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Influence of topography and infiltrability on soil moisture and run-off generation in a small sub-alpine… Williams, Harry F. L. 1982

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INFLUENCE OF TOPOGRAPHY AND INFILTRABILITY ON SOIL MOISTURE AND RUN-OFF GENERATION IN A SMALL SUB-ALPINE WATERSHED by HARRY F.L. WILLIAMS B.Sc.(Hons), Plymouth Polytechnic, Plymouth, England, 1980 A THESIS SUBMITTED IN PARTIAL FULFILMENT OF THE REQUIREMENTS FOR THE DEGREE OF MASTER OF SCIENCE in THE FACULTY OF GRADUATE STUDIES (Geography) We accept this thesis as conforming to the required standard THE UNIVERSITY OF BRITISH COLUMBIA December, 1982 (c)Harry F.L. Williams, 1982 In presenting t h i s thesis i n p a r t i a l f u l f i l m e n t of the requirements for an advanced degree at the University of B r i t i s h Columbia, I agree that the Library s h a l l make i t f r e e l y available for reference and study. I further agree that permission for extensive copying of t h i s thesis for scholarly purposes may be granted by the head of my department or by his or her representatives. It i s understood that copying or publication of t h i s thesis for f i n a n c i a l gain s h a l l not be allowed without my written permission. Department of QBOGCKP\PHY The University of B r i t i s h Columbia 1956 Main Mall Vancouver, Canada V6T 1Y3 Date j2g / Z / 23-DE-6 (3/81) i i ABSTRACT Temporal and spatial variations in soil moisture conditions 1 were studied in a 7500 m sub-alpine watershed, in the Coast Mountains of British Columbia. Measurements of s o i l moisture content were derived from electrical resistance c e l l s , located at 34 sites within the watershed. A series of readings were obtained from the upper and lower part of the s o i l profile at each site, during the snow-free period. The resulting s o i l moisture behaviour patterns were examined in terms of response to r a i n f a l l . Study findings suggest that a model of s o i l moisture behaviour based on watershed topography is not supported by observed s o i l moisture conditions. Re-grouping and re-examination of the data in terms of soil-vegetation complexes, resulted in the emergence of more consistent patterns of s o i l moisture behaviour. These results indicate that s o i l 'heterogeneity was more influential than topography, in determining so i l moisture conditions. The likely implications of this finding for run-off generation are discussed briefly. i i i TABLE OF CONTENTS Page Abstract • • • i i List Of Tables v i i List Of Figures v i i i List Of Symbols xi Acknowledgements x i i i CHAPTER 1 Introduction 1 .0 Purpose 1 1 .1 Objectives 2 1.2 The Hydrologic Significance Of Vegetation 3 1.3 The Hydrologic Characteristics Of Soils 4 1.3.1 The Water Retention Characteristic 4 1.3.2 The Water Flow Characteristic 6 1.3.3 Hysteresis 8 1.4 Factors Influencing Soil Moisture Content 9 1.4.1 Infiltration 9 1.4.2 Antecedent Conditions 10 1.4.3 Infiltration Into Layered Soils 12 1.4.3.a Coarse Over Fine 12 1.4.3.b Fine Over Coarse 13 1.4.3.C Crust-topped And Water Repellent Soils ... 15 1.4.4 Evapotranspiration 16 1.4.5 Soil Moisture Redistribution And Storage Following Infiltration 17 1.5 Topography And Soil Moisture 18 1.6 Theories Of Run-Off Generation 22 iv 1.6.1 I n f i l t r a t i o n E x c e s s O v e r l a n d F l o w 22 1.6.2 S a t u r a t i o n O v e r l a n d F l o w 23 1.6.3 S u b - S u r f a c e S t o r m F l o w 25 1 .7 Summary 26 CHAPTER 2 D e s c r i p t i o n Of S t u d y A r e a 27 CHAPTER 3 E x p e r i m e n t a l Framework And D a t a C o l l e c t i o n 3.0 A p p r o a c h To The P r o b l e m 36 3.1 M e a s u r e m e n t Of S o i l M o i s t u r e C o n t e n t 37 3.2 P r e c i p i t a t i o n M e a s u r e m e n t 39 3.3 S u r f a c e And S u b - s u r f a c e F l o w 39 3.4 S o i l S a m p l e s 41 3.4.1 C e l l C a l i b r a t i o n S a m p l e s 41 3.4.2 O r g a n i c M a t t e r C o n t e n t D e t e r m i n a t i o n s 42 3.4.3 G r a v i m e t r i c D e t e r m i n a t i o n s Of S o i l M o i s t u r e C o n t e n t 45 3.5 B u l k D e n s i t y And P o r o s i t y 45 CHAPTER 4 A n a l y s i s Of D a t a 4.0 The S p a t i a l And T e m p o r a l V a r i a t i o n s I n S o i l M o i s t u r e C o n t e n t 46 4.1 Wa t e r R e t e n t i o n C h a r a c t e r i s t i c C u r v e s 46 4.2 B u l k D e n s i t y , O r g a n i c M a t t e r C o n t e n t And S o i l M o i s t u r e S t o r a g e C a p a c i t y 46 4.3 P r e c i p i t a t i o n 50 4.4 Surface And Sub-surface Flow 54 CHAPTER 5 Analysis Of Errors 5.0 Purpose 55 5.1 Method 55 5.2 Analysis Of Error Based On Water Retention Curves 57 5.2.1 Variability Within A Soil-vegetation Complex 57 5.2.2 Variability Within A Sample 59 5.3 Analysis Of Error Based On Calibration Curves .. 61 5.4 Field Check On Soil Moisture Values Determined By Electrical Resistance Cells 66 5.5 Discussion 66 5.6 Summary 7 6 CHAPTER 6 Analysis Of Results And Discussion 6.0 The Influence Of I n f i l t r a b i l i t y And Topography On Soil Moisture 82 6.1 Run-off Generation 98 Bibliography 101 Appendix A: Major Vegetation Associations 106 Appendix B: Soil Moisture Content Data 107 Appendix C: P r e c i p i t a t i o n Record 109 Appendix D: S o i l M o i s t u r e Response 29/8 - 9/9 111 v i i LIST OF TABLES Page 2.1 Soil-Vegetation Complexes 32 4.1 Calculated Bulk Densities, Porosities And Soil Moisture Storage Capacities 51 4.2 Organic Matter Contents 51 5.1 Mean Deviation Values For Sedge Complex, • Lower Level 58 5.2 Mean Deviation Values For Site 9L 62 5.3 Mean Deviation Values For Site 25Uf 63 5.4 Mean Deviation: Resistance Against Soil Moisture 67 5.5 Bias Correction In The Range 0 - 400 K<fb 76 5.6 Bias Correction In The Range 1000 - 1750 KJL 77 6.1 Median Polish - Upper Soil 90 6.2 Median Polish - Lower Soil 91 6.3 Summary Of Mann-Whitney U Tests 94 v i i i LIST OF FIGURES. Page 1.1 Idealized water Retention Curves 5 1.2 Idealized Water Flow Curves 7 1.3 Relationship Between Infiltration Rate, Rainfall Rate, I n f i l t r a b i l i t y And Time 11 1.4 Appearance Of Finger-Like Protrusions During Infiltration Into A Profile Containing A Fine Layer Overlying A Coarse Layer (Schematic) 14 1.5 Contours And Flow Lines At Right Angles To Contours 19 2.1 Location Of Study Area 28 2.2 Site Topography 29 2.3 Soil Profile In A Sediment Trap 31 2.4 Soil-Vegetation Complexes 33 3.1 Instrumentation 40 3.2 Soil Moisture Calibration Curves For 25U, 14U, 6U 2L, 6L And 1 4L 43 3.3 Soil Moisture Calibration Curves For 31L, 18L, 13L, 9L And 6L 44 4.1 Moisture Retention Curves For 14L, 6L And 2L 47 4.2 Moisture Retention Curves For 31L, 18L, 13L, 9L And 6L 48 ix 4.3 Moisture Retention Curves For 25U, 14U And 6U 49 4.4 Daily Precipitation Totals 52 4.5 Precipitation Of August 31 On An Hourly Basis 53 5.1 Between Samples Mean Deviation 60 5.2 Within Sample Mean Deviation - Cell 9Ld And 25Uf 64 5.3 Mean Deviation About 1:1 Line 68 5.4 Distinction Between Bias And Random Fluctuations .'.71 5.5 Deviations In 0 Of Calibrated Cells From 6 Of Representative Cell 6L In The Lower Sedge 72 5.6 Cell Deviation In The Range 0 - 400 KJl 73 5.7 Cell Deviation In The Range 1000 - 1750 KJi 74 5.8 Correction Of Bias For R1 < 550 Ko7> And R2 > 550 KSL 79 6.1 Upper And Lower Soil Moisture Behaviour - Sedge 83 6.2 Upper And Lower Soil Moisture Behaviour - Heath 84 6.3 Upper And Lower Soil Moisture Behaviour - Debris 85 6.4 Upper And Lower Soil Moisture Behaviour - Hollows 86 X 6.5 Upper And Lower Soil Moisture Behaviour - Straight-Contoured Areas 87 6.6 Upper And Lower Soil Moisture Behaviour - Spurs 88 xi SYMBOLS a Area drained per unit contour length (L z) b A constant D Discharge per unit width of hillslope in a downslope direction (L 3/T) d^/dz Downslope total potential gradient (L/T z) dYm/dz Matric potential gradient (L/T Z) e Rainfall excess intensity (L/T) I Cumulative i n f i l t r a t i o n (L) i Infiltration rate (L/T) K Hydraulic conductivity (L/T) Ks Saturated hydraulic conductivity (L/T) q Rainfall rate (L/T) r Radius of curvature of contours (L) S Sorptivity (L/T1'2) t Time (T) x Distance from the divide (L) z Depth (L) 9 Soil moisture content on a volume basis (L 3/L 3 %) 9i I n i t i a l s o i l moisture content ©m Soil moisture content on a mass basis (M/M %) 9s Saturated so i l moisture content 9t Soil moisture content of transmission zone 9tq Moisture content of transmission zone in equilibrium with r a i n f a l l rate q 9w Soil moisture content at wilting point -j Soil moisture tension (L /T ) Mean value Constant for any given s o i l , related pore size distribution xi i i ACKNOWLEDGEMENTS Thanks are due to my supervisor - Dr. H. Olav Slaymaker and my second reader - Dr. Mike Church, for providing personal encouragement, academic guidence, financial support and for c r i t i c a l l y reviewing earlier versions of this thesis. I also thank the third member of my committee - Dr. R. Allan Freeze, for providing some essential equipment and the use of a laboratory, during the calibration phase of this work. I thank the U.B.C. Natural, Applied and Health Sciences Grants Committee and the Natural Sciences and Engineering Research Council for support of my f i e l d work through operating grants 26-9601 and 67-7073, which were awarded to Dr. Slaymaker. I thank Cathy Lewis for her assistance in the f i e l d . Al and Marti Shaehli are thanked for their kind hospitality. I also thank Dr. J. Ross MacKay for the loan of some equipment, and Thomas Gallie, Gary Barrett and Dr. Jan DeVries, for helpful discussions and suggestions. 1 CHAPTER 1.  Introduction.  1.0 Purpose. In the Coast Mountains of British Columbia, a large proportion of run-off is derived from alpine and sub-alpine areas, yet there has been relatively l i t t l e research into the hydrologic behaviour of watersheds at high altitudes. It is of considerable interest and importance to be able to determine the hydrology of these watersheds. Increased understanding of watershed response will result in a more effective modelling capability and thus enhance run-off prediction (Betson 1964). Established theories of run-off generation are based in large part on assumptions about, and hypothetical relationships between, the hydrologic properties of soil-vegetation complexes and watershed topography (Horton 1933, 1945; Kirkby and Chorley 1967; Dunne and Black 1970b; Hewlett and Hibbert 1967). So i l -vegetation complexes act as regulators on throughput within a watershed. The manner in which a soil-vegetation complex acts on a given input is determined by its hydrologic characteristics, the nature of the input and antecedent conditions (Calver, Kirkby and Weyman 1972; Carson and Kirkby 1972). The topography of a watershed also exerts some control over hydrologic throughput; f i r s t l y , by influencing the nature of moisture flows both within, and between, soil-vegetation complexes; and secondly, by influencing i n f i l t r a t i o n into the soil as a major control on depression storage (Horton 1935). 2 T h e r e f o r e , the response of a watershed to h y d r o l o g i c events depends in l a r g e part on the topography of the watershed, and on the h y d r o l o g i c c h a r a c t e r i s t i c s , d i s t r i b u t i o n and i n t e r a c t i o n of the s o i l - v e g e t a t i o n complexes w i t h i n that watershed. The i n t e r a c t i o n of topography and s o i l - v e g e t a t i o n complexes o f t e n r e s u l t s i n a h y d r o l o g i c response which i s both s p a t i a l l y and temporally v a r i a b l e . T h i s i s i n d i c a t e d by s p a t i a l l y and temporally v a r i a b l e changes i n s o i l moisture content. 1 . 1 O b j e c t i v e s . T h i s study i s based on the p r o p o s a l that the key to understanding watershed response l i e s i n the i n t e r a c t i o n of s o i l - v e g e t a t i o n complexes and topography. I t i s acknowledged that other f a c t o r s , such as g e n e r a l c l i m a t e and man-induced changes, w i l l a l s o i n f l u e n c e watershed response over l a r g e r r e g i o n s . Snowmelt i s by f a r the major component of annual r u n - o f f i n a s u b - a l p i n e environment; t h i s , coupled with the w e l l developed groundwater system of h e a v i l y f r a c t u r e d bedrock, l a r g e l y determines the annual r u n - o f f regime i n - t h i s a r e a. T h i s study, however, i s concerned only with the snow-free p e r i o d . Inputs d u r i n g t h i s time are i n the form of r a i n f a l l and r u n - o f f response i s r e g u l a t e d by the h y d r o l o g i c and topographic p r o p e r t i e s of the s u r f i c i a l m a t e r i a l s . Consequently, i t i s d u r i n g t h i s p e r i o d that the i n f l u e n c e s of topography and s o i l v a r i a b i l i t y on s o i l moisture c o n d i t i o n s and run-off g e n e r a t i o n , are most pronounced. 3 The aims of the study are: 1. to determine the temporal and spatial variations of so i l moisture content within a small sub-alpine watershed in the Coast Mountains of British Columbia during the snow-free period; 2. to determine the relative importance of soil-vegetation complexes and topography, in terms of their influence on surface and sub-surface s o i l moisture conditions; 3. to briefly describe the probable consequences of the study findings for run-off generation. 1.2 The Hydrologic Significance Of Vegetation. The role of vegetation in the watershed hydrologic cycle, involving such processes as interception, transpiration, and so on, is described adequately in many standard texts (see for example Ward, 1967), and no further general review is deemed necessary here. In the present study, vegetation is recognized as a major factor in determining hydrologic behaviour, as is indicated by its incorporation into the term "soil-vegetation complex" (for definition, see Chapter 2), which, i t is proposed, is a basic hydro-geomorphic unit, forming a suitable basis for hydrologic study (Chapter 3 - see also Gallie and Slaymaker 1983). It is anticipated that, for this study, the effect of vegetation on i n f i l t r a b i l i t y w ill be of particular importance. Vegetation cover is directly related to i n f i l t r a b i 1 i t y by the conditioning effect that organic matter has on soil structure, bulk density and porosity. Dead and decaying root material contributes to the development of macro-porosity, promoting a more open structure and increasing the inherent permeability of 4 the root zone. In modifying the soil's i n f i l t r a b i l i t y , vegetation may influence both s o i l moisture behaviour and run-off generation (sections 1.4.5, 1.6) - two central factors in this study. 1.3 The Hydrologic Characteristics Of Soils. The hydrologic behaviour and response of a soi l depends in part on two important hydrologic characteristics of that s o i l . 