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An Overview of Tools for Assessing Groundwater-Surface Water Connectivity Ross Brodie, Baskaran Sundaram, Robyn Tottenham, Stephen Hostetler and Tim Ransley

© Commonwealth of Australia 2007 This work is copyright. Apart from any use as permitted under the Copyright Act 1968, no part may be reproduced by any process without prior written permission from the Commonwealth. Requests and inquiries concerning reproduction and rights should be addressed to the Commonwealth Copyright Administration, Attorney General’s Department, Robert Garran Offices, National Circuit, Barton ACT 2600 or posted at http://www.ag.gov.au/cca.

The Australian Government acting through the Bureau of Rural Sciences has exercised due care and skill in the preparation and compilation of the information and data set out in this publication. Notwithstanding, the Bureau of Rural Sciences, its employees and advisers disclaim all liability, including liability for negligence, for any loss, damage, injury, expense or cost incurred by any person as a result of accessing, using or relying upon any of the information or data set out in this publication to the maximum extent permitted by law.

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Preferred way to cite this publication: Brodie, R, Sundaram, B, Tottenham, R, Hostetler, S, and Ransley, T. (2007) An overview of tools for assessing groundwater-surface water connectivity. Bureau of Rural Sciences, Canberra.

Tools for Assessing Groundwater-Surface Water Connectivity 1

Foreword

Integrated management of surface water and groundwater is critical in ensuring sustainability of the water resource and for meeting the objectives of the National Water Initiative. Water issues such as over-allocation, environmental flows and river salinity are all influenced by the connectivity between streams and aquifers. This means that groundwater-surface water interactions need to be assessed and incorporated into the management response to a range of water quantity and quality issues.

The assessment of stream-aquifer connectivity can be difficult and complex and there is a wide variety of approaches that can be taken. This includes conventional approaches such as interpreting water chemistry or stream hydrographs, as well as other methods which are not routinely used in Australia such as temperature monitoring and seepage meters. Each of these methods have their strengths and weaknesses and measure stream-aquifer connectivity at different scales in time and space.

This report outlines the different approaches available for assessing groundwater-surface water interactions and encourages combining different methods in an overall strategy. Field work has been undertaken in two trial catchments by Bureau of Rural Sciences (BRS) to evaluate some of these assessment methods and to help develop a conceptual understanding of water flow in and between streams, wetlands and aquifers.

This report is part of the Managing Connected Water Resources project, a collaboration between BRS, Australian Bureau of Agricultural and Resource Economics (ABARE), the Australian National University and State agencies. The project objective is to progress a more coordinated approach to the management of surface water and groundwater resources. The project has developed a comprehensive information package on connectivity issues, including assessment methods, at www.connectedwater.gov.au.

Colin Grant

Executive Director

Bureau of Rural Sciences

April 2007



Executive Summary Groundwater and surface water resources are hydraulically connected in many regions of Australia and better understanding of this connectivity is critical for effective water resource management. Assessing groundwater-surface water interactions is often complex and difficult. However, there are a range of methods available as documented in this report, including: (i.) Seepage Measurement, the direct measurement of water flow at the surface

water-groundwater interface using seepage meters or similar devices; (ii.) Field Observations, where an initial reconnaissance can highlight hotspots

where groundwater is interacting with surface water features; (iii.) Ecological Indicators, mapping of specific vegetation communities or biota

that indicate groundwater discharge to surface water features; (iv.) Hydrogeological Mapping, to define the hydrogeology surrounding a surface

water feature including specific geological features such as faults, facies changes or river morphology that can control groundwater flow;

(v.) Geophysics and Remote Sensing, the use of geophysical and remote sensing technologies such as airborne electromagnetics (AEM), radiometrics, seismic waves, electrical charge, or satellite imagery;

(vi.) Hydrographic Analysis, the use of techniques such as recession analysis or baseflow separation to analyse the monitoring record of water levels or flows;

(vii.) Hydrometric Analysis, investigating the hydraulic gradient between groundwater and surface water systems and the hydraulic conductivity of the intervening aquifer and bed material;

(viii.) Hydrochemistry and Environmental Tracers, the interpretation of the chemical constituents of water such as major ions, isotopes, radon and chlorofluorocarbon (CFC);

(ix.) Artificial Tracers, the monitoring of the movement of an introduced tracer such as a fluorescent dye;

(x.) Temperature Studies, the use of time series monitoring of temperature in both the surface water and groundwater systems;

(xi.) Water Budgets, approaches such as river reach water balances; (xii.) Modelling, the use of analytical or numerical modelling techniques based on

governing mathematical equations to predict water movement. Simple methods such as field observations, field chemistry surveys or stream flow measurements can give valuable information in terms of providing a catchment-scale perspective on connectivity as well as targeting areas for more detailed investigation. Site specific investigations using simple tools such as seepage meters, mini-piezometers, temperature loggers or environmental tracers provide more detail in terms of understanding and quantifying key processes. There is a need to use a combination of assessment methods rather than relying on any particular one. This is necessary to not only confirm any interpretation but also to extrapolate any findings in time and space.



Contents Executive Summary.......................................................................................................................... 3

Contents ............................................................................................................................................ 5

Figures............................................................................................................................................... 9

Table ................................................................................................................................................ 11

1. Introduction................................................................................................................................. 13

2. Assessment of Connectivity ..................................................................................................... 15 2.1 Available Assessment Methods for Connectivity .................................................................. 16 2.2 Comparison of Methods ........................................................................................................ 18 2.3 Assessment Strategy............................................................................................................. 23 2.4 Data Collation ........................................................................................................................ 24 2.5 Desktop Analysis ................................................................................................................... 26 2.6 Field Survey........................................................................................................................... 27 2.7 Site Investigations ................................................................................................................. 28

3. Seepage Measurement............................................................................................................... 29 3.1 Seepage Meter Design.......................................................................................................... 30 3.2 Seepage Meter Operation ..................................................................................................... 33 3.3 Automated Seepage Meters.................................................................................................. 35 3.4 Advantages and Disadvantages............................................................................................ 36 3.5 Data Availability ..................................................................................................................... 37 3.6 Relevant Links ....................................................................................................................... 37

4. Field Observations ..................................................................................................................... 39 4.1 Advantages and Disadvantages............................................................................................ 39 4.2 Data Availability ..................................................................................................................... 39

5. Ecological Indicators ................................................................................................................. 43 5.1 Advantages and Disadvantages............................................................................................ 43 5.2 Data Availability ..................................................................................................................... 43

6. Hydrogeological Mapping.......................................................................................................... 45 6.1 Advantages and Disadvantages............................................................................................ 49 6.2 Data Availability ..................................................................................................................... 49

7. Geophysics and Remote Sensing............................................................................................. 51 7.1 Advantages and Disadvantages............................................................................................ 53 7.2 Data Availability ..................................................................................................................... 53

8. Hydrographic Analysis .............................................................................................................. 57


8.1 Baseflow Separation.............................................................................................................. 59 8.2 Graphical Separation Methods.............................................................................................. 59 8.3 Filtering Separation Methods ................................................................................................ 60 8.4 Frequency Analysis Methods ................................................................................................ 62 8.5 Recession Analysis Methods ................................................................................................ 64 8.6 Advantages and Disadvantages............................................................................................ 68 8.7 Data Availability ..................................................................................................................... 70 8.8 Relevant Links ....................................................................................................................... 70

9. Hydrometric Investigations ....................................................................................................... 73 9.1 Piezometers........................................................................................................................... 73 9.2 Head Difference Measurement ............................................................................................. 77 9.3 Hydraulic Conductivity Measurement.................................................................................... 78 9.4 Flow Net Analysis .................................................................................................................. 80 9.5 Advantages and Disadvantages............................................................................................ 81 9.6 Data Availability ..................................................................................................................... 82 9.7 Relevant Links ....................................................................................................................... 82

10. Hydrochemistry ........................................................................................................................ 85 10.1 Field Water Quality Parameters .......................................................................................... 85 10.2 Major Ion Chemistry ............................................................................................................ 86 10.3 Stable Isotopes.................................................................................................................... 87 10.4 Radioactive Isotopes ........................................................................................................... 89 10.5 Industrial Chemicals ............................................................................................................ 90 10.6 Advantages and Disadvantages.......................................................................................... 90 10.7 Data Availability ................................................................................................................... 91 10.8 Relevant Links ..................................................................................................................... 91

11. Artificial Tracers ....................................................................................................................... 95 11.1 Advantages and Disadvantages.......................................................................................... 97 11.2 Relevant Links ..................................................................................................................... 98

12. Temperature Studies................................................................................................................ 99 12.1 Advantage and Disadvantages ......................................................................................... 101 12.2 Data Availability ................................................................................................................. 102 12.3 Relevant Links ................................................................................................................... 102

13. Water Budgets ........................................................................................................................ 105 13.1 Stream Flow Measurement ............................................................................................... 105 13.2 Volumetric Analysis ........................................................................................................... 106 13.3 Velocity-Area method ........................................................................................................ 106 13.4 Slope-Area Method............................................................................................................ 108


13.5 Dilution Gauging................................................................................................................ 111 13.6 Thin Plate Weirs ................................................................................................................ 111 13.7 Advantages and Disadvantages........................................................................................ 112 13.8 Data Availability ................................................................................................................. 112 13.9 Relevant Links ................................................................................................................... 112

14. Acknowledgements................................................................................................................ 115

15. References .............................................................................................................................. 117

16. Appendix 1 .............................................................................................................................. 128



Figures Figure 2.1: Examples of different methods of assessing stream-aquifer connectivity ..................... 19

Figure 2.2: Components of a strategy for investigation and assessment of connectivity ................ 24

Figure 2.3: A framework for conjunctive water management........................................................... 25

Figure 3.1: Basic design of a seepage meter with inverted open chamber ..................................... 32

Figure 3.2: Basic components of the seepage meter including seepage chamber and collection bag.................................................................................................................................... 32

Figure 4.1: Field indicators of discharge of shallow acid groundwater into a coastal drainage network ............................................................................................................................................. 40

Figure 4.2: Lawn Hill Creek at Lawn Hill........................................................................................... 41

Figure 6.1: Different scale groundwater flow systems within a catchment....................................... 45

Figure 6.2: Groundwater flow systems operating within an alluvial riverine valley .......................... 46

Figure 6.3: Schematic cross section of the hydrogeology of the Alstonville Plateau ....................... 47

Figure 6.4: Extent of traditional published hydrogeological maps at 1:250,000 scale or more detailed ............................................................................................................................................. 48

Figure 7.1: A seismic gun used to fire a shotgun charge into the ground to generate a shockwave. ....................................................................................................................................... 53

Figure 7.2: A 144m long floating electric .......................................................................................... 54

Figure 7.3: EC ribbon images........................................................................................................... 54

Figure 7.4: EC ribbon images from geo-electric surveys ................................................................. 54

Figure 7.5: Airborne electromagnetics image showing groundwater recharge zone. ...................... 55

Figure 7.6: Paleochannels detected from airborne magnetics survey in Honeysuckle Creek subcatchment ................................................................................................................................... 55

Figure 8.1: Components of a typical flood hydrograph..................................................................... 58

Figure 8.2: Graphical baseflow separation techniques .................................................................... 60

Figure 8.3: Flow distribution curves for examples of (2a) high baseflow and (2b) low baseflow streams ............................................................................................................................................. 63

Figure 8.4: Procedure for recession curve displacement method.................................................... 68

Figure 8.5: Annual baseflow indices for unregulated streams in the Murray-Darling Basin............. 71

Figure 8.6: Annual volume of Q90 percentile for available stream gauges in the Richmond River catchment................................................................................................................................ 72


Figure 9.1: Configuration of minipiezometer and stilling well for hydrometric measurement of seepage flux. .................................................................................................................................... 74

Figure 9.2: Example of monitoring bore construction....................................................................... 75

Figure 9.3: Stages in the installation of minipiezometer and stilling well for hydrometric investigations of seepage flux .......................................................................................................... 76

Figure 9.4: Watertable contour patterns around streams................................................................. 81

Figure 9.5: Plan view of example groundwater flow net towards a gaining surface water feature............................................................................................................................................... 81

Figure 9.6: Groundwater and surface water level measurements for Yellow Creek site during March, 2005...................................................................................................................................... 83

Figure 9.7: River levels versus nearby bore water levels in key sites in the Border rivers catchment ......................................................................................................................................... 84

Figure 10.1(a) Field electrical conductivity (uS/cm) and (b) pH of Gum Creek and nearby springs, July 2004............................................................................................................................. 92

Figure 10.2. Deuterium versus oxygen-18 concentrations for river water and groundwater in the Border Rivers Catchment ........................................................................................................... 93

Figure 10.3 Chloride versus deuterium concentrations for river water and groundwater in the Border Rivers Catchment ................................................................................................................. 93

Figure 11.1: Dye tracer technique for assessing groundwater and surface water interaction in the field ............................................................................................................................................. 96

Figure 12.1: Common temperature sensors used to measure sediment and stream temperatures..................................................................................................................................... 99

Figure 12.2: Temperature variation of groundwater and stream under gaining (a) and losing stream (b) conditions ...................................................................................................................... 100

Figure 12.3 Observed stream and sediment temperatures downstream of Goondiwindi Weirs.... 103

Figure 13.1: Stream flow measurements taken in November 2004 on the Alstonville Plateau ..... 114


Tables Table 2.1: Summary of tools to assess stream-aquifer connectivity ................................................ 20 Table 2.2: Spatial scales in stream-aquifer connectivity .................................................................. 23 Table 2.3: Time scales in Stream-Aquifer Connectivity.................................................................... 23 Table 2.4: Typical catchment hydrology datasets and sources ....................................................... 26 Table 3.1: Comparison of automated seepage meters .................................................................... 36 Table 3.2: Results of seepage meter trials in Meershaum Drain, Tuckean Swamp ........................ 38 Table 7.1: Different Ground-based electromagnetic techniques...................................................... 53 Table 7.2: Common geophysical tools used in borehole logging..................................................... 51 Table 8.1: Recursive digital filters used in base flow analysis.......................................................... 62 Table 8.2: Different storage-outflow models used in recession analysis ......................................... 69 Table 9.1: Commonly used units for hydraulic conductivity (K) ....................................................... 78 Table 9.2: Indicative hydraulic conductivities of some rock types.................................................... 79 Table 10.1: Australian Standards related to water sampling............................................................ 87 Table 10.2 Relative abundances of the oxygen and hydrogen isotopes.......................................... 88 Table 10.3: Decay constants and half-lives of selected radioactive isotopes with application to

hydrology .................................................................................................................................. 90 Table 13.1: Different procedures for determining mean velocity at a vertical ................................ 107 Table 13.2: Mannings n values for small natural streams.............................................................. 109 Table 13.3: Calculation of Mannings n from Field Observations.................................................... 110 Table 16.1: Inventory of published hydrogeological maps in Australia .......................................... 128



1. Introduction

Groundwater and surface water have been historically managed as isolated components of the hydrologic cycle, even though they interact in a variety of physiographic settings (Sophocleous, 2002). In many catchments, groundwater and surface water are hydraulically connected. For example, surface water features such as rivers, lakes, dams and wetlands can receive groundwater from underlying aquifers (Winter et al, 1998). These interactions can have significant implications for both water quantity and quality. Seepage of fresh groundwater into a river can be important in maintaining flows during extended dry periods. This can be critical for supplying the needs of surface water users such as irrigators as well as for aquatic ecosystems. Pumping from an aquifer near a river can dramatically change the amount of this baseflow to the river. In contrast, if the groundwater is salty or contaminated, increased groundwater discharge can have a negative effect on river water quality. Hence, effective management of water quantity and quality issues requires an understanding of these surface water-groundwater interactions. Assessing groundwater-surface water interactions is often complex and difficult. Commonly, groundwater level measurements are used to define the hydraulic gradient and the direction of groundwater flow. Flow measurements at various points along the stream are used to estimate the magnitude of gains or losses with the underlying aquifer. Other tools used to investigate groundwater-surface water interaction include seepage meters (Lee and Hynes, 1978; Cherkauer and McBride, 1988; Brodie et al, 2005), river bed piezometers (Baxter et al, 2003), time-series temperature measurements (Stonestrom and Constanz, 2003) and environmental tracers (Crandall et al, 1999; McCarthy et al, 1992; Herczeg et al, 2001; Baskaran et al, 2004). In most cases the limited number of data collection points results in a lack of detailed understanding of groundwater-surface water interactions in the field. Numerical modelling approaches on the other hand can provide a valuable tool for developing a framework by combining information obtained from the other field methods. This report documents the tools available to assess connectivity between surface water and groundwater systems in a catchment. This report also gives examples of trials of some of these assessment tools to better understand the nature of connectivity in the two catchments, the Border Rivers in the Murray-Darling Basin and Lower Richmond on the north coast of New South Wales.



2. Assessment of Connectivity Traditionally, surface water and groundwater resources have been independently assessed. An important addition in taking an integrated approach is that connectivity is also assessed. The nature and level of this assessment will depend on the: (i.) Key water management issues within the catchment;

(ii.) Significance of the water resource in terms of social, economic and environmental values;

(iii.) Relative development of the water resource in terms of the ratio between use (and allocation) and sustainable limits;

(iv.) Risk assessment of the likely magnitude of impacts associated with the management issue, such as loss of economic productivity, land and water degradation or poor ecosystem health;

(v.) Availability of resources such as data, budget and expertise, and; (vi.) Management and policy timeframes. Hence, water resource assessment includes investigation of: (i.) Surface water features including streams, reservoirs, wetlands and estuaries.

This includes such aspects as flow duration and dynamics, water storage capacity, water quality, aquatic ecosystems, land use impacts, climate variability and water extraction regimes;

(ii.) Groundwater systems, covering aspects such as aquifer geometry, geological and stratigraphic configurations, hydraulic properties such as transmissivity and storativity, water sources and sinks such as recharge, abstractions and discharge mechanisms, environmental dependencies and the impacts of land use, and;

(iii.) Surface water-groundwater interactions, involving the analysis of the dynamics of water flow between aquifers and surface water features, and the impacts of this interaction in terms of water quantity, quality and ecology.

Hence, the focus is to acquire the baseline information to describe the characteristics of surface water and groundwater systems of the catchment, and their interactions, both spatially and temporally.


2.1 Available Assessment Methods for Connectivity A wide range of tools are available to assess the nature and degree of the connectivity (Figure 2.1). A summary of these methods is outlined below, with more detailed information provided in the following chapters of this report. Seepage Measurement The direct measurement of seepage flux at the stream-aquifer interface can be undertaken using seepage meters or similar devices. The basic concept is to cover and isolate the stream bed with an inverted open chamber and measure the change in volume of water contained in a bag attached to the chamber over a measured time interval. Additional water in the bag over the time of operation indicates gaining stream conditions. Several modifications have been made to the design and operation of the seepage meter to address potential sources of measurement error and to handle logistical issues. Automated versions using different technologies to enable real-time monitoring of seepage flux have been developed. Field Observations Visual evidence of seepage flux can be observed in certain catchments and settings. An initial reconnaissance can highlight hotspots where groundwater is discharging to streams; provide guidance to useful parameters to measure and to identify management issues that are impacted by connectivity. Examples of field indicators include direct observation of water flow from springs at the margins or within the stream bed, water vapour or ice-free conditions around springs during winter, mineral precipitates or iron-bacteria accumulations, or changes in water colour or odour. Ecological Indicators Specific vegetation communities or biota can indicate groundwater discharge to surface water features. Changes in the composition and accumulated biomass of submerged aquatic plants can relate to groundwater seepage. The near-stream presence of phreatophytic plants, which are deep-rooted and can access groundwater, can indicate a shallow watertable. The extent and composition of biota that inhabit the hyphoreic zone can also indicate the processes of near-stream groundwater and surface water mixing. Hydrogeological Mapping Knowledge of the hydrogeology surrounding a surface water feature is critical in understanding connectivity. This involves mapping the configuration and characteristics of the groundwater flow systems within the catchment. This covers aspects such as aquifer geometry, host geology and stratigraphy and hydraulic properties (such as transmissivity and storativity). Also included are specific geological features such as faults, facies changes or river geomorphology that can locally control groundwater flow.


Geophysics and Remote Sensing Geophysical and remote sensing technologies such as airborne electromagnetics (AEM), radiometrics, seismic waves, electrical charge, or satellite imagery can be used to interpret connectivity. These surveys can map the variation in parameters such as groundwater salinity, vegetation types or soil moisture that can be secondary indicators of groundwater discharge. They can also be used to identify geological features that control seepage flux. Mapping of landscape parameters (such as soil type, land use and vegetation cover) that can have an impact on seepage flux can also be supported by geophysical or remote sensing technologies. Hydrographic Analysis The stream hydrograph can be processed and analysed to characterise the magnitude and timing of groundwater discharge to streams. Baseflow separation techniques use the time-series record of stream flow to derive a baseflow hydrograph. Of these techniques, recursive filters are the most commonly applied. Frequency analysis takes a different approach by deriving the relationship between the magnitude and frequency of stream flows. Recession analysis focuses on recession curves which follow stream flow peaks. These curves are fitted using storage-outflow models to characterise the natural storages that feed the stream. Hydrometric Analysis Hydrometric methods are based on Darcy’s Law so focus on the hydraulic gradient between groundwater and surface water systems and the hydraulic conductivity of the intervening aquifer and bed material. Piezometers are used to measure groundwater levels which are compared with the elevation of the stream stage. Pump (or slug) tests can be undertaken on these piezometers to estimate the transmissivity of the aquifer material. Hydrochemistry Studies Interpretation of the chemical constituents of water can provide insights into stream-aquifer connectivity. Dissolved constituents can be used as environmental tracers to track the movement of water. For example, a particular characteristic of the groundwater chemistry (such as high radon levels) can be used as an indicator of groundwater discharge when measured in the surface water. Environmental tracers can occur naturally or have been released into the general landscape by human activities. Some of the commonly used environmental tracers include field parameters such as: EC or pH; the major anions and cations such as calcium, magnesium, sodium, chloride and bicarbonate; stable isotopes in the water molecule of oxygen-18 (18O) and deuterium (2H); radioactive isotopes such as tritium (3H) and radon (222Rn); and industrial chemicals such as chlorofluorocarbons (CFC) and sulphur hexafluoride (SF6). Artificial Tracers Artificial tracer tests are used to evaluate the extent to which aquifers interact with streams, providing information on groundwater flow paths, travel times, velocities, dispersion, flow rates and the degree of hydraulic connection. These tests involve the


introduction of a tracer material or chemical and subsequent monitoring of its movement. This differs from environmental tracer methods which rely on the measurement and interpretation of background concentrations. Fluorescent dyes (such as Rhodamine WT), conservative major ions (such as chloride or bromide), organic compounds (such as ethanol or fluorinated benzoates), isotopes (such as selenate or deuterium) and non-pathogenic micro-organisms or colloidal material (such as clubmoss spores) have been used in tracer studies. Temperature Studies Heat can also be used as a tracer to characterise seepage flux. Time series monitoring of temperature in both the surface water and groundwater systems is used. Stream temperatures have a characteristic diurnal pattern overprinting seasonal trends, whilst regional groundwater temperatures tend to be relatively constant at the daily scale. Temperature monitoring at varying depths in the stream bed can indicate the relative influence of groundwater and surface water processes. Numerical models of heat flow (such as VS2DH and SUTRA) can be used to quantify seepage flux. Water Budgets A common approach to investigating seepage flux between a stream and underlying aquifer is to measure stream flow at specific points. These measurement sites subdivide the stream into reaches and a water budget is estimated for each reach, accounting for inputs such as tributary flows and outputs such as evaporative losses and diversions. The difference between inflows and outflows is then attributed to the seepage flux. The method relies on accurate measurement of stream flow and appropriate accounting of the other gains and losses. 2.2 Comparison of Methods Table 2.1 presents a summary comparing these different assessment methods. These tools are described in the context of: (i.) Spatial Scale, classified in terms of local (ie at a point or site), intermediate (at

the scale of a feature such as a stream reach) and regional (at the catchment scale), refer Table 2.2;

(ii.) Temporal Scale, classified in terms of short-term (over the timeframe of days to months such as tidal, evapotranspiration or discrete episodic processes), medium-term (at the seasonal to yearly scale) and long-term (exceeding the decadal timeframe such as influences of climate change), refer Table 2.3;

(iii.) Cost, associated with collection, analysis and interpretation of data; (iv.) Ease of Use, focusing on the accessibility of technology and the extent of prior

expertise required; (v.) Advantages, the inherent benefits of applying the methodology;

(vi.) Limitations, the potential constraints and limiting assumptions; (vii.) Application, outlining the extent that the method has been used in Australia.


