Seepage area and rate of bedrock groundwater discharge at a granitic unchanneled hillslope

Transcription

1 WATER RESOURCES RESEARCH, VOL. 39, NO. 1, 1018, doi: /2002wr001298, 2003 Seepage area and rate of bedrock groundwater discharge at a granitic unchanneled hillslope Taro Uchida, Yuko Asano, Nobuhito Ohte, and Takahisa Mizuyama Graduate School of Agriculture, Kyoto University, Kyoto, Japan Received 11 March 2001; revised 17 July 2002; accepted 17 July 2002; published 22 January [1] Recent studies have demonstrated the importance of water movement through the bedrock in the rainfall-runoff process on steep hillslopes. However, quantitative information on this process is still limited. The objective of this study was to address the following questions: (1) How large is the area where bedrock groundwater seeps into the soil layer, and (2) what is the rate of water flow out of the bedrock? To address these questions, detailed hydrological, hydrochemical, and thermal measurements were conducted at a forested steep unchanneled granitic concave slope in the Tanakami Mountains, central Japan. The relationship between the amplitude of annual soil temperature variation and the measurement depth showed that in a normal low-flow period, the seepage area ranged between 14 and 21 m 2 and the ratio of this area to that of the whole catchment was about 2.0%. In a drought period the seepage area ranged between 3.5 and 5.5 m 2, and the ratio to the whole catchment was around 0.5%. The variation in the area of seepage was controlled both by the short-term precipitation pattern during the preceding several weeks and by the long-term pattern over several preceding months. A two-component geochemical hydrograph separation indicated that the ratio of bedrock groundwater to streamflow was about 0.82 for the normal low-flow periods and 0.90 for the drought period. The rate of flow out of the bedrock into the soil layer ranged from 0.5 to 3.3 m 3 d 1. That is, although the seepage area was small ( % of the catchment), the contribution of bedrock groundwater was considerable (50 95% of streamflow). INDEX TERMS: 1860 Hydrology: Runoff and streamflow; 1866 Hydrology: Soil moisture; 1829 Hydrology: Groundwater hydrology; KEYWORDS: runoff, hillslope hydrology, thermal response, bedrock groundwater Citation: Uchida, T., Y. Asano, N. Ohte, and T. Mizuyama, Seepage area and rate of bedrock groundwater discharge at a granitic unchanneled hillslope, Water Resour. Res., 39(1), 1018, doi: /2002wr001298, Copyright 2003 by the American Geophysical Union /03/2002WR SBH Introduction [2] Most of the models of hydrological behavior at unchanneled concave or planar hillslopes have ignored the effect of water flow through bedrock fractures [Grayson et al., 1992; Wu and Sidle, 1995]. The Maimai research catchment in New Zealand, where the bedrock is unfractured and extremely watertight, has been the site of ongoing research since the late 1970s [McGlynn et al., 2002]. An early conceptual model of hillslope flow at this catchment featured new water infiltrating rapidly into permeable forest soils via vertical preferential flow paths, or soil matrix structures, to a soil-bedrock interface where a perched groundwater zone formed. This produced the mixing of event water with large volumes of old stored soil matrix water; the water in the saturated zone was then displaced rapidly downslope via soil pipes along the soil-bedrock interface [McDonnell, 1990]. Woods and Rowe [1996] showed that subsurface flow varied widely across a hillslope section, and McDonnell [1997] hypothesized that bedrock surface topography controlled the variation in subsurface flow. The conceptual models applied to the Maimai catchment were corroborated by studies on hillslopes in the Panola Mountains, USA [McDonnell et al., 1996], Tatsunokuchi, Japan [Tani, 1997], and Plastic Lake, Canada [Peters et al., 1995]. These conceptual models agreed with previous numerical models, indicating that subsurface flow along the soilbedrock interface controls the runoff processes on steep hillslopes. [3] However, not all studies showed that water flow along the soil-bedrock interface was dominant contributor of lateral water movement. Several studies of lateral water movement in hillslope showed the importance of lateral flow in the shallow soil layer, including transmissivity feedback [e.g., Bishop et al., 1990] and pipe flow [e.g., Elsenbeer et al., 1995]. In contrast, recent studies, the occurrence of lateral water flow in bedrock have been demonstrated by piezometric observations [Wilson and Dietrich, 1987; Onodera, 1990; Montgomery et al., 1997], tracer approaches [Anderson et al., 1997; Noguchi et al., 1999], and complementary hydrological and hydrochemical measurements [Mulholland, 1993; Hirose et al., 1994; Burns et al., 1998; Tsujimura et al., 1999; Miyaoka et al.,

