13 0.13 0.13 5.89
FF 0.1 10.6 0.67 40.3 51.70 71172 BF 0.03 0.26 1.45 1.41 0.93 1.13 0.81 0.55 0.Q3 0.03 0.03 0.Q3 6.68
FF 0.09 37.9 3.89 7.26 2.70 0.05 0.55 52.39 72173 BF 0.03 0.29 1.79 1.27 0.35 0.37 0.41 0.27 0.12 0.11 0.11 0.10 5.22
FF 0.63 2.79 3.41 73174 BF 0.08 0.42 0.16 0.72 0.68 0.70 0.43 0.40 0.31 0.23 0.19 0.18 4.50
FF 2.89 3.54 2.06 8.49 74175 BF 0.18 0.44 0.58 0.51 0.58 0.52 0.27 0.13 0.05 0.05 0.05 3.37
FF 1.95 0.06 20.9 0.17 23.12 75176 BF 0.54 0.52 0.54 0.54 0.45 0.56 0.50 0.35 0.26 0.27 0.21 0.21 4.95
FF 13.0 13.00 76177 BF 0.24 0.32 0.45 0.62 0.46 0.34 0.46 0.43 0.32 0.23 0.20 0.23 4.31
FF 0.00 77178 BF 0.24 0.26 0.31 0.36 0.35 0.40 0.45 0.39 0.33 0.27 0.22 0.22 3.79
FF 0.37 8.48 0.32 9.17 78179 BF 0.24 0.27 0.37 0.62 0.51 0.63 0.47 0.30 0.10 0.05 0.03 0.Q3 3.62
FF 0.32 0.26 0.58 79/80 BF 0.58 0.87 0.78 0.76 0.43 0.10 0.01 0.01 0.01 0.00 3.56
FF 1.43 10.8 9.45 1.92 23.60 80/81 BF 0.01 0.30 0.36 0.42 0.58 0.15 0.10 0.02 0.D2 1.96
FF 16.7 1.11 17.81 81/82 BF 0.01 0.11 0.16 0.04 0.04 om 0.01 om 0.04 1.08
FF 0.14 0.19 18.1 18.46 82/83 BF 0.01 0.20 0.52 0.69 0.93 0.50 0.16 0.01 om 3.04
FF 0.D2 0.D2 0.49 0.04 0.13 0.71 83/84 BF 0.14 0.13 0.43 0.43 0.68 1.20 0.96 0.62 0.45 0.34 0.21 0.16 5.76
FF 0.00 84/85 BF 1.44 1.42 1.54 1.44 1.23 0.58 0.27 0.21 0.13 -
FF 0.97 0.44 -85/86 BF 0.21 0.27 0.37 0.43 0.42 0.48 0.45 0.43 0.29 0.21 0.13 0.13 3.82
FF 0.02 0.32 0.05 0.15 0.54 86/87 BF 0.14 0.28 0.31 0.26 0.41 0.46 0.38 0.30 0.24 0.19 0.18 -
FF 0.55 0.21 0.01 3.89 -87/88 BF 0.29 0.13 0.14 0.44 0.44 0.47 0.28 4.63 0.08 0.02 0.02 0.33 7.27
FF 15.7 17.6 3.03 24.6 0.03 60.97 88/89 BF 0.24 0.31 0.32 0.37 0.34 0.38 0.34 0.32 0.25 0.24 0.24 0.19 3.54
FF 1.85 8.04 0.20 0.05 10.14 Ave. BF 0.25 0.34 0.56 0.66 0.58 0.65 0.52 0.55 0.20 0.16 0.14 0.14 4.45
FF 0.30 1.30 2.90 1.90 1.70 2.20 10.0 0.10 20.42 Source: WAJ records. Jordan
406
Appendix (B4.6) mean annual observed flood flow of Hasa River at Tannur in MeM.