1. The water retention characteristic describes the relationship between so i l moisture content and s o i l moisture tension. This gives information about the soil's a b i l i t y to store or retain water, for example, against the force of gravity. 2. The water flow characteristic describes the relationship between hydraulic conductivity and soil moisture tension. This relationship yields information about the ab i l i t y of a s o i l to transmit water, for example, during the process of i n f i l t r a t i o n . 1.3.1 The Water Retention Characterstic. The a b i l i t y of a soil to retain water is expressed by the water retention characteristic, which is the functional relationship between the soi l moisture content and the corresponding matric potential. The latter, also known as capillary potential, is a negative pressure potential (tension) resulting from the capillary and adsorptive forces due to the soil matrix ( H i l l e l 1971, pp. 50 and 57).This relationship is most commonly expressed as a curve called the water retention curve ( f i g . 1.1). Tension is applied to a saturated s o i l until the air-entry 5 S o i l M o i s t u r e T e n s i o n F i g : 1.1 I d e a l i z e d Water R e t e n t i o n Curves 6 tension is reached: at th i s point the largest pores empty and the s o i l moisture content is reduced r a p i d l y . A gradual increase in tension w i l l resu l t in the emptying of progress ive ly smaller pores. The amount of moisture retained at r e l a t i v e l y low tensions (up to one bar) depends p r i m a r i l y upon the c a p i l l a r y effect and the pore s ize d i s t r i b u t i o n , and hence is strongly affected by the s tructure of the s o i l . Moisture retent ion at higher tensions is due increas ingly to adsorption and thus influenced less by the structure and more by the texture and spec i f i c surface area of the s o i l m a t e r i a l . Coarse-textured s o i l s with more uniform pore s izes may exhib i t c r i t i c a l a i r - e n t r y phenomena, such as a rapid reduction in s o i l moisture content upon reaching the a i r - e n t r y value of matric tension ( f i g . 1.1). Clayey s o i l s , in general , have a higher moisture content at any p a r t i c u l a r tension and exhib i t a more gradual slope with increasing tens ion, due to a more uniform pore s ize d i s t r i b u t i o n and greater amounts of absorbed moi s ture . 1.3.2 The Water Flow C h a r a c t e r i s t i c . A s o i l ' s water flow c h a r a c t e r i s t i c curve expresses the r e l a t i o n s h i p between hydraul ic conduct iv i ty and s o i l moisture tension ( f i g . 1.2). The a b i l i t y of a s o i l to conduct water has a maximum value at sa turat ion , and decreases as the s o i l water content and corresponding matric po tent ia l decrease. In a saturated s o i l , a l l the pores are f i l l e d and conducting, so the conduct iv i ty i s at a maximum value, Ks. When the s o i l becomes unsaturated, some of the pores become a i r f i l l e d and the conductive port ion of the s o i l ' s c ros s - s ec t iona l area 7 Ks sand S o i l Moisture Tension F i g : 1.2 Idealized Water Flow Curves 8 decreases correspondingly. At saturation, the most conductive soils are those in which large and continuous pores constitute most of the overall pore volume. Thus, a sandy s o i l conducts water more rapidly than a clayey s o i l . However, with the application of tension, the relatively large pores of a sandy s o i l quickly empty and become non-conductive, thus steeply decreasing the i n i t i a l l y high conductivity. In a s o i l with small pores, many of the pores remain f u l l and conductive even at appreciable tensions, so that the hydraulic conductivity does not decrease as steeply and may actually be greater than that of a soil with large pores subjected to the same tension. 1.3.3 Hysteresis. Both the water retention characteristic and the water flow characteristic are subject to the phenomenon of hysteresis. The relationships described above are, in general, not unique and single-valued ones. The causes of hysteresis have been described adequately elsewhere (Baver, Gardner and Gardner 1972; H i l l e l 1971) and will not be discussed here. The consequence of hysteresis is that the equilibrium s o i l moisture content (and thus hydraulic conductivity) at a given matric potential is greater in desorption (drying) than in sorption (wetting). The change in s o i l moisture content which results when sorption changes to desorption or vice visa, is described by scanning curves ( f i g . 1.1). Hysteresis may be particularly important in a f i e l d situation, where so i l moisture changes tend not to be monotonic (i.e.wetting or drying), but tend to be composite (e.g. 9 simultaneous wetting and drying during moisture redistribution) and sequential (e.g. during intermittent ra i n f a l l events). 1.4 Factors Influencing Soil Moisture Content. 1.4.1 In f i l t r a t i o n . Infiltration is the term applied to the process of water entry into the s o i l . Infiltrating water has two components, a transmission component which is constant, representing a steady flow through the s o i l ; and a diffusion component, which is an i n i t i a l l y rapid, and then increasingly slow, filling-up of a i r f i l l e d pore spaces as i n f i l t r a t i o n continues. Both components are directly related to s o i l moisture content. The transmission component is equal to K(Ot), the hydraulic conductivity of the transmission zone at s o i l moisture content Ot. The diffusion component is a function of the sorptivity of the s o i l , which depends largely on the matric potential gradient between the surface and the wetting zone. As i n f i l t r a t i o n continues, the depth to the wetting zone (dz) increases and the matric potential gradient (d m/dz) approaches zero; thus the diffusion component becomes negligible at large values of t. Consequently, the flux that the s o i l profile could absorb through it s surface if i t were maintained in contact with water at atmospheric pressure - termed the soil i n f i l t r a b i l i t y decreases as i n f i l t r a t i o n progresses. The general form of Philip's (1957) equation describes cumulative i n f i l t r a t i o n thus; 10 l=St*+bt 1 Where b=a constant S=sorptivity Differentiating gives the i n f i l t r a t i o n rate; i= ist^+b 2 At large values of t, 1 St approaches zero and the i n f i l t r a t i o n rate becomes equal to the constant b, where b is equal to K(9t), the hydraulic conductivity of the transmission zone. The s o i l moisture content of the transmission zone thus adjusts i t s e l f to the r a i n f a l l rate, such that K(0t), the hydraulic conductivity of the transmission zone, is equal to the r a i n f a l l rate q. If the r a i n f a l l rate is less than the saturated hydraulic conductivity of the transmission zone, i n f i l t r a t i o n is flux-controlled, as the i n f i l t r a t i o n rate is equal to the ra i n f a l l rate. The i n f i l t r a t i o n process becomes profile-controlled i f the ra i n f a l l rate exceeds the saturated hydraulic conductivity of the transmission zone for a time sufficient for the i n f i l t r a t i o n rate to f a l l to Ks (fig. 1.3). The i n f i l t r a t i o n rate then equals Ks, the transmission zone is saturated (assuming negligible air entrapment) and excess water ponds on the surface, giving rise to a potential for surface run-off. 1.4.2 Antecedent Conditions. The i n i t i a l soil moisture content prior to i n f i l t r a t i o n will affect the soil's i n i t i a l i n f i l t r a b i l i t y , the rate of advance of the wetting front and the shape of the s o i l moisture 1 ) . R a i n f a l l r a t e < Ks Ks 11 I r i f i l t r a b i l i t y R a i n f a l l r a t e I n f i l t r a t i o n r a t e 2 ) . R a i n f a l l r a t e = Ks \ I n f i l t r a b i l i t y R a i n f a l l r a t e Ks I n f i l t r a t i o n r a t e R a i n f a l l r a t e Ks 3 ) . R a i n f a l l r a t e > Ks \ I n f i l t r a b i l i t y E x c e s s w a t e r I n f i l t r a t i o n r a t e F i g : 1.3 R e l a t i o n s h i p Between I n f i l t r a t i o n R a t e , R a i n f a l l R a t e , I n f i l t r a b i l i t y And Time. 12 profile after i n f i l t r a t i o n has ceased. A dry s o i l will have a higher i n i t i a l i n f i l t r a b i l i t y than a moist s o i l , because the matric potential gradient, and thus sorptivity, is greater for a dry s o i l than a moist s o i l . The rate of advance of the wetting front will be greater in a moist s o i l , since less moisture has to be diverted into storage in order to increase the i n i t i a l soil moisture content Gi to 9tq, the s o i l moisture content of the transmission zone in equilibrium with a r a i n f a l l rate q. Considering continuity, i t follows that; dz/dt=i/(0tq-0i) 3. Where z=Depth to wetting front. At large values of t,i=q (or i=Ks,if q>Ks), therefore; dz/dt=q/A9 4. Where Ae=0tq-9i Therefore, the higher the i n i t i a l moisture content, the lower the value of Ae and the more rapid the advance of the wetting front. The rate of advance of the wetting front has important implications for soil moisture changes at depth within the s o i l , especially, for example, where perched water tables develop above less permeable layers within the s o i l profile. 1.4.3 Infiltration Into Layered Soils. 1.4.3.a Coarse Over Fine. Ultimately the maximum hydraulic conductivity of any soil profile is determined by the saturated hydraulic conductivity of the less permeable layer within the profile. 13 In the case of a coarse layer of higher saturated hydraulic conductivity overlying a finer-textured layer, the i n f i l t r a t i o n rate is at f i r s t controlled by the coarse layer, but wi l l then be determined by the finer layer when the wetting front reaches i t . If the i n f i l t r a t i o n rate in the coarse layer is higher than the saturated hydraulic conductivity of the fine layer, a saturated zone can build up above the fine layer, resulting in the development of a perched water table. 1.4.3.b Fine Over Coarse. In a profile consisting of a fine layer overlying a coarse layer, i n i t i a l i n f i l t r a t i o n is again controlled by the upper layer. The advance of the wetting front may be halted at the boundary with the coarse layer since moisture at the wetting front is normally under suction, which may be too high to permit entry into the relatively large pores of the coarse layer. Infiltration will continue into the fine layer until the pressure head at the interface builds up sufficiently to penetrate the coarse layer. When the wetting front does penetrate the coarse layer, i t often takes the form of saturated, finger-like protrusions, rather than a uniform wetting front ( f i g . 1.4). The fingering phenomenon is s t i l l incompletely understood, but i t has been recognized and studied by H i l l and Parlange (1972), Raats (1973) and Philip (1975). Fingering has important implications for so i l moisture changes at depth within the profile in that i t provides a 14 q Fine W e t 9 increasing with depth in fine layer Coarse Dry Fig: 1.4 Appearance Of Finger-Like Protrusions During Infil t r a t i o n Into Profile Containing A Fine Layer Overlying A Coarse Layer (Schematic). 15 mechanism f o r r a p i d conductance of water to lower l e v e l s i n the s o i l p r o f i l e . C o n d u c t i o n of water in the s a t u r a t e d f i n g e r s i s much f a s t e r than the u n s a t u r a t e d flow through the c o a r s e l a y e r , which would r e s u l t were the w e t t i n g f r o n t to p e n e t r a t e un i f o r m l y . 1.4.3.C C r u s t - t o p p e d And W a t e r - r e p e l l e n t S o i l s . I n f i l t r a t i o n i n t o c r u s t - t o p p e d and w a t e r - r e p e l l e n t s o i l s i s a s p e c i a l case of i n f i l t r a t i o n i n t o l a y e r e d s o i l s , where a c o n t r a s t i n g l a y e r of very low p e r m e a b i l i t y , perhaps o n l y a few m i l l i m e t r e s t h i c k , o c c u r s a t , or n e a r , the top of the s o i l p r o f i l e . The r e d u c t i o n of p e r m e a b i l i t y at the top of the p r o f i l e may be due to the f o r m a t i o n of a c r u s t of f i n e m a t e r i a l formed by the detachment and m i g r a t i o n of p o r e - b l o c k i n g p a r t i c l e s under the b e a t i n g a c t i o n of r a i n d r o p s (Horton 1933); or to the a c c u m u l a t i o n of hydrophobic o r g a n i c compounds near the s o i l s u r f a c e . The e f f e c t of the c r u s t or hydrophob ic l a y e r i s to g r e a t l y reduce i n f i l t r a b i l i t y , due to i t s v e r y f i n e t e x t u r e or w a t e r - r e p e l l e n t n a t u r e . At l a r g e t i m e s , the maximum i n f i l t r a t i o n r a t e of the p r o f i l e can be no g r e a t e r than the s a t u r a t e d h y d r a u l i c c o n d u c t i v i t y of the impeding l a y e r . The s u r f a c e s of such s o i l s may become e a s i l y s a t u r a t e d due to t h e i r low c o n d u c t i v i t y . S u r f a c e ponding may thus o c c u r , g i v i n g r i s e to a p o t e n t i a l f or i n f i l t r a t i o n - e x c e s s o v e r l a n d f l ow , or f i n g e r i n g in the s u b - s o i l i f the s u r f a c e p r e s s u r e head i s of s u f f i c i e n t magni tude . 