2 SBH 9-2 UCHIDA ET AL.: SEEPAGE AREA AND RATE OF BEDROCK GROUNDWATER 1999]. One of the most detailed piezometric studies was conducted at the CB1 catchment in the Oregon Coast Range, USA, where the bedrock is fractured sandstone [Montgomery et al., 1997]. Observations there showed that bedrock groundwater contributed to the formation and variation of the transient saturated groundwater zone in a steep unchanneled concave slope. In the same catchment, Anderson et al. [1997] demonstrated that vertical infiltration and water emerging from bedrock combined in a small area near the channel head. This area was considered a variable source area, as the size of this saturated zone controlled the discharge rate, the runoff water chemistry, and the age of the water. Recently, several conceptual models describing spatial aspects of flow on steep hillslopes have followed the flow path dynamics proposed by Anderson et al. [1997], in that both preevent soil water and bedrock groundwater contribute to the formation of a saturated zone in the area adjacent to a stream [Onodera, 1990; McGlynn et al., 1999; Bowden et al., 2001; Uchida et al., 2002a]. From these studies, it can be concluded that lateral flow in steep hillslopes can take a variety of forms, including movement as a thin saturated layer above unfractured bedrock [Peters et al., 1995], pipe flow at the base of the soil profile [McDonnell, 1990], macropore flow in the shallow soil layer [Elsenbeer et al., 1995] and transport via fractures in bedrock [Anderson et al., 1995]. [4] The quantitative information about water flow in soil layer was obtained by monitoring of subsurface flow at the excavated trenches [e.g., Sidle et al., 2000] and hydrograph separation using geochemical and isotopic tracers [e.g., Burns et al., 2001]. However, except for a limestone catchment where rapid bedrock flow might be expected, most of the hydrograph separations did not consider the contribution of groundwater in the fractured bedrock to the runoff water. Although quantitative information about the characteristics of bedrock groundwater flow is necessary for more realistic hillslope hydrological modeling, this quantitative information is still limited. The lack of data and poor understanding of the area where groundwater in the fractured bedrock seeps into the soil layer is one of the major obstacles for hillslope hydrological modeling using the Richards equation [e.g., Cloke et al., 2001]. The objective of this study was to address the following questions. (1) How large is the area where bedrock groundwater seeps into the soil layer, and (2) what is the rate of water flow out of the bedrock? [5] We have carried out detailed hydrological, hydrochemical, and isotopic measurements in the Fudoji catchment, a steep forested unchanneled granitic concave slope in the Tanakami Mountains of central Japan [Asano et al., 2002a, 2002b; Uchida et al., 2002b]. Tensiometers showed that a saturated area formed and a downward hydraulic gradient existed continuously in the area near a spring, as similar to the Anderson conceptual model [Uchida et al., 2002b]. Moreover, hydrochemical and isotopic measurements indicate that the spatial distribution of flows at Fudoji agrees with the Anderson conceptual model [Asano et al., 2002a, 2002b; Uchida et al., 2002b], indicating that in the small perennially saturated area near the spring, water percolates through the vadose zone mixed with water emerging from the bedrock. However, during peak runoff, the transient groundwater in the upper slope flowed to the spring via lateral preferential paths. This result did not accord with the studies in CB1 [Anderson et al., 1997; Montgomery et al., 1997], demonstrating that the fractured bedrock flow was a minor source of storm runoff [Uchida et al., 2002b]. In this study, we used detailed hydrological, hydrochemical, and thermal data to address questions (1) and (2) above. A preliminary analysis indicated that during base flow periods, the contribution of bedrock groundwater was relatively large, compared with storm flow periods [Uchida et al., 2002b]. Moreover, according to the observation from 1998 to 2000, the total period of base flow was about 11 months per year and the total base flow is about 93 % of total streamflow. The direct runoff was calculated by simple hydrograph separation with a straight line connected from the point of initial runoff to the inflection point on the recession limb on a semilogarithmic graph scale. Therefore, as the first step in quantifying the bedrock groundwater contribution, we focused mainly on hydrological behavior during the base flow period. 