year Rainfall Runoff runoff Coefficient (%) 68/69 158.3 2.42 1.6 69170 114.3 5.94 5.3 70/71 162.7 13.20 8.2 71172 296.7 9.45 3.1 72173 87.9 2.20 2.5 73174 277 7.03 2.5 74175 281.3 10.33 3.7 75176 105.5 - -76177 136.3 0.66 0.5 77178 167.1 1.32 0.8 78179 118.7 1.98 1.7 79/80 294.5 38.47 13.1 80/81 206.6 14.73 7.2 81182 164.9 7.25 4.4 82/83 246.2 5.06 2.0 83/84 189.0 0.88 0.5 84/85 151.7 4.62 3.1 85/86 195.6 11.0 5.6 Ave. 186.8 8.03 4.3
Source: WAJ records, Jordan
407
Appendix (B4.7) Observed runoff discharge of Hasa River at Ghor Saft in MeM.
year Flow Oct Nov Dec Jan Feb Mar Apr May Jun Jul Aug Sep Ann. 63/64 BF 2.14 2.03 2.66 2.98 3.03 3.44 3.08 2.44 1.92 2.29 2.17 1.63 29.82
FF 8.61 8.61 64/65 BF 1.98 1.89 1.93 2.16 2.46 2.93 2.53 2.40 2.00 2.02 2.04 1.97 26.30
FF 0.00 65/66 BF 2.17 2.18 2.28 2.28 2.13 2.36 2.13 2.04 1.87 1.87 1.85 2.02 25.16
FF 2.06 0.54 2.60 66/67 BF 2.20 2.66 3.42 3.83 4.70 8.33 2.40 2.56 1.99 2.12 2.12 2.05 38.37
FF 5.44 14.9 0.52 0.77 0.94 9.22 8.57 40.32 67/68 BF 1.01 5.30 3.44 3.81 3.98 2.62 2.68 4.74 3.89 3.76 3.79 3.56 42.58
FF 2.31 26.3 1.65 2.38 0.38 0.79 0.35 34.21 68/69 BF
FF 69170 BF
FF 70171 BF
FF 71172 BF
FF 72173 BF 2.17 2.20 2.28 2.28 2.06 2.42 2.15 2.17 2.29 2.26 2.14 1.94 26.35
FF 1.51 0.14 1.65 73174 BF 2.25 2.25 1.87 2.27 2.50 2.87 2.11 2.13 2.05 1.79 1.79 1.66 25.56
FF 0.45 1.23 2.99 1.03 5.70 74175 BF 2.41 2.50 3.07 2.76 2.77 2.99 2.37 2.04 1.87 1.87 1.85 2.02 28.51
FF 1.07 10.5 11.60 75176 BF 2.09 2.07 2.21 2.33 2.21 2.32 1.89 1.65 1.43 1.47 \.34 \.30 22.31
FF 0.00 76177 BF \.34 1.43 1.66 2.00 2.22 2.16 2.13 2.22 \.32 1.21 1.08 1.26 20.03
FF 0.08 0.94 1.02 77178 BF 1.87 2.05 2.47 2.15 1.95 2.11 2.21 2.25 1.77 1.77 1.80 1.79 24.19
FF 0.87 0.04 0.91 78179 BF 1.88 1.91 2.06 2.13 1.91 2.37 2.23 2.29 2.12 2.01 2.01 1.94 24.87
FF \.34 0.23 1.58 79/80 BF 1.96 2.05 2.57 2.28 2.11 3.28 2.56 2.36 2.20 2.30 1.96 1.62 27.25
FF 3.74 2.98 23.5 8.15 1.26 39.67 80/81 BF 1.82 1.94 3.41 2.28 2.00 2.63 2.43 2.23 2.07 2.14 2.14 2.07 27.18
FF 0.45 15.0 0.82 0.07 16.37 81182 BF 2.41 2.73 2.62 2.96 2.87 2.77 2.92 2.90 2.27 2.35 2.24 2.16 31.20