Figure 2.1: Examples of different methods of assessing stream-aquifer connectivity (a) direct measurement of flux using seepage meters (b) hydrometric studies using minipiezometers in the stream bed (c) monitoring of groundwater levels and stream levels/flows (d) temperature monitoring in the stream and shallow bed sediments (e) run-of-river geophysical survey (f) water sampling for hydrochemistry

Field Assessment Tools for Stream-Aquifer Connectivity

a

b

c

d

e

f


Table 2.1: Summary of tools to assess stream-aquifer connectivity

Method Spatial Scale

Temporal Scale Cost Ease of Use Advantages Limitations Application

Desktop Tools

Hydrographic Analysis Processing of time-series stream flow monitoring to define baseflow (groundwater discharge) component

Intermediate to regional Hydrograph represents water balance for subcatchment above gauge

Medium to long-term Depends on length of monitoring record

Low High Many analysis techniques and software tools available. Stream flow data routinely collected

Uses existing flow monitoring data. Can be undertaken as a desktop study prior to detailed field investigations. Provides information of seepage changes through time

Applicable to gaining stream conditions only. Assumption that baseflow is groundwater discharge may not be valid. Baseflow effected by water use and management activities (eg regulation) does not provide spatial distribution of groundwater input along stream

Commonly applied method for unregulated Australian catchments

Hydrogeological Mapping Mapping of groundwater systems including flowpaths, groundwater quality, aquifer structure and properties and geomorphology.

Intermediate to Regional Typical mapping scales of 1:100,000 to 1:250,000

Short to Medium- term Usually ‘average’ conditions at time of mapping Some parameters such as aquifer transmissivity or structural contours are time-insensitive

Medium to High Depends on data availability. Expensive if drilling required to supplement existing data

Low to Medium Knowledge of hydrogeological principles required

Provides conceptual understanding of groundwater systems around stream and hydrogeological controls on connectivity

Compiling and interpreting hydrogeological data can be time consuming and complex. Limited borehole data can lead to misinterpretation.

Groundwater flow system, surface geological and hydrogeological mapping available at a coarse scale for many groundwater management areas across Australia.

Modelling Simulate water flow regime around stream using mathematical equations

Intermediate to Regional Typical models are 2D profiles or 3D grids

Medium to Long-term Used to predict future events

Low to High Depends on data availability and model complexity

Low to Medium Requires good conceptual understanding of hydrological processes and modelling expertise

Useful predictive tool for management and policy. Helps define information gaps. Transient 3-D models can estimate changes in seepage through time and space.

Oversimplified models may not be adequately robust. Over-complex models can be data hungry, costly and time-consuming

Commonly, surface water models for a catchment are developed in isolation to groundwater models.




Field Tools

Field Indicators Visual indications of seepage such as water clarity, springs, aquatic plant species, chemical precipitates etc

Local Site specific observation of seepage indicators

Short-term Current at time of observation

Low

Medium to High Easily incorporated into field work. Depends on familiarity with indicators.

Can identify seepage hotspots quickly. Return visits can provide information on seasonal changes in seepage flux. Field indicators can form basis for mapping (eg airphoto interpretation)

Limited in quantifying seepage flux. Effectiveness varies with observer’s knowledge of field indicators (eg plant or aquatic biota).

Used in specific settings such as acid groundwater (eg iron precipitates, lilies) and karstic streams (eg travertine deposits). Assessment of groundwater-dependent ecosystems not routine

Artificial Tracers Monitoring movement of introduced tracers such as fluorescent dye to track water flow

Local to Intermediate Short to Medium term Typical tracer studies over days to weeks

Medium Need to establish monitoring network

Medium Conceptually simple but needs expertise in field measurement and data interpretation

Can provide direct evidence of water movement between stream and aquifer. Aquifer parameters and fluid transport properties can be quantified.

Tracer studies require careful planning including meeting environmental regulatory controls. Processes such as degradation, precipitation or sorption can affect tracer performance.

Not routinely applied in connectivity studies in Australia. Overseas focus on karstic aquifers or investigations of contaminated sites.

Geophysics and Remote Sensing Use of geophysics (eg resistivity, EM, radiometrics) or remote sensing (eg Landsat) to map landscape features that indicate or control connectivity

Local to Regional Range from site specific (eg downhole surveys) to intermediate (eg run-of-river EC imaging), to catchment scale (eg satellite imagery).

Short-term Measures conditions at the time of survey. Multiple surveys can provide trends through time.

Medium Per hectare cost depends on technology and platform (eg ground, airborne)

Low Needs technical expertise in field equipment operation and data interpretation

Allows rapid, non-invasive mapping of landscape parameters with good spatial resolution. Some techniques provide information at depth.

Requires specific equipment, technical expertise and logistical support. Can require complex data processing and calibration with other datasets. Ground surveys can encounter obstacles such as rough terrain, vegetation cover etc.

Opportunities exist to use geophysical data collected for other purposes eg. mineral exploration. Satellite imagery commercially available, some free in public domain.

Hydrochemistry and Environmental Tracers Use of chemical constituents of water (such as major ions, stable isotopes, radon) to track water flow

Local to Regional Depends on scope of water sampling survey.

Short to Medium-term Defines chemistry at time of sampling. Time-series monitoring (eg EC, pH) possible.

Medium to High Can be expensive due to sampling logistics and cost of analyses

Low Requires expertise in appropriate sampling and data interpretation

Useful in quantifying seepage flux and defining key hydrological processes (such as groundwater recharge and discharge).

Can have long lead times between sample collection and final analytical results.

Commonly used in Australia to identify hydrogeological processes including groundwater seepage to streams.




Hydrometrics Measurement of hydraulic gradient between aquifer and stream and the hydraulic conductivity of intervening aquifer material. Based on Darcy’s Law.

Local to Regional Can range from in-stream studies, to borehole transects to regional flow net analysis

Short to Medium-term Possible to compare hydrographs of stream and groundwater levels

Low to Medium Can use existing data but costly if drilling of bores is required

Medium to High Comparison of groundwater and stream levels simple. Estimation of hydraulic conductivity more difficult.

Comparison of stream and groundwater levels a simple guide to seepage direction. Installation of minipiezometers in stream bed allows direct local measurement of potential seepage direction.

Relies on reasonable estimate of hydraulic conductivity to quantify seepage flux. Assumption of simple groundwater flow conditions may not be valid. Point measurement. Need to correct for density effects.

Comparison of stream levels with nearby groundwater levels commonly used to define direction of potential seepage.

Seepage Measurement Direct measurement of water flow between stream and aquifer using seepage meters

Local Point measurement of seepage. Many measurements required to map spatial variations.

Short-term Meters typically installed over days/weeks. Measures aggregate seepage over time of operation.

Low to Medium Can be time consuming if measuring at multiple sites.

Low Simple concept with meters easy to use and no prior technical knowledge required.

Direct measurement of seepage flux. Meters are simple and inexpensive to construct and provide a semi-quantitative measurement.

Potentially significant measurement errors due to meter design and operation. Unsuitable for high stream flow, gravel and heavy clay sediment beds

Main application to date in Australia has been investigating leakage from irrigation channels or studying aquatic ecosystems

Temperature Monitoring Monitor variations in stream and sediment temperatures to trace seepage.

Local Multiple measurements required to map spatial variability in seepage

Short-Medium term Temperature can be included in time-series monitoring.

Low Temperature loggers are cheap and widely available.

Medium to High Temperature simple to measure. Heat transfer modelling to quantify seepage more difficult.

Temperature loggers are simple, robust and cheap. Heat transfer models that can compliment flow models to quantify seepage are available.

Only measures at a point. Interpretation of monitoring requires confirmation using other assessment methods.

Not specifically applied to study stream-aquifer connectivity in Australia to date. Opportunities to incorporate real-time temperature monitoring into existing hydrographic network

Water Budgets Quantification of stream reach water balance to define seepage component

Intermediate to Regional Does not provide spatial variability of seepage along reach being investigated

Short to Medium Term Possible to use time-series monitoring of stream flow at multiple stations

Low to Medium Can be expensive if data collection required for estimating water balance components

Medium to High Conceptually simple using existing monitoring data. Water balance components such as extraction or diversions can be difficult to quantify

Simple water balances estimated rapidly using existing stream flow monitoring. Provides estimate of aggregate seepage along reach.

Measurement errors in stream flow data can be significant, hence more suited to long reaches. Can be misleading if water balance component (eg extraction) is not adequately accounted for.

Routinely applied, particularly for regulated rivers or irrigation channels.


Table 2.2: Spatial scales in stream-aquifer connectivity

Scale Typical Units Relevance Catchment-scale Regional

>100 km2 Hydrogeological setting Water management areas Catchment management targets Catchment monitoring and reporting

Feature-scale Intermediate

1-100 km Water management decisions Environmental Planning

Site-Scale Local

<100 m Process studies Ecosystem dependencies Water quality protection

Table 2.3: Time scales in Stream-Aquifer Connectivity

Scale Typical Units Relevance Long-term Decades-

centuries Climate variation Land use change Groundwater extraction

Medium-term Seasons-years Water management cycle Allocation and planning Water quality protection

Short-term Days-months Episodic events Evapotranspiration Tidal effects Ecosystem dependencies

2.3 Assessment Strategy Table 2.1 highlights the diversity of the methods available to characterise stream-aquifer connectivity. These variations, particularly in terms of differences in spatial and temporal scale can be used to advantage in an overall assessment strategy (Figure 2.1). Figure 2.2 outlines such a strategy that fits within the overall conjunctive water management framework, as described in Brodie et al. (2007) and summarised in Figure 2.3. The components of the assessment strategy are data collation, desktop analysis, field survey and site investigations. The understanding of connectivity at different scales both in time and space brought about by this assessment strategy is bundled into the conceptual model developed for the groundwater and surface water systems of the catchment. In turn, this conceptual model can be translated into a predictive model, as detailed in Brodie et al. (2007). This process in fact is iterative, as the predictive modelling can highlight information gaps which can spur on additional data collection and assessment.


Figure 2.2: Components of a strategy for investigation and assessment of connectivity 2.4 Data Collation The initial step is to collate the existing baseline data useful in characterising the surface water and groundwater systems of the catchment, and their connectivity. This can be time-consuming and resource-intensive as data requirements are comprehensive and need to be sourced from multiple agencies. Table 2.2 summarises the main data themes and their common sources, and include: (i.) Catchment Properties such as boundaries and topography, remote sensing

imagery; (ii.) Hydrogeology such as existing geological, soils, regolith or hydrogeological

mapping, or borehole databases or geophysical surveys; (iii.) Surface Water Features, including mapping of drainage and waterways; (iv.) Hydrology, including climate data (rainfall, evaporation), stream gauging,

groundwater monitoring and water quality databases; (v.) Ecosystems, such as wetlands mapping, vegetation mapping and

rare/endangered species databases; (vi.) Catchment Use and Management, such as land use mapping, water

infrastructure, water metering and water allocation.

Data Collation

Desktop Analysis

Field Survey

Site Investigations

Conceptualisation

Prediction


Figure 2.3: A framework for conjunctive water management (Brodie et al, 2007). The tools described in this report relate to the water resource assessment component of the framework Without adequate datasets, analysis of stream-aquifer connectivity cannot be carried out with any degree of confidence. Data needs both spatial and temporal distribution to allow proper interpretation of catchment hydrogeology and hydrology to be made. The quality of data available varies with catchments, and this will reflect the type and accuracy of analysis that can be done. The data required to develop a conceptual understanding of connectivity should be determined as early as possible in the planning process. All data sets should meet agreed quality criteria relating to accuracy and temporal and spatial variability. Brodie et al. (2007) provides further information on these key catchment datasets and their data sources in Australia.

Assess Water

Resources

Conjunctive Water Management Framework

Understand and

Predict

Set Management

Targets

Develop and Implement

Management Options

Monitor and Review

Performance

Identify Management

Setting


Table 2.4: Typical catchment hydrology datasets and sources

Date Type Data Sources

Catchment Properties

Topography Topographic maps and DEMs Hydrogeology

Geology and stratigraphy, aquifer extent and thickness, confining units, bedrock configuration Aquifer properties (eg hydraulic conductivity, storativity, anisotropy)

Geological maps Geological databases Hydrogeological maps Remote sensing State agency groundwater databases Scientific literature (journal and conference papers, student theses) Unpublished reports (eg consultant reports, drilling programmes, geophysical surveys)

Surface Water Features

Drainage network, lakes, wetlands, estuaries Topographic maps Bathymetric maps

Hydrology

Rainfall and evapotranspiration Run-off and stream flow Groundwater recharge and discharge Stream-aquifer connectivity Water quality (eg. salinity, acidity)

Climate databases Stream gauging data Groundwater monitoring databases Water quality databases

Ecosystems

Aquatic ecosystems Wetlands Rare/endangered species Vegetation

Water management and environmental protection agencies Vegetation mapping Scientific papers Unpublished reports (eg environmental impact statements)

Catchment Use and Management

Land use Water infrastructure (dams, channels, irrigation, extraction bores, flood mitigation and drainage works, interception or injection schemes) Water allocation and use Community requirements and expectations Legal, regulatory and policy setting

Topographic maps Land use mapping Remote sensing State agency databases Catchment authorities (eg CMAs, councils, water authorities) Unpublished reports

2.5 Desktop Analysis The collation of existing datasets provides an opportunity to undertake an initial desktop analysis of connectivity. Such desktop analysis can provide preliminary insights into seepage flux without any additional investment in data gathering. Depending on budget and time constraints, this may be the extent to which an assessment can be made. Such a desktop analysis can include the approaches of: (i.) Hydrogeological Mapping, where available data such as borehole information,

pump tests, groundwater monitoring, geophysics and geology or soils mapping are combined to compile maps such as groundwater potentials, flow directions, salinity, aquifer structural contours and hydraulic conductivity


distribution. These are necessary in providing the hydrogeological setting for the stream;

(ii.) Hydrographic analysis, where various methods can be applied to the available time-series of stream flow to characterise the baseflow component;

(iii.) Water Balance, where the stream flow record from multiple stream flow gauges can be used to derive the water balance for the intervening stream reaches, or for the overall catchment;

(iv.) Hydrochemistry, where any existing water quality monitoring (such as EC, pH, major ions, nutrients) is used to define variations in surface water and groundwater chemistry that reflects changes in seepage flux in time and space;

(v.) Geophysics and Remote Sensing, where available imagery and data can be processed to provide information on catchment parameters that either control or indicate connectivity;

(vi.) Hydrometrics, most notably when the available time series of stream water levels are compared with nearby monitoring of watertable elevation, to determine changes in the potential direction of groundwater flow. Historic pump tests can indicate the magnitude of aquifer transmissivity;

(vii.) Temperature Monitoring, acknowledging that stream temperature is commonly monitored (in conjunction with stream level and also for ecological purposes) and can potentially also be used in assessing seepage flux.

2.6 Field Survey Additional field surveys can be undertaken to support the initial desktop analysis. These surveys are used to provide greater spatial resolution by interpolating along the stream between existing monitoring sites and to infill areas in the catchment with insufficient data. Such surveys are also used to identify key sites that require more intensive investigations. Initial conceptualisation of processes can also be verified. Examples of these survey methods include: (i.) Hydrochemistry, where water samples are taken along the stream network and

analysed for environmental tracers. This is commonly done to highlight trends and hotspots for groundwater discharge to streams. The tracers analysed range from the simple and cheap (such as field EC and pH) to the more sophisticated and expensive (such as stable isotopes and radon);

(ii.) Geophysics, where surveys can be undertaken down the length of the stream or across the extent of the catchment. For example, geo-electric arrays can be towed behind boats to map the electrical conductivity of the water column and underlying sediments. The technique has been particularly effective in mapping seepage of highly saline groundwater. Ground-based seismic traverses can be useful in mapping geological features that constrain or control groundwater flow, such as the geometry and stratigraphy of alluvial aquifers or fault zones;

(iii.) Water Balance, where flow is measured at multiple points along the stream to help target hotspots in terms of gross seepage losses or gains. Simple water balances can be estimated relatively quickly and cheaply to derive an initial rough estimate of the direction and magnitude of seepage on a stream reach basis;

(iv.) Ecological Indicators, involving the reconnaissance survey of indicator species such as specific aquatic plants, phreatophytes or hyporheic biota to


map groundwater discharge hotspots. Other field observations (such as precipitates, water colour) can be included in the reconnaissance;

(v.) Temperature Measurements, where measurements of water temperature along the stream length are used as a screening tool for identifying gaining and losing stream reaches. This is particularly useful if the groundwater discharge has a significantly higher temperature than the ambient stream, either due to geothermal conditions or due to release from deep regional groundwater flow systems.

2.7 Site Investigations More intensive investigations can be undertaken at key sites in the catchment. This is commonly undertaken to confirm key processes, quantify seepage fluxes or provide more information relating to the hydrological, chemical or ecological aspects of connectivity. Such sites are selected on the basis of the initial desktop analysis and field surveys. Investigations at these specific sites can include: (i.) Seepage Measurement, involving the installation of seepage meters to estimate

the direction and magnitude of seepage flux. As it is a direct measurement, seepage meters have the potential to validate indirect methods that involve measuring secondary indicators such as hydraulic head difference, chemical tracers or isotopes;

(ii.) Hydrometric Analysis, where piezometers and stream gauges are installed to allow local comparison of groundwater and surface water levels and so define hydraulic gradients. Pump tests can be undertaken to estimate shallow aquifer transmissivity;

(iii.) Artificial Tracers, by running a tracer test to quantify aquifer parameters and fluid transport properties, particularly in highly variable aquifers (such as fractured rock or karsts) and in solute transport studies (such as contaminants and nutrients). Specific tracers can be used to track pollutants such as human pathogens, where the movement and fate of these pollutants may not match water flow. Tracers can be used to assess the significance of local geological features (such as faults, clay layers or cave systems) on stream-aquifer connectivity;

(iv.) Temperature Studies, with time-series monitoring of temperature fluctuations for the stream and the sediment profile at varying depths to evaluate seepage flux and hydraulic conductivity. Temperature loggers are robust, simple and relatively inexpensive and available for various scales of measurement.


3. Seepage Measurement Most methods of assessing surface water-groundwater interactions are indirect. This means that the nature and magnitude of water flow is inferred or calculated from the measurements of other parameters such as hydraulic head, hydraulic conductivity, temperature, or isotopes. Direct methods use instruments that directly measure the flow of water at the interface between the surface water feature and the aquifer. Seepage meters are the most commonly used devices for the direct measurement of seepage flux. These were initially developed in the 1940’s to measure loss of water from irrigation channels (Israelson and Reeve, 1944) and resurrected in the 1970’s for use in small lakes and estuaries (McBride and Pfannkuch, 1975; Lee, 1977; John and Lock, 1977; Lee and Cherry, 1978). Seepage meters have since been used in numerous studies of seepage fluxes in rivers (Lee and Hynes, 1978; Libelo and MacIntyre, 1994; Cey et al, 1998; Landon et al, 2001), the near-shore marine zone (Bokuniewicz and Pavlik, 1990; Valiela et al, 1990; Cable et al, 1997; Taniguchi et al, 2003), tidal zones (Belanger and Walker, 1990; Robinson et al, 1998), coral reefs (Simmons and Love, 1984, Lewis, 1987), large lakes (Cherkauer and McBride, 1988) and water-supply reservoirs (Woessner and Sullivan, 1984). A constant-head variant of the seepage meter (the Idaho meter) has been used to measure leakage from irrigation channels into aquifers under Australian conditions (ANCID 2000; Byrnes and Webster, 1981). The basic concept of the seepage meter is to cover and isolate part of the sediment-water interface with a chamber open at the base and measure the change in the volume of water contained in a bag attached to the chamber over a measured time interval. The classic design of Lee (1977) consists of a 15 cm end section of a 55 gallon (~200 L) drum, which is inserted into the sediment. A stopper with a tube is inserted into a hole in the top of the drum and a plastic bag is attached to the tube with rubber bands. The time when the bag is connected and when it is subsequently disconnected is recorded, as well as the change in the volume of water in the bag. The seepage flux (Q) is calculated as:

tAVV

Q of )( −=

Equation 3.1

where Vo is the initial volume of water in the bag, Vf is the final volume of water in the bag, t is the time elapsed between when the bag was connected and disconnected, and A is the surface area of the chamber. Additional water in the bag represents upwards (gaining) seepage and water loss from the bag represents downward (losing) seepage. In environments with positive seepage flux, the water in the bag can be collected for chemical analysis.


3.1 Seepage Meter Design Over recent decades, various modifications have been made to the basic seepage meter to address potential sources of measurement error and to handle operational issues. The inverted open drum is still the basis of the chamber. A wide range of options have been used by previous investigators including capped PVC casing (Schincariol and McNeil, 2002), plastic buckets (Cey et al, 1988; Alexander and Caissie, 2003), a purpose built rectangular stainless-steel funnel (Paulsen et al, 2001), fibreglass domes (Shinn et al, 2002) or a cut-down galvanised water tank (Rosenberry and Morin, 2004). A chamber with a relatively large radius is recommended as laboratory tests indicate that variability in seepage measurements decreased with increasing diameter of the chamber (Isiorho and Meyer, 1999). There may be a requirement for the chamber to be robust and stable for use in dynamic flow conditions. In a particular seepage study of Lake Michigan the chamber used was modified by adding a 50-70 kg layer of concrete to the inside of the chamber, with the lower surface of the concrete conically-shaped to direct upwards flow of water (and gas) to the chamber outlets (Cherkauer and McBride 1988). However, such chambers may be too heavy for use in soft sediments. Other modifications to the chamber include incorporating lugs onto the top (Figure 3.1). This allows a rod to be inserted across the chamber top, to facilitate rotation of the chamber during installation. The rod or a lightweight notched steel picket can be hooked onto the chamber (or alternatively ropes attached) to remove it from the sediment. A central fitting can also be incorporated to allow attachment of a rigid vertical pole to help position or remove the chamber. The top of the chamber can be made removable to minimise disturbance of the sediment bed during installation. The top of the chamber can also be painted white to help find and recover the meter, particularly in turbid water. The tube for the collection bag can be placed to the side of the chamber and another tube at the chamber top is extended above the water surface and open to the atmosphere (Figure 3.1). This configuration is useful in shallow water to keep the bag submerged, while allowing venting of any gas. Alternatively, a small pipe with a ball valve can be added to the top of the chamber, and used to vent any trapped air when the chamber is initially placed into the water body (Cherkauer and McBride, 1988). After the air is released the valve is closed, so this approach does not allow release of gas accumulated during the actual operation of the meter. A flexible bag is used rather than a rigid container for the water storage device as the water in the bag needs to be in hydraulic equilibrium with the chamber and surface water body (Figure 3.2). The principle is that any discharge of groundwater across the surface area of the bed should displace water trapped within the chamber into the bag, likewise any recharge of water to the aquifer would be reflected in loss of water from the bag. The selection of an appropriate bag is based on the objective of minimising the energy required to exchange water between the bag and the chamber. Hence, the bag should be robust but flexible, smooth, compliant and thin-walled to reduce head losses. In many studies, hospital dialysis or intravenous bags are used as they are relatively rugged and designed to be attached to tubing. Oven basting bags have also been used (Shinn et al, 2002). Small-volume elastic bags such as balloons or condoms have been trialled previously but are not recommended as the elastic stretch in the bag


can create artificial pressure differences (Harvey and Lee 2000; Schincariol and McNeil 2002). Bladders from wine casks have also been successfully used in Australian trials refer Box 3.1 (Brodie et al. 2005). The bag can be housed in a protective cover such as an open length of PVC pipe or a perforated bucket (Figure 3.1). Field and laboratory studies have shown that surface water movement like waves, currents or streamflow can cause a venturi effect that reduces the hydraulic head in the collection bag and hence the chamber by a centimetre or more. This head loss is significant compared to the natural hydraulic gradient and can induce anomalous upwards groundwater seepage (Libelo et al. 1994). Hence, measurements from seepage meters are more reliable in slow-moving water with velocities less than 0.6m/s (ANCID, 2000). The tubing used to connect the bag to the chamber should be sufficiently rigid to avoid kinking or flexing. Again, the objective is to minimise head losses by using relatively large diameter tubing and avoiding the use of small-diameter fittings that constrict water flow. This is because frictional head loss is inversely proportional to the diameter of the flow conduit. Laboratory tests recommend that tubing diameter should exceed 7.9mm to reduce the hydraulic resistance that can cause measurement error (Fellows and Brezonik, 1980; Rosenberry and Morin, 2004). A valve can be been incorporated between the chamber and the collection bag, located as close as practical to the bag (Figure 3.1). The valve can be opened to commence the test and closed to finish the test. A two-way valve can also be used, one with tubing connected to the bag, and the other being a short length of tubing open directly to the surface water body. The valve can be manually operated, but remotely operated versions, using a solenoid-controlled switch (as used in fuel-lines in trucks) have been applied (Cherkauer and McBride 1988). Initially the valve directs flow to the short open tube to allow equilibration of water pressure between the inside and outside of the chamber. After a period of stabilisation, the valve is switched to allow connection between the chamber and the bag.


