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Temporary hydrologic changes after deforestation for pioneer homesteading
Journal of Hydrology, 53 (1981) 117--133
117
Elsevier Scientific Publishing Company, Amsterdam -- Printed in The Netherlands
[4] TEMPORARY HYDROLOGIC CHANGES A F T E R D E F O R E S T A T I O N FOR PIONEER HOMESTEADING
HANS RYCKBORST
Petro-Canada, Calgary, Alta. T2P 2M7 (Canada) (Received May 7, 1980; accepted for publication September 16, 1980)
ABSTRACT Ryckborst:. H., 1981. Temporary hydrologic changes after deforestation for pioneer homesteading. J. Hydrol., 53: 117--133. Modern.day pioneering farmers, while clearing forest and brush from Central Albertan homestead lands, create a temporarily disturbed soil profile with a slightly increased effective porosity in groundwater recharge areas. The deforestation of recharge areas creates a drop in the groundwater table of 0.2--0.5 m, with a gradual recovery two years later. In contrast, groundwater levels in groundwater discharge areas rise due to a decreased effective porosity associated with mechanized deforestation. Overall, the impact of pioneering farmers on the hydrology of Central Albertan homestead lands appear to be minor and only temporary.
SCOPE
The first Canadian Dominion Lands Act of 1872 provided for free homesteads of a quarter section of land (64ha) for pioneering settlers. After paying a ten dollar filing fee, the homesteader had to comply with various regulations -- including a three-year residence clause -- before earning freehold title to his land. The work of cleating the land in those days was backbreaking, but the deed gave hundreds of thousands of families permanent status in their new land. After the turn of the century, during the last 75 years, new homesteads in the Canadian Western Plains have been made available to prospective pioneer farmers under Alberta's own provincial Public Lands Act. However, it i§ relatively unknown that even in this day and age significant tracts of homestead lands are sold by the Crown, or leased with five-year purchase options under the Public Lands Act. Homestead sales may take place of those parcels of public lands, classified by the Minister as suitable for cultivation 9rid available for settlement. Transactions with the Crown take place on the proportionate condition that a minimum of 4 h a per quarter section (64ha) be cleared by prospective farmers each year for a period of four years. Clearing means that wood and brush must be removed to make the land free from obstructions and suitable for farming or
0022-1694/81/0000--0000/$02.50 © 1981 Elsevier Scientific Publishing Company
118
ranching. From the second year on, the cleared land must be seeded to crop to a minimum of 24 ha after nine years. Clearing operations on Central Albertan forest lands, designated as homestead lands involve deforestation and the seeding of grasses. The impact of deforestation on the hydrology of prospective Central Albertan homestead land is the subject of this paper.
INTRODUCTION
Prospective Central Albertan homestead lands are located in areas characterized by a typical hydrology. The river basins, in which those homestead lands are located, are usually divided into two parts: (1) At higher elevations, a "land-locked" area is typical for the upper half of a basin. This elevated part of a river basin has not as yet developed an integrated river drainage system. Any previous drainage system has been obliterated during the last ice age, 10,000 years ago. Groundwater flow, groundwater runoff and surface runoff have been evolving, but on a local scale only. The entire hydrologic system is almost closed, or land-locked, with the exception of deep groundwater flow systems and evapotranspiration. In such land-locked areas, the surface waters and groundwaters fill up old glacial depressions, which now contain muskeg--peat deposits and lakes. Surprisingly, no surface runoff emerges across the boundaries of such areas. The hydrological cycle is kept in equilibrium by evapotranspiration. During wet periods excess precipitation and snowmelt causes a rise in groundwater levels and in the water levels of depressions, muskeg and lakes. Consequently, lakes expand or contract, sometimes in a spectacular fashion, in annual and seasonal cycles. The volume of water lost by evapotranspiration increases as a function of the percentage of the basin, occupied by openwater surfaces. The increase in evapotranspiration brings the hydrologic process back to equilibrium. (2) The lower half of a river basin consists of areas, which have since the last ice age (± 10,000 years ago), developed surface drainage systems ranging from poor to well-developed. This part of a river basin rhay receive small groundwater contributions from the land-locked areas. The impact of deforestation on the hydrology of Central Albertan lands, that could possibly be designated as homestead lands, has been analyzed during an 8-yr. experiment in a representative basin, Spring Creek {Fig. 1). This small research catchment (0.06km2), which also shows the abovementioned bi-partition, has been cleared from wood and brush in late 1971 and seeded with grass. Hydrological variables have been recorded during a period of eight years; including a period of four years before and four years after deforestation. The precipitation recorded over Spmlg Creek averages 500 mm/yr.; average runoff is ~ 8 0 mm/yr., evapotranspiration 420 mm/yr. After the 1972 deforestation an increase of groundwater runoff has been detected (Holecek and Noujaim, 1975). This detection was made possible by
119
I
,
- vJI~ '
C
-
J'i
,(
l, Fig. 1. Location of Spring Creek representative and experimental watershed in Canada.
means of the radioactive tracer '2sI which was sprayed uniformly on the snow cover of the small research catchme~at from a helicopter, just prior to the eight annual snowmelt seasons. With the radioactive tracer ,2s I, the groundwater portion of the snowmelt runoff has been measured during eight consecutive years; four before and four after deforestation. In addition, weekly groundwater level measurements have been carried out during the same 8-yr. period. The experiments show (Holecek and Noujaim, 1975) that the deforested soil does not retain less of the melt-water than before clearing. Apparently, more snowmelt infiltrates to the groundwater. The radioactive tracer measurements prove also that the non-steady groundwater component of the streamflow after clearing of the forest is augmented by 90%, from 1.0 to 1.9mm/day. After deforestation, the sublimation and evaporation from the snow cover decreases by 40%. The clearing of the forest is also responsible for a 25% shorter duration of the snowmelt runoff period and a 20% increase in overland flow. The decrease in evaporation contributes 40% out of a 50% recorded increase in the total average daily snowmelt runoff from 1.6 to 2.6 mm/day; the remaining 10% is contributed by increased soil drainage during the snowmelt--runoff period.
120
Total snowmelt runoff consists of two components: (1) groundwater runoff which, after deforestation, changes from 1.0 to 1.9mm/day; and (2) overland flow plus steady-state grom~,dwater seepage which changes very little, from 0.6 to 0.7 mm/day. STEADY-STATE GROUNDWATER CONTRIBUTION TO STREAMFLOW
The T-shaped research catchment is 300m long and 350m wide. It is located in a v',dley which drains into the Peace River (Fig. 1). The geology of the research catchment is known from nine drill holes, on the average 75 m deep, from which the lithology, resistivity logs and spontaneous potential logs have been recorded. The geology is shown in a north--south crosssection (Fig. 2). This figure indicates, that an upper layer of some 30m of little permeable lacustrine clay with thin coal seams is underlain by 25 m of more permeable glaciofluvial sand. The sand rests on a thin, but permeable gravel layer (1 m thick) which in turn covers 25 m of little permeable weathered Cretaceous sandstone (Wapiti Group). The 25m of weathered sandstone is underlain by unweathered Wapiti deposits. The geologic formation from the nine drill holes is supplemented by hydrologic data, consisting of hydraulic potential measurements that have been collected for eight piezometer nests, each equipped with three filters and also for fourteen water-table wells (Fig. 3). These measurements, the geologic profile (Fig. 2) and Laplace's equation have been used to compute the steady-state groundwater contributors to streamflow from the hydraulic potentials h with respect go horizontal datum. The geologic data and Fig. 2 show that the aquifers and aquitards are sufficiently similar in three dimensions to warrant a two-dimensional analysis. The two-dimensional representation is: S P(mV)
/Resistance(~m)
I
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0
LdgL:: 25 E
50 - =..:- _ _ -
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DISTANCE (m) Fig. 2. North--south geological cross-section, showing geophysical logs (spontaneous potential and resistivity).
121
P,ezometers
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~0-
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Fig. 3. Location of piezometer nests and water-table wells in the Spring Creek basin.
