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Scaling up fracture properties
It follows from the above paragraphs that fractures exist over a broad range of scales, ranging from the rock grain to the scale of tectonic plates. Interestingly, it has been found that the characteristics that determine groundwater flow in crystalline rock (e.g. fracture density, aperture, connectivity) scale as fractals (Bonnet et al., 2001), which means that fracture patterns at one scale are in many cases relatively similar to patterns at an entirely different scale. These similarities provide a rationale for extrapolating pat-tern data from one scale (preferably at which fractures may be easily described) to an-other (such as a larger scale where fracture networks are difficult to characterize). Many studies have examined how to quantitatively scale fracture data (e.g. Bour et al., 2002; Gillespie et al., 1993; National Research Council, 1996 and references therein). The exist-ing fractal methods are however still regarded as experimental for several reasons, such as the difficulty to test it rigorously, as cataloging all fractures at all scales is not feasi-ble, or because a single technique can yield significantly different spatial distributions of fractures. So far, stochastic methods which blend aspects of deterministic methods and fractal scale methods (e.g. Cacas et al., 1990; Darcel et al., 2003; Molz, 2004) are the most promising (National Research Council, 1996).
Landscape scale heterogeneities (macro)
Regional and catchment-scale conceptualizations of crystalline aquifers pose a signifi-cant challenge, particularly in the assessment of the structural controls on groundwater flow related to the multi-scale heterogeneity of this type of medium (Comte et al., 2012). To do so, two aspects need to be taken into account: the geological setting that defines the geometry of the hydrogeological environment (i.e. structure, the static properties), and the spatial and time variations of available water (i.e. recharge, the dynamic proper-ties) (Krásný, 2002). This part will focus on the former.
As we have seen, the main sources and reservoirs of groundwater in crystalline rock are the weathered layer and fractures. In view of this, the large-scale geometry of the aquifer can be dissected as, on one part, the heterogeneity and anisotropy of fracture systems at the meso- and macro-scale (see above), and, on the other, the heterogeneity of the geological structure and weathering patterns at the macro-scale (Comte et al., 2012). Weathering patterns at the macro-scale shape the landscape. They determine the landform sequences, where each landform type has a general groundwater potential, and they determine the thickness, extent and characteristics of the weathered layer. These, in turn, depend on several factors (i) climate (both past and present), where significant rainfall contributes to thicker weathered layers, (ii) topography, where weathered layers are mainly formed in erosional peneplain areas of low relief, (iii) the lithology and tex-ture of the parent rock, and (iv) the time span involved in weathering.
Landforms and drainage network
The most commonly developed landforms in crystalline rock are structural hills, in-selbergs, pediments, buried pediments, and valley fills. Their characteristic features are [see Singhal, 2010] (FIGURE 1.14): Structural hills These large-scale structures outcrop after combined processes of tectonism and denudation. These rocks are hard and compact, and thus act as run-off zones and have negligible groundwater potential. Some infiltration may take place in fractures and joints, which may then either discharge as springs or as seepages along the valley portions. Inselbergs are small residual hills which stand above the general level of the surrounding erosional plains. Pediments Formed due to the process of denudation, these broad, flat, or gently sloping areas develop at the base of mountain fronts or plateau escarpments.
Groundwater potential of these units is limited due to the thinness of the weathered material. Buried pediments form when the sloping surface of the pediment gets gradually covered with soil and colluvial material. These areas form better groundwa-ter potential zones as they have a better retention capacity and storage volume.
Valley Fills As pedimentation progresses, channel deposits may develop. These areas are characterized by gentler slopes, and better water retention. They are the most important landforms for groundwater development.
The drainage characteristics of a basin, specifically the drainage pattern and density, are also useful in the assessment of groundwater resources. Drainage patterns are the spatial arrangement of streams, which can be studied and mapped on topographic maps, aerial photographs or satellite images. Dendritic drainage patterns (irregular, branching of streams, resembling a tree) are typical of massive crystalline rocks, while rectangu-lar/angular patterns (streams have right-angled bands) are indicative of fractured rocks. Drainage density, Dd, is the ratio of total channel lengths of streams within a basin to the area of the basin. Generally speaking, low densities are characteristic of regions with highly resistant or permeable surfaces and low reliefs, while high drainage densities are found in regions of weak or impermeable rock (Strahler, 1964). Among drainage network characteristics, other factors such as basin shape, stream frequency and bifurcation ratio are of hydrological interest (Singhal & Gupta, 2010).