16 1.4.4 E v a p o t r a n s p i r a t i o n . E v a p o t r a n s p i r a t i o n i s the l o s s of m o i s t u r e from the s o i l t h r o u g h the combined e f f e c t s of e v a p o r a t i o n from the s o i l and t r a n s p i r a t i o n from p l a n t s u r f a c e s . A net t r a n s f e r of water m o l e c u l e s i n t o the a i r o c c u r s o n l y i f t h e r e i s a vapour p r e s s u r e g r a d i e n t between the e v a p o t r a n s p i r i n g s u r f a c e and the a i r and an a v a i l a b l e source of l a t e n t heat energy t o f a c i l i t a t e a change of s t a t e from l i q u i d t o vapour. The e x t e r n a l heat source may be s o l a r r a d i a t i o n or s e n s i b l e heat from the atmosphere or ground. In g e n e r a l , s o l a r r a d i a t i o n i s the p r i n c i p a l energy source f o r e v a p o t r a n s p i r a t i o n . Other f a c t o r s which d e t e r m i n e the e v a p o t r a n s p i r a t i o n r a t e have been c o v e r e d elsewhere ( M o n t e i t h 1973) and w i l l not be d i s c u s s e d h e r e . The depth t o which e v a p o t r a n s p i r a t i o n o p e r a t e s i n the s o i l p r o f i l e v a r i e s a c c o r d i n g t o such f a c t o r s as r o o t zone d e p t h , a e r a t i o n zone d e p t h , the presence and depth of c r a c k s i n the s o i l and so on. G e n e r a l l y , i n u n v e g e t a t e d s o i l s , e v a p o t r a n s p i r a t i o n i s most a c t i v e a t the s u r f a c e where energy r e c e i p t s a r e a t a maximum, and d e c r e a s e s w i t h d e p t h . In v e g e t a t e d s o i l s e v a p o t r a n s p i r a t i o n may have the g r e a t e s t e f f e c t i n the r o o t zone due t o m o i s t u r e uptake by p l a n t s ( t r a n s p i r a t i o n ) . E v a p o t r a n s p i r a t i o n i n the absence of a water t a b l e ( i . e . i n the absence of s t e a d y - s t a t e upward f l o w of water) r e s u l t s i n d r y i n g of the s o i l . The g r a d u a l r e d u c t i o n of s o i l m o i s t u r e c o n t e n t w i l l e v e n t u a l l y r e s u l t i n a d e c r e a s e of the e v a p o t r a n s p i r a t i o n r a t e , as the h y d r a u l i c c o n d u c t i v i t y of the 1 7 s o i l is reduced. In the case of bare s o i l , the evaporation rate in the early stages of drying can remain fa i r l y constant, as -the increasing matric potential gradient towards the surface tends to compensate for the decreasing hydraulic conductivity. However, eventually the surface s o i l becomes air-dry and the suction gradient can no longer increase, in fact the gradient decreases as the lower soil continues to dry. The combination of a decreasing potential gradient and decreasing hydraulic conductivity inevitably results in a reduction of evaporative flux. 1.4.5 Soil Moisture Redistribution And Storage Following Infiltration. When i n f i l t r a t i o n ceases, the downward movement of water in the s o i l often continues, resulting in a redistribution of moisture within the s o i l profile. The downward movement of s o i l moisture results from a matric potential gradient between the upper moist so i l and lower drier s o i l , and the influence of gravity. The rate of redistribution generally decreases in time, for two reasons; 1. Transfer of moisture from the upper to lower zone decreases the suction gradient between them. 2. The hydraulic conductivity of the upper zone decreases as i t loses moisture. The hydraulic properties of the s o i l profile thus affect redistribution. The hydraulic conductivity of a coarse s o i l , for example, f a l l s off rapidly with increasing suction; whereas the reduction of conductivity in a fine s o i l is more gradual. This 18 is reflected in the change of soi l moisture content of fine and coarse soils during redistribution. The dependence of desorption characteristics on so i l texture results in differences in the moisture retention properties of fine and coarse soils. In general, fine soils will contain a higher soil moisture content than coarse soils at any given time after i n f i l t r a t i o n ceases. The desorption curves of both fine and coarse soils eventually flatten off and further redistribution continues at a very low rate. Desorption of an i n i t i a l l y moist soil proceeds via gravity drainage of "free water" (i.e. water which can not be held by matrix and/or adsorptive forces against the force of gravity). The moisture content at which gravity drainage ceases is known as " f i e l d capacity" (Veihmeyer and Hendrickson 1949). Further desorption may continue by evapotranspiration and under the influence of matrix suctions from adjoining s o i l , but the desorption curve w i l l eventually flatten off and further desorption will become negligible, at this stage the soi l is said to be at its "wilting point". 1.5 Topography And Soil Moisture. Topography influences s o i l moisture by affecting the flows of moisture within a watershed and surface depression storage, a function of gradient which can influence i n f i l t r a t i o n and surface run-off. Generally, surface and sub-surface flow lines are at right angles to the contours on a hillslope. Curved contours produce flow lines which are not parallel, but converge or diverge (fig. 1.5). In an area of concave outwards contours, or a hollow, flow Fig: 1.5 Contours (solid lines) And Flow Lines (broken lines) At Right Angles To Contours. (a) Straight contours; a = x (b) Contours concave outwards (hollow) ; a > x (c) Contours convex outwards (spur); a < x Where a = area drained per unit contour length x = distance from divide along flow line (Adapted from Carson and Kirkby 1972) 20 lines converge causing a progressive concentration of moisture downslope (fi g . 1.5.b). Where contours are convex outwards, as on a spur, flow lines diverge and moisture flows diminish at successive points downslope, in comparison with an area of straight contours, (fig. 1.5.c). Considering that hydraulic conductivity is a function of soi l moisture content (i.e. K=K(6)), the water discharge per unit width of homogeneous hillslope measured in a downslope direction is related to soi l moisture content by Darcy's Law: D = K(9) d$/dz 5 Where D = Discharge per unit width in a downslope direction d$/dz = Downslope total potential gradient Although the relationship between K and 0 is not unique (section 1.3.3), Brooks and Corey (1966) found that a reasonable approximation i s : K oc- ee 6 Where £ = A constant for any given s o i l , related to pore size distribution K can also be defined as: K = Ks.S£ 7 Where S = e-0w/(0s-0w) 0<S<1 8 0w= © at wilting point 0s= Saturated moisture content Ks= Saturated hydraulic conductivity 21 Sustituting 7 in 5 gives: D = Ks.Se.d$/dz 9 Under average conditions the mean flow is equal to the product of the mean rainfa l l excess intensity (drainage or "free water" - see section 1.4.5), e, and the area drained per unit contour length, a, (fig. 1.5): Therefore under equilibrium conditions there is a tendency for D to increase downslope as "a" increases. This requires a corresponding increase in 0 downslope, which increases the Ks.S£ term in equation 9 and balances the two sides of the equation. A reduction of gradient downslope w i l l also require an increase in 0 since the total potential gradient is reduced; Ks.S£ is then increased to maintain the discharge D (Whipkey and Kirkby 1978). The tendency for 0 to increase downslope is greater in hollows, since "a" is increasing faster than "x" the distance from the divide (fig. 1.5.b) and consequently D is increased by geometric flow convergence. On spurs the tendency for D, and thus 0, to increase downslope is reduced since the area drained per unit contour length is less than "x" (i.e. the flow is divergent). Thus, © is dependent on contour curvature such that: Where r = Radius of contour curvature (positive in hollows) This expression suggests that hollows have higher basal moisture contents, and increase their moisture contents to a greater D = e.a 10 9<*e. (] + yr ) 1 1 22 extent during storms, than areas with straight contours. This has three important implications; 1. Hollows, having higher than average moisture contents, are more commonly areas of saturated soil in humid environments. 2. Hollows can support higher evapotranspirat ion rates than adjacent straight-contoured areas, due to a lower so i l moisture d e f i c i t . 3. The higher basal moisture content and more rapid response to storms produce saturation, and thus surface water, more frequently in hollows than in other areas. For a soil of constant surface roughness, the capacity of surface depression storage varies in inverse proportion to surface gradient. Increasing the gradient causes a corresponding decrease in depression storage capacity. This has important implications for i n f i l t r a t i o n , since ponded surface water creates a pressure head which increases the in f i l t r a t i o n rate and often constitutes the necessary condition to initiate the process of fingering. The surface depression storage capacity also influences the timing and production of overland flow, since surface run-off cannot commence until the depression storage has been satisfied (Kirkby 1969). 1.6 Theories Of Run-Off Generation. 1.6.1 Infiltration-excess Overland Flow. Infiltration-excess overland flow, described by Horton (1933, 1945), is the flow of water over the s o i l surface which results when the soils i n f i l t r a b i l i t y is exceeded and surface 23 depression storage is satisfied. By definition, the surface of the s o i l becomes saturated, while the underlying s o i l may remain in an unsaturated condition. Thus i t is the hydraulic properties of the soil's surface which control infiltration-excess overland flow production. Crust-topped and water-repellent soils may be particularly prone to infiltration-excess overland flow, since it is the saturated hydraulic conductivity of the surface crust or water-repellent layer which determines the effective i n f i l t r a b i l i t y of the s o i l profile. Antecedent so i l moisture w i l l be important in determining the i n i t i a l i n f i l t r a b i l i t y of the s o i l , but at large times the saturated hydraulic conductivity, and its relation to r a i n f a l l intensity, will become the controlling factor on overland flow product ion. Topography will be important for determining the flow path of the surface run-off produced and whether the flow w i l l be convergent or divergent. Contiguity of flow paths w i l l determine whether the surface run-off flows into the perennial channel system or is in f i l t r a t e d downslope in an area of greater i n f i l t r a b i l i t y . This will be determined by the surface properties of the soil-vegetation complex(es) downslope of the area of overland flow production. 1.6.2 Saturation Overland Flow. Saturation overland flow occurs where the soil becomes saturated up to the surface as a result of the following processes, operating singly or in combination; 1. A rising water table. 2. A concentration of throughflow from upslope. 24 3. A build up of percolating s o i l moisture above a less permeable horizon within the s o i l profile. Thus, saturation overland flow can occur at r a i n f a l l intensities less than the effective saturated hydraulic conductivity of the so i l profile. The c r i t i c a l condition for saturation overland flow production is that s o i l moisture storage of the potential saturated zone must be satisfied. The net volumetric input of moisture thus becomes more important than the rate of input. The nature of the above processes confines the occurrence of such saturated areas to certain locations; these are (Kirkby and Chorley 1967 ); 1. at the base of slopes, immediately adjacent to water f i l l e d channels. 2. in hollows. 3. in slope profile concavities. 4. in areas of thin or less permeable soils. These locations illustrate the key role played by topography in the production of saturation overland flow. Certain configurations of topography, for example slope profile concavities, are conducive to the convergence of lateral seepage in the s o i l profile. This may not only be important for the development and growth of saturated areas during r a i n f a l l , but may also influence redistribution of soil moisture after i n f i l t r a t i o n has ceased. Topography may, therefore, be reflected in the pattern of antecedent soil moisture. The areas of a watershed that produce saturation overland flow may be either "partial source areas", which occupy fixed 25 positions within the watershed, (Betson 1964) or dynamic "variable source areas" (Hewlett and Hibbert 1967, Dunne and Black 1970b), capable of expanding or contracting in size, depending on the hydrologic input. 1.6.3 Sub-surface Stormflow. Sub-surface stormflow consists of rapid throughflow within the so i l profile, capable of contributing to storm run-off (Whipkey 1965). In general, flow through the so i l matrix is too slow to contribute significant amounts of storm run-off. This is especially true of unsaturated flow, when hydraulic conductivities are usually very low. The build up of a saturated layer above a less permeable horizon, may lead to appreciable saturated flow, but only i f the matrix is very coarse textured and has a relatively high saturated hydraulic conductivity. Rapid throughflow may also take place in sub-surface channels or pipes. Several origins have been proposed for pipes, including; decayed root systems (Whipkey 1969), the enlargement of s o i l cracks (Parker 1963, Weyman 1974), mammal burrows (Newson 1969) and flow through interconnecting macro-pores of buried palaeorills (Huggett 1974). Further reviews of pipes and their potential significance have been provided by Jones(l97l, 1975, 1978) and will not be attempted here. A pre-condition for flow to commence is that the surrounding matrix be saturated (Cheng et al 1975). The formation of a perched water table above a less permeable so i l horizon may thus provide the necessary condition for pipe flow to commence. Rapid saturation of the pipe region, and thus rapid pipe response, may be achieved by fingering in the overlying 26 so i l horizons. Saturated flow through vertical fingers could provide rapid conductance of water down to the sub-surface zone of saturation. 1.7 Summary. This chapter draws attention to the importance of developing a better understanding of the hydrologic response of watersheds at high altitudes. Such research has implications for run-off prediction in the high altitude watersheds which constitute an early link in the Coast Mountain drainage systems. Several run-off generating mechanisms are described, a l l of which involve assumptions about, and hypothetical relationships between, the hydrologic properties of soil-vegetation complexes and watershed topography. The purpose of this study then, is to investigate the influence of topography and i n f i l t r a b i l i t y on surface and sub-surface s o i l moisture behaviour, assess their relative importance and briefly describe the probable consequences of the study findings for the assessment of appropriate run-off generating mechanisms. The investigation will be based on measurements of in-situ s o i l properties and the hydrologic response of the watershed. The latter involves consideration of both spatial and temporal variability of s o i l moisture content. 27 CHAPTER 2. Description Of Study Area. The f i e l d site was a small sub-alpine watershed, situated at a height of approximately 1800m amsl in the Pacific Ranges of the Coast Mountains, Southwestern British Columbia-. The site is located on the crest of a ridge, adjacent to the Lillooet River, above Pemberton Meadows, B.C. (fig. 2.1). Annual precipitation is in excess of 1600 mm, with a strong winter maximum. At the elevation of the study area, most of the precipitation f a l l s as snow. Not surprisingly, most run-off from the study area occurs from May through July as summer snowmelt, but rainstorms during the snow-free period can generate significant F a l l run-off events. The watershed covers an area of approximately 7500 m and has an overall relief of 20 m. It forms an elongated cavity oriented in a south-east to north-west direction. Surface gradients range from approximately 0 to 45 . Contour configurations include both pronounced concavities and convexities, as well as a sunken hollow at the northern edge of the watershed. The topography of the site is shown in detail in f i g . 