year (Figure 2). Other characteristics of water fluxes in the catchment are : - volume of mobile water of the component i (Eq.14). - mean transite time of water ...
Hydrology of Mountainous Areas/Proceedings of the Strbske Pleso Workshop, Czechoslovakia, June IAHS Publ. no. 190, 1990.
Estimation of the surface, subsurface and groundwater runoff components in mountainous areas J.GURTZ, R. SCHWARZE, G.PESCHKE, U.GRttNEWALD Department of Water Sciences, Dresden University of Technology, German Democratic Republic ABSTRACT The analysis of the precipitation-runoff processes is an important basis for the modelling and simulation of the runoff components and the water balance in mountainous areas. Long-term series of daily discharges have been used for the continuous identification of flow components. The total discharge is interpreted as sum over the discharges of several parallel linear reservoirs with specific storage coefficients for each flow component. A continuous water balance evaluation is used on a monthly basis for each flow component the relation of land use processes and atmospheric inputs to water quantity and quality processes in mountainous basins in the southern and western parts of the GDR demonstrate the capability of this method. The dynamic model of the soil water balance BOWAM is used for the simulation of the infiltration, percolation, évapotranspiration, changes in soil moisture, formation of overland flow, interflow and groundwater recharge. The BOWAM model should be applied preferably to sloping areas and higher mountain regions. It is suitable not only for the simulation of individual precipitation-runoff events, but also for the continuous, long-term simulation of the soil water balance, and for coupling to water quality models. INTRODUCTION Information about pathways and interactions of water flow from precipitation to various levels of outflows in a basin is valuable in predicting the magnitude and temporal distribution of runoff and groundwater recharge as well as in estimating the probable chemical composition of waters. The separation of these complex interactions is necessary also for the purpose of adequate representation of hydrologie processes in water yield systems used by agriculture, especially the selection of improved techniques for hydrologie survey of nitrate contamination of aquifers and surface water. CONTINUOUS HYDROGRAPH SEPARATION The discharge observed at a river gauge is often the only information on the transformation and storage processes in a drainage area. Investigations in mountainous areas of the GDR demonstrate the capability of continuous 263
/. Gurtz et al.
264
hydrograph separation to solve the problems mentioned above, in principle using the concept of parallel linear reservoirs with different storage coefficients. But the investigations also demonstrate the need for continuous water balance of individual separated runoff components CONCEPT AND PROCEDURE OF CONTINUOUS SEPARATION From recent investigations in mountainous basins it has been concluded that four flow components can be distinguished: - direct surface flow - interflow components with subsurface flow
RO } RH } direct flow components
- fast base flow RG1 interstitial water flow } long term flow components - delayed base flow RG2 } Each subsytem being the source of a flow component can be described as a separate reservoir. The simplest reservoir concept is that of the linear reservoir with the following parameters: - storage-discharge relation with the storage coefficient C S = C . Q
(1)
- depletion equations Q(t) = Q(t Q ) . exp(-t/C)
(2)
or equations describing the reservoir exhaustion In Q(t) = In Q(t Q ) - t/C
(3)
The recession curve must be a straight line in the semilogarithmic scale due to equation (3). The total discharge is the sum of the discharges of i parallel reservoirs N Q (t) = I Q i (tQ) exp (-t/Ci) (4) i=l or in the natural logarithm form N N In Q(t) = In I Q i (t) = In I Q i (t„) exp(-t/C.) i=l i=l where
Q(t) - total flow at time t, Q . ( t ) - flow of the component i at time t
(5)
Estimation of runoff components
265
Q.(t„)
- starting value for Q.
Our investigations in about 3 0 catchments with areas form 0.5 to 300 km have shown, that the components QO,QH QG1,QG2 mentioned above can be observed generally. For each component a specific storage coefficient CO, CH, CGI, CG2 must be estimated. The continuous separation:!is a gradual procedure and it is carried out on the base of a sufficient long-term series (about ten years) of mean daily discharges. It is favourable to start and end with very dry periods. The first step is the separation of QG2 by determining CG2. Therefore it is necessary to search for long dry recession periods with duration of several month. Then the discharge recession is approximated by a straight line in a semilogarithmic scale as a lower envelope of the hydrograph with number of intersections as low as possible. The inclination of the line gives CG2. With a stable value of CG2 it is possible to separate the recession sections of QG2 step ba step. To obtain a meaning-full continuous hydrograph QG2, the rising limbs between the different recession lines must be determined. However, they are only expected during precipitation periods, mainly during snowmelt event (Schwarze, 1985). The separation of QG2 results in a record of differences (Q-QG2). QG1 is the lowest flow component in the semilogarithmic plot of this differencial hydrograph. Now the separation of the next component QG1 is made by the same procedure described for QG2 using semilogarithmic plot (Q-QG2). In many cases separations of base flow, fast base and the sum of direct components are sufficient. In long wet periods the separation is critical due to long distance between supporting points for construction of QG2, and therefore, it is favourable to determine QG2 by using "respective measurements of the groundwater level". A relation between the groundwater level H characterizing the actual water storage within the aquifer and the river flow component QG2 can be expected as the following: QG2 = a . H n
(6)
where a, n - parameters (Busch, Luckner, 1972; Kliner, Knezek, 1974; Schwarze, 1985) The principles how to obtain appropriate groundwater measurements are given by Knezek (1987). Described simple proportionalities allow to find the QG2 rising lines in periods of successive wet years. The relation is used in predicting river flow also during dry periods (Schwarze, 1985).