Hortonian overland flow
Hortonian overland flow (FIGURE 2.5a) is a process of saturation named after Rob-ert E. Horton, a civil engineer and soil scientist who described this process in a series of papers (Horton, 1933, 1945). Hortonian overland flow occurs when water inputs exceed the saturated hydraulic conductivity of the surface; this forms a wetting front which progresses downslope, and if input persists long enough, the surface soil becomes saturat-ed and the water that accumulates on the surface becomes runoff (Dingman, 2015). Hor-tonian flow is typically important in (i) (semi-)arid regions where rainfall tends to be intense and surface conductivities low due to a lack of vegetation, and (ii) regions where surface conductivity is either naturally very low or has been reduced through soil frost of human and animal activity (Dingman, 2015).
Lateral flow in the unsaturated zone
Lateral flow in the unsaturated zone (interflow) can be divided into two categories: lateral unsaturated (sometimes called “Darcian”) flow through the soil matrix (lateral downslope flow), and lateral preferential flow in macropores and fractures which allow water to bypass the unsaturated soil matrix.
Lateral unsaturated flow
Because the unsaturated zone is under tension–i.e. less than atmospheric–pressure (FIGURE 2.8), water cannot easily move from the unsaturated zone directly into a stream. Thus, lateral unsaturated flow is not a significant source of water to streams under most circumstances. However, a downslope component of drainage can emerge following a water-input event [see Jackson, 1992; Lv et al., 2013]. Downslope flow effects are accentuated in anisotropic environments, where the resistance to downslope flow is reduced (Jackson, 1992). In upland watersheds this is believed to be the major source of streamflow (Dingman, 2015).
Lateral preferential flow
Lateral flow through macropores and fractures, also called lateral preferential flow or bypass flow, is usually much quicker than unsaturated or saturated Darcian flow through the soil matrix (Kirkby, 1988), allowing water to be carried at greater speeds and to greater distances though otherwise unsaturated soils. While macropores occur mostly in the upper soil layers, fractured rock featured openings at depth that can extend over tens or hundreds of meters. Preferential flow can occur regardless of the antecedent soil moisture, and sometimes only a small difference between the diameter of a pore and the grain size of the soil matrix will cause bypassing. For example, in sand of 0.5 mm a pore size of 1 mm is enough to provide a preferential pathway (Kirkby, 1988). Because the number, orientation, size and inter-connectedness of macropores are highly dependent on local factors such as geology, soils, vegetation and fauna, it is difficult to generalize about the importance of bypass flow in models. Nevertheless, macropore flow has been found to contribute significantly to event flow in many environments (e.g. Jones, 2010; Newman et al., 1998; see Beven & Germann, 1982, 2013).
Flow in the saturated zone
Water that enters streams promptly in response to rainfall and snowmelt events is called event flow (also termed storm flow, quick flow, and direct flow). This phenomenon is different from baseflow which is water that enters the stream in a persistent and slow manner, and that is maintained in between water-input events.
Baseflow
Baseflow is the portion of streamflow sustained by groundwater discharge between precipitation events as long as the water table remains above the stream bottom. Stream reaches characterized by high baseflow inputs tend to have low temporal flow variability. There exist many methods to identify groundwater contributions to streamflow, either by analysis of stream hydrographs (also called baseflow analysis, baseflow separation or recession analysis) or by analyzing the chemical composition of stream water. The de-tailed approaches can be found in Dingman (2015). However, it is important to note that (i) the concept of baseflow has no scientific basis and (ii) that the source of base flow is seldom known, and may also originate from other sources, such as the drainage of sur-face water or the slow drainage of soils (Hewlett & Hibbert, 1963).
Event flow
Event flow (FIGURE 2.5c) arises from mechanisms that quickly produce steep hydraulic gradients in near-stream areas, and may be a significant component of event response under some circumstances. There are two types of event flow mechanisms, those generat-ed by near-stream groundwater mounds, or through flow from shallow saturated layers. Near-stream groundwater mounds can be formed gradually through local and upslope recharge which causes the streamward gradient to steepen. They may also be formed suddenly if water percolates to the top of the capillary fringe; this causes the menisci that maintain the tension to be obliterated, thus changing the pressure state of the wa-ter from negative to positive (Dingman, 2015). This then causes an almost instantaneous rise in the near-stream water table, forming a groundwater mound which induces streamward groundwater flow (Abdul & Gillham, 1984, 1989; Gillham, 1984). When thin permeable soils overlay relatively impermeable materials a saturated zone may form which is disconnected from the regional water table, this is called a perched water table. Downflow in these zones may thus contribute to event response; this type of situation is often referred to as a sloping slab (FIGURE 2.5d).