FF 0.50 1.28 0.13 0.40 4.54 6.85 82/83 BF 2.05 2.15 2.54 2.34 2.31 2.74 3.21 2.94 2.25 2.14 2.32 2.33 29.33
FF 3.58 3.58 83/84 BF 2.08 2.52 2.24 2.19 2.42 2.75 2.84 2.61 2.31 2.16 2.31 1.99 28.41
FF 0.00 84/85 BF 2.03 2.05 2.20 2.29 2.15 2.47 2.32 2.23 1.99 1.90 1.89 1.98 25.50
FF 0.00 85/86 BF 2.05 2.13 2.34 2.22 2.03 2.28 2.23 2.34 2.29 2.25 2.09 1.86 26.11
FF 0.00 86/87 BF 1.93 2.18 2.46 2.33 2.24 2.53 2.36 2.50 2.24 2.12 1.91 2.05 26.84
FF 0.00 87/88 BF 2.06 2.12 2.04 2.54 2.56 3.07 2.40 2.32 2.09 1.94 1.99 1.87 26.99
FF 0.00 88/89 BF 2.32 1.96 1.94 2.19 2.20 2.47 2.38 2.45 1.84 2.18 2.01 1.91 25.86
FF 0.00 Ave. BF 2.01 2.29 2.44 2.48 2.49 2.91 2.44 2.45 2.09 2.09 2.04 1.95 27.67
FF 0.64 2.49 0.94 1.41 1.73 0.62 0.11 0.65 8.44 Source: WAJ records, Jordan
408
Appendix (B4.8) Mean annual observed flood flow of Wadi Jurdhan in MeM.
year Rainfall Runoff runoff Coefficient (%) 63/64 48.4 1.12 2.3 64/65 57.3 1.38 2.4 65/66 18.0 0.62 3.4 66/67 27.8 0.40 1.4 67/68 16.9 0.04 0.2 68/69 26.4 1.52 5.8 69/70 12.7 0.02 0.2 70/71 24.6 1.15 4.7 71172 25.8 0.26 1.0 72/73 8.0 0.18 2.3 73/74 38.9 0.13 0.3 74/75 31.5 0.73 2.3 75/76 10.9 0.0 76/77 5.8 0.0 77/78 21.3 0.0 78/79 23.5 0.0 79/80 44.8 1.01 2.3 80/81 26.9 0.67 2.5 81182 23.0 0.044 0.20 Ave. 26 0.5 1.9
Source: WAJ records, Jordan
409
Appendix (Cl) Results of pumping test analysis in the B2/A7 aquifer system.
410
Well No. Duration Saturated S.C T-D-down T-Recov. K (hours) Tickness (m) (m3/h/m) (m2/d) (m
2/d) (mid) A93 16 0.1 - 3.5 A 100 16 42 13.7 - 467 0.11 AIOI 16 100 127 - 4320 0.43 A 102 120 100 103 3888 2592 32.4 AI03 120 35 6.5 207 233 6.29 A 104 20 16 165 5616 - 351 AI05 20 31 51 1728 1728 55.74 AI06 20 28 66 2246 - 80.21 AI08 12 80 14 - 475 5.94 AI09 12 0.08 2.6 2 AIIO 20 100 2.5 0.86 - 0.01 Alii 10 100 0.6 II 2.6 0.07 AI12 15 75 2.9 77.8 25.9 0.69 AI13 16 3.1 104 104 AI14 20 I - 34.6 AI15 10 50 1901 1469 AI16 16 65 0.15 6.9 1.7 0.07 AI17 20 56 1.9 86.4 56.2 1.27 AI18 16 25 I - 34.6 1.38 AI19 16 82 I - 34.6 0.42 AI20 16 9 5.6 - 190 21.11 ···.\/ WAPIWWIlA:HWArn·.·. MP:@:AREA; r·<··········· .