Figure 3.1: Basic design of a seepage meter with inverted open chamber (1) with flanges to assist in installation and recovery (2). The chamber has a sloping top with a gas venting tube (3) attached at the most elevated side. A 4 L wine bladder acts as a seepage collection bag (4) which is housed in an open protective housing (5). The connecting hose (6) has fittings (7) to enable quick release and a valve (8) near the bag.

Figure 3.2: Basic components of the seepage meter including seepage chamber and collection bag. (Brodie et al. 2005)

T

Stream Bed

Surface Water Body

1

3

2

4

5

86

7


3.2 Seepage Meter Operation In terms of the actual operation of the seepage meter, the following procedures and practices are suggested: (i.) Place the chamber on the bed of the surface water body open-end down and

rotate slowly (about 1 cm/s) into the sediment until the top is about 2 cm above the sediment surface. The top of the chamber should not stick up too much out of the bed because of upward advection of interstitial water (Bernoulli effect) caused by such positive relief in environments with waves, tides or currents. This process was interpreted to account for anomalous inflows into meters installed in a shallow marine and reef setting (Shinn et al, 2002). Semi-analytical analysis suggests that a seepage chamber set at a depth that is the same as the chamber radius will collect more than 90% of the ambient flow, assuming efficient design of the bag and tubing (Murdoch and Kelly, 2003). The chamber should be installed as deep as possible to limit the ingress of shallow throughflow or recirculated surface water. However, avoid placing the chamber too deep into the sediment so that the lid is directly on the sediment bed. Also avoid pushing the chamber too rapidly into the sediment as this can cause blowouts that become preferred pathways for water flow (Lee, 1977). In reality the depth of installation is largely predicated on the competence of the sediment, and the need to not excessively disturb the sediment profile. The chamber should be tilted slightly so that the vent hole is relatively elevated, as this allows any entrapped gas to escape freely;

(ii.) Minimise the activity around the meter during installation and operation. By monitoring the pressure within a chamber using a transducer, field studies have shown that walking past or stepping near the meter can effect hydraulic pressure and cause artificial inflows into the chamber (Rosenberry and Morin, 2004). Subsequent measurements of seepage in areas of the sediment bed disturbed by previous installations or by repeated foot traffic can return larger seepage rates. This has been attributed to the disturbance of a thin, lower-permeability sediment veneer (Rosenberry and Morin, 2004);

(iii.) Allow sufficient time between initial installation of the chamber and the commencement of measurements so that hydraulic pressures inside the chamber equilibrate with those of the surface water body. Laboratory tests suggest that 80% of this equilibration occurs in the first 10 minutes (Cherkauer and McBride, 1988; Cable et al. 1997) and investigators have used stabilisation times ranging from 10-15 minutes (Landon et al. 2001) to 2-5 days (Shaw and Prepas, 1989; Shaw and Prepas, 1990);

(iv.) The end of the vent tube can be fixed into position on the bank of the surface water feature using a stake or small star picket. This can be flagged to become a useful marker for the location of the seepage meter during its operation;

(v.) Pre-fill the collection bag with a known volume of water before attaching the bag to the meter. Plastic bags have an inherent tendency to expand slightly during operation of the meter, inducing a head loss. This causes an anomalous short-term influx of water into the bag after being attached to the chamber (Shaw and Prepas, 1989; Blanchfield and Ridgeway, 1996). This error was effectively eliminated in field trials when the bag was filled with 1000 mL of water prior to attachment;


(vi.) Before attaching the tubing (and bag) to the chamber ensure that the water in the bag is in hydraulic equilibrium with the surface water body. This is done by slowly lowering the bag into the water with the valve open and the chamber-end of the tubing above the water surface. This will expel any air within the bag through the tubing, however be careful not to lose any water from the bag. When this is completed, turn the valve closed;

(vii.) Attach the tubing to the chamber via the hose fitting. Before attachment, remove any air within the tubing between the hose fitting and the valve. This is done by submerging the bag and tubing, with the hose fitting directed upwards to allow air bubbles in the tubing to escape. Place the bag inside its protective cover and avoid folding, creasing or kinking the bag as these can result in anomalous and erratic head differences (Murdoch and Kelly, 2003). Weights such as concrete blocks can be placed on both the chamber and the protective cover to prevent any lateral movement due to high stream flow;

(viii.) With the bag attached and inside its protective cover, the meter is ready for operation. Seepage measurement commences when the valve is opened – make sure that you record the time that this was done;

(ix.) A control bag can be used to quantify the effects of factors such as waves, wind or currents or the properties of the bag itself (Sebestyen and Schneider, 2004). This is a pre-filled bag and tubing identical to that used in the meter that is submerged and tethered (with valve closed) about 0.15 m above the sediment bed at the same time as the meter is installed and operated. The bag is positioned near the meter but not attached to the meter. Any change in the water volume within the control bag reflects the magnitude of these effects. Field studies in the nearshore coastal environment have also used complete control meters set up in sand-filled plastic swimming pools on the bed, specifically to measure such measurement artefacts (Cable et al. 1997);

(x.) After a period of time the seepage measurement is ended by returning to the meter, turning the valve closed and recording the time that this was done. The duration of the test is based on the local seepage regime and can vary from less than an hour to several days, so a trial and error approach is required. The change in water volume in the bag should exceed 50 ml. Avoid letting the bag fill close to its maximum capacity due to significantly increased head losses, confirmed by laboratory tests (Murdoch and Kelly, 2003). Likewise, avoid completely draining the contents of the bag in the situation of high negative seepage;

(xi.) Remove the tubing from the chamber via the hose fitting and measure the volume of water in the bag. This can simply be done with a measuring cylinder. An alternative approach is to weigh the pre-test and post-test bag, to define the change in water volume (assuming a density of 1, or alternatively correct for temperature effects). Use Equation 3.1 to derive the seepage rate;

(xii.) Investigators have incorporated a meter correction factor to the calculation of seepage rates, taking account of the measurement artefacts due to frictional resistance and head losses within the meter. Laboratory testing indicated a ratio of measured to actual seepage of 0.77 (Belanger and Montgomery, 1992). For negative fluxes involving movement of surface water into the aquifer, correction factors have ranged between 1.11 and 1.74 (Rosenberry and Morin, 2004). Such correction factors would be unique to a particular seepage meter and would require calibration in a laboratory flume;


(xiii.) The meter should be routinely inspected and cleaned. Potential problems include leaky fittings, perforations or split seams in the bag, plugging of tubing by algal growth and kinks in the tubing.

3.3 Automated Seepage Meters The Lee-type manual seepage meter only measures the aggregated seepage gain or loss over a fixed time period. Various types of automated meters have been developed to record the time-series change in seepage flow. The same principle of an inverted chamber to isolate and direct seepage flux is used. However, instead of using a plastic bag or similar collection device, different instruments are used to continually measure the rate of water flow through the outlet tube. Some automated devices are compared in Table 3.1 and include the: (i.) Continuous heat meter, which applies the “Granier” method commonly, used

in sap flow meters that measure water flux in trees. This is based on the effect of water flow velocity on a temperature gradient established along the flow tube – the temperature difference is at a maximum under no-flow conditions and progressively decreases with increasing water flow velocity (Taniguchi and Iwakawa, 2001);

(ii.) Heat pulse meter which is based on the travel time of a heat pulse generated within the flow tube, which is also a function of water flow velocity (Taniguchi and Fukuo, 1993; Krupa et al, 1998);

(iii.) Ultrasonic meter that is based on the relationship between water flow velocity and the travel time of an ultrasonic signal through the flow tube (Paulsen et al, 2001);

(iv.) Dye-Dilution meter, based on the principle that the rate that a dyed solution is diluted is directly proportional to the water flow rate in the flow tube (Sholkovitz et al, 2003). The method involves the timed injection of a water-soluble dye and the subsequent measurement of the absorbance of the dyed solution;

(v.) Electromagnetic meter which is based on Faraday’s law of induction and measures the voltage induced by the movement of a conductive material (water) perpendicular through a magnetic field, which is proportional to the flow velocity (Rosenberry and Morin, 2004);

(vi.) A proximity switch which is tripped each time a collection bag is filled (Reay and Walthall, 1992).


Table 3.1: Comparison of automated seepage meters (after Sholkovitz et al. 2003)

Continuous Heat Heat Pulse Ultrasonic Dye Dilution Electromagnetic Reference Taniguchi and

Iwakawa 2001 Taniguchi & Fukuo 1993

Paulsen et al 2001 Sholkovitz et al. 2003 Rosenberry & Morin 2004

Area of seepage housing (m2)

0.25 (can be changed)



0.29 (can be changed) 0.25 (can be changed)

Diameter of flow tubing (cm)

1.1 0.95 0.85 5.1

Minimum time resolution 5 min 1 s 5 min 1 sec

Maximum time resolution

Adjustable (up to days)

Adjustable (up to days)

Adjustable (up to days) Adjustable

Range of measurable flow rates (cm/d)

2-40 0.9-300 0.5-150 5.6-5600 17-17000

Able to measure forward and reverse flow

Not in current configuration

Yes Yes Yes

Able to operate manually No No Yes No

3.4 Advantages and Disadvantages A seepage meter only measures flux at a point in space. This means that many measurements are required to derive meaningful interpolations, which is labour intensive and time consuming. High spatial and temporal variability in seepage characteristics can result in poor repeatability of measurements. Such variability can be attributed to variations in water levels through time, spatial variations in aquifer hydraulic conductivity, the presence of a thin clogging layer and changes in its hydraulic resistance, or variable seepage velocities across the stream profile, with velocity decreasing with increasing distance from the bank (Kaleris, 1998). Also manual seepage meters only measure the aggregate seepage over the time period, and do not provide any data on how seepage changes during that time period. Significant measurement errors can be introduced with the design and operation of the seepage meter. Processes such as upward advection of interstitial water caused by the chamber having a positive relief in a flowing stream (the Bernoulli Effect), venturi effects of stream flow on the collection bag, hydraulic resistance along the internal boundaries of the meter causing head losses, or accumulation of sediment gas in the chamber can lead to misleading data. This means that measurements from seepage meters are generally not reliable enough to quantify seepage flux in absolute terms. In low-flux environments (such as in heavy clay sediments or where hydraulic gradients are low) measurements may require days for an adequate change in bag water volume to derive a reliable estimate. The fluxes measured may not be entirely groundwater, but include other sources such as shallow throughflow or recirculation of surface water through the sediments. This can be a major issue if the seepage meter is not installed to a sufficient depth into the sediment. The meters are generally unsuitable for hard, gravelly or weedy sediment beds because of the difficulty in providing an effective seal and installation depth. Sand, silt or soft clay are the best sediment material for bedding down the chamber. The installation and effectiveness of seepage meters is problematic in deep or fast-flowing waterbodies, being more suited to less dynamic environments such as drains, shallow lakes or lagoons.


Due to this range of potential sources of error, it is recommended that seepage meter measurements are accompanied by other indirect methods (such as using minipiezometers to measure head difference) to verify the most likely direction and magnitude of seepage flux. Despite these significant limitations, the seepage meter is still the method that is most commonly available for the direct measurement of water flow at the interface between aquifers and surface water features. As it is a direct measurement, seepage meters have the potential to validate indirect methods that involve measuring secondary indicators such as hydraulic head difference, chemical tracers or isotopes. However, this can only be undertaken when there is confidence that any potential measurement errors have been recognised and minimised. Seepage meters are based on a simple concept and are inexpensive to construct, using readily available components. Although not sufficiently reliable to measure seepage in absolute terms, the method can still be useful for defining relative differences in seepage flux. This is useful in identifying seepage hotspots for further investigation. Seepage meters can also be useful in estimating the direction and order of magnitude of seepage flux. They can be easily incorporated into field investigations as they can be installed at the beginning of the trip, regularly monitored and then removed at the end of the trip. The seepage meter can also be used as a valuable educational tool, to raise awareness of the connectivity between groundwater and surface water resources. 3.5 Data Availability Seepage meter measurements of streams are not routinely undertaken in Australia. The most common use of seepage meters is in the measurement of water losses from irrigation supply channels. A list of rural water suppliers (which potentially have undertaken such investigations) is maintained by the Australian National Committee on Irrigation and Drainage (ANCID). 3.6 Relevant Links ANCID Channel Seepage Management Tool http://www.ancid.org.au/seepage/index.html.


Box 3.1: Trialling seepage meters to measure seepage flux of a coastal drain, Tuckean Swamp, northern NSW A seepage meter investigation was undertaken at a site on the Meerschaum Drain in the Tuckean Swamp (Brodie et al. 2005). The focus was the discharge of shallow acidic groundwater into the drainage network. Two meters were installed in the Meershaum Drain over a week in September 2004. The meters were identical except for the gauge of the gas venting tube – Chamber 1 had 4 mm (fine) black polythene tubing whilst Chamber 2 had 13 mm. Two basic types of collection bags were trialled; the bladders from 4 litre wine casks and the bladders from hydration systems (Flexiflask and Caribee Thirst Pak). The results are summarised in Table 3.2. The initial trial with the Caribee hydration bladder proved unsuccessful due to problems with connecting and disconnecting the bag to the chamber. This particular bag has 7 mm tubing incorporated with it. During its removal, the bag was lifted and without any valve, water from the collection bag would have returned back into the chamber. This is reflected in the collected water being less than the pre-filled 100 mL and the anomalous negative (downwards) seepage flux. Seepage fluxes derived when using the Flexiflask hydration bladder were consistently lower (0.06-0.43 L/d/m3) than those relating to the 4 L wine bladder (1.4-3.0 L/d/m3). This was attributed to the relative rigidity of the Flexiflask which would have provided resistance to any filling of the bag in a gaining situation. In comparison, the wine bladder is a lot more flexible and compliant. The chamber with the small-gauge gas venting tube (Chamber 1) measured consistently lower fluxes regardless of the collection bag used. This was attributed to the excessive accumulation of sediment gas in this chamber effecting the readings. This was highlighted after the final test (Test 6) where a large volume of gas was released from the chamber when it was removed from the drain bed. This did not occur when the other chamber was removed. Overall the seepage meter trial suggests that the Meerschaum Drain was receiving a low positive (gaining) flux with the shallow groundwater system. Due to the problems encountered with the use of seepage meters with either the Flexiflask bag or Chamber 1, with the small-gauge venting tube, Test 5 (with the combination of Chamber 2 and the wine bladder) is considered the most realistic. Assuming a drain width of 6m and homogenous conditions, groundwater inflow of about 18 m3/day for every kilometre of drain is estimated (Table 3.2). This is considered a relatively low seepage flux. However it is realistic considering the competent nature of the clay evident at the base of the drain at this site. Soil profile measurements and shallow piezometers suggest that inflow of groundwater into the drain is more from the shallow oxidised zone (<1m) which intercepts the side of the drain rather than the base. A more robust estimate of seepage flux would require measurements at both the sides and the base of the drain.

Table 3.2: Results of seepage meter trials in Meershaum Drain, Tuckean Swamp (Brodie et al, 2005)

Test Bag Chamber Test Duration

(hrs)

Difference (mL)

Seepage Flux

(L/d/m2)

Seepage Flux (m3/d/km)

1 caribee 1 26.0 -58 -0.19 -1.1A

2 flexiflask 2 26.0 131 0.43 2.6

3 flexiflask 2 24.08 59 0.21 1.2

4 wine 1 23.92 394 1.4 8.4

5 wine 2 53.5 1878 3.0 17.9

6 flexiflask 1 53.5 39 0.06 0.4 B A lifted bag before removing - likely to have lost water B lots of gas discharged from chamber when removed


4. Field Observations

Visual indications of the interaction of groundwater and surface water systems can be observed in certain catchments and settings. A field reconnaissance survey is useful in the initial stages of an assessment to identify specific locations (hotspots) that warrant further investigation involving more detailed monitoring and sampling. The survey can also provide guidance to the parameters that could be measured to help quantify connectivity and also to identify the management issues impacted by the connectivity. There is a range of field indicators of direct groundwater discharge into streams, lakes or estuaries: (i.) The direct observation of water flow from seepages and springs at the margins

or within the bed of the surface water feature. Underwater discharge of groundwater can be observed if the flow rates are sufficiently high;

(ii.) In colder times of the year, the water vapour above discharge zones may be observed due to the contrast between the groundwater and air temperatures. Likewise, in alpine areas during winter, seepage areas can remain continually ice or snow-free in contrast with the surrounding landscape;

(iii.) Changes in the groundwater chemistry due to mixing with surface water can result in mineral precipitates such as iron and manganese oxides. These commonly form with the contact of anoxic groundwater with oxygenated surface water. Iron bacteria that oxidise the dissolved ferrous form (Fe2+) to the ferric form (Fe3+) can also occur as filaments and accumulations. This can be accompanied by an oily sheen on the water surface, similar in appearance to a petrol film. A bacterial origin (rather than a petrol spill) is inferred if the film breaks up into clusters rather than swirling together, when a small stick is trailed through it (NCDWQ, 2004);

(iv.) Water colour and odour can be an indicator, particularly if the groundwater is contaminated. This may be the case in catchments with urban, industrial, mining or intensive agricultural development. Discharge of highly acidic groundwater can be indicated by a dramatic increase in the clarity of the surface water; due to the flocculating of clay particles by elevated levels of dissolved aluminium (refer Box 4.1);

(v.) Carbonate precipitates such as tufa or travertine deposits can indicate discharge of groundwater with high levels of dissolved carbon dioxide and calcium carbonate, notably in a karst landscape. These can form spectacular terraces, cascades and dams that can significantly modify stream morphology (refer Box 4.2).

4.1 Advantages and Disadvantages Identifying and mapping field indicators of groundwater discharge can be readily incorporated into a fieldwork programme with little additional costs. However, there is a high dependence on the observer’s knowledge of visual indicators. Field indicators can show where groundwater seepage occurs but are limited in quantifying seepage flux.

4.2 Data Availability Databases specific to field observations are not routinely maintained. However, field observations relating to the collection of water samples may be relevant. Relevant water quality databases are outlined in Brodie et al. (2007).


Box 4.1 Field Indicators of shallow acid groundwater discharge, Tuckean Swamp, North Coast NSW The Tuckean Swamp is a large estuarine back swamp on the north coast of New South Wales, that has been highly modified with construction of drains and a tidal barrage to manage frequent flooding. The drainage network is efficient in reducing the incidence of flooding and waterlogging, but it also tends to lower the shallow watertable. The swamp contains significant acid sulfate soils and the watertable decline has caused the oxidation of pyrite within the previously waterlogged shallow estuarine sediments, a chemical reaction that generates sulfuric acid. Following major rainfall events, the store of acid migrates into the drains and is exported into the estuary. The consequences of this acidity are fish kills, poor water quality, land degradation, reduced agricultural productivity, loss of estuarine fisheries habitat, and degraded vegetation and wildlife values (Hagley, 1996).

The seepage of shallow acid groundwater into the Tuckean Swamp drains is indicated in the field by:

(i.) Unusually clear water, due to dissolved aluminium in the acid water flocculating the clay particles;

(ii.) Extensive deposition of yellow-brown iron precipitates on the drain bed and on vegetation;

(iii.) The establishment of acid tolerant plants such as lilies.

Changes to these characteristics can occur over small reaches of the drain and associated with rapid decreases in the acidity of the drain water (Figure 4.1)

Figure 4.1: Field indicators of discharge of shallow acid groundwater into a coastal drainage network, Tuckean Swamp, northern NSW. Changes in drain appearance from pH ~ 6 to pH ~3 over drain length of about 100 m.


Box 4.2 Field indicators of groundwater discharge from a limestone aquifer, Barkly karst, North Queensland The Barkly karst is a large karstic terrain near the Queensland-Northern Territory border developed in the Middle Cambrian limestones and dolomites of the Georgina Basin. The perennial streams that drain the catchment are groundwater fed and contain deposits of travertine as terraces, cascades and dams. Indarri Falls on Lawn Hill Creek is considered to be the largest travertine dam in Australia with a height of over 13.5m (Drysdale and Gale, 1997). The regional groundwater has very high dissolved carbon dioxide and is close to calcite saturation, reflecting interaction with the limestones and dolomites. Following discharge into the streams, the carbon dioxide progressively degasses downstream driving the stream water to high levels of supersaturation in terms of dissolved calcium carbonate. This means that travertine deposition typically occurs downstream of the point of groundwater discharge, with deposition occurring for significant distances subsequent. The rate at which this process occurs is a function of stream discharge and gradient. High stream gradients cause greater stream turbulence which enhances degassing of carbon dioxide. Investigations of the Barkly travertine-depositing streams suggest that the length of stream reach required for waters to become supersaturated increases with stream discharge, whilst the length of reach where subsequent deposition occurs also increases with stream discharge (Drysdale et al. 2002).

Figure 4.2: Lawn Hill Creek at Lawn Hill (www.bom.gov.au)



5. Ecological Indicators Certain plants and animals can be used to identify the nature and extent of groundwater-surface water interaction. The most common ecological indicators are aquatic plants, phreatophytes and hyporheic biota. Aquatic plants can be indicators of groundwater discharge. For example, cattail plants have been used as indicators of fresh groundwater input to saline prairie lakes in North Dakota (Swanson et al, 1984). Lodge et al. (1989) indicated that submerged aquatic plant biomass was enhanced where groundwater inflow velocity was greater. In a desert stream setting, algae abundances in streambed sediments recovered more rapidly from flash flooding in areas with upward movement of groundwater (Valett et al, 1994). This was attributed to the groundwater providing a source of nutrients to the algae. Mats of green algae and stalked diatoms also coincide with discharge of nutrient-rich groundwater in the generally unproductive Flathead River in Montana (Ward, 1989). Klijn and Witte (1999) discussed the relationship of plants to groundwater flow systems. Groundwater discharge may favour the growth of particular aquatic plants. The predominance of acid-tolerant species such as water lilies can indicate significant discharge of acidic groundwater. Phreatophytes are deep-rooted plants that can access the watertable. Upland phreatophytic plants near a surface water body can indicate the presence of groundwater at shallow depths. The hyporheic zone is the portion of the stream bed characterised by mixing of groundwater and surface water. It can be a zone of intense biogeochemical activity and assessing the extent and composition of hyporheic biota can indicate the nature of groundwater-surface water interaction. This includes mapping the distribution of organisms that only temporarily migrate into the sediment from the stream floor such as insect larvae, compared with organisms that are permanent inhabitants of the substrate, such as choronimids and amphipods (Ward and Stanford, 1989). For example, the oligchaete worm Phallodrilus sp. can be a useful indicator of groundwater exchange (Lafont et al. 1992). Ostracods are especially useful in assessing past groundwater and surface water relationships because they have specific tolerances to water temperature and chemistry. 5.1 Advantages and Disadvantages Ecological indicators can be readily combined at a basic level with other fieldwork with little additional costs, and included in the reconnaissance of sites of groundwater discharge. Monitoring of ecological indicators may be useful in understanding seasonal changes in seepage flux. Such monitoring is important in the evaluation and management of groundwater dependent ecosystems. More detailed studies, such as hyporheic zone investigations require more time, resources and expertise. The method depends on the observer’s expertise in biological identification. Ecological indicators tend to show where groundwater seepage occurs but are limited in providing quantitative information on seepage flux. 5.2 Data Availability Databases are available on the national assessment of river condition, surveys of wetlands and other environmental assets, and collections relating to specific genera (Brodie et al, 2007).