=
0
In eq. 1 x represents the horizontal coordinate, z the vertical coordinate and Kx and Kz are the horizontal and vertical conductivities, respectively. The solution for the hydraulic potential h u is obtained from a four-ooint finite-difference scheme, applying Taylor's expansion to eq. 1 (Smith, 965) and allowing for a leading error of order Ax 2 (if Ax 2 = Az 2 ):
h:,j = .K. .z.'.i.-. l. '.J. .h H ' J + K x , i , j - l h i , j - s + Kz,i,jhl+l,i_ + Kx,i, yhi,j+l ........ Kz,t,i + Kx,i,i + K z , i - l , i + Kx,~,j+l
(2)
The subscript i denotes the rows and j the columns for a square grid in which Ax = Az = 25m. The northern and southern vertical boundaries and theix hydraulic potential have been derived from the piezometer nests (Fig.3). The lower boundary has been taken as an equipotential line of 664 m, based on piezometer data recorded at depths of 125 m below the surface which is also ~ 5 0 m below the base of the weathered bedrock zone. The potentials for the upper boundary have been derived from ten watertable recorders during the end of the snowmelt period on May 9, 1974. Some initial values for hy:lraulic potentials at the internal nodes have been estimated by projecting hydraulic potential measurements from nearby piezometer nests on the profile as shown in Fig. 4. A vertical hydraulic conductivity value of 0.02 m/day for the lacustrine clay has been computed from field water-level measurements, using the auger method of Van Beers (1963). The horizontal hydraulic conductivity of the lacustrine clay, inferred
122 --Creek 650 - -
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Equipotential Line-----J
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Streomline--J 650 0
HORZ IONTALSCALE( m ) 0
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625 I00 .,,
--E
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600 - ~ . ' - Z G r a v e l . - , ' . . - " - ~
1974 r~
575-
550
-
-- ......
-~-:--:
525 ~
::
= =
---
:
: - = : ::-::>--~:
© Q
0
-
ul
"
Fig. 4. Solution of Laplace equation applied to a profile (Fig. 2) showing the steady-state flow field.
from geophysical logs, increases from 0.02 to 0.10 m/day, reflecting an increasing sand content for the clay layer in a southern direction. The vertical conductivity of the clay is one-fifth of this {Fig. 4, anisotropy). The hydraulic conductivities of the sand and gravel deposits are represented by Kz = 1.0 m/ day and Kx = 3.0m/day, also interpreted from geophysical logs. Weathered bedrock is considered homogeneous and represented by K x = K z = 0.01 m/ day, but the unweathered bedrock {Wapiti Formation) is considered anisotropic with Kx = 0.1 m/day and K z = 0.05 m/day. The anisotropic ratio has been derived from oil and gas well drill-stem tests of Cretaceous and Tertiary formations in the Spring Creek region. The results of the Lap!ace ,computations by using data of May 9, 1974, are shown as equipotential lines and streamlines in Fig. 4. Groundwater is seen to flow laterally through the gravel and sand layer and in an upward direction into the local discharge area of the creek, but also in a downward direction as recharging groundwater into the bedrock from where it drains towards springs emanating from bedrock exposures near the confluence of Spring Creek and the nearby Simonette River. The flowlines have been corrected for anisotropy (Van Ouwerkerk and Pette, 1965) by using ellipses with axes K~j2 and K~'2 tangent to the equipotential lines. The resulting profile (Fig. 4) of such two aquitards, separated by a permeable layer, has been described sometime ago as a "Holland polder profile" by Van Royen (1906) and more recentiy as the "prairie profile" by Meyboom {1962). The measured potential distribution is mainly due to" (1) recharge of the permeable layer beyond
123
the extension of the upper aquitard; (2) the small energy losses of the water flowing through the permeable layer; (3) the discharge of seepage through the upper aquitard. Fig. 4 shows that by the end of the snowmelt period on May 9, 1974, potential differences of ~ 1 m occurred over a vertical distance of 13m near ground surface on both sides of the creek. The associated groundwater flow is defined as vertical flow (Ernst, 1962). When the gradient is multiplied by a vertical hydraulic conductivity of (0.02/5)m per day, the upward steady-state groundwater contribution to streamflow amounts to 0.3 mm/day.
NON-STEADY-STATE G R O U N D W A T E R C O N T R I B U T I O N TO S T R E A M F L O W
Application of the steady-state Laplace equation to the Spring Creek data shows that the creek valley in the upper lacustrine clay deposits represents a local groundwater dischargi~ area, superimposed on a regional groundwater recharge area. Well records show that the groundwater levels are not always in a steady state. Apparently, the groundwater levels fluctuate whereas the recharge by rain and snowmelt differs from well to well. Consequently, a pattern of groundwater recharge--discharge areas could be postulated for the Spring Creek basin. Non-steady horizontal groundwater flow would then compensate for local differences in groundwater levels. Horizontal groundwater flow, superimposed on radial and vertical flows, has been analyzed according to terminology of Ernst (1962). In the case of Spring Creek, the analysis of a possible diiferentiation in location of groundwater recharge-discharge area has been restricted to the upper clay layer (30 m). Non-steady groundwater movement with recharge or discharge {positive or negative P) in the clay layer with phreatic groundwater is described by a two-dimensional diffusion equation (Van der Molen, 1972):
a-T =
(a)
where H represents the hydraulic potential, T the time, x and y the horizontal coordinates, Kx and Ky the horizontal hydraulic conductivities and p the volume of water released from or taken into storage per unit area of the aquifer, due to a change of the water level corresponding to unit potential. In the case of the Spring Creek catchment, there is evidence to warrant the assumption of a constant thickness D of the lacustrine clay (Fig. 3}. If permeability Kx = R y in both horizontal directions in the clay, then the nondimensional form of eq. 3 is given by: ah/3t
= O2h/'t)Z2 + (}2hfi)Y2 4 p
(4)
where X = x / L ; Y = y / L ; t = k D T / p L 2 ; h = H / H o ; and p = p L 2 / K D 2 . The application of Taylor's expansion to eq. 4 and the acceptance of a leading
124
error of the order of Ax 2 results in the following explicit finite-difference equation, using a square grid (Ax = Ay)-
g,i+l,/Jr"
hg, i,j+l = hg, i,j -F - ~
h#,l_l, j-'4he,i, s + he+l,i, j + he_l,i, j
-F Atp
(5) The grid, the boundary nodes, the constant open channel nodes and the internal nodes are shown in Fig. 5. For accuracy, the nodal distance zL~ is taken as ,-i~th of a basin width of 350 m. The thickness of the upper clay layer equals 30 m. In the Laplace's analysis the horizontal hydraulic conductivity ranges between 0.02 and 0.10m/day. Assuming a constant value between these extremes, e.g., Kx = ICy = 0.036 m/day, and substitution of K=0.036, D=30, #=0.02 and L = 3 5 0 yields" At=(KD//~L2)T= 0.00044 T and At/Ax ~ = 0.1, which guarantees stability and a time step At of one day. Laboratory measurements on drilled cores, obtained before deforestation yield p-values of 0.02 (2%}. Recharge--discharge patterns in the research catchment have been computed, by applying eq. 5 to two ~oundwater contour maps. One has been prepared from records of February 13, 1974 {Fig. 6A, winter condition~} and the other from records of May 9, 1974 (F~g. 6B snowmelt conditions).
Boundary nodes
/
,~
~(
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.--River node~
"P" K ~
0
.
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.
.
~LI
50
••
-+TZI;IT.'IIIIII ii_~++i ~
!00 m
Fig. 5. Finite-difference grid of tk~e Spring Creek basin for two-dimensional diffusion equation.
125
7
~'/"i@"
"-
~'i
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o
---.
\......
FEBRUARY13, 1974
X ,
,..,o....,n.,.
.
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Im
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654
';5-~
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' --. i
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MAY 9, 1974
/
• ,t
SNOWMELT
G
nest
x
P,e:ometer
o
Water-table well
•- - . e - - .