Tests carried out by Parker (1970) PP80 3.25 89 4.72 43.4 0.50 PP85 48 121 0.48 15 12.9 0.12 S64 49 125 2.49 63 67 0.52 S66 47 96 3.5 86.4 71 0.82 S69 26 25 19.7 838 33.52 S70 34 48 11.04 210 608 8.52 S71 4 184 0.25 13.4 0.Q7 S75 12 153 32.2 1250 8.17 S76 24 121 21.6 805 6.65 S77 160 80 0.19 3.5 2.9 0.04 S78 0.25 116 0.64 9 0.08 S83 24 107 792 56300 526.17 S84 45 148 7.2 216 205 1.42 S93 24 120 13.8 426 1518 8.1 S95 26 63 443.3 14800 15500 240.48 S96 24 100 39.4 790 851 8.21 S97 5 161 2.18 35 73 0.34 SI19 23 122 305.2 4900 18900 97.54 SI31 24 140 108 4460 31.86 SI32 24 112 9.08 423 3.78 SI34 24 134 3.53 32 44 0.28
Tests carried out by BGR (1987) LAI 72 228 1.1 36 45 0.18 LAIA 73 199 2.2 20 0.1 LA2 72 114 30 1440 12.63 LA4 69 119 2.3 447 728 4.94 LA7 72 175 1.6 34.5 0.2 LA8 48 155 1.1 30.6 0.2 LA9 168 199 0.84 23 0.12 LAI3 432 140 0.7 23.6 79 0.37 LAI4 0.45 157 0.05 0.43 0.003 LAI9 II 65 4 18 0.28 LA20 44 66 2.7 60 0.91 LA21 71 172 1.2 202 1.17
411
Well No. Duration Saturated S.C T-D-down T-Recov. K (hours) Tickness (m) (m3/h/m) (m2/d) (m2/d) (m/d) Tests carried out by JICA (1987)
412
Appendix (Dl) Soil moisture balance (mm) for West Amman sub-catchment for the water year 1982/1983.
Illi: P P-R PET £. P-R-£. L (I=P-R- £.-0) 1lli:1~ 0 0 1.67 0.17 '{).I7 -100.17
2 0 0 1.67 0.17 '{).I7 -10033
3 0 0 1.67 0.17 '{).I7 -1005
4 837 7.42 1.67 6.84 057 .lE.93
5 43 3.81 1.67 3.6 021 .lE.71
6 1.16 1.03 1.67 1.00 .{).(Xi .lE.78
7 0 0 1.67 0.17 '{).I7 .lE.94
8 0 0 1.67 0.17 '{).I7 -100.11
9 0 0 1.67 0.17 '{).I7 -10028
10 0 0 1n7 0.17 '{).I7 -100.44
11 0 0 1.67 0.17 .{).17 -100.61
12 0.84 0.74 1.67 0.83 .{).OO -100.7
13 0.63 055 1.67 0.67 '{).I1 -100.82
14 0 0 In7 0.17 '{).I7 -Ioo~
15 3.% 351 1.67 333 018 -100.8
16 0 0 1.67 0.17 '{).I7 -100.97
17 0 0 1.67 0.17 '{).I7 -101.13
18 0 0 1b7 0.17 '{).I7 -1013
19 0 0 1b7 0.17 '{).I7 -101.47
~ 0 0 1.67 0.17 '{).I7 -101.63
21 2:f) 23 In7 224 0.<Xi -10157
22 03 026 1.67 0.4 '{).I4 -101.71
23 0 0 1.67 0.17 '{).I7 -101.88
24 0 0 1.67 0.17 '{).I7 -102.05
25 0 0 In7 0.17 '{).I7 -102.21
26 0 0 1.67 0.17 '{).I7 -102.38
27 0 0 1.67 0.17 '{).I7 -102.55
28 0 0 In7 0.17 '{).I7 -102.71
29 3.67 325 1.67 3.00 0.16 -102.56
30 8.12 7.19 1.67 6.64 055 -102