2.2. The bedrock consists of heavily fractured gneisses and schists, a roof pendant of the Gambier Group, overlying the granodiorite of the Coast Mountains Pluton (McKee 1972). Surficial materials consist of a thin, discontinuous d i o r i t i c basal t i l l of late Pleistocene origin, overlain by 28 Fig: 2.1 Location Of Study Area. (Adapted from Barrett 1981) 29 • Fig.. 2.2 Site Topography. 30 volcanic ash from the Mt. Mazama (6600 B.P.) and Bridge River (?) (2600 B.P.) volcanic events, interspersed with s i l t y loess and organic deposits. The depositional sequence is best preserved in localized sediment traps (fig. 2.3). The upper ash is less weathered and coarser than the lower ash. Both ashes are coarser than the loess deposit, resulting in localized profile sequences of textural changes. The soils developed in these materials are mainly acidic brunisols and regosols. The brunisols are weakly developed soils, located in stable, well vegetated sites, containing distinct s o i l horizons. The regosols are developed on geomorphically active, or recently active, sites. They consist of colluvium overlying bedrock or t i l l . The so i l horizons are poorly developed and the ash layers, although recognizable in places, are highly disturbed and discontinuous. Four major vegetation communities have been identified within the watershed, these are; 1 . tree islands. 2. heath communities. 3. sedge communities. 4. sparsely vegetated, active sites. The species compositions of the associations are listed in appendix A (Gallie, personal communication 1982). On the basis of so i l type, vegetation cover, .slope position and general stability, a number of soil-vegetation complexes can be identified (Table 2.1). The distribution of these s o i l -vegetation complexes within the watershed is shown in f i g . 2.4. Some sedge communities in this area are known to display 31 loess upper ash loess/organic deposits 30 lower ash loess basal t i l l | I Fig: 2. 3 Soil Profile In A Sediment Trap. T a b l e 2.1 S o i l - V e g e t a t i o n Complexes. S t a b l e , w e l l d e v e l o p e d s o i l h o r i z o n s , w e l l v e g e t a t e d . A s h , l o e s s and o r g a n i c m a t t e r o v e r t i l l . U n s t a b l e o r p o o r l y d e v e l o p e d s o i l h o r i z o n s , p o o r l y v e g e t a t e d . C o l l u v i u m o v e r b e d r o c k o r t i l l . COMPLEX PERCENTAGE OF WATERSHED TREE ISLANDS COMPLEX PERCENTAGE OF WATERSHED BEDROCK OUTCROPS 14 HEATH COMMUNITIES 22 SEDGE COMMUNITIES 34 ACTIVE OR RECENTLY ACTIVE DEBRIS 29 ( a d a p t e d f r o m G a l l i e and Sla y m a k e r 1983) 33 Fig. 2.4 SoilHVegetation Complexes. 34 hydrophobicity (Barrett 1981). This has the effect of greatly reducing i n f i l t r a b i l i t y . Therefore, ponding may occur on sedge areas within the watershed under relatively low intensity r a i n f a l l . Surface ponding may lead to fingering, i f a sufficient pressure head can be developed. This will depend upon the capacity of surface depression storage and is therefore more likely to occur on gently inclined areas than on steeper slopes where depression storage is reduced. Conversely, steeper sedge-covered slopes will more readily produce infiltration-excess overland flow than flatter sedge areas. Water-repellency and fingering may also be present in heath communities. In addition, areas of sedge and heath are known to produce saturation overland flow where concavities force sub-surface flow lines to converge.' Saturation builds up above textural contacts, especially loess over t i l l beneath heath communities (Gallie and Slaymaker 1983). The debris soil-vegetation complex consists of very poorly vegetated, coarse textured material, overlying bedrock or t i l l . Values of i n f i l t r a b i l i t y and saturated hydraulic conductivity are likely to be high. Soil-vegetation complexes may thus form suitable surrogates for i n f i l t r a b i l i t y . Sub-surface channels have been observed in active debris sites. These consist of cores of coarser material than the surrounding matrix, occurring along the contact between t i l l and colluvium (Gallie and Slaymaker 1983). It is possible that these pipes are former surface channels that have been buried by mass movements. Similar pipes, or palaeorills, have been described by 35 Huggett (1974). Flow in these pipes may be fed by saturated zones which develop above the t i l l in response to rapid i n f i l t r a t i o n through the overlying colluvium. Macro-pores have been observed in sedge, heath, debris and tree-island communities by Gallie and Slaymaker (1983). These may be conducting disproportionate quantities of i n f i l t r a t i n g surface water, especially down large root systems under tree-islands, as was observed in forest soils by Cheng et. a l . (1975), and around larger cobbles in debris sites. Exposed bedrock wi l l produce infiltration-excess overland flow. This flow will constitute an additional input to adjacent soil-vegetation complexes. Areas downslope from large bedrock slabs may be more prone to saturation for this reason. 36 CHAPTER 3. Experimental Framework And Data Collection. 3.0 Approach To The Problem. The fieldwork was designed to determine the hydrologic response of the watershed to r a i n f a l l events, as indicated by spatial and temporal variations in s o i l moisture content. It is assumed that the response to a r a i n f a l l event of any point within the watershed will be characterized by changes in so i l moisture content over time (i.e. measurements made before, during and after a r a i n f a l l event) and wi l l reflect the influence of so i l properties and topography, as discussed in sections 1.4 and 1.5. The experimental framework was based on the soil-vegetation complexes shown in table 2.1. Tree islands, representing only 1% of the watershed area, were excluded from the data collection programme on the assumption that their contribution to total run-off is negligible. Bedrock was also excluded on the basis of its assumed zero i n f i l t r a b i l i t y and will be treated as a site of infiltration-excess overland flow production. The data collection programme was based on the remaining three s o i l -vegetation complexes; heath communities, sedge communities and debris. Measurements obtained from different locations within the watershed were intended to reflect any differences in the response characteristics of the soil-vegetation complexes, and to provide some indication of localized topographic effects. In 37 addition, the assessment of spatial variation included vertical contrasts within each soil-vegetation complex, as values of soil moisture content were obtained from two depths at each measurement point. A continuous record of r a i n f a l l was obtained from a tipping bucket rain gauge located near the south east divide of the watershed. A number of s o i l samples were obtained for determinations of organic matter contents, bulk density and porosity (sections 3.4.3 and 3.5). This information provided descriptive data on the soil-vegetation complexes and was used in the calculation of so i l moisture storage capacities and water retention curves. Soil samples were also collected for use in the electrical resistance c e l l calibration process and for gravimetric determinations of so i l moisture contents (sections 3.4.1, 3.4.3), to provide a check on the accuracy of electrical resistance c e l l s . Collecting troughs were constructed in each soil-vegetation complex to intercept surface and sub-surface flows. These measurements provided some assessment of the nature, occurrence and magnitude of run-off generation within each complex (section 3.3). 3.1 Measurement Of Soil Moisture Content. Soil moisture content was derived from electrical resistance cells (Type MC-310A, SOILTEST INC:). The resistance of a c e l l is directly related to its moisture content. A change in c e l l moisture content results as the c e l l moisture tension equilibrates to the moisture tension of the surrounding s o i l . 38 Resistance fluctuations due to temperature changes are automatically compensated for in the resistance meter (Type MC-302, SOILTEST INC:). The nature and amount ,of dissolved solids in,the s o i l moisture wi l l also affect c e l l resistance. However, correction of the resistance measurements to account for soil moisture solute variations was considered too complex a task for the present study. Consequently, some error will be present in the resistance readings. It is d i f f i c u l t to ascertain the probable magnitude of such error. There is a lack of comparable data in the literature on electrical resistance c e l l s . Gardner (1965), however, suggests that precision better than ± 2% water content should not be expected and that errors as high as 100% are easily possible. Gardner further suggests that on-the-site checking of the cells appears necessary if confidence is to be developed in the water content inferences to be made. The gravimetric determinations of moisture content (see 3.5.4) should, therefore, provide some indication of error. The results of the gravimetric determinations of soil moisture, and the implications for c e l l accuracy, will be discussed in chapter 5. Cells were positioned at 34 locations within the watershed. At each site, 2 cells were installed; an upper c e l l (U) at a depth of 5 cm, and a lower c e l l (L) at depths varying from 25 to 60 cm. The lower cells were located at the maximum depth obtainable with a soi l auger. This, in effect, coincided with the depth of the compact t i l l layer, which prevented further augering. The cells were located in contrasting topographic positions, ranging from contour convergent sites to contour 39 divergent sites and from fa i r l y level sites to relatively steeply inclined sites ( fig. 3.1). The frequency of measurement was limited due to poor access to the site. In a l l , eight sets of resistance measurements were obtained between August 29 and October 11, 1981. 3.2 Precipitation Measurement. A continuous record of precipitation was provided by a tipping bucket rain gauge, connected to an -electronic event recorder (constructed by R. Leslie, Technician, Department of Geography, U.B.C ). The number of tips in every 10 minute time period was recorded, with each tip representing 0.33 mm of rain. There is a basic uncertainty of one tip per time period; hence, cumulative r a i n f a l l over one hour is known to approximately 0.33 mm per hour and ra i n f a l l intensities derived from a single 10 minute reading are known to within plus or minus 1.98 mm per hour. The rain gauge was located on the south-eastern divide of the watershed ( f i g . 3.1). 3.3 Surface And Sub-surface Flow. A collecting trough system was constructed in each of the three soil-vegetation complexes under study. In each case, a pit was dug down to the t i l l layer and upper and lower collecting troughs were emplaced into the upslope face of the p i t . The upper trough was positioned just below the ground surface (approximately 5 cm deep) to collect surface run-off. The lower trough was placed just below the t i l l layer (the depth to which varied) to collect throughflow. The pits were covered to prevent 40 F i g . 3.1 Instrumentation. 41 direct precipitation onto the collecting apparatus. Water collected by the troughs was tunneled into a collecting bottle to provide a cumulative total. Although the exact timing and source area of the flow cannot be obtained by this method, i t does provide information on whether or not flow has occurred, the magnitude of the flow and the nature of the rai n f a l l event(s) required to generate them. 3.4 Soil Samples. 3.4.1 Cell Calibration Samples. Soil was collected from the upper and lower positions of a number of electrical resistance c e l l sites. These samples were required for the c e l l calibration process. A number of electrical resistance cells were calibrated against s o i l moisture content using porous plate extraction equipment. Each calibration c e l l was embedded in its accompanying so i l sample, within a small metal ring. The soil was packed in the ring to approximate, as closely as possible, the bulk density found in the f i e l d . These were saturated, weighed and resistance readings were taken. The rings were then placed in porous plate extractors and subjected to progressively greater tensions. At each increase in tension, the cells were left to equilibrate for '72 hours and then weighed and resistances measured. Calibration continued until further weight loss or resistance increase became negligible (usually around a tension of 50 metres of water). There was insufficient time to calibrate a l l of the cells 42 used in the study. Therefore, i t was decided to select an upper and lower c e l l from each soil-vegetation complex and use these as "representative" cells for the complex. The cells selected were from sites 25 upper and 2 lower for the debris, 6 upper and 6 lower for the sedge and 14 upper and 14 lower for the heath. In the limited time available i t was possible to calibrate an additional number of lower cells from the sedge area. The cells selected were 31, 18, 13, and 9. These ce l l s were calibrated so that an attempt could be made to determine the variance between cells from the same group. Thus, some assessment may be made of the error involved in using "representative" c e l l s . At the end of the calibration run the samples were oven-dried and weighed so that their moisture content, on a mass basis for each step of the calibration process, could be determined. Thus, the calibration procedure provided calibration curves of moisture content against resistance (figs. 3.2 - 3.3). 3.4.2 Organic Matter Content Determinations. Organic matter contents were determined for samples from the level of the upper and lower cells at sites 25(debris), 14(heath) and 6(sedge). These determinations were made to provide additional descriptive information on the nature of the soil-vegetation complexes and to fac i l i t a t e the calculation of particle density figures, to be used in the calculation of porosity. Variations of organic matter content between s o i l -vegetation complexes, and at depth within complexes, may also be a significant factor in the hydrologic behaviour of the 0 10 20 30 40 50 S o i l M o i s t u r e C o n t e n t (0m) F i g : 3.2 S o i l M o i s t u r e C a l i b r a t i o n C u r v e s F o r 25U, 14U, 6U, 14L, 6L And 2L. S o i l M o i s t u r e C o n t e n t (0m) F i g : 3.3 S o i l M o i s t u r e C a l i b r a t i o n C u r v e s F o r 31L, 18L, 13L, 9L And 6L. 45 watershed. The organic matter content determinations will provide, at least to some extent, a quantitative assessment of these variations. 