2. Study Site [6] The study site was an unchanneled headwater catchment (called Fudoji), located in southeastern Shiga Prefecture, central Japan. The catchment, underlain by Tanakami granite, covers 0.10 ha and has a mean gradient of 37 degrees; the vegetation consists of dense natural forest, predominately Chamaecyparis obtusa. The area is humid and temperate. The mean annual precipitation and runoff in Kiryu Experimental Forest (10 km north of Fudoji) from 1972 to 2001 were 1645 mm and mm, respectively [Katsuyama, 2002]. Two perennial springs contribute to stream water: one from the soil matrix and the other from a crack in the bedrock (Figure 1). These outflows will be referred to as the spring and the bedrock spring, respectively. In addition, soil pipe outlets with diameters ranging from 3 to 10 cm have been found at the bottom of the slope adjacent to the spring. [7] The soils are predominantly cambisols (brown forest soil). The soil depth to the bedrock along the axis of the concavity ranges from 60 to 120 cm. This depth was measured using a cone penetrometer with cone diameter of 9.5 mm, a weight of 1.17 kg, and a fall distance of 20 cm. The results of the cone penetration test are shown in Figure 2. N 4 is the number of blows required for a 4-cm penetration. Since N 4 increased sharply from 20 to 100, bedrock was assumed for N 4 values greater than 100. Anderson et al. [1997] defined bedrock as the layer that cannot be penetrated with a hand auger. The thickness of the layer with an N 4 less than 100 was almost the same as the thickness of the layer that could be penetrated with a hand auger, so our definition corresponded to that of Anderson et al. [1997]. We also measured bedrock topography at the lower end of the concave slope in detail (Figure 1b). [8] The average saturated hydraulic conductivities of the A and B horizons (measured using three 100-cm 3 field cores in the laboratory) were 9480 and 235 mm h 1, respectively.

3 UCHIDA ET AL.: SEEPAGE AREA AND RATE OF BEDROCK GROUNDWATER SBH 9-3 Figure 1. (a) Topographic map of the Fudoji watershed. The contour interval is 2.5 m. Solid circles F1 though F4 show the locations of observation points. (b) Topographic map of the bedrock surface at the lower end of the Fudoji watershed. The contour interval is 1 m. The saturated hydraulic conductivity of the bedrock was also measured using three 100-cm 3 field cores in the laboratory. The two samples were taken from a 0 5 cm depth from the bedrock surface. The average saturated hydraulic conductivity of the bedrock was 12 mm h 1. The saturated hydraulic conductivity of the bedrock at Fudoji was of the same order as those of weathered granite in the Sierra Nevada Range, California [Graham et al., Figure 2. The longitudinal axis of the concave slope (line A B, as noted in Figure 1). The profiles of the N 4 penetrometer values are shown in the inset graph.

4 SBH 9-4 UCHIDA ET AL.: SEEPAGE AREA AND RATE OF BEDROCK GROUNDWATER 1997], and the Idaho batholiths [Megahan and Clayton, 1986]. 3. Methods [9] The rate of streamflow discharge was measured, beginning in May 1997, using a V notch weir and a water level recorder installed at the watershed outlet. Water levels behind the weir were recorded at 10-min intervals throughout the observation period using a Campbell CR10X recorder. The outflow of the bedrock spring has been directed to a 500 cm 3 tipping-bucket recorder (also a Campbell CR10X). The discharge of the total streamflow (L/s) was divided by the area of the watershed (0.10 ha) to obtain the specific runoff (mm/s). Measurement of the outflow from the bedrock spring commenced in April A tipping-bucket rain gauge was located 1000 m west of the watershed. [10] Groundwater levels above the bedrock were monitored using observation wells installed along the bedrock concavity. These wells were located at F1, F2, F3, and F4 and were constructed using 6-cm-diameter PVC pipes perforated with small holes around their peripheries and installed down to the soil-bedrock interface. The groundwater level gauges each consisted of a vertical rod with a number of