31 3.89 3.44 1.67 326 0.18 -101.83
Ihll'ID 2828 25JXi 1.74 1.74 2332 -78.51
2 7.42 657 1.74 1.74 4.83 -73n7
3 036 032 1.74 1.74 -1.42 -75ffJ
4 3200 28.43 1.74 1.74 26.(9 48.4
5 3.8 337 1.74 1.74 1.63 46.78
6 1.76 156 1.74 1.74 '{).I8 46.%
7 0 0 1.74 1.74 -1.74 48.7
8 0 0 1.74 1.74 -1.74 -50.44
9 0 0 1.74 1.74 -1.74 -5218
10 0.78 om 1.74 1.74 -1.05 -5323
11 0 0 1.74 1.74 -1.74 -54.97
12 0 0 1.74 1.74 -1.74 -56.71
13 0 0 1.74 1.74 -1.74 -5845
14 0 0 1.74 1.74 -1.74 -«J.l9
15 3224 28.56 1.74 1.74 26.82 -3337
16 0.12 0.1 1.74 1.74 -1.64 -35.oI
17 0 .0 1.74 1.74 -1.74 -36.75
18 11.41 10.11 1.74 1.74 8.37 -28.38
413
414
lllE P P-R PET I; P-R-I; L (1 P-R- I;-D) 12Ma-1<ID 0 0 3.56 3.56 -3.56 -2136
lJ 0 0 3.56 3.56 -3.56 -2.4.92
14 6.41 5fJ7 3.56 3.56 211 -22.81
15 3.14 278 3.56 3.56 -0.78 -23.58
16 0 0 3.56 3.56 -3.56 -27.14
17 0 0 3.56 3.56 -3.56 -30.7
18 0 0 3.56 3.56 -3.56 -3426
I~ 0 0 3.56 3.56 -3.56 -37JQ
2U 0 0 3.56 3.56 -3.56 41.38
21 5.36 4.75 3.56 3.56 1.19 -40.I~
22 16.79 14.88 3.56 3.56 1132 -28.88
lJ 1.16 lID 356 3.56 -253 -31.41
14 0 0 356 3.56 -3.56 -34.97
L:> U U 3.56 3.56 -3.56 -38.53
26 0 0 3.56 3.56 -3.56 4200
27 U U 356 3.56 -356 45.65
lIS U 0 3.56 3.56 -3.56 4921
~ U 0 3.56 3.56 -3.56 -5277
30 0 0 3.56 3.56 -3.56 -56.33
31 U U 356 3.56 -3.56 -~JN
lAJrl<ID U 0 4.54 4.54 4.54 -64.43
2 U 0 4.54 0.45 .Q.45 -64.88
3 U 0. 4.54 0.45 .0.45 -6534
4 U U 4.54 0..45 -0.45 -65.79
5 0. 0 4.54 0.45 -0.45 .()625
6 0. 0. 4.54 0.45 -0.45 -&7
7 U 0 4.54 0.45 -0.45 -67.15
8 U U 4.54 0.45 -0.45 -67fJl
9 U U 4.54 0.45 -0.45 -68.06
lU 0.61 0.54 4.54 0.45 0.00 -6798
II U U 4.54 0.45 -0.45 -68.43
12 Ib 1.42 4.54 0.45 0.% -67.47
13 U 0 4.54 0.45 -0.45 -6792
14 0.% 0.85 4.54 0.45 0.4 -6752
15 0 U 4.54 0.45 -0.45 -67.CJ8
16 6.f:IJ :J!:J 4.54 0.45 5.45 -62.53
17 0.78 om 4.54 0.45 02.4 .(J229
18 2f:IJ 236 4.54 0.45 1.91 .0039
19 1.7 151 4.54 0.45 1.05 -5934
2) 0 0 4.54 0.45 -0.45 -~.79
21 U U 4.54 0.45 -0.45 .0014
22 0 0 4.54 0.45 -0.45 .(jJ.7
23 0 0 4.54 0.45 -0.45 -61.15
2.4 Ulil 0.72 4.54 0.45 026 .(jJ~
25 U U 4.54 0.45 -0.45 -6134
26 0 U 4.54 0.45 -0.45 -61.8
27 0 0 4.54 0.45 -0.45 {Q15
28 U U 4.54 0.45 -0.45 -6l..7
'}9 U U 454 0.45 .Q.45 -63.16
30 0 0 4.54 0.45 -0.45
IUIAL 561
415