3.4.3 Gravimetric Determinations Of Soil Moisture Content. A set of samples was also taken for gravimetric determinations of moisture content. This was intended as a check on the s o i l moisture content values obtained from the electrical resistance c e l l s . A random sample was obtained, comprised of soil from the vicinity of both upper and lower c e l l s . A total of seventeen samples was collected. The samples were sealed in so i l moisture cans and their moisture contents were determined gravimetrically within 72 hours. 3.5 Bulk Density And Porosity. A core of material of known volume, was taken from each of the soil-vegetation complexes. The moisture content of each core was estimated from values of moisture content obtained from nearby electrical resistance c e l l s . The weight of moisture in each core was calculated and subtracted, enabling determinations of bulk density to be made. The bulk density information was used in the estimation of porosity. 46 CHAPTER 4. Analysis Of Data. 4.0 The Spatial And Temporal Variations In Soil Moisture Content. Soil moisture content was determined using the calibration curves. Values for non-calibrated cells were derived from the "representative" cells of each soil-vegetation complex. The complete set of so i l moisture content values is shown in appendix B. 4.1 Water Retention Characteristic Curves. Water retention characteristic curves were constructed by plotting s o i l moisture content against the tensions applied during the calibration process ( f i g : 4.1 - 4.3). The curves cover the range of tension from 0 to 50 metres of water. This large range was necessary, as many curves did not begin to level off until tensions as high as 40 m of water were reached. This range of tension is also comparable to the range of tensions measured in the f i e l d . 4.2 Bulk Density, Organic Matter Content And Soil Moisture Storage Capacity. A core of known volume was extracted from the upper 14 cm of each soil-vegetation complex. The cores were weighed, enabling bulk density determinations to be made. The calculations included a correction to account for the moisture contents of the cores, which were estimated from readings Fig: 4.1 Moisture Retention Curves For 14L, 6L And 2L 70 S o i l Moisture Content (0m) 50 obtained from nearby electrical resistance c e l l s . Table 4.1 shows the resulting bulk density values. These figures were used in the calculation of porosity. Particle density figures in these calculations were based on the proportion of mineral soil to organic matter within each core. These proportions were based on the organic matter contents shown in table 4.2. An estimated particle density of 2400 kg/m? was used in the sedge and heath calculations, while a particle density of 2600 kg/m3 was used for debris. The calculated porosities are shown in table 4.1.These porosities, i f assumed uniform throughout the profiles, yield the so i l moisture storage capacities shown in table 4.1. 4.3 Precipitation. The raingauge readings were converted into mm of rain per hour. In a l l , six precipitation events were recorded between August 31 and September 20. The maximum event was 21.78 mm, which f e l l over a period of 12 hours on the 31 August. The maximum 10 minute intensity recorded was 2.75 x 10 mm/s, which occurred on August 31. Events recorded after September 18 were in the form of snow. Details of the precipitation record are shown in Appendix C. Fig. 4.4 shows precipitation totals on a daily basis and f i g . 4.5 shows the rai n f a l l of August 31 on an hourly basis. Data from the previous two years suggests that this event is fai r l y large for the snow-free season, the majority of events falling in the 1 to 5 mm range. Events in the range 20 - 25 mm occur about twice a year during the snow-free period. 51 Table 4.1 Calculated Bulk Densities, porosities And Soil  Moisture Storage Capacities. Soil-Vegetation Bulk Density Porosity Soil Moisture Storage Complex kg/m Capacity (cm per m) Sedge 827 66% 66 Heath 688 71% 71 Debris 1469 44% 44 Table 4.2 Organic Matter Contents, Soil-vegetation Complex Debr i s Debris Heath Heath Sedge Sedge Site Organic Matter Content (% by mass) 25 upper 2.4 25 lower 2.2 14 upper 24.2 14 lower 8.3 6 upper 26.1 6 lower 6.0 25 snow 29 30 31 AUGUST 4 5 6 7 SEPTEMBER 18 19 20 21 22 23 24 25 26 27 28 Fig: 4. 4 Daily Precipitation Totals. 2 0 •* I 1 2 3 4 5 6 7 8 9 10 11 12 Hours F i g . 4.5 P r e c i p i t a t i o n Of August 31 On An H o u r l y B a s i s . 54 4.4 Surface And Sub-surface Flow. No attempt was made to determine the source area and timing of surface and sub-surface flows. The run-off generated can, however, be attributed to the precipitation event preceding the time of collection, as collecting bottles were emptied just before and a few days after, each precipitation event. During the study period, only one incidence of run-off was recorded. This was generated in the surface layer of the heath complex. The upper collecting trough located in this area collected a volume of 2260 ml. The run-off was apparently generated by the ra i n f a l l event of August 31. 55 CHAPTER 5.  Analysis Of Errors.  5.0 Purpose. Errors are inherent in the method employed to assess s o i l moisture content. The use of "representative" ce l l s assumes homogeneity of so i l within a soil-vegetation complex and in the response of cells within the complex. Such homogeneity is unlikely to exist in the f i e l d . The calibration technique is also prone to certain errors. Soil structure is destroyed during the calibration process, resulting in changes in certain hydrologic characteristics of the so i l samples. Hysteresis is not accounted for in the calibration process, the calibration curves being derived from single desorption curves. However, given the coarse nature of soils in this watershed (chapter 2), i t is assumed that hysteresis effects will not be pronounced and the resulting error will be small in comparison with errors from other sources. The purpose of this chapter, then, is to identify and quantify errors inherent in the technique. Once the probable magnitudes of such errors are established, the implications for c e l l accuracy, and thus interpretation of the results, can be determined. 5.1 Method. The stati s t i c utilized was the mean deviation from the 56 r e p r e s e n t a t i v e c e l l i n each sample group d e f i n e d t h u s : d = (£ ld| )/(X-1) Where d=Mean d e v i a t i o n |d|=Deviation from r e p r e s e n t a t i v e c e l l £=Summation index x=Number of samples T h i s e x p r e s s i o n was s e l e c t e d as a s u i t a b l e index of the e r r o r r e s u l t i n g from the use of " r e p r e s e n t a t i v e " c e l l s . T h i s measurement has the advantage of b e i n g n o n - p a r a m e t r i c and s u i t a b l e f o r s m a l l sample s i z e s . The sample groups, based on the a v a i l a b l e c a l i b r a t i o n c u r v e s and the water r e t e n t i o n c u r v e s , were d e s i g n e d t o p r o v i d e q u a n t i t a t i v e assessment of t h r e e major a n t i c i p a t e d s o u r c e s of e r r o r ; 1. s o i l h e t e r o g e n e i t y w i t h i n s o i l - v e g e t a t i o n complexes. 2. c e l l h e t e r o g e n e i t y . 3. c a l i b r a t i o n e r r o r s . In some c a s e s , i t was p o s s i b l e t o e x c l u d e c e r t a i n s o u r c e s of e r r o r from the s t a t i s t i c a l a n a l y s i s ; t h e r e f o r e some r e s o l u t i o n of the r e l a t i v e impact of each so u r c e of e r r o r can be at t e m p t e d . In a d d i t i o n t o the c a l c u l a t i o n of e r r o r p e r c e n t a g e s based on mean d e v i a t i o n from a r e p r e s e n t a t i v e c e l l , a f i e l d check was a l s o p erformed, c o n s i s t i n g of g r a v i m e t r i c d e t e r m i n a t i o n s of s o i l m o i s t u r e c o n t e n t a t a number of s o i l m o i s t u r e c e l l s i t e s t h r o u g h o u t the watershed. A comparison of g r a v i m e t r i c a l l y d e r i v e d s o i l m o i s t u r e v a l u e s w i t h c o r r e s p o n d i n g v a l u e s d e r i v e d 57 from the calibration curves will provide an independent assessment of the error introduced by the use of electrical resistance c e l l s . 5.2 Analysis Of Error Based On Water Retention Curves. 5.2.1 Variability Within A Soil-Vegetation Complex. Water retention curves show the relationship between s o i l moisture content (6) and tension ( T ) . Curves based on measurements obtained from samples of a homogeneous s o i l , during a single desorption run, should be identical and therefore the use of a single "representative" sample should be applicable. Fig. 4.2 show the retention curves for samples from sites 6L, 9L, 13L, 18L and 31L, a l l of which are from the lower level of the sedge soil-vegetation complex, which for the purposes of this study is assumed to be a homogeneous s o i l unit. The mean deviation from the representative c e l l 6L was calculated for a number of tensions and converted into a percentage error. The results are shown in table 5.1. (Note: The random selection of c e l l 6L as "representative" c e l l will result in conservative derivations of error since c e l l 6L's behaviour (purely by chance) forms one extreme of the range of sample c e l l behaviour. The error values calculated are therefore the maximum values that could be derived from this data set. It should also be noted that f i g . 5.1 simply illustrates the relative magnitude of mean deviation plotted as an error bar above and below the c e l l 6L curve. The range of mean deviation shown on f i g . 5.1 does not reflect the likely absolute range of soil moisture values, most of which are likely to be "positive" deviations from c e l l 6L. 58 Table 5.1 Mean Deviation Values For The Sedge Complex, Lower Level. Cell OB D 0. 1B D 0.2B D 0.5B D 0.75B D 6L 31 0 26 0 26 0 22 0 22 0 9L 69 38 43 17 • 41 15 34 12 31 9 1 3L 31 0 24 2 23 3 22 0 22 0 18L 51 20 31 5 30 4 28 6 27 5 31L 39 8 28 2 27 1 24 2 24 2 D= 16.5 D = 6.5 D=5.75 D= 5 D=4 %= 53 % = 25 %=22 %= 22.7 %=18.2 Cont • • Cell 1B D 2B D 5B D 6L 1 9 0 1 7 0 8 0 9L 30 1 1 27 10 9 1 1 3L 21 2 20 3 7 1 1 8L 27 8 26 9 1 2 4 31L 23 4 20 3 8 0 D= 6.25 D= 6.25 D=1.5 %= 32.9 %= 36.8 %=18.8 B=Bars of tension D=Deviation From Representative Cell D=Mean Deviation %=Percentage Error 59 The mean deviation figures were used to construct error bars on the water retention curve of sample 6L (fi g : 5.1). Although the error values over the entire range of tensions used in the calibration procedure (i.e. 0 to 5 bars), vary from 18.2% to 53%, tensions typically encountered in the f i e l d were of the order 0.2 to 0.3 bars; thus, a typical error value is of the order of 22%, as indicated in table 5.1. As the electrical resistance c e l l s were not involved in the determination of the water retention curves, error due to c e l l heterogeneity can be discounted. The remaining sources of error can be identified; 1. s o i l heterogeneity. 2. solute variations among the samples. 3. calibration errors. 5.2.2 Variability Within A Sample. Each sample involved in the calibration process consisted of a fraction of a grab sample taken from a f i e l d site. Soil structure is unavoidably destroyed due to the method of collection. In an attempt to reduce the error resulting from this loss of structure bulk density is approximated as closely as possible during calibration. For the purposes of this study i t is assumed that the hydrologic behaviour of the soil samples does accurately reflect the hydrologic behaviour of the soi l in the f i e l d , to within a margin of error due to approximations made during the calibration process. To assess the effect of the calibration process on sample variability, grab samples from sites 9L and 25U were each sub-divided into five sub-samples and a separate water retention curve was determined for each of the 70 60 50 J 4 0 J 0 1 2 3 4 5 Tension (Bars) Fig: 5.1 Between Samples Mean Deviation. 61 resulting ten samples. Mean deviation was calculated with reference to samples 9Ld and 25Uf, which were selected as the representative samples for their respective sites. The results are shown in table 5.2 and 5.3. These data were used to construct error bars on the water retention curves of 9Ld and 25Uf ( f i g . 5.2). Under tensions typical of f i e l d conditions, an error of about 8% is produced for site 9L, while an error of about 6% is typical for site 25U. Once again, error due to c e l l heterogeneity can be excluded, therefore only the following sources of error need be considered; 1. s o i l heterogeneity between sub-samples. 2. solute variations between sub-samples. 3. calibration errors. In view of the fact that each sub-sample is a fraction of a single 500 gram grab sample, variability due to s o i l heterogeneity and solute variations within the grab sample is assumed minimized due to the extremely close proximity of sampling points. Therefore, i t w i l l be further assumed that variability introduced during the calibration process is the major cause of the error outlined above. Thus, the practice of approximating bulk density is the likely source of this v a r i a b i l i t y . 5.3 Analysis Of Error Based On Calibration Curves. The relationship between c e l l resistance and s o i l moisture content is prone to a l l the sources of error previously mentioned, namely; Table 2 Mean Deviation Values For Site 9L. Cell OB D 0.1B D 0.2B D 0.5B D 0.75B D 9Ld 69 0 43 0 41 0 34 0 31 0 9Le 61 8 39 4 38 3 33 1 30 1 9Lf 66 3 44 1 43 2 36 2 33 2 9Lg 60 9 39 4 37 4 29 5 27 4 9Lh 58 1 1 40 3 38 3 34 0 34 3 D= 7. 8 D = 3 D= 3 D=2 D %= 1 1 % = 6.9 %= 7.3 %=5. 9 % Cont • Cell 1b D 2'B D 4B D 5B D 9Ld 30 0 27 0 15 0 9 0 9Le 27 3 24 3 1 1 4 7 2 9Lf 32 2 31 4 15 0 10 1 9Lg 27 3 24 3 1 2 3 7 2 9Lh 28 2 26 1 12 3 7 2 D= 2. 5 D= 2.75 D< = 2. 5 D= 1 .75 % = 8. 3 %= 10.2 %• = 16 .7 % = 19.4 63 Table 5. 3 Mean Deviation Values For Site 25U. Cell OB D 0..1B D 0.2B D 0 .5B D 0.75B D 25Uf 32 0 21 0 20 0 16 0 15 0 25Uc 31 1 20 1 19 1 15 1 1 4 1 25Ub 28 4 20 1 18 2 15 1 15 0 25Ua 30 2 20 1 19 1 16 0 16 1 25Ug 33 1 22 1 20 0 18 2 17 2 D= 2 D=1 D= 1 D=1 D = %= 6 .3 % = 4 .8 % = 5 %=6. 3 Cont • • Cell 1b D 2B D 4B D 5B D 25Uf 15 0 12 0 5 0 3 0 25Uc 14 1 1 2 0 4 1 3 0 25Ub 1 4 1 12 0 4 1 3 0 25Ua 1 6 1 1 3 1 4 1 3 0 25Ug 16 1 13 1 4 1 3 0 D= :1 D=0. 5 5=1 D= 0 %= '6 .7 % = 4. 2 % = 20 % = 0 F i g : 5.2 Within Sample Mean Deviation - C e l l 9Ld And 25Uf. 65 1. s o i l heterogeneity. 2. c e l l heterogeneity. 3. solute variations. 4. c a l i b r a t i o n errors. Thus, the va r i a t i o n between the c a l i b r a t i o n curves may be due to a combination of the following factors; 1. s o i l heterogeneity: s t r u c t u r a l differences between samples would result in a var i a t i o n in the s o i l moisture content which develops in response to an applied tension. 2. c e l l heterogeneity: there may be a v a r i a t i o n in the response of the c e l l s to an applied tension, causing deviation among the response curves, which may be either l i n e a r or non-linear in nature. 