Table of contents :
Introduction
Groundwater resources in fractured crystalline rock
1. Geography of fractured crystalline rocks
2. Geology of fractured crystalline rocks
2.1. Definition
2.2. Crystalline rock fracturing
2.3. Weathering profile
3. Multi-scale heterogeneity
3.1. Fracture heterogeneity (micro to macro)
3.2. Landscape scale heterogeneities (macro)
4. Aquifer conceptualization
Chapter 2 Aquifer recharge and the water cycle
1. The water cycle
1.1. Generalities
1.2. The “Water Budget Myth”
2. Components of the water cycle
2.1. Precipitation
2.2. Evapotranspiration
2.3. Runoff
2.4. Infiltration and the vadose zone
2.5. The role of aquifers in the water cycle
3. Aquifer recharge
3.1. Definition
3.2. Types of recharge
3.3. Recharge controls
3.4. Global recharge distribution
Chapter 3 Quantitative evaluation of recharge processes
1. The basic water-balance equation
2. Recharge estimation methods
2.1. Recharge from infiltration RI (direct recharge)
2.2. Recharge from surface-water RSW (indirect recharge)
2.3. Total recharge (RI +RSW)
2.4. Choosing the appropriate technique
3. Remaining challenges in estimating recharge
3.1. The variability of recharge in time and space
3.2. The assessment of localized and indirect recharge Humid versus arid regions
3.4. Recharge in fractured rock
Research problem
Chapter 4 Managed Aquifer Recharge
1. Generalities
1.1. Definition
1.2. Key issues
2. MAR methods
2.1. Types of methods
2.2. Types of structures
3. Factors determining the effectiveness of MAR
3.1. Hydrological criteria
3.2. Hydrogeological criteria
4. MAR challenges and risks
4.1. Decrease in infiltration potential
4.2. Contamination of groundwater
4.3. Watershed scale impacts
5. Predicting and assessing MAR efficacy
5.1. Soil maps and hydrogeological reports
5.2. Field methods
5.3. Modeling the aquifer response to MAR
5.4. Monitoring infiltration and groundwater mounding
6. MAR in fractured crystalline rock
7. Watershed development and MAR in India
7.1. National scale
7.2. Telangana state-scale: Mission Kakatiya
Research problem
PART II: EXPERIMENTAL AND NUMERICAL INVESTIGATIONS OF NATURAL AND ARTIFICIAL RECHARGE IN FRACTURED CRYSTALLINE ROCK
Chapter 5 Natural recharge heterogeneity in weathered fractured crystalline rock
1. Introduction
2. Study site
2.1. General
2.2. Geological setting
2.3. Hydrological setting
2.4. Soil types
2.5. Land use
2.6. Reference recharge
3. Methodology
3.1. Hydraulic model
3.2. Basin discretization
3.3. Sensitivity analysis
4. Results and discussion
4.1. Threshold runoff
4.2. Diffuse recharge distribution
4.3. Diffuse recharge/rainfall relationship
4.5. Focused recharge
4.6. Sensitivity analysis of the diffuse recharge model
4.7. Conclusion
Chapter 6 Managed Aquifer Recharge in fractured crystalline rock aquifers
1. Introduction
2. Study site
2.1. Geological setting
2.2. Hydrological setting
2.3. Water level variations in response to recharge
3. Methods
3.1. Estimating infiltration rates and vertical hydraulic conductivity
3.2. Modeling the aquifer response to infiltration
3.2.1. Analytical solutions
3.2.2. Numerical modeling for connected basin
4. Results
4.1. Infiltration estimation and relative contributions
4.2. Vertical hydraulic conductivity
4.3. Horizontal hydraulic conductivity and storativity
4.4. Effect of bedrock relief (numerical model)
4.4.1. Synthetic scenarios
4.4.2. Application to the present field case
5. Discussion
5.1. Representativity of hydraulic parameters
5.2. Comparison to inferred interface relief
5.3. Artificial recharge modeling
6. Conclusions
Acknowledgments
Chapter 7 Conclusions & perspectives
1. Conclusions
1.1. Catchment-scale natural recharge processes
1.2. Site-scale artificial recharge processes
2. Perspectives
2.1. Remote sensing and airborne geophysics
2.2. Recharge quantification
2.3. On the representativity of site-scale analysis
2.4. Effects of MAR on groundwater quality
References