small cups attached at regular height intervals (4.5 or 6 cm) apart. Both the highest water levels since the last observation and the present water levels were measured during regular observations (made once every two to three weeks) beginning in May 1997 [Asano, 2001]. Most of the regular observations were made during base flow periods. Five tensiometers in each of five sets were installed together with Daiki DIK-3150 pressure transducers. These sets are referred to as F1, F1.5, F2, F2.5, and F3 (Figure 1a) and the tensiometers in each are denoted as, for example, F1 86, F1.5 55,F2 112, F2.5 73, and F3 66, (where the subscripts refer to the depth in cm). Tensiometers were installed to measure pore pressures just above the bedrock. The pore pressures were measured from April to December 1999 and from April to December [11] Soil temperature was monitored at 8 points (F1 10, F1 40,F1 86, F1.5 55,F2 10,F2 40,F2 108, F2.5 73, and F3 66 ) and was recorded at 10-min intervals throughout the observation period using Campbell CR10X and TABAI Thermo- Recorders. Soil temperatures at F1 10, F1 40, F1 86, F2 10, F2 40, and F2 108 were measured from April 1999 to May 2001 and those at F1.5 55, F2.5 73, and F3 66 from April 2000 to May [12] Through fall was sampled using three plastic bottles equipped with mesh-covered funnels with a diameter of 21 cm. Spring water, bedrock spring water, and stream water were sampled directly during base flow conditions. Soil water samples were collected using PVC pan zero-tension lysimeters (15 20 cm). The lysimeters were installed at 10 and 40 cm depths at F1, F2, and F3. Except for the one at F2 40, the lysimeters were duplicated. Groundwater just above the bedrock was sampled from 6-cm-diameter PVC wells fitted with a 10-cm screen at the bottom. The groundwater sampling wells were installed at F1, F2, F3, and F4. Groundwater samples were collected directly from cups attached to a vertical rod. Samplings were conducted at two- to three-week intervals. Samplings for through fall, groundwater, spring water, and stream water were conducted from May 1997 through May 2001 and those for soil water were conducted from July 1999 through May The SiO 2 concentration was estimated by the molybdenum yellow method using a HITACHI U-1000 spectrophotometer. 4. Results 4.1. Groundwater Level, Pore Water Pressures, and Streamflow [13] The Tanakami Mountains region has two major wet seasons: (1) the monsoon (Baiu) season from mid-june to mid-july and (2) the typhoon season from late August through early October. Thus the discharges of base flow in July, August, and October were relatively large, while those from January through March were small (Figure 3f). The mean annual precipitation and runoff from 1998 to 2000 were 1680 mm and 1470 mm, respectively. That is, the annual water loss ( precipitation minus runoff) of Fudoji is about 200 mm. However, Suzuki [1980] reported that the annual evapotranspiration rate from forests in Tanakami Mountain ranged from 720 to 760 mm. These results indicate that the drainage area (capture area) of this catchment is greater than the surface drainage area. The discharge rate of the bedrock spring was less variable, ranging from 0.91 to 1.48 m 3 d 1 (Figure 3e). During base flow periods, the ratio of outflow from the bedrock spring to total streamflow ranged from 35 to 70%. [14] Except for the period from mid-january to early June in 2000, a saturated area was continuously present at F1 (Figure 3d), while the soil-bedrock interfaces at F2 through F4 were not commonly saturated by groundwater during base flow periods (Figures 3a through 3c). Most storms produced a transient saturated groundwater area above the bedrock at F2 and F3 (Figures 3b and 3c). In contrast, only the heaviest storms produced a saturated area above the bedrock at F4 (Figure 3a). Only during the period from mid- January to early June in 2000 was there no saturated area at F1 (Figure 3d); the base flow rate was also the lowest during this period (Figure 3e), suggesting that the period from mid-january to early June in 2000 was the driest in the four-year observation period. Due to this, base flow periods during the four-year observation were classified into two periods: (1) the normal low-flow periods (except for the period from mid-january to early June in 2000) and (2) the drought period (the period from mid-january to early June in 2000). [15] Pore pressures at F2 112, F2.5 73, and F3 66 were sensitive to rainfall intensity, while pore pressure at F1 86 remained