3. solute variations may exist between the s o i l sample solutions, due to d i f f e r i n g amounts of soluble materials within the s o i l samples. This may be viewed as a special case of s o i l heterogeneity, the heterogeneity existing in the composition of the s o i l , rather than i t s physical structure. 4. c a l i b r a t i o n errors may include poor contact between c e l l s and the surrounding s o i l causing impaired transmission of the applied tension and re s u l t i n g in incomplete e q u i l i b r a t i o n ; i n s u f f i c i e n t time for e q u i l i b r a t i o n , such that s o i l heterogeneity becomes a factor in determining the degree of e q u i l i b r a t i o n ; errors in the measurement of the applied tension and resistance reading; equipment f a i l u r e at higher tensions at tensions above two bars i t was d i f f i c u l t to maintain a constant reading over the three day c a l i b r a t i o n period. Analysis of mean deviation from the representative c e l l , 66 6L, i n t h e sedge complex, y i e l d s t h e r e s u l t s i n t a b l e 5.4. A t t y p i c a l f i e l d r e s i s t a n c e v a l u e s of 100 - 200KJb, an e r r o r of a b o u t 14 - 28% i s p r e s e n t . 5.4 F i e l d Check On S o i l M o i s t u r e V a l u e s D e t e r m i n e d By E l e c t r i c a l R e s i s t a n c e C e l l s . A s e r i e s o f g r a v i m e t r i c a l l y measured s o i l m o i s t u r e c o n t e n t v a l u e s e n a b l e d an i n d e p e n d e n t c h e c k t o be made on t h e v a l u e s d e r i v e d f r o m t h e e l e c t r i c a l r e s i s t a n c e c e l l s . Once a g a i n , mean d e v i a t i o n was u s e d t o c h a r a c t e r i z e t h e e r r o r p r e s e n t i n t h e c e l l r e a d i n g s . O v e r a l l mean d e v i a t i o n f r o m t h e 1:1 l i n e on f i g . 5.3 was c a l c u l a t e d and c o n v e r t e d i n t o a p e r c e n t a g e e r r o r f o r t h r e e m o i s t u r e c o n t e n t s , g i v i n g an e r r o r of 17.3% a t 15% 0, 13% a t 20% 9 and 10.4% a t 25% 0. On t h e a s s u m p t i o n t h a t t h e g r a v i m e t r i c r e a d i n g s c o n t a i n o n l y i n s i g n i f i c a n t e r r o r s i n c o m p a r i s o n t o t h e c e l l r e a d i n g s , t h e e r r o r v a l u e s shown r e s u l t f r o m t h e e l e c t r i c a l r e s i s t a n c e method and a r e due t o ; 1. s o i l h e t e r o g e n e i t y . 2. c e l l h e t e r o g e n e i t y . 3. s o l u t e v a r i a t i o n s . 4. c a l i b r a t i o n e r r o r s - i n c l u d i n g v a r i a t i o n s due t o h y s t e r e s i s e f f e c t s , w h i c h may be p r e s e n t i n t h e f i e l d , b u t were not a c c o u n t e d f o r i n t h e c a l i b r a t i o n p r o c e s s . 5.5 D i s c u s s i o n . B o t h t h e f i e l d c h e c k and t h e s o i l m o i s t u r e c a l i b r a t i o n c h e c k ( s e c t i o n 5.3) g i v e an a s s e s s m e n t o f t h e e r r o r i n t h e s o i l Table 5.4 Mean Deviation: Resistance Against Soil Moisture Content. Cell 2 OK D 50K D 1 00K D 1 50K D 200K D 6L 28 0 25 0 23.7 0 22.4 0 20.5 0 9L 45 17 33 8 31 7.3 30.5 8.1 30 9.5 1 3L 32 4 26.4 1 .4 25.3 1 .6 25 2.6 24.8 4.3 18L 29.3 1.3 28 3 27.7 4 27.5 5.1 27.2 6.7 1 3L 25.5 2.5 23.5 1 .5 23.2 0.5 23 0.6 22.8 2.3 D=6. 2 D=3. 5 D=3. 4 D=4. 1 D=5. % = 22 %=14 %=1 4 %=18 .3 % = 27 Cont: Cell 300K D 500K D 1 000K D 1 500K D 6L 16.7 0 10.8 0 7.5 0 6.5 0 9L 29 12.3 26 15.2 19.5 1 2 16 9.5 1 3L 22.5 5.8 22.2 11.4 21 13.5 17.5 1 1 18L 27. 1 10.4 26.5 15.7 23.3 15.8 19 12.5 31L 24.5 7.8 24.2 13.4 22.8 15.3 20.8 14.3 5=9. 1 D=13. 9 D=14.2 D=1 % = 54 %= 1 29 %= 1 89 %=1 K=1000 ohms. Gravimetric 9m F i g : 5.3 Mean Deviation About 1 : 1 Line. 69 moisture values found via the electrical resistance c e l l method. The former, by comparing the s o i l moisture values from calibration against the assumed "actual" values; the latter by testing the validity of the representativeness assumption in terms of the error in the resulting s o i l moisture data. Overall, the results based on the calibration curves indicate an error in the range of "typical" resistance values of the order 20 - 28% due presumably to so i l and c e l l heterogeneity. Both these sources of- error appear to be of greater significance than the error introduced by structural disturbance during calibration, which, i f the assumptions in section 5.2.2 are correct, is of the order of 6 - 8%. The f i e l d check, based on samples from both levels of the debris soil-vegetation complex, seems to confirm these findings; the error of 10.4 - 17.3% being at least of a similar magnitude. The disparity between the two results may be due to a variation in the error associated with each complex-level combination. While i t is acknowledged that each complex-level combination may have its own particular error value (due, for example, to differing degrees of so i l heterogeneity within complexes), for the purposes of this study, the figure of 20 - 28% will be assumed typical for the whole watershed. This level of error is sufficiently high to prevent meaningful interpretation of much of the data. In many instances the apparent response of the cells is of a similar or smaller magnitude than the calculated error value of 20 - 28% (Appendix B). However, further examination of the behaviour of the calibrated cells suggests much of the apparent error in the data 70 may be due to a consistent bias in the s o i l moisture values derived from the calibration curves, rather than random fluctuations. The presence of bias is evident in figs. 3.3 and 4.2. The behaviour of these curves suggests that much of the deviation from the representative c e l l is due to a f a i r l y consistent offset. While such bias would affect the accuracy of s o i l moisture data derived from the representative c e l l in terms of absolute values, i t is possible to compensate for the bias in such a manner that relative changes in s o i l moisture wi l l be subject only to the error in the precision of the data, which results from random fluctuations. •Fig. 5.4 illustrates the distinction between bias (accuracy) and random (precision) fluctuations. The examples shown illustrate that i f the factor A<R>dB/&JL is known i t is possible to calculate the change in s o i l moisture to within a margin of error ±d which results from only random fluctuations in the data. While such relative changes in s o i l moisture values may be determined, as long as the bias offset "B" remains an unknown factor the absolute values of s o i l moisture can not be calculated. An examination of the deviations of calibrated cells from the representative c e l l 6L in the lower sedge does suggest that the bias in this data is acting in a f a i r l y consistent manner (fig . 5.5). Fig. 5.5 suggests that the factor d9/d<Ji> is constant and positive up to a resistance of about 550 KJband then becomes constant and negative as resistance increases. Figs. 5.6 and 5.7 show the behaviour of c e l l deviation in the 71 + CD < Unknown C e l l 6 Of R e p r e s e n t a t i v e C e l l R l R e s i s t a n c e R2 Case ( i ) C o n s t a n t O f f s e t : In t h i s case 0 o f the unknown c e l l a t R l e q u a l s 0 o f the r e p r e s e n t a t i v e c e l l a t R1+B+ d|. 0 o f unknown c e l l a t R2 e q u a l s 0 o f r e p r e s e n t a t i v e c e l l a t R2+B+ d^ T h e r e f o r e , change i n 0 o f unknown c e l l between R l and R2 e q u a l s 0 o f r e p r e s e n t a t i v e c e l l a t R l - 0 o f r e p r e s e n t a t i v e c e l l a t R2 + d . B = D e v i a t i o n due to b i a s d = D e v i a t i o n s due to random f l u c t u a t i o n s d=Mean d e v i a t i o n due to random f l u c t u a t i o n s + 0 Of R e p r e s e n t a t i v e C e l l R l R e s i s t a n c e R2 Case ( i i ) Non-cons t an t O f f s e t : In t h i s c a se 0 o f the unknown c e l l a t R l e q u a l s 0 o f the r e p r e s e n t a t i v e c e l l a t R l + B + d,. 0 o f the unknown c e l l a t R2 e q u a l s 0 o f the r e p r e s e n t a t i v e c e l l a t R2 + B + A<*d0/dJi. + d2. T h e r e f o r e , change i n 0 o f unknown c e l l between R l and R2 e q u a l s 0 o f r e p r e s e n t a t i v e c e l l a t R l - 0 o f r e p r e s e n t a t i v e c e l l a t R2 - AuidO/dJl + d . B=Constant offset between Rl and R2 due to bias AiAd0/doi=Additional offset at R2, where =R2 - Rl dQ/d =Addition to bias per unit change in F i g . 5 . 4 D i s t i n c t i o n Between B i a s And Random F l u c t u a t i o n s . 18 , 16 J _ 4 J Fig. 5.5 Deviations In 9 Of Calibrated Cells From 9 Of Representative Cell 6L In The Lower Sedge. ro Fig. 5.6 A0 In The Range 0 - 400 KJl. 75 range 0 - 400 KJb and 1000 - 1750 Kcfr respectively. The behaviour of each c e l l does appear to conform to a consistent linear trend which changes from positive to negative in the region of 550 KJL (fig. 5.5). The factor dQ/dJb (equivalent to the gradient of the trend lines) is very similar for each of the calibrated c e l l s . The equations of the deviations of cells 9L, 13L, 18L and 31L can be calculated since unlike "unknown c e l l s " (i.e. uncalibrated cells) the offset component of bias ("B" in f i g . 5.4) of these cells is known for a l l values of resistance "R". The equations have the linear form 0=a + d0/dJb Where a = Offset at R=0 Once these equations are found, bias can be removed from the soil moisture values leaving only the error due to random fluctuations. Tables 5.5 and 5.6 show the equations, the bias correction factors and the corrected s o i l moisture values along with the remaining error, for the ranges 0 - 400 KJfc and 1000 1750 KJb respectively. The results indicate that the error in the data is now considerably reduced. The typical error value is now of the order 2 - 2.5% which is an order of magnitude lower than the error of 20 - 25% present in the data before bias was removed (table 5.1, 5.4). The removal of bias from the uncalibrated cells will thus enable a much greater degree of confidence to be placed in the resulting s o i l moisture data and subsequent interpretation of the data. To remove bias from the uncalibrated cells the assumption 76 Table 5.5 Bias Correction In The Range 0 - 400 Kjl. Cell , Equation Bias Correction Factor 9L 0 = 3.8 + 2.2n -3.8-2.2n 18L 0 = 0.5 + 2.6n -0.5-2.6n 31L 0 =-1.2+ 2.6n 1.2-2.6n 13L 0 =-3.0 + 2.6n 3.0-2.6n n=Resistance/100 Corrected Readings Cell 50K D 1 00K D 1 50K D 200K D 300K D 400K D 6L 21 .3 0 19.6 0 18.5 0 17 0 13.8 0 10.3 0 9L 21.9 0.6 19.6 0 18.1 0.4 16.6 0.4 13.6 0.2 10.1 0.2 13L 21.1 0.2 19.6 0 18.1 •0.4 16.7 0.3 13.8 0 1 1 0.7 18L 21.4 0.1 19.8 0.2 18.3 0.2 16.8 0.2 14.1 0.3 10.9 0.6 31L 21.7 0.4 19.5 0.1 18.0 0.5 16.5 0.5 13.7 0.1 10.9 0.6 D=0. 33 D=0 .08 D=0.38 D=0. 35 D=0 .15 D=0 %=1 . 5 % = 0 .4 %=2. 1 %=2. 1 %=1 . 1 % = 5 Table 5.6 Bias Correction In The Range 1000 - 1 750 Kul. Cell Equation Bias Correction Factor 9L 0 = 13.9 - 0. 4n -13.9+0.4n 1 3L 0 = 15.3 - 0. 4n -15.3+0.4n 1 8L 0 = 16.6 - 0. 4n -16.6+0.4n 31L 0 = 17.3 - 0. 4n -17.3+0.4n n=Resistance/1 00 Corrected Readings Cell 1 000K D 1 250K D 1 500K D 1750K D 6L 6.2 0 5.8 0 5.4 0 5.1 0 9L 6.2 0 5.7 0.1 5.3 0.1 5.3 0.2 1 3L 6. 1 0.1 6.2 0.4 5.2 0.2 5.0 0.1 1 8L 6.3 0.1 5.8 0 5.1 0.3 4.8 0.3 31L 6.1 0.1 5.7 0.1 5.9 0.5 5.0 0.1 D=0 .08 D=0. 15 D=0. 28 D=0. 18 %=1 .2 % = 2. 6 % = 5. 1 %=3. 4 78 must be made that a l l uncalibrated c e l l s in the watershed behave in the same consistent manner as the cells in the lower sedge (fig . 5.6 and 5.7). Therefore i t is assumed that deviation from the representative c e l l increases at the rate of 2.5A0 per 100 Kji< up to a resistance of 550 KJl and thereafter decreases at the rate 0.4A8 per 100 KJl>. Where R1 and R2 f a l l on either side of 550 KSl, the d9/dJl> factor is assumed to change to negative from positive at 550 KSb and the bias is calculated accordingly, as illustrated in f i g . 5.8 (Note: error due to random fluctuations is not included in this figure.). As the offset is unknown for uncalibrated c e l l s absolute values of 9 can not be calculated, however the change in 9 in response to a r a i n f a l l event can be calculated with a high degree of confidence using the methods described in figs. 5.4 and 5.8. Thus, the changes in soil moisture were calculated for a l l the cells in the watershed for the period August 29 to September 9. As these changes are not based on absolute values of s o i l moisture content, the lowest soil moisture value for each c e l l (corresponding to the highe.st resistance reading) was arbitarily set at zero and changes in s o i l moisture were calculated as increases above this datum. The resulting s o i l moisture departures from datum are shown in Appendix D. 5.6 Summary. The results of the error analysis illustrate that the technique employed to determine s o i l moisture content is prone to considerable error. The electrical resistance method and the use of representative c e l l s introduced a level of error into the so i l moisture data which was of a similar or larger magnitude 79 a> < + Offset 0 Of Representative C e l l Rl 550 R2 0 of unknown c e l l at Rl equals 0 of representative c e l l at Rl + o f f s e t . 0 of unknown c e l l at R2 equals 0 of representative c e l l at R2 + o f f s e t + (a - b). Therefore, change i n s o i l moisture between Rl and R2 equals 0 of representative c e l l at Rl - 0 of representative c e l l at R2 - (a - b). Fi g . 5.8 Correction Of Bias For Rl<550M<And R2> 550 K^ . 80 than the apparent magnitude of many soi l moisture changes. Consequently the level of error was sufficient to prevent meaningful interpretation of much of the data. On the basis of data from calibrated cells i t appeared that much of the error was due to a c e l l to c e l l bias rather than random fluctuations in the behaviour of c e l l s . On the assumption that bias operated in the same consistent manner for a l l ce l l s , i t was possible to remove the bias and reduce the error an order of magnitude, although the resulting s o i l moisture data was in the form of relative changes since absolute values cannot be determined by this method. These findings suggest that the electrical resistance cells are not very well suited to the determination of absolute values of s o i l moisture at a large number of sampling points. To calibrate every c e l l would be very time consuming and errors due to hysteresis and the calibration process would s t i l l remain. The use of representative cells and a number of assumptions including s o i l homogeneity within broadly defined s o i l -vegetation complexes and uniformity of c e l l response, clearly introduces a large degree of error. A viable alternative in such situations may be the use of a more intensive network of gravimetric sampling, to provide at-a-point measurements without the necessity for assumptions of soil homogeneity introduced by the use of representative samples. The drawback of gravimetric sampling is the problem of site destruction, but as long as the number of samples required is relatively low (i.e. one sample collected before and after several r a i n f a l l events) disturbance of the site should not 81 constitute a major problem. The advantage of gravimetric sampling is that i t provides very accurate so i l moisture contents, a large number of samples can be processed in a short time and very few assumptions need to be made. 82 Chapter 6. Analysis Of Results And Discussion. 