Pmlte-&fference grid boundary
.-b~~ ......
Oroundwoter contour ( m ) Contour interval : ] m
r 0
~'st50
.
):
~,,
\k i
o
'....653
~ : 1 lOOm
Fig. 6. Groundwater contour map for the Spring Creek basin of: (A) winter freezing co~lditions, February 13, 1974; and (B) snowmelt conditions, May 9, 1974.
126
The two maps represent the basin after deforestation. The boundary and river hydraulic head nodes have been kept constant and the initial levels for the internal nodes have been obtained from the contour maps as shown in Fig. 6. A comparison of the results (Fig. 7) indicates that recharge--discharge patterns occur, and also that an expansion and contraction of different recharge and discharge areas has taken place after a time lapse of three months. Comparison of Fig. 7A and B indicates also some changes in recharge--discharge rates expressed in mm/day. More detailed computations not reported herein, using daily changes in groundwater levels, suggest that recharge--discharge boundaries change locations on a daily basis. In the summer these transition zones may be recognized by botanical communities that are transitional between shallower and deeper groundwater tables of the discharge, respectively recharge areas. Under winter conditions (Fig. 7A) the catchment soil is snow covered and frozen, and the recharging/discharging groundwaters are therefore not visible. However, the spring recharge--discharge conditions can occasionally be recognized in t h e field. For example, Fig. 8 shows an oblique angle photograph of the 6-ha research catchment taken from a helicopter. It shows the southern discharge area of Fig. 7B, while the sun just happened to reflect on the discharging groundwaters. Average non-steady recharge and discharge groundwater velocities have been integrated from Fig. 7, which graphically portray the computed groundwater flow conditions for the wiuter and the end of the snowmelt period in 1974. The recharge areas (48% of basin Fig. 7 A ) c o n t r i b u t e 13.9m3/day whereas the discharge areas (3.1 ha) contribute 14.3 m3/day. Winter groundwater flow velocities in recharge areas average 0.39 vs. 0.37 mm/day for discharge areas. Conditions at the termination of the snowmelt period {Fig. 7B) result in a recharge rate of 21.1m3/day (56% of catchment} whereas the discharge areas (2.6 ha) contribute 11.2 m3/day. Snowmelt groundwater flow velocities in the recharge areas average 0.5 vs. 0.2mm/day for the discharge areas. The net non-steady groundwater flow, computed for winter conditions, is almost negligible {14.3 -- 13.9 = 0.4 m3/day) and directed upwards. By the end of the snowmelt period the calculated non-steady net groundwater flow is directed downward, amounting to 10m3/day, or 0.16 mm/day from 6ha. The calculated rate of 0.16mm/day compares with a measured non-radioactive groundwater runoff of 0.01 ram/day. The discrepancy of 0.15 mm/day is probably due to evapotranspiration losses on May 9, 1974. The steadystate vertical groundwater seepage (0.3 ram/day, see the previous section} is also lost to evapotranspiration. The eight years of experimenting shows that snowmelt groundwater runoff, which has lost its radioactive tracer ~2sI due to the infiltration effects of the soil~ averages ~55% of the total snowmelt runoff volume before deforestation. This increased slightly to 65% after deforestation. Clearing of the forest is also responsible for a 25% shorter duration of the snowmelt runoff period, as well for a 40% decrease in snowcover sublimation and evaporation. The effect of these three factors is that
127 o ,iuiIU,
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RECHARGE {
DISCHARGE
m
o~-~
{~
L_ [ ~ ' . ~
WINTER FREEZING
~o,-, ~~r~-:.:...:~
O-Obmm day-I
J
06-25mm dJy "1
j •
®
RECHARGE - DISCHARGE VELOCITIES (mm day -! ) = 0.02 (2°/o)
[ ,
-¢' --:_i ...........
i_._._._
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,oo~
RECHARGE ~
cz ~ISCHARGE ~
®
~ ~
06-2 5 mm doyl
2 5-7ram da~,1 0-06 mm day "1
06- 25 mm day'l
!
:::: - ~
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I
..... iJ
'
..
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.
::::;
~iii!ii!i!!il :::::1 S.OWMELT PE'~,OD RECHARGE
-
D I S C H A R G E
i!i!!!~i
VELOCITIES (ram day ")
.i::]!i::ii::'~
JJ" = 0.02 ( 2 °/o)
2.5-3 mm day I
Fig. 7. Distribui~on of recharge--discharge velocities in the Spring Creek basin in mm/day during: (A) winter freezing conditions; and (B) spring snowmelt conditions.
Fig. 8. Oblique-angle photo of the Spring Creek basin, showing one discharge area where the sun reflects on the water.
t~ G0
129 on the average, the groundwater contribution to streamflow recorded during the annual snowmelt periods, changes from 1.0 to 1 . 9 m m / d a y after deforestation. As Fig. 7B suggests, such net flow rates are derived from expanding and contracting recharge and discharge areas.
STREAMFLOW, GROUNDWATER LEVELS AND GROUNDWATER STORAGE The main annual streamflow event in the small Spring Creek research catchment is the short-lived snowmelt runoff, which takes place mostly within 4--5 weeks during the spring time. After the snowmelt the creek carries very little flow due to high evapotranspiration rates and usually the creek dries up or reduces to a trickle during the summer. The fall rains bring again a short-lived a m o u n t of runoff, before winter sets in. A one-dimensional version of diffusion equation (3) has been applied to groundwater level records (Fig. 9) and to rainfall--runoff data from the 6-ha research catchment in order to explain the reported changes in deforestation:
~}x 2
~)t
(6)
p(t)
The symbols in eq. 6 have been previously defined in the context of eq. 3. When eq. 6 is applied to symmetrical boundaries and special conditions of the research catchment (consisting of an aquifer underlain by an aquitard), an analytical solution of eq. 6 exists (Wesseling, 1973). This solution yields the groundwater level yt midway between two creeks (at x - 0.5L), where L represents t]~e distance between creeks, the total phreatic groundwater storage Rt; and the groundwater runoff qt. 0 E
! m.
~
iI
" ~
.
' )
Ij~ x
I
,
~t \~'~ \':~'1~...
~
'""'"'
' "
III; v \'V,'I I \k,,.."~JJ I/~:%" k~'~l
J
" "
I
~"'..1 ~'~
~!!
-Deforeslat~on
3 1969
1970
1971
1972
i973
1974
YEARS
Fig. 9. Groundwater levels of four wells before and after deforestation in October, 197 t.
130
An analytical solution of eq. 6 can be applied to Spring Creek, if the Dupuit approximation is applicable to the upper lacustrine clay layer. In this research catchment the properties of the clay appears to warrant the application of Kirkham's (1967)slot--slab model. There is evidence that slabs of clay and till have been fractured during or after last glaciation. Kirkham's slot--slab model follows the Dupuit theory. The analytical solution for qt, yt, Rt is" tl
qt.. = ~ (ut.. + p t r . )
(7)
I
where n counts the number of terms (1, 3, 5, 7, . . . ) in an infinite series ( 2 n - - 1 ) ; as n increases, accuracy increases; ut,. is a carry-over term at time t: Ut.n
-"
Ut-l.ne,
+ PtSr.
en = e x p - - [b(2n
(8)
(9)
-- 1)2/j]
where j (in units of time) represents a four-parameter drainage resistance (Krayenhoff van de Leur, 1962)" J = (1/lr2 )I~L2/KD
(10)
in which L is the distance (m) between two parallel ditches or creeks and:
S. = 8/Tr2[1/(2n- 1)2(1--e,,)]
(11)
rn-1--8/Tr2I~1/(2n--1)21
(12)
since" 81~ 2 [
lira n=I,2,3
{ l l ( 2 n - 1) 2 }] = 1
(13)
.....
The groundwater level Yt is given by" n
2n -1
Yt.. = ~ qt..[Dr/{2p(2n-1)}l +pt[4j/(g~)] 7ra/32 -I
E
(1/v)3 1
v= 1 . - 3 . 5 . - 7 . . . .