266
/. Gurtz et al.
BALANCE OF THE SEPARATION In order to handle and properly interpret the separation of the total hydrographs, and so obtain an information as extensive as possible, a continuous water balance evaluation on a monthly basis is used for each flow component. For the balance following elements are necessary: P RG1 RG2 RD
-
sum of precipitation or snowmelt sum of fast base flow formation sum of delayed base flow formation sum of flow formation of the direct components (QO+QH) W+ETR - sum of residual, containing real évapotranspiration and storage, characterized by a depletion of ETR Because the direct runoff formed during a monthly interval is transformed into discharge (RD=QO+QH) within the same interval, the water balance can be written as follows : P - RG1 - RG2 - RD - (W+ETR) = 0
(7)
In consequence of equation (1) RG2 is calculated by the balance equation (8) A S mm = RG2 mm = CG2.QG2 m 3 /s
.86.4/A^ km2
(8)
where QG2 is distance of consecutive recession lines at the time of the peak of the rising limb. RG1 can be calculated from the hydrograph (Q-QG2) according to equation (8). By estimation of RG1 and RG2 a reminding term REM is also calculated: REM = P - RG1 - RG2
(9)
The discharge component (runoff concentration) for QG2 can be obtained by equations (10) or (11) as follows: mo t=2 8... 31 I QG2 = / QG2(tn) ._ exp(-t'/CG2)dt' -0' t=tQ=0
(-10)
mo I QG2 =[QG2(tQ) .CG2 ( 1-exp ( (28 . . . 31) /CG2)J . 86 . 4/AE (11) where
mo I - sum of monthly values
QG1 is calculated in the same way. Finally the equation (12) can be used for testing equation (13):
267
Estimation of runoff components
RD = (QO + QH) = Q - I QG1 - £ QG2
(12)
REM - RD = W + E T R i 0
(13)
Changes of the storage level A S of the individual flow components, losses from the groundwater mainly induced by évapotranspiration and groundwater inflow and outflow through the aquifer boundary, have to be calculated additionally in order to enable the use of the equations (8) to (13) within a computer program. It is advantageous to compare the results (annual means) of the equation (13) and calculations of évapotranspiration carried out by other producers (Golf, Schleicher, 1988) . The separation method described above is realized in the program DIFGA for ROBOTRON computer PC 1715. Figure 1 gives an example for a semilogarithmic plot of separated hydrograph. RESULTS The obtained mean values of runoff components provide valuable information on the hydrologie regime of different runoff components. The example of the Saidenbach catchment (Table 1) demostrates the utility of this method. The portion of direct surface flow is e.g. less than twenty percent of the total runoff. The part of impermeable and water covered areas with surface runoff nearly 100% of P - precipitation amounts to 2.7% of the drainage area. That means, that 97.3% of the catchment only about 4% of P - precipitation transform into RO. In the same way we cancalculate groundwater recharge components or time-variable runoff formation within the year (Figure 2 ) . Other characteristics of water fluxes in the catchment are : - volume of mobile water of the component i (Eq.14) - mean transite time of water (Eq.15) max S. = max Q. . C i
l
N T = I a± . Ci i=l where
(14)
l
(15)
a. - part of flow component i of the total flow, and ( I a i = 1) .
The subject method allows also information about the water quality processes. Figure 3 demonstrates the weak "relationship" between nitrogen concentration NO., and total discharge Q.A strong connection NO =f(QG2) results using the separated values of QG2 (Fig.47• This indicates, if different runoff components and their origin
268
/. Gurtzetal.
zdcyr—F Separation der 1-t Einzugsgebietsi-' ae Hieder-schlagsatati
^r-
ïrt^dt
ffiponenr e ! : ' - i ; ^ ^ l e t Bill.1
sas = 898.8
Fig. 1
1/s
Example for the separation of the delayed base flow using the computer program DIFGA (see also Table 2)
Estimation of runoff components
269
Table 1
Results of continuous identification and analysis of flow components from long-term series of daily discharges
Drainage basin of the Saidenbach-Reservoir Location: Erzgebirger, southern drainage area: 54.7 km2 height above sea-level:555 m Geology: gneiss, loamy sand
part of the GDR hill slope: 8.6 % channel slope: 2.0 % Land use: 78 % agricultural 18 % forest 4 % urban
Results of the period 1968-1983: summer half-year(V-X) winter half-year(XI-VT) Precipit.: P Runoff: R
512 mm _ 157.5 mm = 31% of P
419 mm _ 299.1 mm =71'.% of P.
with the following runoff components: RO 26.0 mm RH 55.2 mm RG1 30.5 mm RG2 45.8 mm RO 34.6 mm RH 103.8 mm RG1 85.2 mm RG2 75.5 mm Mean transit
16.8 34.8 19.3 29.1
5.1 10.8 6.0 8.9
83.2% of R underground flow!