6.0 The Influence Of I n f i l t r a b i l i t y And Topography On Soil Moisture. It has been proposed (chapter 2) that soil-vegetation complexes form suitable surrogates for in f i l t r a b i 1 i t y and may therefore constitute a major influence on s o i l moisture behaviour. Topographic setting may also represent a major control on soi l moisture conditions by influencing s o i l moisture flows in the manner described in section 1.5. To investigate the relationship between soil-vegetation complexes, topography and soi l moisture behaviour, the cells were sub-divided into twelve categories consisting of an upper and lower level of each of the three soil-vegetation complexes (based on f i g . 2.4) and an upper and lower level of each of three topographic settings; hollows, straight-contoured areas and spurs (based on f i g . 2.2). Figs. 6.1 - 6.6 show the soil moisture changes that occured within each of the soil-vegetation level and topographic setting level combinations, during the period August 29 to September 9 which includes the largest r a i n f a l l event, recorded on August 31. An anticipated source of error in these responses is the possibility that the soi l in some locations may have been at or near saturation at the beginning of a rai n f a l l event and may therefore show only a limited response. The occurence of surface run-off in the vicinity of c e l l 40 n 29 30 31 1 2 3 4 5 6 7 8 9 AUGUST SEPTEMBER Fig. 6.1 Upper And Lower Soil Moisture Behaviour - Sedge. AUGUST SEPTEMBER Fig. 6.2 Upper And Lower Soil Moisture Behaviour - Heath. 85 + 32U 27U 26U 25U, 11U, 16U 24U, 34U, IU, 30U, 15U 4: 29U, 12U, 17U, 21U, 22U, 20, 2U, 8U, 10U o 17L, 22L 23L, 29L 12L, 25L, 27L, 20L, I L , 2L 10L, 11L, 16L, 24L, 32L 29 30 31 1 2 3 4 5 6 7 8 9 AUGUST SEPTEMBER F i g . 6.3 Upper And Lower S o i l M o i s t u r e B e h a v i o u r - D e b r i s . 86 CD S-i PM 29 30 31 1 2 3 4 5 6 7 8 9 AUGUST SEPTEMBER Fig. 6.4 Upper And Lower Soil Moisture Behaviour - Hollows. 87 30 -, 88 u cx & CD < O 30 i 20 J 10 30 20 10 20 J .2 15 id Z 10J & •r-i U 0) U PH + 31U 29 30 31 1 2 AUGUST 4 5 6 7 SEPTEMBER 24U t 34U + 15U | 7U, 14U, 17U 22U, 20U, 29U 17L, 22L 29L, 7L 14L 0 20L, 24L F i g . 6.6 Upper And Lower S o i l M o i s t u r e B e h a v i o u r - S p u r s . 89 14U in the heath with apparently l i t t l e response in the upper heath (fig. 6.2) suggests that the surface soil at this location was saturated and that further interpretation based on the response of the upper heath may be invalid. For this reason and the fact that only two c e l l sites were located on the heath, subsequent analysis will be concerned only with a comparison of the sedge and debris. For the purposes of this study i t will be assumed that the heath is likely to produce s o i l moisture behaviour which is at least similar to that of the sedge, this being based on their similar physical characteristics (i.e. both are well vegetated, stable and contain distinct horizons with similar characteristics). To provide a preliminary indication of the relative influences of i n f i l t r a b i l i t y and topography, a simple two-way analysis of variance technique, median polish, was applied to the data. For each soil-vegetation type, topographic setting and level combination (i.e. upper sedge spur), the response to the rai n f a l l of August 31 was calculated. It is assumed that this response is indicated by the change in soi l moisture content between August 30 and September 1. The median polish method, described by Tukey 1977, simply breaks down the median responses of each soil-vegetation / topographic setting / level combination into its component parts based on the following model: Median response = typical value + soil-vegetation effect + topographic effect + residual value. The results for the upper and lower so i l responses are shown in table 6.1 and 6.2 respectively. The results for the 90 Table 6.1 Median Polish - Upper Soil. Response = Typical value + soil-vegetation effect +topography effect + residual Spur Straight Hollow Sedge 5.45 3.5 7.45 Debris 1 .2 1 .3 1 . 1 Median responses Spur Straight Hollow Sedge 2.125 1 . 1 3.175 Debris -2.125 -1.1 -3. 175 3.325 2.4 4.275 - Row medians taken out Sedge Debris Topo: effect Spur 0 0 0 Straight -1.025 1.025 -0.925 - Column medians Hollow 1.05 -1.05 0.95 taken out Soil-veg: 2.125 -2.125 3.325 - Typical value effect 91 Table 6.2 Median Polish - Lower So i l . Response = Typical value + soil-vegetation effect + topography effect + residual Spur Straight Hollow Sedge 1 .6 1 .5 1 .25 Debris 3.1 3.9 1 .5 - Median responses Sedge Debris Spur -0.75 0.75 2.35 - Row medians Straight -1.2 1.2 2.7 taken out Hollow -0.125 0.125 1.375 Spur Soi1-veg: ef fect Sedge 0 Straight -0.45 Hollow 0.625 0.75 Debris Topog: effect 0 6 0.45 0.35 - Column medians -0.625 -0.975 taken out 0.75 2.35 - Typical value 92 upper so i l suggests that soil-vegetation type does have a greater effect than topographic setting. The response of a c e l l tends to be increased by a sedge location and decreased i f located in debris. The results also suggest that hollows tend to increase response, the response of.cells in straight-contoured areas tends to be decreased and spurs have a zero effect. These results seem to provide some support for the relationships discussed in section 1.5, in that hollows tend to increase so i l moisture response. The results for the lower so i l (table 6.2) are less clear. Location in hollows has the largest effect, tending to reduce so i l moisture response; the soil-vegetation effects are the next largest, sedge causes a decreased response, while debris tends to increase the response; the effect of straight-contoured areas is a small increase in response, while spurs again have a zero effect. The soil-vegetation effects are consistent with an explanation of soi l moisture response based on a contrast in i n f i l t r a b i l i t y between the sedge and debris. On the assumption that the sedge has low i n f i l t r a b i l i t y in comparison to the debris, moisture retention may be greater in the upper sedge relative to the upper debris, thus s o i l moisture response would tend to be higher in the upper sedge. In the lower soil this pattern is reversed due to greater transmission of moisture down through the profile in debris areas. The reversal of the effect of hollows between the upper and lower soil is not consistent with the model of soil moisture behaviour based on topography, since no distinction is made 93 between the upper and lower soil in the relationships proposed in section 1.5. Inconsistencies in the results may be due at least in part to imperfections in the soil-vegetation classification system. The sub-division of the watershed into (in effect) three major soil-vegetation complexes is probably inadequate to cover the f u l l range of soil-vegetation variability within the watershed. The "debris" complex, for example, includes both stable, poorly vegetated, recently active debris and unstable, unvegetated, active debris lobes. As such there may be some significant contrasts in behaviour within each of the broadly defined complexes. The median polish analysis indicates that some differentiation does exist between soil moisture responses divided on the basis of soil-vegetation type and topography. This method does not, however, assign any level of significance to the resulting values. To test for significance in the differences between responses associated with the various combinations, Mann-Whitney U tests were applied to the response data. Based on a rejection level of 95%, the results (table 6.3) indicate that the upper sedge has a significantly greater response than the lower sedge and upper debris, but is not st a t i s t i c a l l y distinct from the lower debris. No s t a t i s t i c a l l y significant differences could be found between the responses of the lower debris, upper debris and lower sedge. A comparison of responses based on topographic setting produced no significant differences. Possible explanations for the distinction between the 94 T a b l e 6.3 Summary Of Mann-Whitney U T e s t s . Upper Sedge : : Lower Sedge S i g n i f i c a n t a t 97.5% Upper Sedge : : Upper D e b r i s S i g n i f i c a n t a t 95% Upper Sedge : : Lower D e b r i s Not s i g n i f i c a n t Upper D e b r i s : : Lower D e b r i s Not s i g n i f i c a n t Upper D e b r i s : : Lower Sedge Not s i g n i f i c a n t Lower D e b r i s : : Lower Sedge Not s i g n i f i c a n t H o l l o w : Spur Not s i g n i f i c a n t H o l l o w : S t r a i g h t Not s i g n i f i c a n t S t r a i g h t : Spur Not s i g n i f i c a n t R e j e c t i o n l e v e l = B e l o w 95% \ 95 response of the upper and lower sedge include the following: 1. Such a pattern could be produced by a time lag in transmission of moisture from the upper to lower so i l (i.e. the lower s o i l may show a more pronounced response, similar to that of the upper s o i l , but at a later time); however, in the debris areas lower cells are installed at similar depths as those of the sedge and yet there is no distinction between upper and lower c e l l response. This suggests that the rate of i n f i l t r a t i o n is lower in the sedge, resulting in a greater proportion of the rainfa l l being retained for a longer period in the upper sedge than in the upper debris. 2. Reduced porosity at depth in the sedge, due to increased overburden pressure and less biotic activity, may limit the potential for response of the lower so i l (i.e. the saturation moisture content of the lower so i l may be considerably less than at the surface). However, this possibility is not supported by the study findings. Data from moisture retention curves for calibrated cells in the lower sedge (fig. 4.2) produce lower soi l porosities ranging from 26 - 57%, while typical lower so i l moisture contents gained from calibrated cells are of the order of 20 - 25%. The calibrated c e l l 9L, for example, indicates a maximum so i l moisture content of 26%, whereas calculated porosity in the vicinity of c e l l 9L is 57%. On the basis of these observations i t appears that the distinction in soil moisture behaviour between the sedge and debris is due to a contrast in i n f i l t r a b i l i t y between the two soil-vegetation complexes. The occurence of overland flow under relatively low r a i n f a l l intensities on the heath (which is 96 assumed to have similar hydrologic characteristics to the sedge) is consistent with this possibility. These fingings suggest that i n f i l t r a b i l i t y may be significantly lower in vegetated sections of the watershed than in unvegetated sections. This contrast may be due to: 1. Overall textural and compositional differences between the soils associated with vegetated and unvegetated areas. The presence of vegetation is i t s e l f a major factor in that vegetation stabilizes the site, traps and retains fine windblown material resulting in the development of loess horizons, and contributes significant amounts of organic matter to the s o i l . 2. The presence of distinct textural changes in well-vegetated, geomorphically stable sites. Sequences of textural changes may act to impede i n f i l t r a t i o n in the manner described in section 1 .4.3. 3. The possibility that a hydrophobic layer exists at a shallow depth within vegetated soils of this watershed. Hydrophobic layers have been found within vegetated soils in areas adjacent to the study watershed (Barrett 1981). The lack of support in the study findings for the model of soi l moisture behaviour based on watershed topography suggests that the proposed relationships between topography and s o i l moisture may be subject to a number of limitations and conditions: 1. The occurence of lateral flows is a necessary condition of these relationships. Yet such flows may only occur under certain r a i n f a l l intensities and/or amounts. During the study period for example only one incidence of lateral flow was 97 recorded (section 4.4). This was surface run-off generated on a spur in the heath area(pit 1, f i g . 3.1), following the rainfa l l of August 31. If the proposed model is to- be applied to s o i l moisture variations at depth within the soil profile, then the occurrence of interflow, in the form of a diffuse matrix flux, is implied, and yet no form of interflow was recorded during the entire study period. Rainfalls insufficient to cause significant lateral seepage may simply be diverted into the available s o i l storage capacity or undergo vertical i n f i l t r a t i o n . Under these conditions, variations in s o i l moisture behaviour within a watershed wi l l be largely a function of soi l heterogeneity, rather than topography. 2. The relationships are based on the assumption that s o i l homogeneity exists in the watershed. Often, as in this study, this may not be the case. The occurrence of soil heterogeneity within a watershed is likely to result in spatial variations in i n f i l t r a b i l i t y . Such variations may constitute a major influence on so i l moisture behaviour, perhaps even to a greater extent than topography. This has implications not only for contrasts between different locations within a watershed, but also for soil moisture variations with depth at a given location. 3. The role of preferred pathways is not considered. However, such pathways may play an important role in many watersheds; f i r s t l y , by providing zones of moisture concentration, which may in some cases be independent of topographic influences, and secondly, by providing a means of by-passing much of the so i l matrix. Fingering, for example, 98 r e p r e s e n t s a l o c a l i z e d t r a n s f e r of moisture from a s a t u r a t e d s u r f a c e l a y e r to a lower s o i l l a y e r , whereby much of the i n t e r v e n i n g s o i l matrix i s by-passed. On the b a s i s of the study f i n d i n g s i t appears that topography had no s t a t i s t i c a l l y s i g n i f i c a n t e f f e c t on s o i l moisture c o n d i t i o n s encountered d u r i n g the study p e r i o d . The evidence i n d i c a t e s that s o i l - v e g e t a t i o n type, which i s assumed to be a sur r o g a t e f o r i n f i l t r a b i l i t y , d i d s i g n i f i c a n t l y i n f l u e n c e the behaviour of s o i l moisture w i t h i n the watershed. I t i s assumed that the c o n t r a s t i n s o i l moisture behaviour a s s o c i a t e d with the v a r i a t i o n i n i n f i l t r a b i l i t y between the sedge and d e b r i s , can be viewed more simply as a c o n t r a s t between w e l l - v e g e t a t e d and bare or p o o r l y - v e g e t a t e d s e t i o n s of the watershed. The former i n c l u d e s both the sedge and the heath areas, the l a t t e r has been d e f i n e d as the " d e b r i s " s o i l -v e g e t a t i o n complex. 