almost constant. The recession curves of soil pore water pressures above the bedrock became steeper away from the spring. Tensiometer results indicated that during the wettest period (from late June to mid-july in 1999), continuously saturated groundwater occurred at F2, but lasted for only about 1 month (Figure 4f ). The saturated area at F3 and F2.5 dissipated two to three days after a heavy storm (Figures 4b and 4c). Pore pressures at F2 112, F2.5 73, and F3 66 decreased dramatically after storms in June through August in 2000 (Figures 4b through 4d), corresponding to the period of dry weather (Figure 4a). However, there was no significant difference in the recession curve of F1 86 between the observation periods in 1999 and 2000 (Figure 4f ). The response of tensiometer F showed a

5 UCHIDA ET AL.: SEEPAGE AREA AND RATE OF BEDROCK GROUNDWATER SBH 9-5 Figure 3. (a d) Temporal variation in groundwater levels at F4, F3, F2, and F1, (e) outflow from the bedrock spring, and (f ) streamflow. Shaded area represents the drought period. Solid circles in Figures 3a 3d represent the groundwater level at routine observations. Open squares in Figures 3a 3d indicate the highest levels since the last routine observations. trend intermediate between those at F1 86 and F2 112 (Figure 4e). The pore pressure of F was sensitive to the rainfall intensity, as was that at F2 112, but the recession curve was not as steep as those at F2 112 through F3 66. The saturated area lasted for more than three months after the rainy season at F1.5 (Figure 4e) Temperature [16] The soil temperature at each observation point showed nearly sinusoidal seasonal variation overlain by short-term fluctuations (Figure 5). As depth increased, the amplitude was reduced and the phase was delayed at each location. The thermal response in the shallow soil layer (10 and 40 cm) at F1 (Figures 5a and 5b) was similar to that at F2 (Figures 5e and 5f ). The amplitude of the annual thermal response at F2 108, F2.5 73, and F3 66 decreased with increasing depth (Figures 5g 5i). [17] However, the amplitude of the annual thermal response at F was smaller than those at F2 108, F2.5 73, and F3 66, although the measurement depth of F was shallower than all the others (Figure 5d). The highest temperature at F1 86 was lower than that at F2 108, although the measurement depth at F1 86 was shallower than that at F2 108 (Figure 5g). The lowest temperature at F2 108 in February 2000 was similar to that in 2001 (Figure 5g), but the lowest temperature at F1 86 in February 2000 was 2.5

6 SBH 9-6 UCHIDA ET AL.: SEEPAGE AREA AND RATE OF BEDROCK GROUNDWATER Figure 4. (a) Hyetograph and (b f) temporal variation in the soil pore water pressure heads at F3 66, F2.5 73,F2 112, F1.5 55, and F1 86. degrees lower than that in 2001 (Figure 5c). Thus the lowest temperature at F1 86 in February 2001 was higher than that at F2 108, while the lowest temperature at F1 86 in February 2000 was lower than that at F Spatial Variability of the Dissolved Silica Concentration [18] The median dissolved silica concentrations in soil water did not vary much, regardless of sampling depth or topographic position (Figure 6). Variations in silica concentration in the 10-cm soil water were greater than those in the 40-cm soil water. Silica concentrations in transient groundwaters at F2, F3, and F4 were similar to those in the soil water, while concentrations in perennial groundwater at F1 were much higher than in the soil water or transient groundwater. The median silica concentration of F1 groundwater was 2.4 times greater than that of F2 groundwater. The highest median concentration of silica was found in the bedrock spring; the variation in silica concentration in this spring was small. Spring water and streamflow silica concentrations were generally intermediate between those of groundwater and the bedrock spring. 5. Discussion 5.1. Area of Bedrock Groundwater Seepage Into the Soil Layer [19] The differences of the lowest temperature at F1 10, F1 40,F2 10,F2 40 and F2 108 between 2000 and 2001 were less than 1.2 degrees. Thus the annual thermal response at F1 10,F1 40, F ,F2 10,F2 40,F2 108, F2.5 73, and F3 66 can be described by a single sinusoidal curve (Figure 5). However, since the lowest temperature at F1 86 in February 2000 was 2.5 degrees lower than that in 2001 (Figure 5c), the thermal response at F1 86 was described by two sinusoidal curves, one for the drought period (from January to May in 2000) and the other for normal low-flow periods (Figure 5).