6. Hydrogeological Mapping Knowledge of the hydrogeological setting is critical in understanding groundwater-surface water interactions. This involves mapping the configuration and characteristics of groundwater flow systems within the catchment, covering aspects such as aquifer geometry, geological and stratigraphic configurations, and hydraulic properties such as transmissivity and storativity, and recharge and discharge mechanisms. A hierarchy of groundwater flow systems of different size and depth can develop in a catchment depending on the combination of surface topography, geology and climate (Figure 6.1). There can be a combination of groundwater flow systems that are local, intermediate or regional in scale (Toth, 1963). Local flow systems are the shallowest and most dynamic, involving short flow paths (mostly < 5 km) with groundwater discharging to the nearest lowland feature. In contrast, regional flow systems have the deepest and longest flow paths (typically exceeding 50 km), with intermediate systems operating between these two end-members. Local flow systems tend to be dominant in areas of high topographic relief, while intermediate-regional systems are more evident in flat-lying areas. Groundwater exchange with surface water features are primarily governed by their location with respect to groundwater flow systems, the geological characteristics of their beds and climatic factors (Winter, 1999). River reaches can receive contributions of groundwater from flow systems of different scales and provenance (Figure 6.2). Figure 6.1: Different scale groundwater flow systems within a catchment (Winter et al, 1998, after Toth, 1963)


Figure 6.2: Groundwater flow systems operating within an alluvial riverine valley (Winter et al, 1998) Groundwater flow system mapping is available at a national level as well as at the catchment level. The National Land and Water Resources Audit (http://www.nlwra.gov.au) provides access to a national coverage of groundwater flow system mapping based on recharge and flow behaviour and using a combination of geology, geological and topographical criteria (Coram et al. 2000). Local, intermediate and regional groundwater flow systems are identified across a range of geological and geomorphological terrains. The focus of the mapping was to predict how groundwater systems respond to changing recharge and for defining appropriate dryland salinity management options. Such mapping is useful in defining the provenance and time-scale of groundwater movement, but is not specific in terms of groundwater-surface water interaction. Traditional hydrogeological maps are available for certain groundwater management regions across Australia and are published at 1:250,000 scale or more detailed, refer Figure 6.4. These are the printed maps that depict various combinations of hydrogeological parameters such as groundwater availability, salinity, potentials and aquifer structure. Such information is useful as context when evaluating the extent and direction of groundwater-surface water exchange. However, this perspective of the extent of hydrogeological mapping across Australia is by no means complete. The published hydrogeological map is only a small subset of a vast repository of mapping that has been undertaken, with mapping embedded in journals, reports, unpublished consultancies, research theses or management plans.


Box 6.1 Hydrogeological mapping of the Alstonville Plateau, NSW to help understand stream-aquifer connectivity By compiling a series of cross sections interpreted from borehole data such as lithological logs, water cut intervals, pump tests and water levels, the general hydrogeological features of the Alstonville Plateau could be interpreted (Brodie and Green, 2002). Figure 6.3 is a generalised cross section of the Tertiary Lismore Basalt sequence that makes up the plateau. This conceptualisation is also based on the facies model characteristic of continental basaltic successions (Cas and Wright, 1995). Two main groundwater systems can be found in the basalt plateau:

(i) A local-scale unconfined groundwater flow system operating in the shallow profile of soil and weathered or highly fractured basalt. Being a weathered mantle, the aquifer geometry is largely controlled by topography. Shallow groundwater flow is also a function of topography and dominated by local-scale flow systems having flowpaths typically less than 5 km. This means that groundwater divides for this shallow aquifer tend to correspond with surface water divides. Groundwater recharged from the hills or ridges, flows down-slope largely constrained by the contact between the weathered material and relatively fresh and unfractured basalt. Groundwater discharge occurs as springs or seepage areas lower in the valley floor. Mid-slope springs can occur where structural benches of unweathered or unfractured basalt impede this down-slope movement.

(ii) A deeper intermediate-scale groundwater flow system operating in interlayered and fractured horizons within the basaltic sequence. These aquifers can range from being semi-confined to confined in nature. Vertical columnar jointing or fracturing can provide, in part, a degree of interconnection between permeable horizons. Deeper aquifers are found in the buried weathered horizons, the vesicular and highly fractured components of basalt flows and interbedded fluvial deposits. These aquifers can be separated and confined by relatively thick sequences of massive, poorly fractured basalt. The deeper groundwater flow in the Lismore Basalt is largely controlled by the dip of the volcanic sequence. The flow path length is determined by the position of the particular aquifer within the basalt sequence and the level of dissection. As indicated in Figure 6.4, the deeper groundwater flow can cross local surface water divides with discharge into neighbouring streams or along the plateau escarpment.

Figure 6.3: Schematic cross section of the hydrogeology of the Alstonville Plateau (Brodie and Green, 2002)


Figure 6.4: Extent of traditional published hydrogeological maps at 1:250,000 scale or more detailed (Brodie, 2002) In the absence of specific hydrogeological mapping, more generic geological information is useful at the stream-reach level as well as the catchment level. Identification of geological structures such as faults or basement highs that control the geometry and hydraulic properties of aquifers is particularly important. Stratigraphic information such as the distribution of low-permeability clay layers or paleochannels that act as preferential pathways is used to map variability of stream-aquifer connectivity. For example, connectivity in the Cudgegong Valley (near Mudgee, NSW) is controlled by geological features (Hamilton 2004). Bedrock constrictions or faulting restricts groundwater throughflow in the alluvial aquifer resulting in shallower watertables and gaining conditions in the river. Away from these geological features, the regulated Cudgegong River is largely a losing stream. Box 6.1 gives the example of how hydrogeological mapping (by constructing a series of cross sections using the available borehole data) was used to develop an understanding of how groundwater systems feed the streams on the Alstonville Plateau (Brodie and Green, 2002). Mapping of geomorphological features have also been used to characterise connectivity. Analysis of the geomorphology of some North American alluvial aquifers inferred a relationship between dominant groundwater direction and parameters such as channel slope, sinuosity, incision and channel width-to-depth ratio (Larkin and Sharp, 1992). Groundwater was dominantly lateral (underflow conditions) in alluvial systems with large channel gradients, small sinuosities, large width-to-depth ratios and low river incisions, with the alluvial systems with opposing attributes dominated by vertical groundwater flow (gaining or losing conditions).


6.1 Advantages and Disadvantages An understanding of the hydrogeological setting is critical in defining connectivity between streams and aquifers. Hydrogeological mapping is an important part of developing a conceptual model showing the broader perspective of the nature and configuration of groundwater systems, the scale and direction of groundwater flow, geomorphological features and the hydraulic properties of aquifers. However, compiling and interpreting hydrogeological data can be time consuming and complex. It can involve interpolation of limited data from a sparse network of bores, so is subject to misinterpretation. A working knowledge of hydrogeological principles is required. 6.2 Data Availability A listing of available hydrogeological maps is provided in Appendix 16.1.



7. Geophysics and Remote Sensing Geophysical or remote sensing imagery can be integrated with other datasets to assist in the interpretation of groundwater–surface water interactions. Geophysical surveys can provide mapping of the spatial (and temporal) variation in important properties such as groundwater chemistry (particularly salinity), soil moisture content and soil/sediment texture. Such surveys can map the geological and geomorphological features that control stream-aquifer connectivity. They can also be used to map indicators of groundwater discharge (such as vegetation types, waterlogged or saline areas) in the landscape. Geophysical and remote sensing has been used to identify hotspots in terms of saline groundwater ingress into streams, pollution plumes or recharge of fresh stream water into alluvial aquifers. Geophysical data collection can be undertaken at different scales using different platforms, including satellite, aircraft, vehicles or boats. Downhole geophysical surveys can also be undertaken on existing bores (Table 7.1). Table 7.1: Common geophysical tools used in borehole logging

Log Parameters Measured Applications Caliper Borehole or casing diameter. Fracture identification, lithologic changes, and well construction. Natural Gamma Natural gamma radioactivity. Lithology and estimation of clay content in overburden. Fluid Temperature Temperature of borehole fluid. Indicates geothermal gradient, and water flow in borehole or between

borehole and fractures. Fluid Resistivity Resistivity of borehole fluid. Indicates water flow within borehole, or between borehole and fractures;

and water quality. Single Point Resistance

Resistance of materials between probe and ground surface electrode.

Lithology, fracture identification, and location of well screens.

Normal Resistivity Apparent resistivity of material.

Lithology and water quality.

Spontaneous Potential (SP)

Electrical potentials between probe and surface electrodes.

Lithology, water quality, and in some cases, fractures in resistive crystalline rock.

EM Conductivity (Induction)

Electrical conductivity in medium surrounding borehole.

Location of contaminant plumes, conductive clay units, or bedrock fractures. Monitor water quality changes over time.

Flowmeter Continuous or point measurements of water flow in borehole.

Identification of permeable zones and apparent vertical hydraulic conductivity and flow direction.

Borehole Video Provides visual record of lithology, fractures, well construction.

Lithologic logging; identification of fractures; examination of casing or well construction.

Acoustic Televiewer

Provides acoustically-generated image of boring walls.

Structural logging; identification and orientation of fractures and foliation; examination of casing or well construction.

Optical Televiewer Provides optically-generated image of boring walls.

Lithologic & structural logging; identification and orientation of structure & lithologic changes; examination of casing or well construction.

(Source: http://www.negeophysical.com/#bore) Various techniques are available including: (i.) Resistivity/Electrical Conductivity, using an array of transmitters to introduce

an electrical current and receiver electrodes to measure subsequent voltage differences. Distortion of the electrical field by conductivity variations due to salinity, texture or moisture content can be imaged vertically at various depths. Ground-based surveys can be undertaken as transects parallel or orthogonal to the surface water feature. Also the geo-electric array, either submersible or floating, can be towed behind boats along rivers and irrigation channels to map


the electrical conductivity of the water column and underlying sediments (Allen and Merrick, 2004), refer Box 7.1. The technique has been particularly effective in mapping seepage of highly saline groundwater and evaluating the effectiveness of salt interception schemes;

(ii.) Electromagnetics, measuring the secondary magnetic field from induced electrical currents. Various instruments are routinely used in ground-based surveys to map soil conductivity to various depths, refer Table7.2. Such surveys (with the instrument either manually carried or vehicle mounted) have been used to map seepage from irrigation channels into surrounding sediments (ANCID, 2000), or to identify shallow saltwater intrusion in coastal settings. Fixed wing or helicopter based electromagnetic surveys can also be undertaken at the catchment scale, refer Box 7.2. These surveys can map the bulk electrical conductivity of the geological material down to depths exceeding 100 metres, and are useful for providing the overall geological framework, for identifying losing stream reaches that recharge aquifers and the distribution of salt stores relative to the surface drainage network;

(iii.) Magnetics, measuring the magnetic field to define geological boundaries or geological structures that can define or constrict groundwater flow, refer Box 7.3;

(iv.) Seismics, where the reflections or refractions of seismic waves from a simple energy source at the surface (such as sledgehammer, explosive, vibrating plate) are detected to map stratigraphy and geological structures, refer Figure 7.1. Ground-based seismic traverses can be useful in mapping features that constrain or control groundwater flow, such as the geometry and stratigraphy of alluvial aquifers, the bedrock structures and preferred pathways such as palaeochannels;

(v.) Radar, measuring the reflectance of transmitted microwaves to interpret moisture content and chemical composition of the shallow soil profile. Ground penetrating radar involves measuring the reflectances from high frequency pulses to map near-surface geological features;

(vi.) Gamma-ray spectrometry (or radiometrics), where gamma radiation emissions are used to derive the concentration of thorium, uranium and potassium within the shallow soil profile. A spectrometer is used to count the number of gamma rays across multiple bands within the energy spectrum, with peaks in particular bands attributed to each of the three radioactive isotopes considered. The technique is used to provide detailed information about the characteristics of the soil and its parent geological material, including surface texture, weathering, leaching, soil depth and clay mineralogy (Bierwirth, 1997). This is useful in mapping landforms such as near-surface palaeochannels or estimate soil hydraulic properties. The method is typically incorporated in airborne geophysical surveys, but portable spectrometers can also be used in ground-based surveys;

(vii.) Multi-spectral imagery, such as Landsat TM imagery that can be processed and classified to map landscape indicators such as moisture content, vegetation type and stress, land use and terrain. These can be integrated with other data to identify seepage areas. The Near Infrared imagery is particularly useful as water is a strong absorber in this part of the spectrum. Hyperspectral imagery involves greater partitioning of the electromagnetic spectrum and has the potential to map more specific soil or land cover properties.


Table 7.2: Different Ground-based electromagnetic techniques

Technique Signal Penetration Depth (m)

EM38 1-1.5

EM31 6

EM34 7.5-30 (depending on coil separation) (Source http://www.stratscan.co.uk/em31.html)

Figure 7.1: A seismic gun used to fire a shotgun charge into the ground to generate a shockwave.

7.1 Advantages and Disadvantages Geophysics and remote sensing provide opportunities for the rapid, non-invasive mapping of landscape parameters that either indicate or control groundwater-surface water interactions (such as groundwater salinity or aquifer texture). Surveys can provide good spatial resolution in the vicinity of surface water features, while multiple surveys can provide information in terms of changes through time. Various platforms from satellite and aircraft to vehicles, boats and downhole loggers can be used to collect data of various resolution and geometry. However, undertaking and interpreting these surveys can be complex, requiring specific equipment, technical expertise and logistical support. The equipment used can be expensive to purchase or hire and can require ongoing maintenance and calibration. The data processing following field collection to remove measurement artefacts or to derive mapped outputs can be complex and may require extensive calibration with other datasets, such as borehole logs and chemical analyses. The cost of commercially available data (such as satellite imagery) can vary significantly. Ground or water-based surveys can have significant logistical problems encountering obstacles such as fallen trees, snags, fences, rough terrain, shallow water levels, thick vegetation or boggy conditions.

7.2 Data Availability There is the opportunity to use existing geophysical datasets generated for geological mapping and mineral exploration for catchment mapping. The Geoscience Portal (http://www.geoscience.gov.au) provides information on historical geophysical surveys as well as other geoscience datasets. The State geological surveys also undertake regional surveys as part of their mapping initiatives.


Box 7.1 Application of electrical conductivity imaging of streams in the Border Rivers catchment Rapid imaging of groundwater salinity and sediment texture beneath the Dumaresq and MacIntyre River at different river reaches was undertaken using geophysical survey as per the method of Allen and Merrick (2004), refer figure 7.2. Water quality in the area downstream of Keetah may suffer from inflow of saline groundwater. The rivers flow across a vast plain of self mulching clay soil of low permeability and moderate salinity. Where the groundwater is in contact with rivers, interaction occurs.

Figure 7.2: A 144 m long floating electric Water course to measure groundwater and streambed electrical conductivity (Allen, 2005)

Figure 7.3: EC ribbon images from geo-electric surveys – Dumaresq River at Glenarbon (Allen, 2005)

Figure 7.4: EC ribbon images from geo-electric surveys

(a) MacIntyre River at Mungindi and (b) South Callandoon (Allen, 2005)

The geophysical survey was undertaken in the Dumaresq River at Glenarbon to investigate river-aquifer interaction. The survey results show that at Glenarbon, upstream of Keetah, groundwater had very low salinity (EC of 100-300 µS/cm) as indicated by blue in the imagery (Figure 7.3). At least 20 m of aquifer filled with fresh water is evident under the river. The results indicate that river recharge is a significant source of recharge to the underlying aquifer in this river reach.

A steady increase in groundwater salinity beneath the rivers can be observed in a downstream direction, particularly below Goondiwindi. This is indicated by the green and red in the imagery (Figure 7.4). The MacIntyre/Barwon River appears to be underlain by uniform sediment in comparison with other MDB rivers, with only minor conductivity variation evident. Most transverse variation is evident in the South Callandoon image where the river intersects, at numerous locations, the sides of a swath of recent river palaeochannels which are recessed into the broad scale floodplain. However, for the MacIntyre/Barwon river reach at Mungindi, salinity increases and the geophysical signature is quite uniform in the sediment.

(a) (b)


Box 7.2 Airborne eletromagnetics (AEM) survey in the Lower Balonne Catchment, Queensland An airborne electromagnetics survey was flown in the Lower Balonne area in Queensland to investigate salt stores in the catchment. The survey maps the electrical conductivity of the geological material with groundwater salinity one of the major contributors to the conductivity response. Saline water is more conductive when compared to fresh water. The geophysics is calibrated using field data from boreholes. Figure 7.5 shows an output from the survey showing a low-conductivity (blue) signal in the northern section of the survey area. This is interpreted as a fresh groundwater plume, caused by leakage from the Maranoa River.

Figure 7.5: Airborne electromagnetics image showing groundwater recharge zone.

Box 7.3 Airborne magnetics survey in the Honeysuckle Creek catchment, Victoria Airborne magnetics have been successfully employed in the Honeysuckle Creek catchment of southern Australia to map buried paleochannels filled with iron-rich gravels (Figure 7.6). Subsequent drilling confirmed these channels were pathways of preferred groundwater flow and were carrying saline groundwater. Where these intersect and are hydraulically connected with the modern stream network is important for salinity management.

Figure 7.6: Paleochannels detected from airborne magnetics survey in Honeysuckle Creek subcatchment



8. Hydrographic Analysis A hydrograph is the time-series record of water level, water flow or other hydraulic properties, and can be analysed to gain insights into the relationships between rivers and aquifers. Typically, a stream hydrograph shows the fluctuations in stream flow through time and is a commonly available dataset routinely measured to support the management of water resources. For a gaining stream, where groundwater is contributing to stream flow, analysis of the stream hydrograph can indicate the magnitude and timing of this contribution. Hydrograph separation is probably the most widely used technique to solve many hydrological problems. This method aims to separate the observed hydrograph into two components: (i.) Quickflow –the direct response to a rainfall event including overland flow

(runoff), lateral movement in the soil profile (interflow) and direct rainfall onto the stream surface (direct precipitation),

(ii.) Baseflow–the longer-term discharge derived from natural storages, mostly assumed to be groundwater discharge from the shallow unconfined aquifer (refer Box 8.1).

Analysing the stream hydrograph to separate out the baseflow component provides information on the characteristics of the natural storages feeding the stream. Groundwater discharge from the shallow unconfined aquifer is commonly assumed to be the main contributor to baseflow. For this to be a significant process, the unconfined aquifer needs to be adequately replenished (typically on a seasonal basis), have a shallow watertable that is higher than the stream water level, and have adequate water storage and transmission properties to maintain flow to the stream (Smakhtin, 2001). For a gaining stream, where the underlying aquifer satisfies this criteria and groundwater contributes to stream flow, analysis of the stream hydrograph can indicate the magnitude and timing of this contribution. However, in certain catchments baseflow may not be dominated by groundwater discharge from the shallow unconfined aquifer. Other storages such as connected lakes or wetlands, snow, glaciers, caverns in karst terrains, or temporary storage within the river bank following the passage of high-flow events (bank storage) can also contribute to the baseflow regime of a stream (Griffiths and Clausen, 1997). Another complication is that baseflow is also influenced by any water losses from the stream. The hydrographic record essentially represents the net balance between gains to and losses from the stream. These losses include direct evaporation from the stream channel or from any connected surface water features such as lakes and wetlands, transpiration from riparian vegetation, evapotranspiration from source groundwater seepages, leakage to the underlying aquifer, or rewetting of stream bank and alluvial deposits (Smakhtin, 2001). These processes are often aggregated into a transmission loss for the reach of the stream.


Box 8.1 The baseflow component of streamflow The relative contribution of quickflow and baseflow components changes through the stream hydrographic record. The flood or storm hydrograph is the classic response to a rainfall event and consists of three main stages (Figure 8.1): (i.) Prior low-flow conditions in the stream consisting entirely of baseflow at the end of a

dry period; (ii.) With rainfall, an increase in streamflow with input of quickflow dominated by runoff

and interflow. This initiates the rising limb towards the crest of the flood hydrograph. The rapid rise of the stream level relative to surrounding groundwater levels reduces or can even reverse the hydraulic gradient towards the stream. This is expressed as a reduction in the baseflow component at this stage;

(iii.) The quickflow component passes, expressed by the falling limb of the flood hydrograph. With declining stream levels timed with the delayed response of a rising watertable from infiltrating rainfall, the hydraulic gradient towards the stream increases. At this time, the baseflow component starts to increase. At some point along the falling limb, quickflow ceases and streamflow is again entirely baseflow. Over time, baseflow declines as natural storages are gradually drained during the dry period up until the next significant rainfall event.

Figure 8.1: Components of a typical flood hydrograph

Also, water use or management activities can significantly affect the baseflow regime. Many streams have highly modified flows due to the development and use of water resources. Over-extraction can mean that streams that were naturally perennial due to prolonged baseflow can become intermittent. Major regulated systems such as the River Murray have artificially high flows during the summer due to releases to supply irrigation and urban users. Specific activities that can influence baseflow include: (i.) Stream regulation where flow is controlled by infrastructure such as dams,

locks or weirs. Releases from surface water storages for downstream users can make up the bulk of streamflow during dry periods. Baseflow analysis should be undertaken in unregulated reaches, or at least the regulated catchment area should be no more than 10% of the catchment area of the streamflow gauge (Neal et al. 2004);


(ii.) Direct pumping of water from the stream for consumptive uses such as irrigation, urban supply or industry;

(iii.) Artificial diversion of water into or out of the stream as part of inter-basin transfer schemes;

(iv.) Direct discharges into the stream, such as from sewage treatment plants, industrial outfalls or mine dewatering activities;

(v.) Seasonal return flows from drainage of irrigation areas; (vi.) Artificial drainage of the floodplain, typically for agricultural or urban

development, which can enhance rapid runoff and reduce delayed drainage; (vii.) Changes in land use, such as clearing, reafforestation or changes in crop type,

which can significantly alter evapotranspiration rates; (viii.) Groundwater extraction, sufficient to lower the watertable and decrease or

reverse the hydraulic gradient towards the stream. Careful consideration of the overall water budget and management regime for the stream is required before the assumption that baseflow equates to groundwater discharge can be made. Analysing the baseflow component of the stream hydrograph has had a long history of development since the early theoretical and empirical work of Boussinesq (1904), Maillet (1905) and Horton (1933). Several useful reviews have been written including Hall (1968), Nathan and McMahon (1990), Tallaksen (1995) and Smakhtin (2001) to map this development. The multitude of methods that have evolved can be conveniently categorised into three basic approaches of baseflow separation, frequency analysis and recession analysis. 8.1 Baseflow Separation Baseflow separation techniques use the time-series record of stream flow to derive the baseflow signature. The common separation methods are either graphical which tend to focus on defining the points where baseflow intersects the rising and falling limbs of the quickflow response, or involve filtering where data processing of the entire stream hydrograph derives a baseflow hydrograph. 8.2 Graphical Separation Methods Graphical methods are commonly used to plot the baseflow component of a flood hydrograph event, including the point where the baseflow intersects the falling limb (Figure 8.2). Stream flow subsequent to this point is assumed to be entirely baseflow, until the start of the hydrographic response to the next significant rainfall event. These graphical approaches to partitioning baseflow vary in complexity and include: (i.) An empirical relationship for estimating the point along the falling limb where

quickflow has ceased and all of the stream flow is baseflow,

D = 0.827A0.2 (Equation 8.1)

where D is the number of days between the storm crest and the end of quickflow, and A is the area of the catchment in square kilometres (Linsley et al. 1975). The value of


the exponential constant (0.2) can vary depending on catchment characteristics such as slope, vegetation and geology; (ii.) The constant discharge method assumes that baseflow is constant during the

storm hydrograph (Linsley et al, 1958). The minimum streamflow immediately prior to the rising limb is used as the constant value;

(iii.) The constant slope method connects the start of the rising limb with the inflection point on the receeding limb. This assumes an instant response in baseflow to the rainfall event;

(iv.) The concave method attempts to represent the assumed initial decrease in baseflow during the climbing limb by projecting the declining hydrographic trend evident prior to the rainfall event to directly under the crest of the flood hydrograph (Linsley et al, 1958). This minima is then connected to the inflection point on the receeding limb of storm hydrograph to model the delayed increase in baseflow;

(v.) Using the trends of the falling limbs before and after the storm hydrograph to set the bounding limits for the baseflow component (Frohlich et al, 1994);

(vi.) Use the Boussinesq equation as the basis for defining the point along the falling limb where all of the streamflow is baseflow (Szilagyi and Parlange, 1998).