11.6 34.7 28.5 25.2
8.3 24.8 28.5 18.0
88.4% of R underground flow!
time of water:
total runoff long term flow components
T
R = 120 d T R g = 195 d
Volume of mobile groundwater: max SG - 280 mm
/. Gurtz et al.
5 Si
Cj «V
s: «0
••1 •Ç
$
S, 4l U 5J
>£n k [^ k+ PI m _I Al J J J K 1=1 ]=1 j=l
t
"
K7)
J
(18)
For the homogenous soil is n,=n.=n, K,=K.=K and ty-.=~ty
/. Gurtzetal.
274
and than equation (18) is reduced to: m
PI
( - ^ - l) (I K
At ± PIi)>^ n i
i=1
(18a)
i
The formation of overland flow depends therefore on the relationship between PI and K during the moistening by rainfall and on the soil moisture deficit at the beginning of the rainfall. By means of further formulas the ponding time within or at the beginning of the mth interval can be calculated. The amount of infiltration up to. ponding yields from equation (19): K-l n
)] k [ \ + P I m i=l .1, A l jJ ( K7-ib k j ,
S
PI K
m - 1
K-l r Al-(n.-n,) j=l
(19)
J
k
The subject process of saturation phase is determined by the submodel SATUR2. Small rain intensities do not saturate the soil surface but they lead to an increase in soil moisture. These new soil moisture distribution is calculated by means of the submodel SATURl. Whereas in the saturation phase the infiltration intensity f. is equal to the precipitation intensity PI, f. decreases nonlinearly indepently of PI in the recession phase and approximates asymptotically to the K-value of the particular soil layer. For the potentially possible amount of infiltration we obtain: F pot (t)=F n _i+A n (t-t n _i)-B n + +2A n C n (t-t n _ 1 ) } where
{(An(t-tn_i)-Bn)a+
1/2
(20)
Fn_i - the amount of water infiltrated into the first (n-1) layers, tn_i - the time after which the wetting front has penetrated the (n-1) layers.
The transition time of the nth layer we calculate from equation (21): nn .1 n tn
= __
r
i
, 2B n n-C- i
[1+ n .1 +C ]
n n n n The abbreviations are described as follows: A n = K n / 2.0, n-1 Bn = nn Kn .L -
Ali.
.„. . (21)
Estimation of runoff components
275
C n
= 2n
n
*n
n-1 + I
i=l
Al
(22)
Supposing that the rainfall intensity is always greater or- equal to the infiltration intensity, the recession phase of infiltration is completely described by these equations. In this case, the water quantity above the irx filtration curve f (t) forms overland flow (see Fig.5) . In the case of precipitation intensities smaller than the infiltration rate, we have to balance the rising moisture deficit and precipitation excess in the following time intervals. In layered soil profiles interflow can be formed and calculated from the difference of infiltration curves belonging to different soil layers (see Fig.5). The calculation of the recession phase including the formation of runoff components is realised by the submodel INFIL. At the end of the recession phase the saturated part of the soil profile is defined by an upper and lower boundaries. This infiltrated plug of moisture can continue to percolate as simulated by submodel PERCO.Th'è. aimL of percolation. modelling is to describe the local and temporal variation of the soil moisture content immediately after infiltration. The share of the soil water between the water contents at saturation and the field capacity percolates under the influence of gravity. Evapot ranspiration In contrast to submodels described so far the calculation of évapotranspiration is carried out only in the daily intervals because of the low degree of process dynamics compared to the infiltration. The value of the evaporation is determined daily by the model. Besides specific site characteristics, the daily values of the air temperature, the relative air moisture constitute the data basis. The potential évapotranspiration ETP and the actual one ETR are determined successively. Whereas by the modelling of ETP, the essential atmospheric influences on evaporation are taken into account. In the calculation of the actual evapotraspiration above all th.e soil-physical and plant-physiological properties are taken into account. The vertical distribution of roots in the soil and the actual soil moisture are the ma in quantities influencing the actual transpiration. These terms are represented by reduction functions f (z) and f, (9). So we obtain for ETR: ETR = ETP - f1
(z) - f 2 (9)
D e t a i l ed i n f o r m a t i o n a b o u t t h e t y p e of t h e s e a r e t o be found by P e s c h k e , Dunger and G u r t z
(23) functions (19 8 6 ) .
/. Gurîz et al.
276
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