7 UCHIDA ET AL.: SEEPAGE AREA AND RATE OF BEDROCK GROUNDWATER SBH 9-7 Figure 5. Temporal variation in the mean daily soil temperature. Shaded area represents the drought period. [20] When water movement does not affect the thermal response and the thermal diffusivity of the soil (K) is homogeneous, thermal diffusion is ¼ 2 where q is temperature and Z is depth. If the amplitude of the annual thermal variation at the soil surface is represented by A 0, the amplitude (A i ) of the annual thermal variation at depth Z = Z i becomes: rffiffiffiffiffiffi p A i ¼ A 0 exp Z i KT ð1þ ð2þ where T is the period of one cycle (=1 year), indicating that A i decreases exponentially with increasing depth. [21] The amplitudes at F1 10, F1 40, F2 10, F2 40, F2 112, F2.5 73, and F3 66 decreased exponentially with increasing depth (Figure 7), suggesting that water movement had little effect on the thermal response at these locations. Using the above relationship between Z and A and equation (2), we obtain a thermal diffusivity of cm 2 s 1. This value is similar to those found by Tani et al. [1979], who reported that the thermal diffusivity of five forest soils ranged from to cm 2 s 1. This agreement supports the assumptions that water movement did not affect the thermal response and the variation in K was small. [22] In normal low-flow periods, the amplitude at F1 86 was smaller than that given by the Z to A relationship for F2 (Figure 7). From June to September 1999, the temperature at F1 86 was lower than that at F2 108, although tensiometer observations indicated little difference in the soil water contents at F1 86 and F This shows that the small amplitude at F1 86 cannot be attributed to spatial variation in the thermal diffusivity of soil caused by spatial variability of the water content. Thus the small change in the thermal response indicates that water flows upward from the bedrock zone. The value of A for F1 86 during the drought period fit the Z to A relationship for F2 well (Figure 7). This indicates that during the drought period, the effect of water flow out of the bedrock on the thermal response at F1 86 was small. In other words, the variation in the saturated groundwater area (Figure 3d) was related to the variation in the area of bedrock groundwater seepage into the soil layer. The annual thermal response at F also had a small amplitude, indicating that groundwater at F1.5 was delivered from both the soil layer and the bedrock. The concentrations of dissolved silica also support the presence of a contribution of bedrock groundwater to the formation of saturated groundwater at F1. Deep groundwater usually has a significantly higher silica concentration than shallow groundwater in headwater catchments [e.g., Rice and Hornberger, 1998; Uhlenbrook et al., 2000]. Spatial variability in the dissolved silica concentration at Fudoji agrees with these studies, indicating that the highest median concentration of silica was in the bedrock spring (Figure 6). During normal low-flow periods, the silica concentration in perennial groundwater at F1 reflected the higher silica concentration in the bedrock groundwater. The dissolved silica concentrations indicated that the contribution of bedrock groundwater increased between F1 and the spring, showing that bedrock groundwater also seeped into the soil layer in the area between F1 and the spring. [23] In comparison, the dissolved silica concentration of spring water during the drought period was significant higher than that of the transient groundwater at F2 through F4, as occurred in the normal low-flow periods (Figure 8). Moreover, the stream water silica concentration during the drought period was greater than that during the normal lowflow periods. This suggests that the bedrock groundwater contribution varied less than other sources of runoff as drought set in. This indicates that water from the pathway in the bedrock seeped into the soil layer between F1 and the spring, despite the drought conditions. These indicate that when the area of bedrock outflow was greatest, the upper boundary of the area was between F1.5 and F2, while the

8 SBH 9-8 UCHIDA ET AL.: SEEPAGE AREA AND RATE OF BEDROCK GROUNDWATER Figure 6. Box plots of the dissolved silica concentration. The boundary of the box closest to zero indicates the 25th percentile; the line within the box marks the median; and the box farthest from zero indicates the 75th percentile. The whiskers left and right of the box indicate the 10th and 90th percentiles. Abbreviations are TF, through fall; SW, soil water; GW groundwater; Br, bedrock. The numbers in parentheses indicate the numbers of samples. boundary was between F1 and the spring, when the area was smallest. [24] Previous piezometric observations showed a spatially continuous pattern of the groundwater table along the axis of concavity, although the groundwater table intersected the soil-bedrock interface [Wilson and Dietrich, 1987; Onodera, 1990; Anderson et al., 1997]. Therefore we assumed that the boundary of the area of bedrock groundwater emergence was along the contour of the bedrock topography (Figure 9). Consequently, during the normal low-flow periods, the area where the bedrock groundwater discharged ranged from 14 to 21 m 2 (Figure 9a), while the area in the drought period Figure 7. Relationship between the amplitude of thermal responses and measurement depth. Open and solid squares show the result at F1 86 during the normal low-flow periods and the drought period, respectively. Figure 8. Differences in the dissolved silica concentration of spring water and streamflow between normal low-flow periods and the drought period. The rules used for box plotting are the same as those for Figure 7. Abbreviations are SP, spring water; ST, streamflow.