(14) and the groundwater storage Rt equals:
Rt.. = ~ qt.nj[1/(2n -- 1)] + Pt(8j/lr 2) 1r4/96 -- ~ 1/(2n -- 1) 4 1
1
(15)
Analytical solutions (7}, (14) and (15) have been used to determine the possible effects of deforestation from changing parameter values, through matching of: (1) The recorded monthly groundwater levels with computed levels from 1969--1975 (Fig. 10).
131 --
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15% Prectp j : 4.5 days U. : 2% "~ - - D e f o r e s t a t i o n
iII
1969
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If
III
1 I,
1970
t l | I T t
IIr
III
IT
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1 1 1
1971
I
it
i
1972
t
l
WELL '5611 l
,
l
l
l
t
t
it
I
1973
rT
T ' | t
IT
l i t 1
~ I ~ T
1974
YEARS
Fig. 10. Measured vs. computed groundwater levels before and after deforestation.
(2) The computed groundwater contribution to streamflow during the snowmelt period vs. the groundwater contribution to streamflow measured with the radioactive tracer from 1968--1975 (see Table I). Groundwater levels, groundwater storage and groundwater runoff have been matched before deforestation and after deforestation, using recorded snowmelt and two sets of parameters; j = 45 days and # = 2% before deforesttation, vs. j = 60 days and # = 3% after deforestation. The j-value represents a combination of four catchment characteristics lumped in one parameter: j = (1/Ir)2(I~L2/KD) (dimension: time). Fig. 10 shows that the combination j = 45 days and /~ = 2% matches the groundwater levels of well 5611 (recharge area) slightly better before deforestation. The combination j = 60 days and ~ = 3% appears to fit better the first years after deforestation. Thereafter the combination j = 45 days and # = 2% fits somewhat better. This indicates that deforestation in recharge areas created a temporarily and only slightly disrupted soil profile with an increased effective porosity, thus explaining the measured increase in groundwater runoff. Figs. 9 and 10 show also that deforestation in the recharge areas created a drop of the groundwater table of 0.2--0.5 m, with a gradual recovery two years later (wells 5611 and 5608). Ho~fever, in the discharge areas, groundwater tables rose (Fig. 9, wells 5605 and 5596). The rise of the water tables in discharge areas is associated with decreased effective porosity/~ in those wet and lowlying areas that are compacted during mechanical deforestation operations. Table I shows a comparison of the computed groundwater runoff vs. the groundwater runoff qt measured wi~h the 12sI isotope. The snowmelt in 1974, the last year of simulation, was exceptional (150ram) due to an unusually thick snow cover preceded by a wet fall season in 1973. The average snowwater runoff is usually only 75 ram/yr. Table I shows that j = 45 days and/~ = 2% fits better, whereas the combi-
132 TABLE I
Total groundwater runoff, q t, computed during snowmelt q t, mm per snowmelt period
measured with radioactive tracer
computed with
computed with
j = 45 days, p -- 2%
j = 60 days,/~ = 3% 12.9 15.0 14.9 17.9 ( d e f o r e s t a t i o n ) 24.4 19.0 27.3
1968 1969 1970 1971 1972 1973 ~974
14.6 28.2 17.1 18.7 18.7 22.0 84.6
14.8 24.6 17.0 19.8 27.9 21.1 30.4
Total
203.9
155.6
13! .4
Total excluding 1974
119.3
125.2
104.1
nation j = 60 days and p = 3% proves slightly better shortly after deforestation. The conclusion, drawn from the 8-yr. experiment is, that deforestation of Central Albertan prospective homestead lands has only minor and temporary (1--2 yr.) effects on the hydrology of these lands.
CONCLUSIONS
Clearing operation on Central Albertan forest lands, designated as modernday pioneering farmer homestead lands, involve deforestation and the seeding of grasses An 8-yr. experiment on the impact of deforestation on the hydrology of a research catchment, located in the prospective homestead lands, shows that only minor and temporary impacts can be expected. This conclusion was reached by the use of an 12sI isotope in the experiment. The main changes are a slight increase in total snowmelt runoff from 60 to 75 mm/yr.; a shortening of the snowmelt period from 5 to 4 weeks; an increase of snowmelt groundwater runoff from 1.0 to 1.9 mm/day and a reduction of snow-cover sublimation and evapQration by 40%. Deforestation in groundwater recharge areas creates a temporarily disturbed s3il profile with a slightly increased effective porosity, thus explaining the recc,rded increase in groundwater runoff. In the recharge areas, deforestation creates a drop in the groundwater table of 0.2--0.5 m, with a gradual recovery two years later. However, in the remaining groundwater discharge areas, groundwater tables rise due to a decreased effective porosity associated with mechanized deforestation operations.
133 ACKNOWLEDGEMENTS
Assistance of Mr. V. Da Silva, R. Bohmer and the liason of Dr. J. Toth and computer programming of Mr. A. Litviak is acknowledged. George Holecek, Head Alberta River Basin Research, functioned as project leader.
REFERENCES
Ei'nst, L.F., 1962. Groundwater flow in the saturated zone and its calculation when horizonal parallel open conduits are present. Versl. Landbouwk. Onderz., Wageningen, No. 65-15, pp. 1--189. Holecek, G.F. and Noujaim, A.A., 1975. Separation of the surface and subsurface flow from snowmelt by the use of radioactive tracers. Tokyo Symp. Proc., Int. Assoc. Sci. Hydrol., Publ. 117, pp. 497--505. Kirkham, D., 1967. Explanation of paradoxes in Dupuit--Forchheimer seepage theory. Water Resour. Res., 2(3): 609--622. Krayenhoff van de Leur, D.A., 1962. A study of non-steady groundwater flow. Ingenieur, 2o: B285--B292. Meyboom, P., 1962. Patterns of ground- ter flow in the Prairie Profile (with discussion). Proc 3rd Can. Hydrol. Syrup., Calgary, Alta., pp. 5--33. Smith, G.D., 1965. Numerical Solution of Partial Differential Equations. Oxford University Press, London, 175 pp. Van Beers, W.F.J., 1963. The auger hold method. Int. Inst. Land Reclam. Improv., Wageningen, Bull. No. 1, pp. 1--32. "Van der Molen, W.H., 1972. Non-steady groundwater flow. Lect. M.Sc. Course on Soil Sciences and Water Management, Agric. Univ., Wageningen, Intern. Rep. (unpublished). Van Ouwerkerk, J.H. and Pette, D., 1965. A technique to compute stationary groundwater flow, H. Tijdschr. K. Ned. Heidemaatsch., pp 444--457. Van Royen, R.P., 1906. Reply to discussion of H.E. de Bruyn on The flow of water in the underground. Ingenieur, 26:480--482. Wesseling, J., 1973. Subsurface flow into drains. In: Drainage Principles and Applications, Vol. If, Theories of Field Drainage and Watershed Runoff. Int. Inst. Land Reclam. Improv., Wageningen, Publ. No. 16, pp. 1--31.
117
Elsevier Scientific Publishing Company, Amsterdam -- Printed in The Netherlands
[4] TEMPORARY HYDROLOGIC CHANGES A F T E R D E F O R E S T A T I O N FOR PIONEER HOMESTEADING
HANS RYCKBORST
Petro-Canada, Calgary, Alta. T2P 2M7 (Canada) (Received May 7, 1980; accepted for publication September 16, 1980)
ABSTRACT Ryckborst:. H., 1981. Temporary hydrologic changes after deforestation for pioneer homesteading. J. Hydrol., 53: 117--133. Modern.day pioneering farmers, while clearing forest and brush from Central Albertan homestead lands, create a temporarily disturbed soil profile with a slightly increased effective porosity in groundwater recharge areas. The deforestation of recharge areas creates a drop in the groundwater table of 0.2--0.5 m, with a gradual recovery two years later. In contrast, groundwater levels in groundwater discharge areas rise due to a decreased effective porosity associated with mechanized deforestation. Overall, the impact of pioneering farmers on the hydrology of Central Albertan homestead lands appear to be minor and only temporary.