Figure 8.2: Graphical baseflow separation techniques including (1a) constant discharge method (1b) constant slope method and (1c) concave method (Linsley et al. 1958) 8.3 Filtering Separation Methods The baseflow component of the streamflow time series can also be separated using data processing or filtering procedures. These methods tend not to have any hydrological basis but aim to generate an objective, repeatable and easily automated index that can be related to the baseflow response of a catchment (Nathan and McMahon, 1990). The baseflow index (BFI) or reliability index, which is the long-term ratio of baseflow to total streamflow, is commonly generated from this analysis. Other indices include the mean annual baseflow volume and the long-term average daily baseflow (Smakhtin, 2001). Examples of continuous hydrographic separation techniques based on processing or filtering the data record include:

1a

1b

1c


(i.) Increasing the base flow at each time step, either at a constant rate or varied by

a fraction of the runoff (Boughton, 1988); (ii.) The smoothed minima technique which uses the minima of 5 day non-

overlapping periods derived from the hydrograph. (Institute of Hydrology 1980; FREND, 1989). The baseflow hydrograph is generated by connecting a subset of points selected from this minima series. The HYSEP hydrograph separation programme (http://water.usgs.gov/software/hysep.html) uses a variant of this called the local-minimum method (Sloto and Crouse, 1996);

(iii.) The fixed interval method discretises the hydrographic record into increments of fixed time (Pettyjohn and Henning, 1979). The magnitude of the time interval used is calculated by doubling (and rounding up) the duration of quickflow calculated empirically from Equation 8.1. The baseflow component of each time increment is assigned the minimum streamflow recorded within the increment;

(iv.) The sliding-interval method assigns a baseflow to each daily record in the hydrograph based on the lowest discharge found within a fixed time period before and after that particular day (Pettyjohn and Henning, 1979);

(v.) Recursive digital filters, which are routine tools in signal analysis and processing, are used to remove the high-frequency quickflow signal to derive the low-frequency baseflow signal (Nathan and McMahon 1990). Table 8.1 outlines some of the digital filters that have been applied to smooth hydrographic data. Box 8.2 outlines an assessment of unregulated streams in the Murray Darling Basin using a recursive digital filter. Eckhardt (2005) has developed a general formulation that can devolve into several of the commonly used one-parameter filters:

max

max)1(max)( 1

)1()1(aBFI

qBFIaaqBFIq iib

ib −

−+−= − (Equation 8.2)

Where qb(i) is the baseflow at time step i, qb(i-1) is the baseflow at the previous time step i-1, qi is the stream flow at time step i, a is the recession constant and BFImax is the maximum value of the baseflow index that can be measured; (vi.) The streamflow partitioning method uses both the daily record of streamflow

and rainfall (Shirmohammadi et al. 1984). Baseflow equates to streamflow on a given day, if rainfall on that day and a set number of days previous, is less than a defined rainfall threshold value. Linear interpolation is used to separate the quickflow component during high rainfall events.


Table 8.1: Recursive digital filters used in base flow analysis (Grayson et al. 1996; Chapman, 1999; Furey and Gupta, 2001)

Filter Name Filter Equation Source Comments One-parameter algorithm

)()1()( 21

2 iibib qkkq

kkq

−−

+−

= − Chapman and Maxwell (1996)

qb(i) ≤q(i) Applied as a single pass through the data.

Boughton two-parameter algorithm

)()1()( 11 iibib qC

CqC

kq+

++

= − Boughton (1993) Chapman and Maxwell (1996)

qb(i) ≤q(i) Applied as a single pass through the data Allows calibration against other baseflow information such as tracers, by adjusting parameter C

IHACRES three-parameter algorithm )(

11 )1()()1()( −− ++

++

= iqiibib qqC

CqC

kq α Jakeman and Hornberger (1993)

Extension of Boughton two-parameter algorithm

Lyne and Hollick algorithm

21)( )1()()1()(

αα +−+= −− iiifif qqqq

Lyne and Hollick (1979) Nathan and McMahon, (1990)

qf(i) ≥0 α value of 0.925 recommended for daily stream data filter recommended to be applied in three passes Baseflow is qb = q - qf

Chapman algorithm )(

32

313

)1()()1()( −− −−

+−−

= iiifif qqqq ααα

α

Chapman (1991) Mau and Winter (1997)

Baseflow is qb = q - qf

Furey and Gupta filter )()1( )1()1(

1

3)1()( −−−−− −+−= dibdiibib qq

ccqq γγ

Furey and Gupta (2001) Physically-based filter using mass balance equation for baseflow through a hillside

q(i) is the original streamflow for the ith sampling instant qb(i) is the filtered baseflow response for the ith sampling instant qf(i) is the filtered quickflow for the ith sampling instant q(i-1) is the original streamflow for the previous sampling instant to i qb(i-1) is the filtered baseflow response for the previous sampling instant to i qf(i-1) is the filtered quickflow for the previous sampling instant to i k is the filter parameter given by the recession constant α, αq are filter parameters C is a parameter that allows the shape of the separation to be altered γ, c1, c3 are physically based parameters 8.4 Frequency Analysis Methods Frequency analysis takes a different approach in characterising baseflow by deriving the relationship between the magnitude and frequency of streamflow discharges from the hydrographic record. In its most common application, a flow duration curve (FDC) is generated. Instead of plotting as a time series, a flow duration curve shows the percentage of time that a given flow rate is equalled or exceeded. The FDC is constructed from flow data of fixed time period (eg daily, monthly, and annual) by: (i.) Sorting the flow data in order of decreasing flow;

(ii.) Assigning a unique ranking number m to each flow, starting with 1 for the maximum flow to n for the minimum flow, where n is the number of flow measurements;

(iii.) The probability P that a given flow will be equalled or exceeded is defined by:

1100

+=

nmP (Equation 8.3)


(iv.) The flow-probability relationship is typically presented as a log-normal plot (Fetter, 1994).

Flow duration curves can be constructed for the entire record of flow measurement, or for specific time periods such as similar calendar months or seasons. The FDC provides information on the baseflow component of stream flow. The median flow (Q50) is the discharge which is equalled or exceeded 50% of the time. The part of the curve with flows below the median flow represents low-flow conditions. Baseflow is interpreted to be significant if this part of the curve has a low slope, as this reflects continuous discharge to the stream. A steep slope for these low-flows suggests relatively small contributions from natural storages like groundwater (Figure 8.3). These streams may cease to flow for relatively long periods. In this way, the shape of the FDC can indicate the hydrogeological characteristics of a catchment (Smakhtin, 2001).

Figure 8.3: Flow distribution curves for examples of (2a) high baseflow and (2b) low baseflow streams Various indices are used to represent the characteristics of the low-flow regime for a stream. The ratio of the discharge which is equalled or exceeded 90% of the time, to that of 50% of the time (Q90/Q50) is commonly used to indicate the proportion of streamflow contributed from groundwater storage (Nathan and McMahon, 1990). Box 8.3 outlines a study where the Q90 percentile was used as an indicator of the importance of groundwater inputs. Other low-flow indices include: (i.) One or n-day discharges that are exceeded at defined percentages of time, say

75, 90 or 95% eg. Q75(7), Q75(10), Q95(10); (ii.) The percentage of time the stream is at zero-flow conditions;

(iii.) The longest recorded period of consecutive zero-flow days (Smakhtin, 2001). A Low-flow Frequency Curve (LFFC) shows the proportion of years when a low-flow rate is exceeded. This depicts the recurrence interval which is the average interval (in years) that the stream discharge falls below a given rate, and can also be used to represent baseflow conditions. The curve is generated from the series of annual minimum flow values extracted from the stream monitoring data. Like the flow duration curve, various indices can be used to indicate baseflow conditions including: (i.) The slope of the LFFC, as the larger the slope indicates more variability in

low-flows;

2a

2b


(ii.) Breaks in the curve near the modal value have been interpreted as representing when streamflow is exclusively from groundwater storage;

(iii.) Lowest average flows that occur over a set number of consecutive days (eg 3, 7 days) at defined recurrence intervals (eg 2, 10 years), for example 7day 10 year low flow (7Q10) or 7 day 2 year low flow (7Q2);

(iv.) The average of the annual series of minimum 7day average flows, MAM7 or also known as dry weather flow;

(v.) Indices of seasonal low flows such as mean 30 day summer low flows (Smakhtin, 2001).

8.5 Recession Analysis Methods The recession curve is the specific part of the flood hydrograph after the crest (and the rainfall event) where streamflow diminishes, refer Box 8.1. The slope of the recession curve flattens over time from its initial steepness as the quickflow component passes and baseflow becomes dominant. A recession period lasts until stream flow begins to increase again due to subsequent rainfall. Hence, recession curves are the parts of the hydrograph that are dominated by the release of water from natural storages, typically assumed to be groundwater discharge. Recession segments are selected from the hydrograph and can be individually or collectively analysed to gain an understanding of these discharge processes that make up baseflow. Graphical approaches have traditionally been taken but more recently analysis has focussed on defining an analytical solution or mathematical model that can adequately fit the recession segments. Each recession segment is often considered as a classic exponential decay function as applied in other fields such as heat flow, diffusion or radioactivity, and expressed as:

tt eQQ α−= 0 or cT

t

t eQQ−

= 0 (Equation 8.4) where Qt is the stream flow at time t, Q0 is the initial stream flow at the start of the recession segment, α is a constant also known as the cut-off frequency (fc) and Tc is the residence time or turnover time of the groundwater system defined as the ratio of storage to flow. The term e-α in this equation can be replaced by k, called the recession constant or depletion factor, which is commonly used as an indicator of the extent of baseflow (Nathan and McMahon, 1990). The typical ranges of daily recession constants for streamflow components, namely runoff (0.2-0.8), interflow (0.7-0.94) and groundwater flow (0.93-0.995) do overlap (Nathan and McMahon, 1990). However, high recession constants (eg > 0.9) tend to indicate dominance of baseflow in streamflow. Another parameter interpreted from the recession segment is the recession index (K) which is the time (in days) required for baseflow to recede by one log-cycle ie Q0 to 0.1Q0. A similar index called the half-flow period or half-life, which is the time (in days) for flow to halve, can also be calculated. For streams with low baseflow inputs the half-life may be in the range of 7-21 days, while discharge from large stable natural storages can result in a half-life exceeding 120 days (Smakhtin, 2001). The integrated form of the classic recession function of Equation 8.4 is


tt SQ α= (Equation 8.5)

where St is the storage in the reservoir that is discharging into the stream at time t. This relationship is called a linear storage-outflow model and implies that the recession will plot as a straight line on a semi-logarithmic scale. However, semi-logarithmic plots of individual recessions are commonly curved rather than linear. This is because other natural storages (eg bank storage, wetlands, deeper confined aquifers) can also contribute to baseflow, and these have different regimes of water release to the stream than that of the groundwater stored in the shallow aquifer (Sujono et al, 2004). The recession curve is effectively a composite of water discharged into the stream from multiple natural storages. This coincides with the concept that a catchment is a series of interconnected reservoirs (such as rainfall, snow, aquifers, soil, biomass etc), each having distinct characteristics in terms of recharge, storage and discharge (Smakhtin, 2001). A curved semi-logarithmic plot for recessions means that the storage-outflow relationship is non-linear. For groundwater discharge from a shallow unconfined aquifer, there are three main reasons for this non-linearity (Van de Griend et al, 2002): (i.) A falling watertable continually decreases the effective thickness of the

aquifer and decreases the ability to drain. Declining watertables can also be attributed to other processes other than stream discharge, such as evapotranspiration or groundwater extraction;

(ii.) The hydraulic conductivity tends to decrease with depth. This is attributed to increased compaction with depth in unconsolidated sediments, and decreased fracturing with depth in hard rock formations;

(iii.) With prolonged drainage, the lower order stream channels can run dry; leaving only the highest order reaches receiving baseflow.

Another complication is that the recession behaviour for a stream can change through time. This is reflected in variations in the shape of the recession segments found in a stream hydrograph. This is due to variability in such factors as the areal distribution of rainfall, residual storage in connected surface water bodies, catchment wetness, saturated aquifer thickness or depth of stream penetration into the aquifer. Baseflows are also influenced by seasonal effects such as variations in rainfall and evapotranspiration. High evapotranspiration rates during warm weather or active growing seasons can significantly reduce the baseflow component, particularly in shallow watertable areas. Different approaches have been used in recession analysis to address this non-linearity and variability in recession: (i.) Approximating the semi-logarithmic plot of the recession curve as three

straight lines of different slope (Barnes, 1940). The gradients of these three lines are inferred to be the recession constants for the main streamflow components of runoff, interflow and groundwater flow. The plotting of the three lines is difficult because of the gradual nature of the change in curvature in the recession;


(ii.) Plotting flow ratios (Q0/Qt) instead of flow (Qt) on the semi-logarithmic plot (Hino and Hasebe, 1984) to facilitate better interpretation of the recession;

(iii.) Using a double logarithmic plot of streamflow against time (Hewlett and Hibbert, 1963). Any abrupt change in slope is interpreted to mark the transition from quickflow to baseflow;

(iv.) The correlation method where the current flow Q is plotted on a natural scale against the flow (Qt) at some fixed time interval t previously (eg 2 days before) for each of the recession curves evident in the hydrograph (Langbein 1938). A line enveloping the traces of these multiple recessions is drawn through the origin to derive the master recession curve. By rearranging the exponential decay (Equation 8.4) the recession constant k can be derived from the slope of this master recession curve and the lag time interval t.

t

QQk /1

0

)(= (Equation 8.6)

(v.) The matching strip method involves plotting multiple recession curves derived from the hydrograph on the one semi-logarithmic plot in order of increasing minimum discharge (Toebes and Strang, 1964). Each recession curve is superimposed and adjusted horizontally to produce an overlapping sequence. The master recession curve is interpreted as the envelope to this sequence, and the recession constant k derived from its slope (Equation 8.6);

(vi.) The tabulation method where data from the multiple recession curves are used to derive the master recession curve and average discharges calculated for the period of the hydrographic record (Johnson and Dils, 1956). Recession periods are tabulated and sorted and mean discharges calculated for each timestep. This is either done computationally (Boughton, 1995) or by an analytical solution (Singh, 1989);

(vii.) The recession ratio method which analyses the ratios of current flow (Q) to the flow (Qt) at some fixed time interval t previously (eg 2 days before). A cumulative frequency diagram is plotted to estimate indices such as the median recession ratio (REC50) as a substitute for the recession constant, k (Smakhtin, 2001);

(viii.) The parameter averaging method where the recession function (Equation 8.4) is fitted for each of the recession segments in the hydrograph. The recession constants that are derived are then averaged (James and Thompson 1970);

(ix.) Wavelet transform analysis is a technique to break down a signal into its components and applied in such fields as image processing and geophysics. The technique can also be used in hydrograph recession analysis in terms of separating out the low frequency signature of the baseflow. Plots of frequency against time called mean-square wavelet maps are used to derive recession constants (Sujono et al. 2004);

(x.) Using different storage-outflow models or combinations of storage-outflow models to obtain a better fit to the recession curve. The classic exponential decay function (Equation 8.4) represents a linear relationship between storage and outflow. Other equations have been developed to model discharge from different types of natural storages (Table 8.2). By combining these equations, discharge from the various natural storages can be better accounted for. For example, a simple option is to add a constant (b) to the linear reservoir equation:


beQQ t += −α0 (Equation 8.7)

(xi.) This provides a better fit to recession curves that stabilise to a constant streamflow over time. This constant flow may represent discharge from large groundwater storage or from ice or snow reserves. A model based on combining two linear storages has also been used to provide a better fit to the recession curves for a small forested catchment (Moore 1997). These two storages were interpreted to represent different residence times for water in footslope and upslope zones in the catchment;

(xii.) The Meyboom method uses stream hydrograph data over two or more consecutive years (Meyboom, 1961). The baseflow is assumed to be entirely groundwater discharged from the unconfined aquifer. An annual recession is interpreted as the long-term decline during the dry season following the phase of rising streamflow during the wet season. The total potential groundwater discharge (Vtp) to the stream during this complete recession phase is derived as:

3.20KQVtp = (Equation 8.8)

where Q0 is the baseflow at the start of the recession and K is the recession index, the time for baseflow to decline from Q0 to 0.1Q0.;

(xiii.) The recession-curve-displacement method is based on the upward displacement of the recession curve during the rainfall event (Rorabaugh, 1964; Rutledge and Daniel, 1994; Rutledge, 1998). The method assumes that baseflow is entirely groundwater discharge from an unconfined aquifer of uniform thickness and hydraulic properties, with the stream fully penetrating the aquifer. On the basis of the algorithms developed, the total recharge to the groundwater system during the rainfall event has been shown to be about twice the total potential discharge to the stream at a critical time (Tc) after the hydrographic peak. Hence, the total volume of groundwater recharge due to the rainfall event (R) can be estimated from the stream hydrograph by:

3026.2)(2 12 KQQR −

= (Equation 8.9)

where Q1 is the baseflow at the critical time (Tc) extrapolated from the pre-event recession curve, Q2 is the baseflow at the critical time (Tc) extrapolated from the post-event recession curve, and K is the recession index (Figure 8.4).


Figure 8.4: Procedure for recession curve displacement method (after Rutledge and Daniel, 1993)

(1) Estimate the recession index (K) from the stream hydrograph record (2) Calculate the critical time (Tc), using the relationship KTc 2144.0= (3) Locate the time on the hydrograph which is Tc days after the peak, where streamflow recessions will be

extrapolated to (4) Extrapolate the pre-event recession curve to derive Q1 (5) Extrapolate the post-event curve to derive Q2 (6) Calculate total potential groundwater recharge using these parameters

8.6 Advantages and Disadvantages Baseflow analysis of stream hydrographs can provide valuable insights into how the groundwater contribution to stream flow changes through time. A distinct advantage of the approach is that it uses stream flow data that is regularly collected and placed in the public domain. This means that recession analysis can be readily undertaken as a desktop study prior to any detailed field investigations. However, baseflow analysis is only applicable for gaining stream conditions, although frequency analysis can provide insights into losing conditions. The assumption that baseflow is entirely groundwater discharge which may not always be valid, as water can be released to the stream from different storages such as connected lakes or wetlands over different timeframes. Baseflow can be affected by water use and management activities, so that the method cannot be applied in rivers that are regulated or have significant diversions or extractions. Careful consideration of the overall water budget and management regime for the stream is required. The hydrograph represents the water balance for the catchment above the stream gauge. This means that baseflow analysis can provide information on the temporal changes but not the spatial distribution of groundwater inputs along a stream between gauging stations.


Table 8.2: Different storage-outflow models used in recession analysis (Moore, 1997; Griffiths and Clausen, 1997; Dewandel et al. 2003)

Conceptual Model

Storage-Outflow Relation

Recession Function Storage Types

Source Comments

Linear Reservoir

kSQ = kteQQ −= 0 General storage Boussinesq

(1877) Maillet (1905)

Linearised Depuit-Boussinesq equation. Approximation for short time periods

Horton Double Exponential

mteQQ 20

α−= General storage Horton (1933) Transformation of

linear reservoir model

)1(00 ))1(1( nntnQQ −−+= α

Coutagne (1948)

nnc tnQQQ +−+−= − )1(

00 ))1(1( α

Karstic aquifers Padilla et al. (1994)

Qc is discharge from low-transmissivity components of karst

Channel Bank Storage

kteQ −=α Channel banks Cooper and

Rorabaugh, (1963)

Variant of linear reservoir. Also used to model evapotranspirative losses

Exponential Reservoir

SDBeQQ φ−= )1/( 00 tQQQ φ+= Throughflow in

soil hydraulic

conductivity assumed to exponentially decrease with depth

Power-law Reservoir

βαSQ = ptQQ )1(0 µ+=

)1/( ββ −=p βββ βαµ )1(

0/1 )1( −−= Q

Springs and unconfined aquifers (p = -2) Soil moisture

Hall (1968) Brutsaert and Nieber (1977)

Recessions modelled using p ~ 1.67 (Wittenberg 1994)

Depuit-Boussinesq Aquifer STorage

230 )1( −+= taQQ

Shallow unconfined aquifer

Boussinesq (1904)

Special case of power-law reservoir for Depuit-Boussinesq aquifer model

Depression Storage Detention Storage

321 )1/( tQ αα +=

Surface depressions such as lakes and wetlands, Overland flow

Griffiths and Clausen (1997)

variant of power-law reservoir

Two parallel Linear Reservoirs

2211 SkSkQ += tktk eQeQQ 2121

−− += Independent aquifers

Barnes (1939)

Two Serial Linear Reservoirs

22SkQ =

2212 1 SkeQ

dtdS tk −= −

)( 212

12

120

tktktk eekk

QkeQQ −−− −−

+=

Cavern Storage

tQ 21 αα −= Underground caverns in karst terrain

Griffiths and Clausen (1997)

Hyperbola Reservoir

btQ r += −1α

Ice melt, lakes Toebes and Strang (1964)

Constant Reservoir

α=Q Permanent snow and ice pack, large groundwater storages

Constant stream flow over a finite time period

Q - discharge S,S1,S2 – reservoir storages SD – catchment storage deficit t– time since beginning of recession Q0 – discharge for t = 0 QB, Q1, Q2, k, k1, k2, α,β,ϕ - parameters to be determined by calibration


8.7 Data Availability Hydrographic analysis has the advantage that it uses pre-existing data in the form of historic stream flow monitoring. In terms of data availability, Brodie et al, (2007) outlines the significant surface water monitoring databases in Australia, highlighting the State and Territory agencies involved in water management as the main data custodians. As a guide to the location of existing stream gauging sites, the Water Resources Station Catalogue (http://www.bom.gov.au/hydro/wrsc) developed by the Bureau of Meteorology provides a national inventory of the river gauging stations, as well as rainfall and evaporation stations, across the country. 8.8 Relevant Links ASTHyDA Project (http://www.geo.uio.no/drought/) – European project focussing on tools for assessing low surface water and groundwater flows Australian Hydrographers Association (http://www.aha.net.au/) BFI (http://www.usbr.gov/pmts/hydraulics_lab/twahl/bfi/index.html) US Bureau of Reclamation software for determining a Base Flow Index using a local minimum approach CRC Catchment Hydrology River Analysis Package including baseflow analysis HYSEP (http://water.usgs.gov/software/hysep.html) – USGS hydrographic separation program based on the fixed-interval, sliding-interval and local-minimum methods Low Flows 2000 (http://www.nwl.ac.uk/ih/www/products/iproducts.html)- Decision support tool for estimating low flows in the United Kingdom PART (http://water.usgs.gov/ogw/part/) - USGS program for estimating baseflows using a stream partitioning method PULSE http://water.usgs.gov/ogw/pulse/: - USGS analytical solutions to estimate groundwater discharge and baseflow of stream based on recession-curve-displacement method RAP (http://www.toolkit.net.au/cgi-bin/WebObjects/toolkit.woa/wa/productDetails?productID=1000002) – RECESS (http://water.usgs.gov/ogw/recess/) USGS method for analysing streamflow recession to determine the master recession curve RORA (http://water.usgs.gov/ogw/rora/): USGS program for estimating groundwater recharge using the recession-curve-displacement method


Box 8.2 Murray-Darling Basin Baseflow Survey A recursive digital filter was used to separate the baseflow component of hydrographs from 178 gauging sites across the Murray-Darling Basin (Neal et al, 2004). The filter applied was that used by Nathan and McMahon (1990) with a consistent filter parameter value of 0.925 adopted. Seasonal and annual baseflow indices were then calculated using the filtered baseflow. The criteria for selecting sites for analysis were that the flow was unregulated, the flow record covered at least the period 1990 to 1999 inclusive, and that no more than 5% of the raw data record was missing. This resulted in most of the stations analysed being in upland fractured rock catchments. Annual baseflow indices ranged between 0.04 and 0.76, with a mean of 0.34. The median baseflow indices tended to be higher for the Victorian sites (0.58), compared with those in New South Wales (0.23) and Queensland (0.11), refer Figure 8.5. This trend was inferred to be due to climatic effects, rather than hydrogeological differences.