9 UCHIDA ET AL.: SEEPAGE AREA AND RATE OF BEDROCK GROUNDWATER SBH 9-9 Figure 9. The size of the area of bedrock groundwater seepage into the soil layer. The contour lines show the bedrock surface topography. Dark and light shaded areas represent the minimum and maximum predicted area, respectively. ranged from 3.5 to 5.5 m 2 (Figure 9b); that is, the area of bedrock groundwater seepage into the soil layer was 0.4 to 2.1% of the total watershed area Contribution of Water Seeping From Bedrock to Total Streamflow [25] The dissolved silica in natural water in headwaters is a good index of chemical weathering, since precipitation is assumed to contain little or no dissolved silica, whereas water seeping through mineral soil gains silica mainly through chemical weathering reaction [Kennedy, 1971; Schlesinger, 1997]. Moreover, recent studies showed that deep groundwater usually has significantly higher silica concentrations than shallow groundwater in headwater catchments, because of a difference in weathering availability between shallow and deep layers [Scanlon et al., Figure 10. (a) The ratio of bedrock groundwater to streamflow and (b) observed streamflow, predicted bedrock groundwater component, and observed runoff from the bedrock spring. Shaded area represents the drought period. When C s is assumed to be equal to the mean value of 40-cm soil waters and transient groundwaters, the predicted ratio is indicated by bold line in Figure 10a. When C s is assumed to be equal to 10th and 90th percentiles of 40-cm soil water and transient groundwater, the predicted ratios are indicated by fine lines in Figure 10a.

10 SBH 9-10 UCHIDA ET AL.: SEEPAGE AREA AND RATE OF BEDROCK GROUNDWATER Figure 11. period Temporal variation in the antecedent precipitation index (API). Area represents the drought 2001], indicating that dissolved silica is a useful tracer for differentiating shallow groundwater from deep groundwater [e.g., Rice and Hornberger, 1998; Uhlenbrook et al., 2000]. [26] The spatial variation in the dissolved silica concentration of shallow water (soil water and transient groundwater above the bedrock) was relatively small, compared with the difference in the silica concentration between the shallow water and the bedrock spring (Figure 6), indicating that the dissolved silica concentration is a useful tracer for differentiating shallow water from bedrock groundwater. Therefore, to quantify the water flux from the bedrock into the soil layer, hydrograph separation was conducted using the dissolved silica concentration as the basis of separation. Since this study focuses on the base flow period, it can be assumed that the direct contribution of through fall to runoff water is negligible. Therefore the streamflow was separated into two components, assuming mass conservation, as follows Q r ¼ Q b þ Q s Q r C r ¼ Q b C b þ Q s C s where Q is the discharge, C is the dissolved silica concentration, and the subscripts r, s, and b refer to streamflow, water traveling through the soil layer, and water seeping from the pathway in the bedrock, respectively. ð3þ ð4þ [27] During storm runoff, shallow soil water was sometimes delivered to the stream via preferential flow paths, bypassing the normal process of mixing with groundwater through the soil matrix [Hagedorn et al., 2000; Uchida et al., 2001]. During base flow periods, however, these preferential flow paths play little part in the rainfallrunoff process on forested hillslopes [e.g., Tsuboyama et al., 1994], and the direct contribution of the 10-cm soil water to streamflow was small. Therefore we assumed that C s was equal to the mean value of the 40-cm soil waters and transient groundwaters (F2 through F4). To examine the effects of variability in silica concentration on the calculation, we used three values of C s : the mean value and the 10th and 90th percentiles of the 40-cm soil water and transient groundwater. In contrast, the value of C b was assumed to be the mean value from the bedrock spring, because the variation in the dissolved silica concentration of bedrock spring water was very small (Figure 6). [28] The percentage of the discharge of water emerging from the bedrock to the total streamflow ranged from 50 to 95% (Figure 10a), although the area of bedrock groundwater seepage into the soil layer was 0.4 to 2.1% of the watershed area. The mean ratio of the bedrock groundwater to streamflow was about 0.82 in normal low-flow periods, and 0.90 in the drought period. The percentage of the discharge of water emerging from the bedrock to the total streamflow increased as streamflow declined, with the exception of the results of days 378 and 813. The stream