SCOPE
The first Canadian Dominion Lands Act of 1872 provided for free homesteads of a quarter section of land (64ha) for pioneering settlers. After paying a ten dollar filing fee, the homesteader had to comply with various regulations -- including a three-year residence clause -- before earning freehold title to his land. The work of cleating the land in those days was backbreaking, but the deed gave hundreds of thousands of families permanent status in their new land. After the turn of the century, during the last 75 years, new homesteads in the Canadian Western Plains have been made available to prospective pioneer farmers under Alberta's own provincial Public Lands Act. However, it i§ relatively unknown that even in this day and age significant tracts of homestead lands are sold by the Crown, or leased with five-year purchase options under the Public Lands Act. Homestead sales may take place of those parcels of public lands, classified by the Minister as suitable for cultivation 9rid available for settlement. Transactions with the Crown take place on the proportionate condition that a minimum of 4 h a per quarter section (64ha) be cleared by prospective farmers each year for a period of four years. Clearing means that wood and brush must be removed to make the land free from obstructions and suitable for farming or
0022-1694/81/0000--0000/$02.50 © 1981 Elsevier Scientific Publishing Company
118
ranching. From the second year on, the cleared land must be seeded to crop to a minimum of 24 ha after nine years. Clearing operations on Central Albertan forest lands, designated as homestead lands involve deforestation and the seeding of grasses. The impact of deforestation on the hydrology of prospective Central Albertan homestead land is the subject of this paper.
INTRODUCTION
Prospective Central Albertan homestead lands are located in areas characterized by a typical hydrology. The river basins, in which those homestead lands are located, are usually divided into two parts: (1) At higher elevations, a "land-locked" area is typical for the upper half of a basin. This elevated part of a river basin has not as yet developed an integrated river drainage system. Any previous drainage system has been obliterated during the last ice age, 10,000 years ago. Groundwater flow, groundwater runoff and surface runoff have been evolving, but on a local scale only. The entire hydrologic system is almost closed, or land-locked, with the exception of deep groundwater flow systems and evapotranspiration. In such land-locked areas, the surface waters and groundwaters fill up old glacial depressions, which now contain muskeg--peat deposits and lakes. Surprisingly, no surface runoff emerges across the boundaries of such areas. The hydrological cycle is kept in equilibrium by evapotranspiration. During wet periods excess precipitation and snowmelt causes a rise in groundwater levels and in the water levels of depressions, muskeg and lakes. Consequently, lakes expand or contract, sometimes in a spectacular fashion, in annual and seasonal cycles. The volume of water lost by evapotranspiration increases as a function of the percentage of the basin, occupied by openwater surfaces. The increase in evapotranspiration brings the hydrologic process back to equilibrium. (2) The lower half of a river basin consists of areas, which have since the last ice age (± 10,000 years ago), developed surface drainage systems ranging from poor to well-developed. This part of a river basin rhay receive small groundwater contributions from the land-locked areas. The impact of deforestation on the hydrology of Central Albertan lands, that could possibly be designated as homestead lands, has been analyzed during an 8-yr. experiment in a representative basin, Spring Creek {Fig. 1). This small research catchment (0.06km2), which also shows the abovementioned bi-partition, has been cleared from wood and brush in late 1971 and seeded with grass. Hydrological variables have been recorded during a period of eight years; including a period of four years before and four years after deforestation. The precipitation recorded over Spmlg Creek averages 500 mm/yr.; average runoff is ~ 8 0 mm/yr., evapotranspiration 420 mm/yr. After the 1972 deforestation an increase of groundwater runoff has been detected (Holecek and Noujaim, 1975). This detection was made possible by
119
I
,
- vJI~ '
C
-
J'i
,(
l, Fig. 1. Location of Spring Creek representative and experimental watershed in Canada.
means of the radioactive tracer '2sI which was sprayed uniformly on the snow cover of the small research catchme~at from a helicopter, just prior to the eight annual snowmelt seasons. With the radioactive tracer ,2s I, the groundwater portion of the snowmelt runoff has been measured during eight consecutive years; four before and four after deforestation. In addition, weekly groundwater level measurements have been carried out during the same 8-yr. period. The experiments show (Holecek and Noujaim, 1975) that the deforested soil does not retain less of the melt-water than before clearing. Apparently, more snowmelt infiltrates to the groundwater. The radioactive tracer measurements prove also that the non-steady groundwater component of the streamflow after clearing of the forest is augmented by 90%, from 1.0 to 1.9mm/day. After deforestation, the sublimation and evaporation from the snow cover decreases by 40%. The clearing of the forest is also responsible for a 25% shorter duration of the snowmelt runoff period and a 20% increase in overland flow. The decrease in evaporation contributes 40% out of a 50% recorded increase in the total average daily snowmelt runoff from 1.6 to 2.6 mm/day; the remaining 10% is contributed by increased soil drainage during the snowmelt--runoff period.
120
Total snowmelt runoff consists of two components: (1) groundwater runoff which, after deforestation, changes from 1.0 to 1.9mm/day; and (2) overland flow plus steady-state grom~,dwater seepage which changes very little, from 0.6 to 0.7 mm/day. STEADY-STATE GROUNDWATER CONTRIBUTION TO STREAMFLOW
The T-shaped research catchment is 300m long and 350m wide. It is located in a v',dley which drains into the Peace River (Fig. 1). The geology of the research catchment is known from nine drill holes, on the average 75 m deep, from which the lithology, resistivity logs and spontaneous potential logs have been recorded. The geology is shown in a north--south crosssection (Fig. 2). This figure indicates, that an upper layer of some 30m of little permeable lacustrine clay with thin coal seams is underlain by 25 m of more permeable glaciofluvial sand. The sand rests on a thin, but permeable gravel layer (1 m thick) which in turn covers 25 m of little permeable weathered Cretaceous sandstone (Wapiti Group). The 25m of weathered sandstone is underlain by unweathered Wapiti deposits. The geologic formation from the nine drill holes is supplemented by hydrologic data, consisting of hydraulic potential measurements that have been collected for eight piezometer nests, each equipped with three filters and also for fourteen water-table wells (Fig. 3). These measurements, the geologic profile (Fig. 2) and Laplace's equation have been used to compute the steady-state groundwater contributors to streamflow from the hydraulic potentials h with respect go horizontal datum. The geologic data and Fig. 2 show that the aquifers and aquitards are sufficiently similar in three dimensions to warrant a two-dimensional analysis. The two-dimensional representation is: S P(mV)
/Resistance(~m)
I
//,. /
0
LdgL:: 25 E
50 - =..:- _ _ -
75 ¸
I 0 0 -
o
so
,
,
,oo
,50
,. . . . .
,
200
. . . .
,
250
.....
,
,
300
,oo
DISTANCE (m) Fig. 2. North--south geological cross-section, showing geophysical logs (spontaneous potential and resistivity).
121
P,ezometers
Well--~ O~
-
--
,0-
:-=4
~0-
::~:::;
3o-
--_=~-=
//
°
50-
de{o
_
.
Wo,*er-{,oble we{{ +rec
(~
P,ezometer nest
/65
/opograph,c contours
60-
~ _ _
70-
~ ; =-3',--~ Wop,,,~o,mo,,on
80--
_--_____+--,~_ -:-_ -~-- "+-'~-~--::--:
--'-'_~-
,r-.~--- T 0
\
+r
50
~,~i,
100
....
o
~.~ ~,' ~
l
lSO(m )
Fig. 3. Location of piezometer nests and water-table wells in the Spring Creek basin.