Figure 8.5: Annual baseflow indices for unregulated streams in the Murray-Darling Basin (Neal et al. 2004)


Box 8.3: Analysis of stream flow percentiles in the Richmond catchment, NSW

The available stream flow monitoring in the Richmond River catchment was analysed to investigate differences in groundwater inputs to streams across the catchment. It was assumed that discharge from the shallow unconfined aquifer dominated low flow conditions. The long term Q90 percentile (the flow rate for which the stream exceeds for 90% of the time) was calculated for the gauges relating to unregulated subcatchments. The annual volume of Q90 was calculated by summing the 90th percentile of flow for each individual day. Values can depend on the length of the streamflow record, but the resulting Q90 streamflow volume was considered to be a semi-quantitative indicator of groundwater discharge for each subcatchment. Within the Richmond River Catchment, the groundwater component of baseflow (as indicated by the Q90 volume) varies depending on the aquifer system (Figure 8.6). Towards the north, the aquifers of the Alstonville Basalt and North Coast Fractured Rocks are composed of fractured basalts which are known to maintain perennial streams. In the southern part of the catchment, the geology consists of tight sandstones of the Clarence-Morton Basin. The difference in Q90 volume between the two geological systems is over two orders of magnitude. When normalised by area, the groundwater contribution for the northern basaltic subcatchments decreases westwards from 126 ML/d/km2 to 37 ML/d/km2. This probably relates to the inland decreasing trend in average rainfall. In contrast, the southern sedimentary subcatchments have much lower values of 0.05-0.6 ML/d/km2.

This simple analysis would be useful to gauge the importance of groundwater discharge to streams, particularly when assessing the sustainable yield of groundwater systems. Figure 8.6: Annual volume of Q90 percentile for available stream gauges in the Richmond River catchment


9. Hydrometric Investigations

Hydrometric methods are based on Darcy’s Law for fluid movement in a porous medium;

KdldhAQ = Equation 9.1

where Q is the flux of water (volume per unit time), A is the cross-sectional area of the porous medium through which flow occurs, dh/dl is the hydraulic gradient where dh is the change in hydraulic head along the distance dl of the groundwater flow line, and K is the hydraulic conductivity of the material. In its simplest form, Darcy’s Law can be applied to the situation where the flow between the stream and shallow aquifer are assumed to be entirely vertical (Figure 9.1). Hence, estimating the rate and direction of vertical seepage flux involves measuring: (i.) The head difference (dh), which is the difference in the stream level and the

groundwater level; (ii.) The vertical distance between the measuring point in the aquifer and the

stream bed (dl); (iii.) The vertical hydraulic conductivity (Kv) of the material along the vertical flow

path between the groundwater measuring point and the stream bed. In reality, groundwater flow paths near streams can be significantly more complex than simple vertical movement as represented in Figure 9.1. Even at the metre-scale of the near-stream environment, the groundwater flow in channel deposits can be a complex pattern of upward, downward and lateral movement (Vaux, 1968; Woessner, 2000). This complexity of groundwater flow direction becomes more evident when the level of investigation broadens to the kilometre-scale. However, it is still useful to compare stream levels to nearby groundwater levels to determine if the reach is likely to be gaining or losing. This requires access to the aquifer to be able to measure the groundwater level. This is usually done by constructing a piezometer, which is a cased hole with a screen at a fixed interval down the hole. Also called monitoring bores or observation bores, piezometers allow groundwater conditions at a single point within the aquifer to be measured. 9.1 Piezometers Piezometers are constructed to enable measurement of groundwater levels to characterise the groundwater flow regime surrounding a stream. Construction requirements for such monitoring bores are outlined in LWBC (2003) with an example given in Figure 9.2. Different configurations of piezometers can be used to gain an understanding of groundwater flow directions and aquifer characteristics. Options include: (i.) A piezometer nest, where a number of piezometers of varying depth are

constructed at one location. This is used to gain information on the hydraulic gradient in the vertical direction at a particular point;


(ii.) A piezometer transect, where piezometers are constructed in a line relative to the stream. This gives information on the groundwater flow in a vertical cross section, including whether losing or gaining conditions are evident. Typically, transects are constructed perpendicular to the stream. However, these transects should ideally be aligned along the groundwater flow path, and this may not be necessarily perpendicular to the stream (Woessner, 2000). Longitudinal sections are transects that are constructed down the axis of the stream, to map downstream changes in the vertical profile;

(iii.) A piezometer network, where piezometers are sited at different points in the landscape to provide information on the groundwater flow in the horizontal dimension as a plan view. This is done by contouring the elevation of the groundwater potentials measured from the piezometers – this depicts the potentiometric surface of the aquifer being monitored. All the piezometers need to access the same aquifer because of the complication that different aquifers can have different groundwater levels at the same geographical location.

Figure 9.1: Configuration of minipiezometer and stilling well for hydrometric measurement of seepage flux. In this example the groundwater level is lower than the surface water level and seepage flux is negative and downwards.

dh

Minipiezometer Stilling well

Surface water body

Sediment bed

Kv

dl

QQ


Figure 9.2: Example of monitoring bore construction (LWBC, 2003) However, more detailed information is often required in the immediate vicinity of the surface water feature where actual interaction takes place. For example, the magnitude of the flow of water between stream and aquifer may not be defined by the hydraulic conductivity of the aquifer proper (as measured via the piezometer network) but that of the relatively thin riverbed deposits. Minipiezometers are scaled-down versions of piezometers designed to monitor conditions in this near-stream environment. Minipiezometers monitor shallow groundwater conditions in the stream bed, typically at depths of less than 2 metres, refer Box 9.1. The basic construction is a small-diameter pipe stoppered at the base with gauzed holes to serve as an inlet. Different designs and materials have been trailed including: (i.) A closed metal or PVC tube sealed at the base, with holes drilled near the tip

and a small wad of fibreglass inserted inside the tip to act as a filter (Winter et al. 1988);

(ii.) A flexible plastic tube closed at the end with a perforated tip wrapped in fibreglass or nylon mesh (Lee and Cherry, 1978);


(iii.) Commercially available PVC or stainless steel probes with fitted drive point, inlet holes and porous filters.

There are two general methods of installing the minipiezometer. Robust versions such as those constructed of stainless steel with a pointed tip are driven directly into the sediment either manually, using a hammer or hydraulic ram. Alternatively, casing is driven into the sediment either directly (Lee and Cherry, 1978) or with the aid of an internal solid driver rod (Baxter et al, 2003), refer Figure 9.3. The piezometer is then lowered down inside the casing and held in place while the casing is slowly raised and removed. The sediment is allowed to collapse around the installed minipiezometer. Water can be removed from the minipiezometer by using a hand-pump or a syringe connected to plastic tubing. This is done to check if the inlet is not clogged and that the piezometer refills to an equilibrium level.

Figure 9.3: Stages in the installation of minipiezometer and stilling well for hydrometric investigations of seepage flux (modified from Baxter et al, 2003)

(1) Driver mechanism consisting of solid steel driver rod (C) and steel outer casing with flange (A) hammered into sediment to suitable depth using a cap fitting (B)

(2) Driver rod (C) removed with the steel outer casing retained (3) Minipiezometer inserted into the outer steel casing (4) Outer steel casing removed with minipiezometer held in position and sediment was manually tamped

around the minipiezometer. Bentonite clay can also be used to seal the annulus between minipiezometer and hole above the inlet.

(5) Stilling well fitted and secured using a star picket

ooo

o o o

o o o

1 2 4 3 5

A

C D

Stream

Sediment

B


9.2 Head Difference Measurement Understanding the hydraulic gradient between the stream and aquifer requires measurement of the water levels in these two systems. Installing minipiezometers directly into the stream bed allows the opportunity for direct measurement of the vertical hydraulic gradient (dh/dl). The length of the vertical flow path (dl) is the depth from the top of the sediment bed to the uppermost opening of the minipiezometer (Figure 9.1). The head difference (dh) between the groundwater level in the minipiezometer and the stream level can be directly measured by using: (i.) A manometer, where a clear plastic tube making up the minipiezometer is

connected to a vertical manometer board, as is a similar plastic tube feeding into the stream (Lee and Cherry, 1978; Winter et al, 1998). The groundwater and surface water are drawn up into the board by using the vacuum pump, and their relative levels are manually read from the meter stick;

(ii.) A stilling well, which is an open hollow tube connected vertically to the casing of the minipiezometer, along the side perpendicular to stream flow (Baxter et al. 2003). The top of the tube above the surface water level and the base is located just above but not embedded within the sediment bed (Figure 9.1). The purpose of the stilling well is to provide a stable surface water level by reducing the effects of streamflow. The two water levels can be measured by using a chalked wire or wooden dowel, or with the various water level recorders commercially available.

However, it is more common for the water levels in surface water and groundwater systems to be measured at different locations. For example, river levels are taken at gauging stations and groundwater levels may be taken at nearby monitoring bores. Water level measurements taken at a site are usually relative to an arbitrary benchmark such as the base of a staff gauge (for stream measurements) or the top of the piezometer casing (for groundwater measurements). When measuring the head difference it is important that the measurements of water level in the surface water body and aquifer use the same reference datum (such as metres AHD), or that the height difference between the benchmarks used is known. Another complication is that corrections need to be made to the water level measurements if there is a significant density difference between the surface water and the groundwater. This may be due to contrasts in water temperature and salinity. These corrections are made by converting the water level measurements to freshwater heads (hf) by:

pfpf hh )/( ρρ= Equation 9.2

where ρp is the observed water density, ρf is the density of fresh water, and hp is the height of the water level relative to the reference datum (Lusczynski, 1961). The water level in a stream is called the stage and is measured relative to an arbitrary datum. The most common methods of taking stage measurements are: (i.) a staff gauge which is a vertical graduated marker established to visually

estimate the water level;


(ii.) Automatic water level recorders that are installed in a stilling well to minimise the effects of turbulence or wave action. Different options include a float recorder, where the vertical movements of a float on the water surface is recorded, a bubbler gauge where the pressure required to force out nitrogen gas bubbles out of a tube is correlated with water level, and pressure transducers, which measure hydrostatic pressure of the water column.

Different techniques can be used to measure the groundwater level within the piezometer, including: (i.) Plopper, where a concave metal casting attached to the graduated tape makes a

plopping noise when it hits the groundwater surface; (ii.) Electrical sounder, where the insulated wires for a pair of electrodes are

incorporated into a graduated flat tape. A circuit is completed when the electrodes come into contact with the groundwater surface, which activates a light and/or buzzer;

(iii.) Wetted-tape, where a weighted tape that is rubbed with coloured chalk is lowered down the piezometer until it is submerged. The water level is indicated by where the chalk has been removed;

(iv.) Bubble tube, where a length of plastic tubing marked with depth increments is lowered down the piezometer. Contact with the standing water is distinguished by blowing into the tube and listening for the sound of bubbles;

(v.) Automatic water level recorders, similar to that used in surface water bodies such as pressure transducers, or capacitance probes.

9.3 Hydraulic Conductivity Measurement Hydraulic conductivity (K) defines the rate of movement of water through the aquifer. It is the constant of proportionality in Darcy’s Law and as such is defined as the flow volume per unit cross-sectional area of porous medium under the influence of a unit hydraulic gradient. This translates to SI units of m3/m2/day or m/d, but other measurement units are commonly used (Table 9.1) Table 9.1: Commonly used units for hydraulic conductivity (K)

Description metres/day (m/d)

metres/second (m/s)

millimetres/day (mm/d)

millimetres/hour (mm/hr)

Extremely slow 0.000001 1.5741x10-11 0.001 0.000041667 Very Slow 0.0001 1.5741x10-9 0.1 0.0041667 Slow 0.01 1.5741x10-7 10 0.41667 Moderate 1 1.5741x10-5 1000 41.667 Fast 10 1.5741x10-4 10000 416.667 Very Fast 100 1.5741x10-3 100000 4166.667 Measurement of hydraulic conductivity is problematic, considering the parameter can differ over several orders of magnitude across the spectrum of sediments and rock types, as indicated in Table 9.2. The parameter can also vary markedly in space, even with apparently minor changes in sediment characteristics. Hydraulic conductivity is influenced by the properties of the fluid being transmitted (such as viscosity) as well as the porous medium. Hydraulic conductivity is also scale dependent, so that measurements taken at the core sample level may not be directly extrapolated to the aquifer scale. It is also direction dependent, so that hydraulic conductivity can be markedly different in the vertical from the horizontal. The hydraulic conductivity in


the vertical direction (Kv) in aquifers can be several orders of magnitude lower than that in the horizontal direction (Kh). This is partly because the interlayering of finer-grained clays and silts impedes vertical water movement, while laterally extensive sand and gravel deposits enable high rates of horizontal flow. Hydraulic conductivity cannot be directly measured but inferred from field, laboratory or modelled data. Table 9.2: Indicative hydraulic conductivities of some rock types (Bouwer, 1978; Brassington, 1988)

Rock Type Grain size (mm) Hydraulic Conductivity K (m/d) Clay 0.0005-0.002 10-8-10-2 Silt 0.002-0.06 10-2 - 1 Fine Sand 0.06 –0.25 1-5 Medium Sand 0.25-0.50 5-20 Coarse Sand 0.50-2 20-100 Gravel 2-64 100-1000 Shale small 5x10-8 – 5x10-6 Sandstone medium 10-3-1 Limestone variable 10-5-1 Basalt small 0.0003-3 Granite large 0.0003-0.03 Slate small 10-8-10-5 Schist medium 10-7-10-4

Different approaches have been taken to estimate hydraulic conductivity, including: (i.) Using seepage meters (refer Chapter 3) which directly measure the flux (Q) at

the interface between the stream and aquifer. The basic method is to isolate part of the sediment-water interface with a chamber open at the base (with surface area A) and measure the change in water volume contained in a bag attached to the chamber over a measured time period. When combined with head gradient measurements (dh/dl) between the sediment bed and the stream from minipiezometers (Lee and Cherry, 1978), the vertical hydraulic conductivity can be derived from Darcy’s Law:

dhdl

AQKv = Equation 9.3

(ii.) Infiltration Tests, where infiltrometers (also known as permeameters) are used to measure the rate that water can infiltrate downwards through the sediment/soil profile which is a function of vertical hydraulic conductivity. Two basic methods can be employed. In falling-head tests water is added to reach a target level in the infiltrometer, after which the subsequent decline in water level is recorded as infiltration occurs. In constant-head tests, the target level is maintained during the test duration by adding increments of water of known volume. These tests are commonly used for unsaturated soils, but can also be carried out in calm, shallow water bodies (McMahon et al. 1995; Duwelius, 1996; Lindgren and Landon, 2000; Rosenberry, 2000). Various instruments are available that are based on measuring infiltration, including well, disc and ring (double and single) permeameters (ANCID, 2000). Similar tests using constant-head or falling-head configurations can be undertaken in the laboratory on core samples taken from the field site;

(iii.) Pump Tests, involving pumping groundwater from the piezometer and monitoring the pumping rate, as well as the groundwater level in the piezometer or in nearby piezometers. The pump test (also called aquifer test)


indicates how the aquifer responds to groundwater withdrawals, with the data used to estimate aquifer characteristics such as hydraulic conductivity. A wide variety of formulas are available for the analysis of pump test data, based on differences in aquifer type and geometry, boundary conditions, and underlying assumptions. Slug Tests are particular tests where the rate of groundwater recovery is measured after a small volume (slug) of water is suddenly displaced (Duwelius, 1996; Cey et al, 1998; Springer et al, 1999);

(iv.) Grainsize analysis, involving determining the distribution of grainsizes within the sediment using standard sieves. Empirical relationships are used to estimate hydraulic conductivity from standard grainsize parameters (Vukovic and Soro, 1992). The grainsize diameter at which 10% of the sediment is finer (d10) is applied in a commonly used empirical formula, initially developed by Hazen (1893):

210CTdAK H= Equation 9.4

where AH is a dimension coefficient (= 1.0 for m/d), C is an empirical constant (=860) and T is a temperature correction factor (=1 at 10° C). Another empirical relationship developed by Alayamani and Sen (1993) uses the slope and intercept (I0) of the grainsize distribution curve between d10 and the median grainsize (d50):

Equation 9.5

9.4 Flow Net Analysis Contours of the watertable elevation as interpreted from a piezometer network can be used in flow net analysis to estimate groundwater seepage rates. This is a hydrometric method that can be applied to water level data either in the horizontal or vertical plane. As a first step, if the network is of sufficient density, particularly near the stream or lake, contours of the potentiometric surface can indicate the relationship between the aquifer and surface water feature. Contours that curve in towards the upstream end of the stream indicates a gaining stream (Figure 9.4a) and contours that curve towards the downstream end indicate a losing stream (Figure 9.4b). A flow net is generated by the combination of the potential contours and groundwater flow lines that by definition are perpendicular to these contours. The objective in constructing the flow net is to generate cells that are equidimensional (Figure 9.5). A flow tube is defined as being bounded by the groundwater flow lines. In the example given in Figure 9.5, total groundwater discharge to the stream (Q = q1 + q2 +q3 + q4) is derived from Darcy’s Law as:

nhbKpQ /∆= Equation 9.6

where K is the average hydraulic conductivity, p is the number of flow tubes (=4), ∆h is the difference in head between the two bounding potentiometric contours (=6-0), b is the aquifer thickness and n is the number of head drops in the flow net (n=3, 6 to 4, 4 to 2 and 2 to 0) (Loaiciga and Zekster, 2003).

210500 ))(025.0(1300 ddIK −+=


Figure 9.4: Watertable contour patterns around streams (a) Contours pointing upstream for gaining streams (b) contours pointing downstream for losing streams (Winter et al, 1998) Figure 9.5: Plan view of example groundwater flow net towards a gaining surface water feature (after Loaiciga and Zekster, 2002) 9.5 Advantages and Disadvantages Hydrometric methods are based on fundamental principles for groundwater flow in porous medium. Hydrometric investigations in stream beds give valuable information of the water/sediment interface. Many minipiezometers can be quickly and cheaply installed to provide information on the areal distribution of seepage. Comparison of surface water levels and groundwater levels are a simple guide to the direction of potential seepage, and can be used in quickly targeting losing or gaining conditions. Comparison of hydrographs from stream gauging sites with piezometers can indicate temporal changes in hydraulic gradient and therefore potential seepage direction, (refer Box 9.2.).

q1 q2

q3 q4

Equipotential Contour

Streamline

h =6m

h =4m

h =2m

h =0m


However, hydrometric methods rely on a reasonable estimate for hydraulic conductivity (K) which has a potential range over several orders of magnitude. Measurement of hydraulic conductivity tends to be point estimates and may not reflect the average conditions for the groundwater flow path considered. This is particularly true if thin clay veneers in bed deposits are the main control for water flow. Also, assumptions of simple groundwater flow conditions (ie vertical) may not reflect actual conditions. This assumption is typically invalid when using piezometers a long distance from the surface water feature. The method can be expensive if monitoring bores need to be constructed to obtain information on groundwater levels and hydraulic conductivity. Flow nets can be generated for various geometrical and physical configurations to define key processes. This approach was used to generate a framework for determining groundwater-surface water interactions for lakes and other surface water bodies (Nield et al, 1994). This allows flux conditions (i.e. losing, gaining, through-flow) to be predicted using geometrical and physical parameters, and flux boundary conditions. Flow net analysis can provide a simple, quick and cost-effective way of estimating first-order approximation of seepage flux. However, the method cannot account for spatial variability and other local groundwater factors. 9.6 Data Availability Hydrometric analysis of seepage flux in streams requires data on stream water levels, nearby groundwater levels and estimates of aquifer properties (mainly hydraulic conductivity). The State and Territory agencies involved in water management are the main custodians for monitoring of stream water levels. The major hydrographic databases containing this data record are outlined in Brodie et al. (2007). These agencies also tend to be involved in the collection of groundwater information (such as borehole databases). 9.7 Relevant Links Aquifer Test Forum http://www.aquifertest.com/forum/ Minimum construction requirements for water bores in Australia http://www.iah.asn.au/pdfs/mcrwba.pdf


Box 9.1 Use of minipiezometers and stilling wells in the Lower Richmond catchment, NSW Simple hydrometric investigations can be undertaken by installing a stilling well to limit the effects of winds and current on the surface water level measurement coupled with a minipiezometer to measure the groundwater level in the underlying sediment bed. This was undertaken at a site on Yellow Creek on the floodplain at the base of the Alstonville Plateau escarpment (Brodie et al, 2005).

The minipiezometers were constructed using 2 m lengths of 25 mm OD electrical conduit stoppered at the base with a series of 2 mm holes drilled to form a 15 cm long perforated inlet. These were then installed into the stream bed using an outer steel casing and driver mechanism. The minipiezometers were then developed by moving a length of solid steel rod up and down within the conduit, to dislodge any smeared clay and ensure hydraulic connection at the inlet. A length of 50 mm PVC casing was attached to a star picket and used as a stilling well to stabilise stream water levels. Once established, readings were taken of the depth to groundwater relative to the top of the minipiezometer and of the depth to the stream water level relative to the top of the stilling well. The vertical difference between these two reference points are taken into account when calculating the head difference between the stream and the shallow groundwater system. Density effects due to contrasts in water salinity or temperature were considered negligible.

At the Yellow Creek site, the minipiezometer was installed to a depth of 1.23 m within heavy competent clay, with the stream about 0.7 m deep. The stream flow rate was very low at the time of installation. Figure 9.6 plots the water level measurements taken from the stilling well and minipiezometer at this site. The groundwater level in the minipiezometer took about two days to re-equilibrate following installation, due to the dominance of clay in the stream bed sediment. The groundwater level stabilised to about 0.16 m below the surface water level. This suggests that the potential direction of seepage flux at this site was downward during the period of measurement.

A slug test was undertaken using the minipiezometer by filling the conduit with water and monitoring the subsequent decline in water level. The resulting horizontal hydraulic conductivity (Kh) estimate of 0.004 m/d using the standard Hvorslev method is within the typical range for clays. A seepage flux of 0.16 m3/d/km can be estimated from Darcy’s Law by deriving a vertical hydraulic conductivity (Kv) by assuming anisotropy of 0.1 and assuming an average stream width and homogenous conditions. Although indicative, this very low seepage rate is as anticipated in such a heavy clay stream bed.

-2

-1.8

-1.6

-1.4

-1.2

-1

-0.8

-0.6

-0.4

-0.2

02/03 3/03 4/03 5/03 6/03 7/03 8/03 9/03 10/03 11/03

Dep

th to

Wat

er (m

)

Groundw ater Depth

Stream Level depth

Figure 9.6: Groundwater and surface water level measurements for Yellow Creek site during March, 2005 (Brodie et al. 2005). Water level measurements are relative to the top of the stilling well.


Box 9.2 Comparison of groundwater levels with river stage at key sites in the Border Rivers catchment Monitoring of water levels in the stream and aquifer can be used to assess the hydraulic gradient between these two systems. This was done by analysing the relationship between river stage data and groundwater hydrographs along transects made up of existing gauging stations and appropriately located monitoring bores in the Border Rivers catchment. River stages at gauging stations/weirs were compared with nearby groundwater levels at three key sites (Bonshaw, Goondiwindi and Terrewah). The potential direction of flux between the river and aquifer could be inferred.

Figure 9.7: River levels versus nearby bore water levels for key sites in the Border Rivers catchment (Baskaran et al, 2005)

A comparison of groundwater levels of three NR&M monitoring bores (41630062, 41630063 and

41630071) and the Dumaresq River stage at Bonshaw weir (Figure 9.7a) indicates that the Dumaresq River stage downstream of Bonshaw weir is higher than the local watertable, indicating a downward hydraulic gradient and potential for downward leakage of river water into the aquifer.

Groundwater levels in three NR&M monitoring bores were compared to the MacIntyre River stage at Goondiwindi weir (Figure 9.7b). Water table fluctuations appear to be influenced by changes in river stage at this site and groundwater levels are within a few metres of the river level. The results shows that the river stage at the Goondiwindi section is slightly higher than the local water table, indicating that this reach has the potential to be a losing stream.

Groundwater levels in two NR&M (41620019 and 41620020) and two DNR (GW036699 and GW036963) monitoring bores were compared to the MacIntyre River stage near Terrewah gauging station (Figure 9.7c). The water table in both the QLD and NSW bores was more than 10 m below the river stage and did not appear to be influenced by changes in river stage. The river reach at the Terrewah section is composed of heavy clay sediments underneath the river. The MacIntyre river stage is higher than the local water table in this river reach indicating that the flux direction should be from the river to the aquifer.