6 .1 Run-Off Generation. The study f i n d i n g s i n d i c a t e that i n f i l t r a b i 1 i t y was more i n f l u e n t i a l than topography i n determining s o i l moisture behaviour d u r i n g the study p e r i o d . I t should be noted that under c o n d i t i o n s of higher r a i n f a l l i n t e n s i t y and/or amount, topography may exert a g r e a t e r i n f l u e n c e on s o i l moisture c o n d i t i o n s due to an a n t i c i p a t e d i n c r e a s e i n the occurrence and magnitude of l a t e r a l moisture flows, both w i t h i n and over the s o i l s of the watershed. However, f o r the purposes of t h i s study i t w i l l be assumed that c o n t r a s t s in i n f i l t r a b i l i t y between vegetated and unvegetated p a r t s of the watershed w i l l continue to be the 99 dominant influence on so i l moisture behaviour during r a i n f a l l events which generate significant run-off. On the basis of the findings of this and other hydrologic studies in this area, the probable consequences for the assessment of appropriate mechanisms of run-off generation will now be examined. The findings of Barrett (1981) suggest that the vegetated areas of the watershed, consisting primarily of the sedge and heath complexes, may contain a hydrophobic layer of low hydraulic conductivity at a shallow depth below the surface. The findings of this study, which indicate relatively low i n f i l t r a b i l i t y in the sedge and the occurrence of surface run-off on the heath, are consistent with this possibility. Whatever the cause of the apparently low i n f i l t r a b i l i t y on vegetated sections of the watershed, the consequence is that i n f i l t r a t i o n -excess overland flow may occur on relatively large areas under relatively low r a i n f a l l intensities. The debris areas, which cover most of the remainder of the watershed, are apparently capable of more rapid moisture transmission down through the profile, resulting in an increase of s o i l moisture content at the colluvium-ti11 interface. Such behaviour is consistent with the possibility that pipeflow is generated in debris areas. The development of a saturated zone above the t i l l during larger r a i n f a l l events, could provide the necessary source of "feed water" for pipeflow generation along pipes which have- been observed beneath active debris lobes. The possibility that saturated zones develop above the t i l l -colluvium interface would presumably be enhanced in locations downslope from large bedrock slabs. It is assumed that these 100 slabs would contribute surface run-off to adjacent downslope areas during rainf a l l events. Both these proposed run-off mechanisms contribute rapid storm run-off via preferred pathways, rather than as a form of diffuse matrix flux. If the preceeding assumptions are correct and such mechanisms are indeed the dominant form of run-off generation in the watershed, this suggests a need for a shift in emphasis away from conceptual analyses based on soi l homogeneity and topographic controls on s o i l moisture conditions, towards a conceptual framework based on soi l heterogeneity and the influence of spatial variations in i n f i l t r a b i l i t y , on s o i l moisture behaviour and patterns of run-off generation. This would presumably be accompanied by a shift in the basis of theoretical models of watershed behaviour, whereby moisture flow in the form of a diffuse, uniform matrix flux is rejected, in favour of flow via preferred pathways, whereby much of the soi l matrix may be by-passed. A similar shift in emphasis would be anticipated in the analysis of other high altitude watersheds. 101 Bibliography Barrett, G.E., 1981, Streamflow Generation In A Sub-Alpine Basin  In The Coast Mountains Of British Columbia, Unpublished M.Sc Thesis, University Of British Columbia, p. 89. Baver, L.D., Gardner, W.H., and Gardner, W.R., 1972, Soil Physics, Wiley, New York, p. 498. Betson, R.P., 1964, What Is Watershed Run-Off? Journal Of  Geophysical Research, V.69, pp. 1541 - 1551. Calver, A., Kirkby, M.J., and Weyman, D.R., 1972, Modelling Hillslope And Channel Flow, in Chorley, R.J., (ed.) Spatial  Analysis In Geomorphology, Methuen, London, pp. 197 - 218. Carson, M.A., and Kirkby, M.J., 1972, Hillslope Form  And Process, Cambridge University Press, London, p.475 Cheng, J.D., Black, T.A., and Willington, R.P., 1975, The Generation Of Stormflow From Small Forested Watersheds In Coast Mountains Of Southwestern British Columbia. Proceedings Of The Canadian Hydrology Symposium, Winnipeg, August 11 - 14, 1975, pp. 542 - 551. 1 02 Dunne, T., and Black, R.D., 1970b, Partial Area Contributions To Storm Run-Off In A Small New England Watershed. Water  Resources Research, V.6, pp. 1296 - 1311. Gallie, T.M., and Slaymaker, H.O., 1983, Variable Solute Sources And Hydrologic Pathways In A Coastal Subalpine Environment, in Walling, D.E., and Burt, T.P., (eds) 1983, Watershed  Experiments, Published By Geo Books, (in press). Gardner, W.H., 1965, Soil Moisture Content, in Black, C.A., 1965, Methods Of Soil Analysis, American Society Of Agronomy, New York, p. 1572. Hewlett, J.D., and Hibbert, A.R., 1967, Factors Affecting The Response Of Small Watersheds To Precipitation In Humid Areas, Proceedings Of The International Symposium On Forest  Hydrology, Pennsylvania State University, pp. 275 - 290. H i l l , D.E., and Parlange, J.Y., 1972, Wetting Front Instability In Layered Soils, Soil Science Society Of America Proceedings, V.36, pp. 697 - 702. H i l l e l , D., 1971, Soil And Water: Physical Principles And  Processes, Academic Press, New York, p. 288. Horton, R.E., 1-933, The Role Of Infiltration In The Hydrological Cycle, Transactions Of American Geophysical Union, V.14 pp. 446 - 460. ' 103 Horton, R.E., 1935, Surface Run-Off Phenomena - Part 1. Analysis Of The Hydrograph, Horton Hydrological  Laboratory Publication, No. 101, New York, p. 73. Horton, R.E., 1945, Erosional Development Of Streams And Their Drainage Basins: Hydrophysical Approach To Quantitative Morphology, Bulletin Of The Geological Society Of America, V.56, pp. 275 - 370. Huggett, R.A., 1974, Percolines Or Palaeorills? Area, V.6(3) , pp. 238 - 239. Jones, J.A.A., 1971, Soil Piping And Stream Channel Initiation, Water Resources Research, V.7(3), pp. 602 - 610. Jones, J.A.A., 1975, Soil Piping And The Subsurface Initiation  Of Stream Channels, Unpublished ph.D Thesis, University Of Cambridge, p.476. Jones, J.A.A, 1978a, Soil Pipe Networks: Distribution And Discharge, Cambria, V.5, pp. 1-21. Kirkby, M.J., 1969, Erosion By Water On Hillslopes, in Chorley, R.J., 1969, Water, Earth And Man, Methuen, London, p.588 Kirkby, M.J., and Chorley, R.J., 1967, Throughflow, Overland Flow And Erosion, Bulletin Of International Association  For Scientific Hydrology, V.12, pp. 5 - 21. 1 04 Liakopoulos, A.C., 1965b, Retention And Distribution Of Moisture In Soils After Infiltration Has Ceased, Bulletin Of The  International Association Of Scientific Hydrology, V.10, pp. 58 - 69. McKee, B, 1972, Cascadia, McGraw-Hill, New York, p.394. Monteith, J.L., 1973, Principles Of Environmental Physics, Edward Arnold, London, p.241. Newson, M.D., 1976, Soil Piping In upland Wales: A Call For More Information, Cambria, V.3(1), pp. 33 - 39. Parker, G.G., 1963, Piping, A Geomorphic Agent In Landform Development In The Drylands, International Association  Of Scientific Hydrology, Publication No. 65, pp.103 - 113. Philip, J.R., 1957, The Theory Of Infiltration, Soil Science, V.83, pp. 345 - 357 and 435 - 448, V.84, pp. 163 - 177, 257 - 264 and 329 - 339, V.85, pp. 278 - 286 and 333 - 337. Philip, J.R., 1975, Stability Analysis Of Infiltration, Proceedings Of The Soil Science Society Of America, V. 39, pp. 1042 - 1049. 105 Raats, P.A.C., 1973, Unstable Wetting Fronts In Uniform And Non-Uniform Soils, Proceedings Of The Soil Science Society  Of America, V. 37, pp. 681 - 685. Tukey, J.W., 1977, Exploratory Data Analysis, Addison-Wesley, Reading, Mass. p.506. Veihmeyer, F.J., and Hendrickson, A.H., 1949, Methoods Of Measuring Field Capacity And Permanent Wilting Percentage Of Soils, Soil Science, V. 68(2), pp. 75 - 94. Ward, R.C, 1967, Principles Of Hydrology, McGraw-Hill, London, p. 403. Weyman, D.R., 1975, Run-Off Processes And Streamflow Modelling, Oxford University Press, London, p.54. Whipkey, R.Z., 1965, Subsurface Stormflow From Forested Slopes, Bulletin Of The International Association Of Scientific  Hydrology, V.10, pp. 74 - 85. Whipkey, R.Z., 1969, Storm Run-Off From Forested Catchments By Subsurface Routes, in Floods And Their Computation,  Volume 2, I.A.S.H., Publication 85, pp. 773 - 779. Whipkey, R.Z., and Kirkby, M.J., 1978, Flow Within The Soil, in Kirkby, M.J., (ed) 1978, Hillslope Hydrology, Wiley, London, pp. 121 - 144. 106 Appendix A: Major V e g e t a t i o n A s s o c i a t i o n s . Heather W i t h Dwarf T r e e s C a s s i o p e m e r t e n s i a n a P h y l l o d o c e e m p e t r i f o r m i s , i n t e r m e d i a , g l a n u l i f o r a A b i e s l a s i o c a r p a L u e t k e a p e c t i n a t a Lycopodium s i t c h e n s e G a u l t h e r i a humifusa C l a o n i a sp. Heather, Sedge, Forb P h y l l o d o c e e m p e t r i f o r m i s  C a s s i o p e m e r t e n s i a n a  L u e t k e a p e c t i n a t a  Juncus p a r r y i  P i n u s a l b i c a u l i s  A b i e s l a s i o c a r p a  L u p i n u s l a t i f o l i u s  Anemone o c c i d e n t a l i s  Deschampsia a t r o p u r p u r e a  Carex s p e c t a b i l i s Sedge Carex n i g r i c a n s  Juncus drummondi i  D e s c h a m p s i a a t r o p u r p u r e a  Carex s p e c t a b i l i s  P o l y t r ichum s e x a n g u l a i e Sedge, F o r b , Moss Carex n i g r i c a n s , s p e c t a b i l i s  Juncus drummondii, m e r t e n s i a n u s  E p i l o b i u m a l p i n u m , l a t i f o l i u m  P h i l o n t i s f o n t a n a  Deschampsia a t r o p u r p u r e a Sedge, Sphagnum Carex n i g r i c a n s , s p e c t a b i l i s Sphagnum o t h e r mosses 107 Appendix B: Soil Moisture Content Data. 29 / 8 30 / 8 1 / 9 2 / 9 Cell L U L U L U L U 1 5.5 18.2 6.5 18.2 5.5 22 5.5 24 2 19 10 19 10.8 19.5 10.8 19.5 1 2 3 21 90 23.5 90 26 90 26 90 4 12.5 55.5 14.8 60 23.7 73 26.5 73 5 25 43.5 25 46.5 27 80 28 85 6 11.3 45 1 2 48.5 17.5 60 23 60 7 27 78 27 78 29.4 75 29.5 75 8 25 11.5 24 12.7 32 12.5 32 12.5 9 30.5 53 31 60 31.3 70 31.3 64 10 18 14.5 18.2 14.5 18.8 1 5 18.8 1 5 1 1 19.3 25 19.3 24 19.6 20.5 19.6 20.5 1 2 19.5 15.2 19.5 1 6 40 17 20 16.3 13 21 90 21 .6 90 23 90 23.2 90 14 24.2 66.5 24.2 67 26.5 69 26.5 67 1 5 6.6 18 8.3 18.5 11.5 19 9.5 19.5 1 6 16.5 18.2 16.7 18.5 18.4 24 18.7 26 17 3 15.1 4 15.5 18 17.7 17.7 17.7 18 1 4 60 19.3 60 27.5 57.5 1 4 54.5 19 29 75 26 75 30 82 30 84 20 1 6 14.5 16.8 14.7 19.4 15.5 19.4 15.5 21 4 13.5 4 1 4 16.7 14.7 9.5 14.7 22 3 1 5 4 15 19.3 18 18.8 17.5 23 26 34 26.5 34 29 84 30 84 24 18.5 17.5 19.4 18.2 20 26 19.6 24 25 1 5 6.5 15 6.5 16 8.5 18.2 7.6 26 8.4 3.5 8.4 3.5 17 3.5 19 3.5 27 18.6 3.8 19 4 19 6.5 19.2 8.5 28 4 15.5 4 15.5 4 19.5 6.5 19.5 29 18.8 18.2 19.3 18.2 21 19 23 19 30 8.5 19 9.5 19.5 19 26 18.8 26 31 1 2 46 22.8 49 26.8 64 26.3 69 32 1 3 16.5 13.8 18 16.5 31.5 1 5 28 33 8 64 8.5 60 8.5 69 5 69 34 3 24 3 28 3 31 .5 4 28 108 Cont inued: 9 / 9 24 / 9 25 / 9 11 / 10 Cell L U L U L U L U 1 6.5 22 5.5 18.2 3 18 5.5 24 2 19.5 1 4 19.2 10.8 19.4 11.5 19.2 10.8 3 25.5 90 24.6 45.5 24.6 90 24.5 80 4 28.3 83 17.5 80 17.5 80 28.3 90 5 28 75 27.5 67 27.5 65.5 28 75 6 23.7 60 22.5 50 22.3 50 22.3 49 7 29.4 77 28.5 69 28.5 68 28.2 67 8 32 12. 5 23 7.5 23 9.5 23 1 1 9 31 51 30.7 48.5 30.7 51 31 .3 58 10 18.6 15. 2 18.2 15.2 18.2 1 5 18 15 1 1 19.7 26 19.4 22 19.4 22 19.6 22 1 2 19.9 15. 6 19.5 15.3 19.7 15.2 19.8 15.2 1 3 23.2 90 23. 1 78 23. 1 78 23.3 75 14 26.8 69 26 57 25.8 55.5 30.5 52 15 9.9 20. 5 9.5 19 9.2 19 1 4 22 1 6 18.6 22 18.5 18.2 18.5 18.2 19 17.2 1 7 16.3 16 16.7 15.5 16 15.2 16.7 15.2 18 19.3 60 17 55 19.3 55 28 58 19 29 80 28.5 75 28.5 75 29 73 20 19.4 15. 5 19 15.2 19 15.2 18.8 15.4 21 11.5 14. 7 9.5 13 9.5 13 1 3 13.2 22 17.8 16. 7 18.3 16.4 17.8 16 17.6 16.2 23 29 75 28 69 28 66 27 65 24 19.5 22 19.5 20.5 19.3 18.2 19.4 22 25 19.3 1 1 19 9 19.2 9.5 20 16.2 26 19.3 6. 5 19 3.5 19 4.7 19.3 16 27 19.2 14. 8 18.8 15.5 19.1 9.4 -- — 28 9.2 17. 8 8.5 19.5 8.5 19 8.5 17.5 29 21 .5 19. 5 21 .5 18 21 .5 18 20 19.5 30 17.8 22 18 26 17.8 19.5 18 19.5 31 25.6 90 25.4 75 25.8 72 25.4 76 32 16.5 25 16.5 20.5 17 20.5 17.5 19.5 33 20 66 24 60 24 60 26 59 34 19 28 19.2 24 19 24 -- — A l l values are soi l moisture content % (on a mass basis) L=Lower c e l l U=Upper c e l l 109 Appendix C: Precipitation Record. Event #1 31 /8 10.49 a.m. Hours mm 1 1 .65 2 0.66 3 2.31 4 3.3 5 3.63 6 3.63 7 2.97 8 0.33 9 1.65 10 0.99 11 0.33 12 0.33 Total=21.78 1=2.75 Event #2 3 /9 5.05 p.m. mm 0.99 0.55 0.44 0 0.66 1 .54 0.44 0 0.33 Total=4.95 1 = 1 .65 Event #3 9 /9 2 a.m. mm 1 .65 0.33 Total=1.98 1=1.65 Event #4 18 /9 4.55 a.m. mm 1 • V 0.99 1 .43 2.31 Total=5.83 1=1.65 1 10 Cont: Event #5 Event #6 20 /9 20 /9 3.49 a.m. 1.28 p.m. Hours mm mm 1 0.99 4.95 2 0.22 1.65 3 0.11 0.33 Total=6.93 1=2.75 A l l values are mm precipitation I=Maximum 10 minute intensity in mm/s x 10 Total=1.32 1 = 1 .1 111 Append ix D: Soil Moisture Response 29/8 • - 9/9. 29/8 30/8 1/9 2/9 9/9 e l l U L U L U L U L U L 1 0 0 0 3.4 5.5 0 8.5 0 5.5 3.4 2 0 0 0.7 0 0.7 1 0 1 5.5 1 3 0 0 0 0.3 0 1 0 1 0 0.8 4 0 0 6.4 0.6 13.1 3.7 14 4.8 22. 1 6 5 0 0 1.6 0 27.7 0.7 31.7 1 .5 23.5 1.5 6 0 0 4.3 1 .4 12.4 6.8 12.4 8.8 12.4 9.4 7 4. 1 0 2.4 0 0 1 . 1 0 1 . 1 2.4 1 . 1 8 0 1 .4 0.2 0 0.4 11.6 0.4 11.6 0.4 11.6 9 1 .4 0 6.8 0.4 13.6 0.4 10 0.4 0 0.4 10 0 0 0 0 0 1 . 1 0.2 1 . 1 0 1 .4 1 1 3.8 0 3.1 0 0 2 0 1 .5 6 1 .8 1 2 0 0 0.7 0 2 29.8 1 0.6 0.2 0.4 1 3 0 0 0 0.5 0 1 .6 0 1 .6 0 1 .8 1 4 0 0 0 0 1 .7 1 .6 0 1 .6 1.7 1 .8 1 5 0 0 0.6 4.5 1 .2 9.8 2.2 7.3 3.7 8 1 6 0 0 0 0 8 0.9 10.8 1 5 0.9 17 0 0 0.3 5.5 1 .5 21.1 2.5 21.2 0.3 21.4 18 4.2 0 4.2 6.4 2.9 9.4 0 0 4.2 6.4 19 0 1 .9 0 0 5.7 2.6 7.3 2.6 4.1 1 .9 20 0 0 0 0 0 0 0 0 0 0 21 0 0 0.3 0 0.5 15.7 0.5 13.1 0.5 15.6 22 0 0 0 5.5 2.7 21.1 2.1 21 .3 1 21 .3 23 0 0 0 0.6 35.6 2 35.6 2.8 28.2 1 24 0 0 0.9 0 12.3 0.2 9.4 0 6.4 0 25 0 0 0 0 2 1 .9 1 .7 1 .2 6.7 1 26 0 0 0 0 0 3.9 0 3.6 9 3.6 27 0 0 1 .5 0.2 7.7 0.2 6.9 0.1 9.9 0.1 28 0 0 0 0 23. 1 0 23. 1 2.9 19.1 7.1 29 0 0 0 0.4 0.3 8.5 0.3 4.5 0.9 2.4 30 0 0 0.7 2.5 10.2 5.2 10.2 5.2 4.3 5.2 31 0 0 1 .8 9.5 15.1 11.6 17.5 11.6 34.5 11.3 32 0 0 2 0 21.7 0.4 16.5 0.4 12.2 0.4 33 3.2 7.3 0 7.9 8.3 7.9 8.1 0 4.8 8.5 34 0 0 5.3 0 8.7 0 5.3 4.8 5.3 10.7 

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