=
0
In eq. 1 x represents the horizontal coordinate, z the vertical coordinate and Kx and Kz are the horizontal and vertical conductivities, respectively. The solution for the hydraulic potential h u is obtained from a four-ooint finite-difference scheme, applying Taylor's expansion to eq. 1 (Smith, 965) and allowing for a leading error of order Ax 2 (if Ax 2 = Az 2 ):
h:,j = .K. .z.'.i.-. l. '.J. .h H ' J + K x , i , j - l h i , j - s + Kz,i,jhl+l,i_ + Kx,i, yhi,j+l ........ Kz,t,i + Kx,i,i + K z , i - l , i + Kx,~,j+l
(2)
The subscript i denotes the rows and j the columns for a square grid in which Ax = Az = 25m. The northern and southern vertical boundaries and theix hydraulic potential have been derived from the piezometer nests (Fig.3). The lower boundary has been taken as an equipotential line of 664 m, based on piezometer data recorded at depths of 125 m below the surface which is also ~ 5 0 m below the base of the weathered bedrock zone. The potentials for the upper boundary have been derived from ten watertable recorders during the end of the snowmelt period on May 9, 1974. Some initial values for hy:lraulic potentials at the internal nodes have been estimated by projecting hydraulic potential measurements from nearby piezometer nests on the profile as shown in Fig. 4. A vertical hydraulic conductivity value of 0.02 m/day for the lacustrine clay has been computed from field water-level measurements, using the auger method of Van Beers (1963). The horizontal hydraulic conductivity of the lacustrine clay, inferred
122 --Creek 650 - -
650
o54
625
625
6 2 0 ~ ,,oI
669
_.._.____._._~600
.....
. . . . . . .
m *MSL525
/
/
•/ /
,
525 m * MSL
/
Equipotential Line-----J
,/ /
Streomline--J 650 0
HORZ IONTALSCALE( m ) 0
50
MAY
9,
625 I00 .,,
--E
Sand z
600 - ~ . ' - Z G r a v e l . - , ' . . - " - ~
1974 r~
575-
550
-
-- ......
-~-:--:
525 ~
::
= =
---
:
: - = : ::-::>--~:
© Q
0
-
ul
"
Fig. 4. Solution of Laplace equation applied to a profile (Fig. 2) showing the steady-state flow field.
from geophysical logs, increases from 0.02 to 0.10 m/day, reflecting an increasing sand content for the clay layer in a southern direction. The vertical conductivity of the clay is one-fifth of this {Fig. 4, anisotropy). The hydraulic conductivities of the sand and gravel deposits are represented by Kz = 1.0 m/ day and Kx = 3.0m/day, also interpreted from geophysical logs. Weathered bedrock is considered homogeneous and represented by K x = K z = 0.01 m/ day, but the unweathered bedrock {Wapiti Formation) is considered anisotropic with Kx = 0.1 m/day and K z = 0.05 m/day. The anisotropic ratio has been derived from oil and gas well drill-stem tests of Cretaceous and Tertiary formations in the Spring Creek region. The results of the Lap!ace ,computations by using data of May 9, 1974, are shown as equipotential lines and streamlines in Fig. 4. Groundwater is seen to flow laterally through the gravel and sand layer and in an upward direction into the local discharge area of the creek, but also in a downward direction as recharging groundwater into the bedrock from where it drains towards springs emanating from bedrock exposures near the confluence of Spring Creek and the nearby Simonette River. The flowlines have been corrected for anisotropy (Van Ouwerkerk and Pette, 1965) by using ellipses with axes K~j2 and K~'2 tangent to the equipotential lines. The resulting profile (Fig. 4) of such two aquitards, separated by a permeable layer, has been described sometime ago as a "Holland polder profile" by Van Royen (1906) and more recentiy as the "prairie profile" by Meyboom {1962). The measured potential distribution is mainly due to" (1) recharge of the permeable layer beyond
123
the extension of the upper aquitard; (2) the small energy losses of the water flowing through the permeable layer; (3) the discharge of seepage through the upper aquitard. Fig. 4 shows that by the end of the snowmelt period on May 9, 1974, potential differences of ~ 1 m occurred over a vertical distance of 13m near ground surface on both sides of the creek. The associated groundwater flow is defined as vertical flow (Ernst, 1962). When the gradient is multiplied by a vertical hydraulic conductivity of (0.02/5)m per day, the upward steady-state groundwater contribution to streamflow amounts to 0.3 mm/day.
NON-STEADY-STATE G R O U N D W A T E R C O N T R I B U T I O N TO S T R E A M F L O W
Application of the steady-state Laplace equation to the Spring Creek data shows that the creek valley in the upper lacustrine clay deposits represents a local groundwater dischargi~ area, superimposed on a regional groundwater recharge area. Well records show that the groundwater levels are not always in a steady state. Apparently, the groundwater levels fluctuate whereas the recharge by rain and snowmelt differs from well to well. Consequently, a pattern of groundwater recharge--discharge areas could be postulated for the Spring Creek basin. Non-steady horizontal groundwater flow would then compensate for local differences in groundwater levels. Horizontal groundwater flow, superimposed on radial and vertical flows, has been analyzed according to terminology of Ernst (1962). In the case of Spring Creek, the analysis of a possible diiferentiation in location of groundwater recharge-discharge area has been restricted to the upper clay layer (30 m). Non-steady groundwater movement with recharge or discharge {positive or negative P) in the clay layer with phreatic groundwater is described by a two-dimensional diffusion equation (Van der Molen, 1972):
a-T =
(a)
where H represents the hydraulic potential, T the time, x and y the horizontal coordinates, Kx and Ky the horizontal hydraulic conductivities and p the volume of water released from or taken into storage per unit area of the aquifer, due to a change of the water level corresponding to unit potential. In the case of the Spring Creek catchment, there is evidence to warrant the assumption of a constant thickness D of the lacustrine clay (Fig. 3}. If permeability Kx = R y in both horizontal directions in the clay, then the nondimensional form of eq. 3 is given by: ah/3t
= O2h/'t)Z2 + (}2hfi)Y2 4 p
(4)
where X = x / L ; Y = y / L ; t = k D T / p L 2 ; h = H / H o ; and p = p L 2 / K D 2 . The application of Taylor's expansion to eq. 4 and the acceptance of a leading
124
error of the order of Ax 2 results in the following explicit finite-difference equation, using a square grid (Ax = Ay)-
g,i+l,/Jr"
hg, i,j+l = hg, i,j -F - ~
h#,l_l, j-'4he,i, s + he+l,i, j + he_l,i, j
-F Atp
(5) The grid, the boundary nodes, the constant open channel nodes and the internal nodes are shown in Fig. 5. For accuracy, the nodal distance zL~ is taken as ,-i~th of a basin width of 350 m. The thickness of the upper clay layer equals 30 m. In the Laplace's analysis the horizontal hydraulic conductivity ranges between 0.02 and 0.10m/day. Assuming a constant value between these extremes, e.g., Kx = ICy = 0.036 m/day, and substitution of K=0.036, D=30, #=0.02 and L = 3 5 0 yields" At=(KD//~L2)T= 0.00044 T and At/Ax ~ = 0.1, which guarantees stability and a time step At of one day. Laboratory measurements on drilled cores, obtained before deforestation yield p-values of 0.02 (2%}. Recharge--discharge patterns in the research catchment have been computed, by applying eq. 5 to two ~oundwater contour maps. One has been prepared from records of February 13, 1974 {Fig. 6A, winter condition~} and the other from records of May 9, 1974 (F~g. 6B snowmelt conditions).
Boundary nodes
/
,~
~(
x
x
i
x
,
I
'~
K
K
x
K
x
I+
X
X
X
x
internal ~odes
x
x
x
x
X
X
X
"x
X
x
t(
X
x
x
I
]~
x
.--River node~
"P" K ~
0
.
.
.
.
~LI
50
••
-+TZI;IT.'IIIIII ii_~++i ~
!00 m
Fig. 5. Finite-difference grid of tk~e Spring Creek basin for two-dimensional diffusion equation.
125
7
~'/"i@"
"-
~'i
b~>
o
---.
\......
FEBRUARY13, 1974
X ,
,..,o....,n.,.
.