11 UCHIDA ET AL.: SEEPAGE AREA AND RATE OF BEDROCK GROUNDWATER SBH 9-11 water samplings of days 378 and 813 were conducted immediately after the end of direct runoff (<12 h). Thus it can be thought that the relatively low percentage of bedrock groundwater to the total streamflow might be caused by the residual effects of previous direct runoff, such as contribution of shallow soil water and through fall. [29] Although the runoff rate from the bedrock spring remained almost constant, ranging from 0.91 to 1.48 m 3 d 1, the predicted bedrock groundwater rate ranged from 1.7 to 5.2 mm d 1 (m 3 d 1 ). The mean predicted bedrock groundwater runoff rate was 2.8 mm d 1 (m 3 d 1 ) for normal low-flow periods, and 1.8 mm d 1 (m 3 d 1 ) for the drought period (Figure 10b). The water flow rate from the pathway in bedrock into the soil layer can be computed by subtracting the discharge rate of bedrock spring from Q b. During the drought period, the water flow rate from bedrock into the soil layer was small, from 0.5 to 1.3 m 3 d 1, while the flow rate in ordinary low-flow periods ranged between 1.1 and 3.3 m 3 d 1. This indicates that the water flow rate out of the bedrock into the soil layer was related to the area of bedrock groundwater seepage, indicating that the temporal variation in the rate of water out of the bedrock per unit area was small Influence of the Preceding Precipitation Pattern on the Contribution of Bedrock Groundwater to Streamflow [30] Here we discuss the effects of temporal precipitation patterns on the size of the area of bedrock groundwater seepage into the soil layer. The antecedent precipitation index (API) is commonly used to model the residual effect of previous precipitation on current soil moisture or runoff [e.g., Ziemer and Albright, 1987]. Therefore we can argue that if the variation in the size of bedrock seepage area was controlled by the preceding precipitation pattern, this size should vary with the API. To test this hypothesis, the API at any time was calculated as: API ¼ X500 0:5 i=t P i ð5þ i¼1 where T is the half-life representing the decay characteristic of a particular recession, and P i is the total rainfall amount i days beforehand. The values of T that were tested ranged from 20 to 120 days. For periods when the local precipitation data was not available, data from Kiryu Experimental Watershed (10 km north of our measurement site) were used to calculate the API. [31] The API for T = 100 days in normal low-flow periods exceeded 500 mm, while the API in the drought period was less than 500 mm (Figure 11c). This indicates that when T was greater than 100 days, the API was useful as an indicator of size of bedrock seepage area, which was controlled by the preceding precipitation pattern. In comparison, when T was smaller than 80 days, the API from January to March in 1999 was similar to the API from January to May in 2000 (Figures 11a and 11b). In other words, when T was smaller than 80 days, the API alone did not describe the variation in the size of seepage area, indicating that the size of this area was influenced not only by the precipitation pattern over the preceding few weeks (short-term effects), but also by the precipitation pattern over several preceding months (long-term effects). 6. Summary and Conclusions [32] Based on the data from detailed hydrologic, hydrochemical, and thermal observations, we determined the area where the bedrock groundwater emerged into the soil layer and the water flux out of the bedrock on a steep unchanneled concave hillslope. The area where the bedrock groundwater emerged into the soil layer became small when conditions were extremely dry. The size of the seepage area was controlled by both the rainfall intensities during the preceding several weeks, and those during the preceding several months. In normal low-flow periods, the area ranged between 14 and 21 m 2 and the ratio of the area to the whole area of the catchment was about 2.0%. In the drought period, the area ranged between 3.5 and 5.5 m 2, and the ratio to the area of the whole catchment was around 0.5%. The ratio of bedrock groundwater to streamflow was about 0.82 for normal low-flow periods, and 0.90 for the drought period. The rate of water flow out of the bedrock into the soil layer ranged from 0.5 to 3.3 mm d 1. That is, although the area where the bedrock groundwater emerged into the soil layer was small ( % of the whole catchment), the contribution of bedrock groundwater was considerable (50 95% of total streamflow). [33] Recent studies have demonstrated that bedrock geology has a strong influence on the runoff characteristics of bedrock groundwater flow [e.g., Tsujimura et al., 1999; Onda et al., 2001]. This suggests that it is necessary to elucidate the runoff characteristics of bedrock groundwater flow in a variety of geological conditions for a better understanding of hillslope hydrological processes. From the results of this study, we concluded that, in the temperate region, hydrometric, thermal, and geochemical measurements need to be coupled in order to quantify the bedrock groundwater flow processes. [34] Acknowledgments. This study was supported by a grant from the Fund of the Japanese Ministry of Education and Culture for Science Research. We are also grateful to Yasunori Nakagawa and Masatoshi Kawasaki for assistance in the field, and Naoko Tokuchi for assistance in conducting the laboratory experiments. 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