Dumaresq River @ Bonshaw weir - QLD

270

275

280

285

290

295

300

1/01

/198

5

16/0

5/19

86

28/0

9/19

87

9/02

/198

9

24/0

6/19

90

6/11

/199

1

20/0

3/19

93

2/08

/199

4

15/1

2/19

95

28/0

4/19

97

10/0

9/19

98

23/0

1/20

00

6/06

/200

1

19/1

0/20

02

2/03

/200

4

AHD

(m)

river levelbore 41630062bore 41630063Bbore 41630071B

MacIntyre River @ Goondiwindi weir - QLD

205207209211213215217219

4/05

/200

0

1/10

/200

0

28/0

2/20

01

28/0

7/20

01

25/1

2/20

01

24/0

5/20

02

21/1

0/20

02

20/0

3/20

03

17/0

8/20

03

14/0

1/20

04

12/0

6/20

04

9/11

/200

4

8/04

/200

5

AH

D (m

)

river levelbore 41620021bore 41620022bore 41620023

MacIntyre River @ Terrewah

164168172176180184188192196200

1/01

/198

5

21/0

1/19

87

9/02

/198

9

1/03

/199

1

20/0

3/19

93

9/04

/199

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28/0

4/19

97

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99

6/06

/200

1

26/0

6/20

03

AH

D (m

)

GW036699GW036963river levelbore 41620019bore 41620020

a

b

c


10. Hydrochemistry Analysing and interpreting the chemistry of water can provide valuable insights into groundwater-surface water interactions. Dissolved constituents can be used as environmental tracers to track the movement of water. For example, a particular characteristic of the groundwater chemistry can be used as an indicator of groundwater discharge when measured in the surface water. Such tracers can be used to determine source areas of water and dissolved chemicals in catchments, calculate hydrologic and chemical fluxes between groundwater and surface water, calculate water ages that indicate the length of time water and dissolved chemicals have been present in the catchment (residence times), and determine average rates of chemical reactions that take place during transport (Winter et al, 1998). Geochemical mass balance models have been used to estimate mixing ratios of river water and groundwater (Cook et al, 2003). Chemicals or materials can be introduced specifically to study groundwater-surface water interactions and are referred to as artificial tracers (Chapter 11). Environmental tracers can occur naturally or have been released into the general landscape by human activities. Some of the commonly used environmental tracers include: (i.) Field parameters such as electrical conductivity or pH;

(ii.) The major anions and cations such as calcium, magnesium, sodium, chloride and bicarbonate;

(iii.) Stable isotopes in the water molecule of oxygen-18 (18O) and deuterium (2H); (iv.) Radioactive isotopes such as tritium (3H) and radon (222Rn); (v.) Industrial chemicals such as chlorofluorocarbons (CFC) and sulphur

hexafluoride (SF6). Several studies have used a combination of these tracers (eg major ions, stable and radioactive isotopes) to assess groundwater-surface water interactions (Crandall et al, 1999; McCarthy et al, 1992; Herczeg et al, 2001; Cook et al, 2003; Baskaran et al, 2004). 10.1 Field Water Quality Parameters Water quality parameters such as electrical conductivity (EC), pH or redox (Eh) that can be readily measured in the field have been used to investigate groundwater discharge to streams. This is particularly the case if the shallow groundwater is relatively saline and is a significant contributor to the salt load of the stream. Run of river salinity surveys have been used to define the key reaches of the River Murray in South Australia with large groundwater-induced salt accessions (Porter, 2001). The groundwater contribution to stream flow can be estimated, based on measurements of groundwater salinity and stream flow. Box 10.1 provides an example of this, using a simple mass balance equation. Monitoring of stream salinity in conjunction with stream flow can provide useful information on the temporal variability of salt loads and, by inference, groundwater discharge.


Field surveys using other water quality parameters may also be useful in characterising groundwater discharge. In coastal catchments where acid sulfate soils are an issue, surveys of stream acidity can be a simple and effective technique of mapping key reaches receiving shallow acidic groundwaters. 10.2 Major Ion Chemistry Cations (such as calcium, magnesium, sodium and potassium) and anions (such as chloride, bicarbonate, sulfate and bromide) have been used as tracers to determine groundwater input to a stream during high flow and low flow periods. Groundwater can have a chemistry that is distinctly different to the connected stream, and these characteristics can be used as indicators of groundwater discharge. Solutes that are present in groundwater are derived from two main sources: (i) input from rainfall, which have their origin from both marine salts and continental dust, and (ii) acquisition during weathering and water-rock interactions. Processes that affect hydrochemistry include (i) acid-base reactions, (ii) precipitation and dissolution of minerals, (iii) sorption and ion exchange, (iv) oxidation-reduction reactions, (v) biodegradation and (vi) dissolution and exsolution of gases. Hydrochemistry can be interpreted to understand the key processes that have occurred during the movement of water through aquifers and streams. Representative samples of the stream and the groundwater system, as well as other possible inputs such as rainfall, need to be taken. Protocols are available for the appropriate collection, field preparation and storage of these water samples (Table 10.1). Major anions (chloride, bromide and sulphate) can be determined by ion chromatography and major cations (calcium, magnesium, sodium and potassium) by atomic absorption spectrophotometry (AAS) or by inductively coupled plasma - atomic emission spectrometry (ICP-AES). These are routinely undertaken in most analytical laboratories. The accuracy must be demonstrated both by the use of appropriate standards and also by use of the ionic balance to check electrical neutrality. Major ions data may be presented in graphical format, of which the most useful plots are the trilinear diagram that show the total major anion or cation composition on separate or combination (Piper) diagrams (Hem, 1989). These diagrams have the advantage for tracer work of showing a large number of analyses in one plot to define distinct populations or trends. The relationships and evolution between dominant compositions (eg Ca-HCO3 to Na-Cl) usually indicate trends along flow paths or mixing between water bodies. Other types of diagrams include the mixing plot (Lawrence et al, 1976), Shoeller type diagrams and the Durov diagram (Howard and Lloyd, 1983; Petalas and Diamantis, 1999).


Table 10.1: Australian Standards related to water sampling (http://www.standards.com.au) Australian Standard Title Content

AS 4276.1-1995 Water microbiology - General information and procedures

Information on the microbiological examination of water.

AS/NZS 5667.1:1998 Water quality - Sampling - Guidance on the design of sampling programmes, sampling techniques and the preservation and handling of samples

General principles to be applied in the design of sampling programmes, general guidance on sampling techniques and guidance on the procedures to be taken to preserve and transport samples for the physical, chemical and radiological analysis of waters and wastewaters, including bottom sediment and sludges, for the purposes of process control, quality characterisation, identification of sources of pollution, compliance with water quality guidelines or standards, and other specific reasons.

AS/NZS 5667.10:1998 Water quality - Sampling - Guidance on sampling of waste waters

Detailed guidance on the design of sampling programmes, sampling techniques and the handling and preservation of samples of waste water. It is identical with and has been reproduced from ISO 5667-10:1992.

AS/NZS 5667.11:1998 Water quality - Sampling - Guidance on sampling of groundwaters

Part of ISO 5667 that provides guidance on the design of sampling programmes, sampling techniques and the handling of water samples taken from groundwater for physical, chemical and microbiological assessment.

AS/NZS 5667.12:1999 Water quality - Sampling - Guidance on sampling of bottom sediments

Part of ISO 5667 that provides guidance on the sampling of sedimentary materials from inland rivers and streams; lakes and similar standing bodies; and estuarine and harbour areas.

AS/NZS 5667.4:1998 Water quality - Sampling - Guidance on sampling from natural and man-made lakes,

Detailed guidance on the design of sampling programmes techniques and the handling and preservation of samples of water from natural and man-made lakes. It is technically equivalent to and has been reproduced from ISO 5667-4:1987.

AS/NZS 5667.6:1998 Water quality - Sampling - Guidance on sampling of rivers and streams

Part of ISO 5667 that sets out the principles to be applied to the design of sampling programmes, sampling techniques and the handling of water samples from rivers and streams for physical, chemical and microbiological assessment.

10.3 Stable Isotopes Isotopes are forms of a given chemical element that have different atomic masses. For a particular element, the isotopes have the same numbers of protons, and so have the same atomic number. However, each isotope has a different number of neutrons and therefore has a different atomic mass. Stable isotopes are those isotopes that do not undergo radioactive decay; so their nuclei are stable and their masses remain the same. However, they may themselves be the product of the decay of radioactive isotopes. In hydrological studies, the stable isotopes of interest generally relate to H, C, N, O, S, B, and Li. In terms of the water molecule itself, oxygen has three stable isotopes, 16O, 17O, and 18O; and hydrogen has two stable isotopes, 1H and 2H (deuterium). The relative abundances of these stable isotopes of hydrogen and oxygen are given in Table 10.2. The stable isotopes of 18O (oxygen-18) and 2H (deuterium) are


used to provide information on hydrological processes, including groundwater-surface water interactions, refer Box 10.2. Table 10.2 Relative abundances of the oxygen and hydrogen isotopes

Hydrogen Oxygen

Isotope Abundance Isotope Abundance 1H 0.99985 16O 0.99757 2H 17O 0.00038

18O 0.00205

Numerous papers or books have been published over the past 40 years that deal with applications of environmental isotopes in hydrological investigations. Some of the most comprehensive information on isotopes appears in the Handbook of Environmental Isotope Geochemistry series edited by Fritz and Fontes in the 1980s as well as reviews by Fontes and Edmunds (1989), Coplen (1993) and Gat (1996). Textbooks recently published by Mazor (1997), Clark and Fritz (1997) and Cook and Herczeg (2000) are excellent reference works on isotopes application in hydrology. Water samples collected for isotopic analysis should be stored in bottles with tight closures, such as caps with conical plastic inserts. Bottles for archived samples should be glass. End-member samples (representing the key components of the hydrological system) should be analysed prior to undertaking a detailed study to determine if sufficient isotopic discrimination exists in the hydrologic system. Background samples can also be very important. Often these will be samples of shallow local groundwater, near to, but outside, the area of investigation. Oxygen and hydrogen stable isotopic ratios are measured by isotope mass spectrometry. Hydrogen analysis is done on hydrogen gas obtained through high-temperature reduction of water on metal (Kendell and Coplen, 1985). Oxygen analyses are done on carbon dioxide that has equilibrated with water at a constant temperature (Epstein and Mayeda 1953). Oxygen and hydrogen isotope compositions are commonly reported relative to an agreed sample of ocean water, referred to as the Standard Mean Ocean Water (SMOW), representing the largest and most equilibrated water body. Stable isotope ratios of deuterium/hydrogen (2H/1H) and 18O/16O of water are conventionally expressed as units of parts per thousand (per mil, ‰) deviation from SMOW. Isotopic fractionation of water molecules due to evaporation of seawater and subsequent precipitation in rainfall was recognised by Craig (1961). Based on about 400 water samples from rivers, lakes and precipitation, a linear relationship between deuterium and oxygen-18 was established for average global meteoric waters. This relationship (δD=8δ18O+10) is known as the Global Meteoric Water Line (GMWL) and provides a useful benchmark against which regional or local waters can be compared and their isotopic composition interpreted. The slope of this curve represents Rayleigh fractionation due to repeated evaporation and precipitation, the intercept (termed the deuterium excess) is largely a function of the mean relative humidity of the atmosphere above the ocean water (Merlivat and Torzel, 1979). Local meteoric water lines can be established from isotopic analysis of local precipitation events.


Comparison of the stable isotope data for surface water and groundwater samples relative to the global or local meteoric water lines can provide information on processes. For example, isotopically light water molecules evaporate more efficiently than isotopically heavy water molecules. Due to this variability in isotopic vapour pressures, evaporation produces residual water enriched in the heavier isotopes relative to the initial isotopic composition. Therefore water that has undergone evaporation lies to the right of the local meteoric water line due to this enrichment (Coplen, 1993). The trend line for evaporation from surface water tends to have a slope between 4 and 6, with a slope less than 4 indicating evapotranspiration of soil water in the unsaturated zone (Allison, 1982). 10.4 Radioactive Isotopes Radioactive isotopes have unstable nuclei that decay, emitting alpha, beta, and sometimes gamma rays. Such isotopes eventually reach stability in the form of non-radioactive isotopes of other chemical elements, termed radiogenic daughters. Decay of a radionuclide to a stable radiogenic daughter is a function of time measured in units of half-lives. The decay constants (λ) and half-lives (t1/2) of radioactive isotopes that are frequently used as environmental tracers in the field of hydrology are listed in Table 10.3. Radioactive isotopes are useful indicators of the time that water has spent in the groundwater system. For example, tritium (3H) is a well-known radioactive isotope of hydrogen that had peak concentrations in precipitation in the mid-1960s as a result of above-ground nuclear bomb testing conducted at that time. Radon-222 (222Rn) is a radioactive daughter isotope of radium-226 that has a half-life of only 3.8 days. It is produced naturally in groundwater as a product of the radioactive decay of 226Ra in uranium-bearing rocks and sediments. Radon concentrations in groundwater depends on the presence of these radioactive isotopes in the aquifer matrix, and can vary from <2 Bq/L within clastic sediments to >200 Bq/L in igneous and metamorphic rocks (Lee and Hollyday, 1993).Several studies (Ellins et al, 1990; Crandall et al, 1999; Pritchard et al, 2000; Cook et al, 2003) have demonstrated that radon can be used to identify locations of significant groundwater input to a stream. Radon was also used in a study in France to determine stream water loss to groundwater as a result of groundwater withdrawals (Bertin and Bourg, 1994). Radon is a gas, and natural radon concentrations in the atmosphere are so low that natural waters in contact with the atmosphere will continually lose radon by volatilisation. Hence, groundwater has a higher concentration of 222Rn than surface water. Any significant concentration of radon in a stream or river is a sensitive indicator of local inputs of ground water. Kraemer and Genereux (1998) provide a detailed discussion of 222Rn mixing models and the use of 222Rn to determine areas of ground water discharge to streams.


Table 10.3: Decay constants and half-lives of selected radioactive isotopes with application to hydrology (adopted from Browne and Firestone, 1999)

Isotope Decay Constant Half-life

(year-1) (day-1) (year) (day)

Rubidium (87Rb) 1.46 x 10-11 4.00 x 10-14 4.75 x 1010 1.73 x 1013

Uranium (238U) 1.55 x 10-10 4.24 x 10-13 4.468 x 109 1.63 x 1012

Iodine (129I) 4.41 x 10-8 1.21 x 10-10 1.57 x 107 5.73 x 109

Chlorine (36Cl) 2.3 x 10-6 6.30 x 10-9 3.01 x 105 1.10 x 108

Krypton (81Kr) 3.03 x 10-6 9.03 x 10-9 2.29 x 105 8.36 x 107

Carbon (14C) 1.21 x 10-4 3.31 x 10-7 5730 2.09 x 106

Radium (226Ra) 4.33 x 10-4 1.19 x 10-6 1600 5.84 x 105

Argon (39Ar) 2.58 x 10-3 7.06 x 10-6 269 9.83 x 104

Silicon (32Si) 4.95 x 10-3 1.36 x 10-5 140 5.11 x 104

Strontium (90Sr) 0.0241 6.65 x 10-5 28.78 1.05 x 104

Hydrogen (3H) 0.0558 1.53 x 10-4 12.43 4540

Krypton (83Kr) 0.0644 1.77 x 10-4 10.756 3929

Radium (228Ra) 0.121 3.31 x 10-4 5.75 2100

Sulphur (35S) 2.89 7.92 x 10-3 0.240 87.51

Argon (37Ar) 7.23 1.98 x 10-2 0.0959 35.04

Radon (222Rn) 66.0 0.181 0.0105 3.82

10.5 Industrial Chemicals Chlorofluorocarbons (CFCs), which are industrial chemicals that are present in groundwater less than 50 years old, can be used to calculate groundwater age in different parts of catchments. CFCs are useful tracers in hydrological studies as they are non-reactive and resistant to degradation and have low toxicity (Davis et al, 1980). Likewise, sulfur hexafluoride (SF6) can also be used as an environmental tracer, being a persistent and stable industrial chemical with emissions commencing in the 1950s. SF6 is detectable at low concentrations (10-16 mol/L) and non-toxic (Kass, 1998). It has been used in groundwater dating (Busenberg and Plummer, 2000) and for monitoring septic tank effluent (Dillon et al, 1999). 10.6 Advantages and Disadvantages Environmental tracers have been used for the past few decades to quantify seepage flux and for defining key hydrogeological processes. Hydrochemistry is a valuable tool in developing a conceptual understanding of groundwater flow near a stream and has been used to provide information on groundwater evolution, residence times or mixing ratios that would otherwise be difficult to determine. The range of


hydrogeological processes that can be investigated under field conditions is probably the great strength of this method. Measurements of an environmental tracer along the stream can be a powerful tool to map the spatial distribution of groundwater inflows. This can be more rapid or cheaper than physically based methods (such as seepage meters or hydrometric studies) particularly if field chemistry parameters such as EC or pH can be used. Time series monitoring of environmental tracers can provide information on the changes in seepage flux at a stream site. Such water chemistry monitoring is commonly undertaken to complement hydrographic data collection and analysis. Stable and radioactive isotopes can be used as a preliminary investigation tool or as an independent means of confirming the results of other methods. Deuterium, oxygen-18 and radon-222 are the most commonly used isotopes to investigate groundwater-surface water interactions. Environmental tracers can be expensive due to the logistics of sampling or the cost of laboratory analysis. A high level of expertise may be required for their sampling and interpretation. Tracers such as deuterium, oxygen-18 or tritium can have long lead times between sample collection and the final analytical results. This is in contrast with parameters such as water level, temperature or electrical conductivity where real-time monitoring and data access is possible. Models used to quantify seepage flux from hydrochemical data can require estimates of parameters that are difficult to measure in the field. 10.7 Data Availability Water quality monitoring is undertaken by a range of Commonwealth, State or Local Government agencies, private companies, research groups and community-based groups, and over a range of settings, the most significant being rivers and creeks, industrial effluent and reservoirs and lakes. Such monitoring is undertaken for a number of reasons including compliance with health or environmental regulations, operational or process control, or for environmental or catchment health (Atech, 2000). Such data has the potential to be interpreted on the basis of being environmental tracers. Brodie et al. (2007) outlines some of the agencies that undertake water sampling and analysis. 10.8 Relevant Links Australian Nuclear Science and Technology Organisation http://www.ansto.gov.au/ Australian Radiation Protection and Nuclear Safety Agency Radioanalytical Services http://www.arpansa.gov.au/rad_serv.htm Clark I and P. Fritz, 1997 Environmental Isotopes in Hydrogeology. http://www.science.uottawa.ca/%7Eeih/ch1/ch1.htm#cocah CSIRO Isotope Analysis Service http://www.clw.csiro.au/services/isotope/ International Atomic Energy Agency http://www-naweb.iaea.org/na/index.html Murray-Darling Basin Groundwater Quality Sampling Guidelines http://www.mdbc.gov.au/__data/page/127/GroundwaterqualityguideliesReport.pdf National Association of Testing Authorities, Australia (NATA) http://www.nata.asn.au/ The National Water Quality Management Strategy (NWQMS) http://www.deh.gov.au/water/quality/nwqms/ Sustainability of Semi-Arid Hydrology and Riparian Areas (SAHRA) Isotopes and Hydrology http://www.sahra.arizona.edu/programs/isotopes/ USGS Resources on Isotopes http://wwwrcamnl.wr.usgs.gov/isoig/res/


Box 10.1 Use of stream salinity surveys to identify groundwater discharge, Alstonville Plateau NSW Differences in the salinity of the shallow and deeper groundwater systems was used as a basis for investigating groundwater discharge to streams on the Alstonville Plateau, northern NSW (Brodie et al, 2005). A field chemistry survey of the plateau streams was undertaken, involving the measurement of parameters such as electrical conductivity, temperature, pH and Eh. Rainfall during the survey period was minor (< 5 mm) and the stream flow was considered to reflect baseflow conditions.

Figure 10.1 shows the survey results for the Gum Creek catchment on the plateau. The stream shows a slight increase in salinity as it flows down the plateau escarpment, inferred to be due to contribution from the deeper, more saline aquifers. This results in stream salinity at the base of the escarpment that is about 25 µS/cm higher than at the top. Numerous springs have been mapped along the middle and base levels of the escarpment. The relative contribution from the deeper groundwater system (QG) can be estimated using a mass balance equation of:

)/()( GGSGSSSG ECECECECQQ −−= ++ Equation 10.1

(Oxtobee and Novakowski 2002)

where QS is the ambient stream discharge, ECS is the measured ambient electrical conductivity of the stream (75 µS/cm at the plateau top), ECG is the measured electrical conductivity of the discharging groundwater (350 µS/cm from a spring measurement), and ECS+G is the electrical conductivity of the stream resulting from mixing with the groundwater input (100 µS/cm at the plateau base). This gives an estimate of the relative contribution of the deeper groundwater system down the plateau escarpment as being in the order of 10% of stream flow under these low-flow conditions.

The field stream salinity data for the eastern tributary of Gum Creek can be used to support this estimate. At the plateau top, this tributary was not flowing at the time (ECS =92 µS/cm), so it is assumed that there is no contribution from the shallow aquifer. At its junction with Gum Creek, the tributary has a small flow and higher salinity (ECS+G = 232 µS/cm), refer Figure 10.1. If this stream flow is assumed to be entirely baseflow from the deeper groundwater system then QG = QS and the mass balance equation (10.1) can be rearranged to:

SGSG ECECEC −= +2 Equation 10.2

Using the salinity data for the tributary gives an estimate of the salinity of the groundwater discharge (ECG) of 372 µS/cm, similar to the spring discharge measurement (350 µS/cm) used in the calculation of the relative contribution of deeper groundwater discharge to overall stream flow for Gum Creek.

Figure 10.1(a) Field electrical conductivity (µS/cm) and (b) pH of Gum Creek and nearby springs, July 2004 (Brodie et al. 2005)


Box 10.2 Use of stable isotopes to define hydrological processes in the Border Rivers catchment, Murray-Darling Basin A combined major ions, stable isotopes (deuterium and oxygen-18) and a radioactive isotope (radon-222) sampling method was trialled in the Border Rivers catchment to investigate stream-aquifer connectivity (Baskaran et al. 2005). Groundwater and river water samples were collected during November-December 2004, coinciding with the main irrigation season. The δD and δ18O compositions for the groundwater that is close to the river are different from the isotopic composition of the groundwater that is farthest from the river (Figure 10.2). The near-stream groundwaters upstream of Keetah have a relatively enriched isotopic signature (less negative) that is similar to that of the river water. This indicates that infiltrating river water is the main source for the groundwater that is close to the river. Under low or average flow conditions, river water tends to be isotopically enriched relative to rainfall because of surface water evaporation (Simpson and Herczeg, 1991). The more depleted signature for groundwater that is further from the river suggests that this water may not originate from infiltrating river water. However, recharge following large flood events could have such a depleted isotopic signature because heavy rain tends to have a relatively negative isotopic composition and there is relatively less evaporation. This is supported by the fact that many of the groundwaters distant from the river upstream of Keetah cluster about the average (more depleted) signature for rainfall from large events, exceeding 200 mm/month (Figure 10.2).

-60

-50

-40

-30

-20

-10

0

10

-10 -8 -6 -4 -2 0 2 4

Oxygen-18 (o/oo, SMOW)

Deu

teriu

m (

o / oo,

SM

OW

)

River water

GW-upstream of Keetah(far stream)GW-upstream of Keetah(near stream)GW-downstream ofKeetah

Local Meteoric Water Line (Barakula)

Global Meteoric Water Line

Evaporation trendslope = 4.7

>200 mm rain

Figure 10.2. Deuterium versus oxygen-18 concentrations for river water and groundwater in the Border Rivers Catchment The chloride-deuterium plot (Figure 10.3) suggests that three types of groundwater occur in the study area namely: (i) some groundwaters upstream of Keetah, characterised by low chloride and relatively enriched δD, that is most frequently recharged by river water (ii) some groundwaters upstream of Keetah and a few groundwaters that are close to the MacIntyre River reach near Goondiwindi Weir, with low chloride and depleted δD, representing areas that are recharged less frequently by the river and more frequently by high rainfall and (iii) highly saline groundwaters further downstream of Goondiwindi, with very high chloride and lower δD, that never or rarely receive recharge from surface water.

1

10

100

1000

10000

100000

-50 -40 -30 -20 -10 0 10Deuterium (o/oo, SMOW)

Chl

orid

e (m

g/L)

GW-upstream of Keetah River water GW-downstream of Keetah

most frequently recharged by river

less frequentlyrecharged by river

infrequently or neverrecharged by river

Figure 10.3 Chloride versus deuterium concentrations for river water and groundwater in the Border Rivers Catchment The results of hydrochemical and environmental isotope sampling from the Border Rivers catchment indicate that the river and the shallow alluvial aquifers close to the river in the area upstream of Keetah have a close hydraulic relationship. In this upper catchment area, the streams are dominantly losing and recharge the shallow aquifers. The environmental isotope data demonstrated that recharge of the alluvial aquifers by surface water occurs by bank infiltration as well as diffuse recharge during high rainfall events.