_.-
/
'i I
\..,o ~, :
Im
:
654
';5-~
W ~/:',, ? o , J
o.ou.~,.o~..,o.,., Contour interval
/ !
~ ~
~ ~
,,i x
')/
0
/. • ..ll~l
I
II, /
/
j,/f ."//
/
0 /"
"'o
/t
~"
o
A"
\ o
o
,l~.
d54.
~mmmm~
' --. i
J
I
A,
! /%
MAY 9, 1974
/
• ,t
SNOWMELT
G
nest
x
P,e:ometer
o
Water-table well
•- - . e - - .
Pmlte-&fference grid boundary
.-b~~ ......
Oroundwoter contour ( m ) Contour interval : ] m
r 0
~'st50
.
):
~,,
\k i
o
'....653
~ : 1 lOOm
Fig. 6. Groundwater contour map for the Spring Creek basin of: (A) winter freezing co~lditions, February 13, 1974; and (B) snowmelt conditions, May 9, 1974.
126
The two maps represent the basin after deforestation. The boundary and river hydraulic head nodes have been kept constant and the initial levels for the internal nodes have been obtained from the contour maps as shown in Fig. 6. A comparison of the results (Fig. 7) indicates that recharge--discharge patterns occur, and also that an expansion and contraction of different recharge and discharge areas has taken place after a time lapse of three months. Comparison of Fig. 7A and B indicates also some changes in recharge--discharge rates expressed in mm/day. More detailed computations not reported herein, using daily changes in groundwater levels, suggest that recharge--discharge boundaries change locations on a daily basis. In the summer these transition zones may be recognized by botanical communities that are transitional between shallower and deeper groundwater tables of the discharge, respectively recharge areas. Under winter conditions (Fig. 7A) the catchment soil is snow covered and frozen, and the recharging/discharging groundwaters are therefore not visible. However, the spring recharge--discharge conditions can occasionally be recognized in t h e field. For example, Fig. 8 shows an oblique angle photograph of the 6-ha research catchment taken from a helicopter. It shows the southern discharge area of Fig. 7B, while the sun just happened to reflect on the discharging groundwaters. Average non-steady recharge and discharge groundwater velocities have been integrated from Fig. 7, which graphically portray the computed groundwater flow conditions for the wiuter and the end of the snowmelt period in 1974. The recharge areas (48% of basin Fig. 7 A ) c o n t r i b u t e 13.9m3/day whereas the discharge areas (3.1 ha) contribute 14.3 m3/day. Winter groundwater flow velocities in recharge areas average 0.39 vs. 0.37 mm/day for discharge areas. Conditions at the termination of the snowmelt period {Fig. 7B) result in a recharge rate of 21.1m3/day (56% of catchment} whereas the discharge areas (2.6 ha) contribute 11.2 m3/day. Snowmelt groundwater flow velocities in the recharge areas average 0.5 vs. 0.2mm/day for the discharge areas. The net non-steady groundwater flow, computed for winter conditions, is almost negligible {14.3 -- 13.9 = 0.4 m3/day) and directed upwards. By the end of the snowmelt period the calculated non-steady net groundwater flow is directed downward, amounting to 10m3/day, or 0.16 mm/day from 6ha. The calculated rate of 0.16mm/day compares with a measured non-radioactive groundwater runoff of 0.01 ram/day. The discrepancy of 0.15 mm/day is probably due to evapotranspiration losses on May 9, 1974. The steadystate vertical groundwater seepage (0.3 ram/day, see the previous section} is also lost to evapotranspiration. The eight years of experimenting shows that snowmelt groundwater runoff, which has lost its radioactive tracer ~2sI due to the infiltration effects of the soil~ averages ~55% of the total snowmelt runoff volume before deforestation. This increased slightly to 65% after deforestation. Clearing of the forest is also responsible for a 25% shorter duration of the snowmelt runoff period, as well for a 40% decrease in snowcover sublimation and evaporation. The effect of these three factors is that
127 o ,iuiIU,
o itn,iI,
I, • ,
, • i
•
~b B
j
I
~
¸
o
I
O
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I 50
10b
o
0-6 mm day-1
RECHARGE {
DISCHARGE
m
o~-~
{~
L_ [ ~ ' . ~
WINTER FREEZING
~o,-, ~~r~-:.:...:~
O-Obmm day-I
J
06-25mm dJy "1
j •
®
RECHARGE - DISCHARGE VELOCITIES (mm day -! ) = 0.02 (2°/o)
[ ,
-¢' --:_i ...........
i_._._._
i
f ~o
,oo~
RECHARGE ~
cz ~ISCHARGE ~
®
~ ~
06-2 5 mm doyl
2 5-7ram da~,1 0-06 mm day "1
06- 25 mm day'l
!
:::: - ~
• --.
I
..... iJ
'
..
j~ ' •
.
::::;
~iii!ii!i!!il :::::1 S.OWMELT PE'~,OD RECHARGE
-
D I S C H A R G E
i!i!!!~i
VELOCITIES (ram day ")
.i::]!i::ii::'~
JJ" = 0.02 ( 2 °/o)
2.5-3 mm day I
Fig. 7. Distribui~on of recharge--discharge velocities in the Spring Creek basin in mm/day during: (A) winter freezing conditions; and (B) spring snowmelt conditions.
Fig. 8. Oblique-angle photo of the Spring Creek basin, showing one discharge area where the sun reflects on the water.
t~ G0
129 on the average, the groundwater contribution to streamflow recorded during the annual snowmelt periods, changes from 1.0 to 1 . 9 m m / d a y after deforestation. As Fig. 7B suggests, such net flow rates are derived from expanding and contracting recharge and discharge areas.
STREAMFLOW, GROUNDWATER LEVELS AND GROUNDWATER STORAGE The main annual streamflow event in the small Spring Creek research catchment is the short-lived snowmelt runoff, which takes place mostly within 4--5 weeks during the spring time. After the snowmelt the creek carries very little flow due to high evapotranspiration rates and usually the creek dries up or reduces to a trickle during the summer. The fall rains bring again a short-lived a m o u n t of runoff, before winter sets in. A one-dimensional version of diffusion equation (3) has been applied to groundwater level records (Fig. 9) and to rainfall--runoff data from the 6-ha research catchment in order to explain the reported changes in deforestation:
~}x 2
~)t
(6)
p(t)
The symbols in eq. 6 have been previously defined in the context of eq. 3. When eq. 6 is applied to symmetrical boundaries and special conditions of the research catchment (consisting of an aquifer underlain by an aquitard), an analytical solution of eq. 6 exists (Wesseling, 1973). This solution yields the groundwater level yt midway between two creeks (at x - 0.5L), where L represents t]~e distance between creeks, the total phreatic groundwater storage Rt; and the groundwater runoff qt. 0 E
! m.
~
iI
" ~
.
' )
Ij~ x
I
,
~t \~'~ \':~'1~...
~
'""'"'
' "
III; v \'V,'I I \k,,.."~JJ I/~:%" k~'~l
J
" "
I
~"'..1 ~'~
~!!
-Deforeslat~on
3 1969
1970
1971
1972
i973
1974
YEARS
Fig. 9. Groundwater levels of four wells before and after deforestation in October, 197 t.
130
An analytical solution of eq. 6 can be applied to Spring Creek, if the Dupuit approximation is applicable to the upper lacustrine clay layer. In this research catchment the properties of the clay appears to warrant the application of Kirkham's (1967)slot--slab model. There is evidence that slabs of clay and till have been fractured during or after last glaciation. Kirkham's slot--slab model follows the Dupuit theory. The analytical solution for qt, yt, Rt is" tl
qt.. = ~ (ut.. + p t r . )
(7)
I
where n counts the number of terms (1, 3, 5, 7, . . . ) in an infinite series ( 2 n - - 1 ) ; as n increases, accuracy increases; ut,. is a carry-over term at time t: Ut.n
-"
Ut-l.ne,
+ PtSr.
en = e x p - - [b(2n
(8)
(9)
-- 1)2/j]
where j (in units of time) represents a four-parameter drainage resistance (Krayenhoff van de Leur, 1962)" J = (1/lr2 )I~L2/KD
(10)
in which L is the distance (m) between two parallel ditches or creeks and:
S. = 8/Tr2[1/(2n- 1)2(1--e,,)]
(11)
rn-1--8/Tr2I~1/(2n--1)21
(12)
since" 81~ 2 [
lira n=I,2,3
{ l l ( 2 n - 1) 2 }] = 1
(13)
.....
The groundwater level Yt is given by" n
2n -1
Yt.. = ~ qt..[Dr/{2p(2n-1)}l +pt[4j/(g~)] 7ra/32 -I
E
(1/v)3 1
v= 1 . - 3 . 5 . - 7 . . . .
(14) and the groundwater storage Rt equals:
Rt.. = ~ qt.nj[1/(2n -- 1)] + Pt(8j/lr 2) 1r4/96 -- ~ 1/(2n -- 1) 4 1
1
(15)
Analytical solutions (7}, (14) and (15) have been used to determine the possible effects of deforestation from changing parameter values, through matching of: (1) The recorded monthly groundwater levels with computed levels from 1969--1975 (Fig. 10).
131 --
u,.u
.
I
.
.
Well 5611
0~ Lu
.
.
•
.
.
.
.
.
.
.
•
,
'x
rL
~
,
3
l
It
,gj
i'll
x--x--x
15% Prec,p j :60 days U : 3'/o
~_.,_,
15% Prectp j : 4.5 days U. : 2% "~ - - D e f o r e s t a t i o n
iII
1969
IV11
If
III
1 I,
1970
t l | I T t
IIr
III
IT
I i
1 1 1
1971
I
it
i
1972
t
l
WELL '5611 l
,
l
l
l
t
t
it
I
1973
rT
T ' | t
IT
l i t 1
~ I ~ T
1974
YEARS
Fig. 10. Measured vs. computed groundwater levels before and after deforestation.
(2) The computed groundwater contribution to streamflow during the snowmelt period vs. the groundwater contribution to streamflow measured with the radioactive tracer from 1968--1975 (see Table I). Groundwater levels, groundwater storage and groundwater runoff have been matched before deforestation and after deforestation, using recorded snowmelt and two sets of parameters; j = 45 days and # = 2% before deforesttation, vs. j = 60 days and # = 3% after deforestation. The j-value represents a combination of four catchment characteristics lumped in one parameter: j = (1/Ir)2(I~L2/KD) (dimension: time). Fig. 10 shows that the combination j = 45 days and /~ = 2% matches the groundwater levels of well 5611 (recharge area) slightly better before deforestation. The combination j = 60 days and ~ = 3% appears to fit better the first years after deforestation. Thereafter the combination j = 45 days and # = 2% fits somewhat better. This indicates that deforestation in recharge areas created a temporarily and only slightly disrupted soil profile with an increased effective porosity, thus explaining the measured increase in groundwater runoff. Figs. 9 and 10 show also that deforestation in the recharge areas created a drop of the groundwater table of 0.2--0.5 m, with a gradual recovery two years later (wells 5611 and 5608). Ho~fever, in the discharge areas, groundwater tables rose (Fig. 9, wells 5605 and 5596). The rise of the water tables in discharge areas is associated with decreased effective porosity/~ in those wet and lowlying areas that are compacted during mechanical deforestation operations. Table I shows a comparison of the computed groundwater runoff vs. the groundwater runoff qt measured wi~h the 12sI isotope. The snowmelt in 1974, the last year of simulation, was exceptional (150ram) due to an unusually thick snow cover preceded by a wet fall season in 1973. The average snowwater runoff is usually only 75 ram/yr. Table I shows that j = 45 days and/~ = 2% fits better, whereas the combi-
132 TABLE I
Total groundwater runoff, q t, computed during snowmelt q t, mm per snowmelt period
measured with radioactive tracer
computed with
computed with
j = 45 days, p -- 2%
j = 60 days,/~ = 3% 12.9 15.0 14.9 17.9 ( d e f o r e s t a t i o n ) 24.4 19.0 27.3
1968 1969 1970 1971 1972 1973 ~974
14.6 28.2 17.1 18.7 18.7 22.0 84.6
14.8 24.6 17.0 19.8 27.9 21.1 30.4
Total
203.9
155.6
13! .4
Total excluding 1974
119.3
125.2
104.1
nation j = 60 days and p = 3% proves slightly better shortly after deforestation. The conclusion, drawn from the 8-yr. experiment is, that deforestation of Central Albertan prospective homestead lands has only minor and temporary (1--2 yr.) effects on the hydrology of these lands.
CONCLUSIONS
Clearing operation on Central Albertan forest lands, designated as modernday pioneering farmer homestead lands, involve deforestation and the seeding of grasses An 8-yr. experiment on the impact of deforestation on the hydrology of a research catchment, located in the prospective homestead lands, shows that only minor and temporary impacts can be expected. This conclusion was reached by the use of an 12sI isotope in the experiment. The main changes are a slight increase in total snowmelt runoff from 60 to 75 mm/yr.; a shortening of the snowmelt period from 5 to 4 weeks; an increase of snowmelt groundwater runoff from 1.0 to 1.9 mm/day and a reduction of snow-cover sublimation and evapQration by 40%. Deforestation in groundwater recharge areas creates a temporarily disturbed s3il profile with a slightly increased effective porosity, thus explaining the recc,rded increase in groundwater runoff. In the recharge areas, deforestation creates a drop in the groundwater table of 0.2--0.5 m, with a gradual recovery two years later. However, in the remaining groundwater discharge areas, groundwater tables rise due to a decreased effective porosity associated with mechanized deforestation operations.
133 ACKNOWLEDGEMENTS
Assistance of Mr. V. Da Silva, R. Bohmer and the liason of Dr. J. Toth and computer programming of Mr. A. Litviak is acknowledged. George Holecek, Head Alberta River Basin Research, functioned as project leader.
REFERENCES
Ei'nst, L.F., 1962. Groundwater flow in the saturated zone and its calculation when horizonal parallel open conduits are present. Versl. Landbouwk. Onderz., Wageningen, No. 65-15, pp. 1--189. Holecek, G.F. and Noujaim, A.A., 1975. Separation of the surface and subsurface flow from snowmelt by the use of radioactive tracers. Tokyo Symp. Proc., Int. Assoc. Sci. Hydrol., Publ. 117, pp. 497--505. Kirkham, D., 1967. Explanation of paradoxes in Dupuit--Forchheimer seepage theory. Water Resour. Res., 2(3): 609--622. Krayenhoff van de Leur, D.A., 1962. A study of non-steady groundwater flow. Ingenieur, 2o: B285--B292. Meyboom, P., 1962. Patterns of ground- ter flow in the Prairie Profile (with discussion). Proc 3rd Can. Hydrol. Syrup., Calgary, Alta., pp. 5--33. Smith, G.D., 1965. Numerical Solution of Partial Differential Equations. Oxford University Press, London, 175 pp. Van Beers, W.F.J., 1963. The auger hold method. Int. Inst. Land Reclam. Improv., Wageningen, Bull. No. 1, pp. 1--32. "Van der Molen, W.H., 1972. Non-steady groundwater flow. Lect. M.Sc. Course on Soil Sciences and Water Management, Agric. Univ., Wageningen, Intern. Rep. (unpublished). Van Ouwerkerk, J.H. and Pette, D., 1965. A technique to compute stationary groundwater flow, H. Tijdschr. K. Ned. Heidemaatsch., pp 444--457. Van Royen, R.P., 1906. Reply to discussion of H.E. de Bruyn on The flow of water in the underground. Ingenieur, 26:480--482. Wesseling, J., 1973. Subsurface flow into drains. In: Drainage Principles and Applications, Vol. If, Theories of Field Drainage and Watershed Runoff. Int. Inst. Land Reclam. Improv., Wageningen, Publ. No. 16, pp. 1--31.