11. Artificial Tracers Artificial tracer tests can be used to evaluate the extent to which aquifers interact with streams, providing information on groundwater flow paths, travel times, velocities, dispersion, flow rates and the degree of hydraulic connection (Flury and Wai, 2003;Otz et al, 2003). These tests involve the introduction of a tracer material or chemical and subsequent monitoring of its movement. This differs from environmental tracer methods which rely on the measurement and interpretation of background concentrations of the chemical constituents of water (such as major ions, stable or radioactive isotopes, refer Chapter 10). In artificial tracer tests, a substance is introduced and monitored to track the movement of water. Hence, the movement of the tracer should match that of the water flow regime, so it should not be effected by sorption onto geological material, changes in chemistry (such as pH or salinity), or degradation by physical or biological processes. The tracer should not affect the water flow regime, by changing fluid density or viscosity. The tracer should have low background levels and be able to be measured simply and cheaply to low detection levels. As tracer tests involve an introduction of a substance into the environment, the tracer should have low toxicological or pathogenic impacts. The tracer should be stable for the duration of the test but not be retained as residual material in the longer term. Dyes have had a long tradition of use as tracers, commencing with tracing the source of typhus epidemics in Europe (des Carrieres, 1883). Sulforhodamine B, Rhodamine B, Sodium Fluorescein and Rhodamine WT are popular due to low cost, easy detection to low limits with a fluorometer, and the potential for visualisation (Flury and Wai, 2003). Major ions such as chloride and bromide have been used as they behave conservatively and rarely sorb onto geological material. A lithium chloride solute tracer was used to estimate seepage fluxes in a headwater stream (Harvey et al, 1996). Organic compounds such as ethanol, benzoate and fluorinated benzoates have been proposed as tracers (Malcolm et al, 1980; Bowman and Gibbens, 1992). However, retardation and degradation is an issue in low pH conditions and with the presence of abundant clay, iron oxides or organic material (McCarthy et al, 2000; Jaynes, 1994). Flourescent polyaromatic sulfonates have been trialled in geothermal groundwater studies due to their resistance to thermal decay (Rose et al, 1998). The use of isotopes as artificial tracers tends to be limited due to radiation risks and the complexity of chemical analysis. Short lived isotopes such as selenate (75S) as well as deuterium (2H) are considered the most useful (Flury and Wai, 2003). In some studies, a particular pollutant of water is investigated, requiring the tracer to follow the fate of the pollutant rather than water flow. This is the case when non-pathogenic microorganisms are used to trace the transport of human pathogens. Solid and colloidal particles such as clubmoss spores have been used in studies in karstic aquifers, but analysis requires filtering and microscopic examination (Drew, 1968). The use of nanotechnology in hydrogeochemical studies, especially the application of chemical-specific nano-scale tracers developed by the biomedical industry has been suggested (Divine and McDonnell, 2005).


Various strategies can be used to undertake an artificial tracer test. The constant injection rate technique is commonly employed in the field to assure complete mixing of the injected tracer. A dye such as Rhodamine WT is released upstream of the study reach (see Figure 11.1). Using medical devices for controlling intravenous fluid injection, the dye injection can be started several hours before the beginning of the investigation. Water samples are collected regularly at different locations and the dye concentration analysed with a fluorometer. In this way, very low concentrations can then be measured far downstream. The samples are taken from the centre of the channel at each location. By analysing the dye concentration in the sample, the amount of water required to dilute the injected dye solution to the sample concentration can be determined. The amount of water includes the amount of stream flow passing the injection site, plus any groundwater contributions the stream has received between the injection and sample sites.

Figure 11.1: Dye tracer technique for assessing groundwater and surface water interaction in the field (Otz et al, 2003) A tracer injection trial can provide valuable insights into the rate and direction of groundwater movement near a surface water feature (Dahm and Valett, 1996). This involves establishing monitoring sites both within the aquifer and along the stream. This network can include:

(i.) Existing boreholes or piezometers in the vicinity of the stream; (ii.) Sampling pits dug near the stream and accessing the shallow watertable (these

should be refilled after the experiment); (iii.) Any springs or seepage areas evident in the area; (iv.) Specific sites to monitor the stream itself; (v.) Constructing minipiezometers both within and adjacent to the stream bed.

A tracer is then injected into a centrally located pit or piezometer and the time recorded. Each of the groundwater and stream monitoring sites is subsequently sampled to measure if the tracer is detected, and the time when this occurs. This is used with the distance between the injection site and the monitoring site to calculate groundwater velocity and direction.


An even simpler approach is to inject a small bolus of dye about 5-15 cm within the stream sediment using a syringe with a long cannula, or via a minipiezometer (Grimm and Fisher, 1984). The time of injection and also the time and position that the dye first appears at the sediment surface is recorded. The distance between the injection point and emergence point is used to calculate groundwater velocity and direction. If the dye does not emerge, dig up the injection site or check the minipiezometer for the status of the dye bolus. If the dye has disappeared this suggests that lateral or downward seepage conditions prevail, rather than upward. In a point dilution test, the tracer is added to a piezometer and the subsequent rate of dilution of the tracer within the piezometer is monitored. This data is used to estimate the groundwater velocity at the piezometer (Halevy et al, 1967; Gaspar, 1987). Electrical conductivity measurements were used to record the dilution of a KCl tracer added to piezometers at three contrasting Australian riparian and estuarine sites (Lamontagne et al, 2002). Displacement of the tracer by its improper release into the piezometer or due to the subsequent recirculation process was found to be the main technical difficulty. 11.1 Advantages and Disadvantages Artificial tracers are a useful investigation tool as the application and monitoring of the tracer can be designed and implemented in a controlled way. The amount of tracer introduced is known, allowing quantification of aquifer parameters and fluid transport properties. Specific processes in particular hydrogeological settings can be investigated by using appropriate tracers. The method is particularly useful in characterising groundwater flow in highly variable aquifers (such as fractured rock or karsts, refer Box 11.1) and in solute transport studies (such as contaminants and nutrients). Specific tracers can be used to track pollutants such as human pathogens, where the movement and fate of these pollutants may not match water flow. Tracers can be used to assess the significance of local geological features (such as faults, clay layers or cave systems) on stream-aquifer connectivity. Seepage can be assessed either qualitatively (such as visual inspection of the presence of dye or the use of charcoal-based detectors) or quantitatively (such as modelling of time-concentration breakthrough curves to derive travel time characteristics). Tracers can provide direct evidence for the movement of water between one point to another, which is easily understood by the public, regulatory agencies or the courts (Mull et al, 1988). Appropriate application of tracers requires careful planning and design with some pre-test knowledge of hydrogeology. Unanticipated short travel times can lead to high tracer concentrations being released to watercourses and potentially into public water supplies. This was the case when a borefield tracer test in 2003 unintentionally dyed red the water supply for about a million people (SS Papadopulos and Associates, 2004). Part of this planning is meeting any regulatory controls on the release of chemicals in the environment for public health or ecosystem protection. The performance of the tracer in matching water movement can vary with the hydrogeological setting. Dyes can have complex chemical interactions which tend to be pH-dependent or can be selectively sorbed with geological material. Sometimes, the dye mixture viscosity can be dramatically affected by variations in the ambient temperature, complicating the determination of flow rates. Tracer tests can have


overheads in terms of cost and time, particularly when investigating longer or slower groundwater flow paths. 11.2 Relevant Links USGS Rhodamine WT Reader http://smig.usgs.gov/SMIG/rhodamine_reader.html Joint International Isotopes in Hydrology Program (JIIHP) http://www.iaea.or.at/programmes/ripc/ih/jiihp_home.htm Box 11.1: Dye tracing tests of the Piora aquifer, Switzerland (Otz et al, 2003) In Switzerland, the Piora Valley is bounded on the north by the Cadimo Valley, to the southwest by the Leventina Valley, to the northwest by the Canaria Valley, and to the east by the St. Maria Valley. The Piora Valley is noted for losing streams and leaking lakes. Since construction of the hydroelectric power plant and dam at Lake Ritom over 80 years ago, dye tracing tests have been used to evaluate the hydrologic flow paths in aquifers of the Piora Valley. These dye tracing tests provided local information on the hydrogeology of the Piora Zone, but not a comprehensive understanding of the overall hydrodynamics of the Piora aquifer. The objective of the study by Otz et al, (2003) was to clarify and quantify water loss from major streams and lakes using additional dye tracing tests and to determine whether flow paths are regional in scale.

A variety of organic fluorescent dyes were used to evaluate the hydrogeologic flow system in the Piora Valley and the adjacent valleys in Switzerland. The nature and chemical characteristics of the dyes and standard tracing procedures are described in Kass (1988). Briefly, multiple dyes were placed in different parts of the Pioral flow system and the dye breakthrough curves were monitored at possible discharge points. From the breakthrough curves, maximal flow velocities for the dyes to move from places of introduction to re-emergence were calculated.

Results of seven dye tracing tests done from 1993-1997 showed that the direction of groundwater flow in the Piora aquifer is from the Pioral Valley to the Ri di Lareccio springs in the Santa Maria Valley, and even further east to the di Campo Valley. Dye tracing tests show that a major sinkhole in the Piora Valley, Calderoni sinkhole, is located precisely on the water divide where subsurface flow in the Piora Valley and surface water diverge and move in the opposite direction. The dye tracing results also showed no hydraulic connection between surface water in the Piora Valley and the famous Pertusio spring, located in the upper Santa Maria Valley. Only a small amount of dye from the two dye tracing tests done in 1993 and 1997 entered an exploratory gallery built to test the variability of the AlpTransit tunnel, being built in competent rock under the Triassic Piora aquifer, effectively perched above.


12. Temperature Studies In a connected system, the exchange of water between the stream and shallow aquifer plays a key role in influencing temperature not only in streams, but also in their underlying sediments. As a result, analysis of subsurface temperature patterns can provide information about seepage flux. Studies, notably in North America, have used temperature monitoring in the stream and underlying sediments as a screening tool for identifying gaining and losing reaches (Silliman and Booth, 1993; Stonestrom and Constanz, 2003). Recently, heat as a tracer has been demonstrated to be a robust method for quantifying surface water-groundwater exchanges in a range of environments, from perennial streams in humid regions (Lapham, 1989; Silliman and Booth, 1993) to ephemeral channels in arid locations (Stonestrom and Constanz, 2003). Logging devices that measure temperature at specific time intervals and store the information in memory can be installed both within the stream and at different depths in the sediments below the stream bed. The sensors most often employed are thermocouples, thermistors, resistance temperature devices and integrated circuit sensors (Figure 12.1). The characteristics of each type of temperature logging equipment along with their advantages and disadvantages and installation methods are

presented in Stonestrom and Constantz (2003). Figure 12.1: Common temperature sensors used to measure sediment and stream temperatures (Stonestrom and Constanz, 2003) The hydraulic transport of heat enables its use as a tracer with temperature monitoring especially suited for delineating fine-scale flow paths. The heat tracer method has been used to estimate groundwater velocity and aquifer hydraulic properties, and to identify areas of recharge and discharge (e.g. Bouyoucos, 1915; Suzuki, 1960; Lee, 1985; Lapham, 1989; Silliman and Booth, 1993; Brewster Conant Jr, 2004). One way of using heat tracing in stream-aquifer studies is to compare the temporal patterns evident in stream and shallow sediment temperature. Stream temperatures have a characteristic diurnal pattern overprinting seasonal trends, being influenced by changes in solar radiation, air and ground temperature, rainfall and stream inflows that include groundwater discharge (Sinokrat and Stefan, 1993). These diurnal variations in temperature in the near-stream environment are often large and rapid, providing a clear thermal signal that is easy to measure. In contrast, the temperature of regional groundwater tends to be relatively constant at the daily scale. The movement of heat between surface water and groundwater systems is both advective (associated with


fluid movement) and conductive (through the static solid/liquid phase). Ignoring the effect of insitu sources of thermal energy (such as from biological activity), the temperature pattern in the shallow stream sediment profile can be used to evaluate seepage flux. The temperature signatures for three potential forms of stream-aquifer connectivity (gaining, losing and neutral) have been hypothesised (Silliman and Booth, 1993). In gaining stream reaches, the hydraulic gradient is upward as indicated by a groundwater level in piezometers that is higher than the stream stage (Figure 12.2a). Although the stream has a large diurnal temperature variation, the shallow sediment has only a slight or no diurnal variation. The downward propagation of any surface temperature effects is moderated by water that is flowing up from depths where temperatures are constant at the daily time scale. At any given depth beneath the streambed, higher flows of groundwater to the stream lead to smaller variations in sediment temperature while smaller flow leads to larger variations. Consequently, shallow installation of temperature equipment is necessary to characterise gaining stream reaches, in order to detect significant temperature variations. In losing stream reaches, the stream stage is higher than the groundwater level so that the potential hydraulic gradient is downward. This downward flow of water transports heat by advection from the stream, resulting in deeper propagation of diurnal temperature fluctuations into the sediment profile (Figure 12.2b). As a consequence, deeper installation of temperature equipment (inside the piezometer or beneath the stream bed) is necessary for losing streams to be characterised. Losing streams also tend to have larger daily temperature fluctuations than gaining reaches, due to the absence of any moderating effect from groundwater inflow (Constanz, 1998).

Figure 12.2: Temperature variation of groundwater and stream under gaining (a) and losing stream (b) conditions (adopted from Stonestrom and Constanz, 2003) In neutral reaches of the stream, thermal conduction will control stream sediment temperatures. This means that sediment temperature can vary due to changes in surface water temperatures, and will have an average that is between that of the surface water and groundwater. The distinct temperature signal of episodic infiltration associated with ephemeral streams has also been characterised (Ronan et al, 1998; Constanz et al, 2002).


Estimating the water exchange between the stream and shallow aquifer requires knowledge of the hydraulic and thermal conductivity of the material and the hydraulic gradient as defined by the stream stage and groundwater level. Numerical models of heat flow, such as VS2DH (Healy and Ronan, 1996) and SUTRA (Voss, 1984) can be used to quantify seepage fluxes. These can supplement and help calibrate more traditional groundwater flow models. In particular, temperature modelling can help constrain estimates of hydraulic conductivity which can vary over several orders of magnitude, as the thermal conductivity of sediments has a much smaller range of potential values. Stream sediments composed of sand and gravel can have a hydraulic conductivity six orders of magnitude higher than clay. In contrast, the thermal conductivity of porous materials depends upon the composition and arrangement of the solid phase with the potential range in thermal conductivity between coarse grained sand (2.2 W/m oC) and clay (1.4 W/m oC) being much smaller than that for hydraulic conductivity (Stonestrom and Constantz, 2003). The work of Bravo et al (2002) is a recent example of using temperature data to constrain estimates of boundary fluxes and hydraulic conductivity in a groundwater flow model for a wetland system. Fluctuations in temperature can also directly influence seepage rates due to its influence on water density. The hydraulic conductivity of the stream bed is both a function of the porous medium and the water itself, the latter in terms of density and dynamic viscosity. Hence, transmission rates through the sediment bed can increase with increased water temperature. This process was used to explain diurnal variations in seepage flux in losing reaches of a small alpine stream (Constanz, 1998). 12.1 Advantage and Disadvantages The use of temperature as a hydrologic tracer has several advantages over other field methods. Temperature logging devices are robust, simple and relatively inexpensive and available for various scales of measurement. The temperature signal arrives naturally and the temperature data are immediately available for inspection and interpretation. Temperature monitoring can be used as a screening tool for identifying gaining and losing stream reaches. Such a screening method can be valuable both as a rapid investigative tool for small studies and as a precursor to more detailed studies such as the design/installation of a groundwater monitoring network. Once installed, loggers can also provide useful time-series data that can provide information on seasonal changes in seepage flux. Temperature studies are particularly useful in defining small-scale flow paths, such as associated with stream banks or sand bars (Stonestrom and Constanz, 2003). Despite these advantages, this method has some potential limitations. Interpretation of the temperature data can be ambiguous when viewed in isolation. It is recommended that temperature monitoring be used in conjunction with other methods such as minipiezometers, seepage meters or hydrographic analysis when interpreting stream-aquifer connectivity. Also, the temperature measurement is at a point in space and many measurements may be required to obtain information on spatial variability. It can be difficult to separate localised effects (such as associated with weirs or shallow throughflow) from the broader seepage domain.


It is suggested that temperature loggers can be readily and cheaply incorporated into existing hydrographic networks to provide a supplementary dataset for understanding stream-aquifer connectivity. This is because the water level data can indicate the potential seepage direction and the temperature data can help estimate the magnitude of the seepage. It is also recommended that the existing temperature logging used to calibrate pressure transducers for monitoring water levels be upgraded to sufficient accuracy for heat transfer studies. Temperature monitoring would be particularly useful in estimating infiltration rates in Australian ephemeral streams, where conventional water level recording and hydrographic analysis is problematic. Routine recording of temperature data also has relevance to the investigation and management of aquatic ecosystems, notably within the hyporheic zone. 12.2 Data Availability Stream temperature data can be routinely collected as part of the stream gauging network, as pressure transducers that measure stream level require such data for calibration. Brodie et al (2007) outlines the key hydrographic databases across Australia, mostly maintained by State and Territory water management agencies. However, this temperature data is not commonly made available and may not be of adequate resolution for heat transfer studies. Groundwater temperature can also be recorded in the same way, as pressure transducers are a common logging device for groundwater levels. It would be rare for both groundwater and surface water temperatures to be monitored in close enough vicinity at a site for application in heat transfer studies. For instance, temperature monitoring in the shallow stream bed sediments at stream gauging stations is not routinely done. 12.3 Relevant Links US Geological Survey Circular 1260 Heat as a tool for studying the movement of groundwater near streams http://pubs.water.usgs.gov/circ1260


Box 12.1 Trialling of temperature monitoring in the Border Rivers Monitoring of temperature in the stream sediment (0.25 – 1.2 m depth) as well as the stream itself was used to investigate groundwater-surface water interactions in the Border Rivers catchment in the Murray-Darling Basin (Baskaran et al, 2005). When interpreted with hydrographic and hydraulic conductivity data, the temperature monitoring provided useful insights into the spatial and temporal variability of stream-aquifer connectivity. At one site, sediment temperatures fluctuated with the diurnal temperature variation of the stream, reflecting river leakage (Figure 12.3). No diurnal signal was detected in the sediment temperatures at other sites, which is a typical indicator of gaining conditions. However, with water level measurements indicating negative gradients and the stream sediments dominated by clay at these sites, this lack of sediment temperature variability is interpreted to reflect very low rates of downward seepage. At one site, a transition from gaining to losing conditions was observed through time. In the field trials, operational issues such as timing the monitoring to coincide with reasonable diurnal variations of stream temperature, the requirement of understanding the shallow stratigraphy of the stream bed, and separating out localised effects (such as from weirs) were highlighted. The trials also highlighted that interpretation of the temperature data can be ambiguous when viewed in isolation.

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Figure 12.3 Observed stream and sediment temperatures downstream of Goondiwindi Weir during (a) high flow and (b) low flow seasons (Baskaran et al, 2005)

Stream and sediment temperatures recorded at the Goondiwindi site during high flow (summer) and low flow (winter) seasons are depicted in Figure 12.3. The sediment temperature shows regular fluctuations that relate to the diurnal stream signal. Such fluctuations were evident at different depths in the sediment during the November 2004 high flow conditions (Figure 12.3a). The sediment temperature measured at shallow depth (0.25 m) recorded a diurnal signal varying between 24.3-25.8oC, with the peaks lagging by 4 hours from the corresponding maxima in daily stream temperature. The dramatic change in shallow sediment temperatures after 13 November 2005 was due to premature removal of the logger from the stream bed, thereafter effectively measuring the stream temperature. The diurnal pattern deeper in the profile (0.5m) is more subtle in amplitude with a time lag of about 6 hours. This reflects the time taken to transfer heat by both conduction and advection downwards through the sediments.

It should be noted that the diurnal variation in sediment temperature (0.5 m depth) is significantly greater during high flow than in the low flow season. This is most likely due to the fact that the stream temperature diurnal fluctuation was high (6-7oC) during the high flow summer but very low (0.5-2oC) during the low flow winter. Temperature fluctuations evident in the sediment profile indicate that the stream is losing water to the groundwater system, confirmed by measurement of the head gradient between the river stage and shallow groundwater level.

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13. Water Budgets A common approach to investigating seepage flux between a stream and underlying aquifer is to measure stream flow at specific points. These measurement sites subdivide the stream into reaches and a water budget is estimated for each reach, accounting for inputs such as tributary flows and outputs such as evaporative losses and diversions. The difference between inflows and outflows is then attributed to the interaction between the stream and the underlying aquifer. When applied to a defined reach, the groundwater flux (Qgw) is estimated from:

∑ ∑−+−= inoutupdngw QQQQQ Equation 13.1

where Qdn is the flow at the downstream end of the reach, Qup is the flow at the upstream end, Qout are outputs from the reach (such as distributaries, evaporation and extraction) and Qin are inputs to the reach (such as direct rainfall, runoff, tributaries, irrigation drainage and sewage outfall). This follows the convention that a positive Qgw indicates a net input of groundwater to the reach. A negative Qgw indicates a net loss of surface water to the groundwater system and is commonly termed a transmission loss. Although the method is simple, involving the quantification of all fluxes to, or from, the river, it is difficult to apply in many cases. The method relies on the accurate measurement of surface water flow, as well as appropriately accounting for all the other gains and losses evident for the reach. The uncertainties associated with the flow measurements and estimates for water balance components such as unmetered extraction, evaporation, ungauged tributary flows, overbank flooding losses and flood return flows can often exceed the magnitude of the seepage flux being estimated. Stream flow measurement errors can be +25% during high flow conditions, and from –50 to +100% for flash floods in semiarid catchments (Lerner et al, 1990). This means that the reach must be relatively long so that the cumulative volume of seepage exceeds the errors in the water balance. A specific type of water balance technique called a pondage test is commonly used for man-made structures such as irrigation supply channels. A reach of the channel is isolated by placing embankments at each end, and filled with water to (or higher than) the operating level. After correcting for rainfall and evaporation, the subsequent decline in water level is attributed to seepage losses in the underlying aquifer. Alternatively, water is added to maintain a constant water level, and the added volume used in the seepage calculations. 13.1 Stream Flow Measurement The water budget method relies on accurate measurement of stream flow at the end points of the reach being considered. Stream flow or discharge is the rate at which a volume of water passes through the cross section of the stream per unit time, and as such has SI units of cubic metres per second (m3/s) or cumecs. Other common units used in Australia include megalitres per day (ML/d) or litres per second (L/s). Water budgets are commonly estimated for a specified river reach defined between two gauging


stations where stream flow is routinely measured. This typically involves the monitoring stream level and defining a relationship between level and flow. The stage or the height of the water surface from a benchmark is monitored as a time-series whilst stream discharge (Q) is measured at a less regular basis, particularly at various flow regimes. The relationship between stream stage and flow is called the flow rating curve and typically takes the form:

bZhaQ )( −= Equation 13.2

where h is the measured gauge height, Z is the gauge height at zero-flow conditions, and a and b are best-fit coefficients using non-linear regression. In this way, a time-series of stream flow called a hydrograph can be generated from any continuous monitoring of stream level at the site. There are also a range of methods for the direct measurement of stream flow. These can be used in detailed surveys along the extent of the stream reach being considered. The common methods for measuring stream flow are volumetric analysis, the velocity-area method, the slope-area method, dilution gauging and thin-plate weirs. 13.2 Volumetric Analysis Volumetric analysis or the ‘bucket and stopwatch method’ involves the measurement of the time taken for a container of known capacity to be filled. This is a simple method for measuring small streams where all of the flow can be concentrated (such as naturally within a rock cascade or by constructing a temporary dam with a pipe). The method can also be used in larger streams which have flow concentrated or partitioned by culverts, pipes or weirs. A short, wide container is more suitable to fit under the falling water, than a tall, narrow one. The container volume should be such that it takes at least 3 seconds to fill. It is important that the bucket is held upright and that multiple (>3) readings are taken to reduce measurement error. An alternative approach is to hold down and open a heavy-gauge plastic garbage bag on the stream bed, time its filling and empty the bag contents into a measuring container (Hauer and Lamberti, 1996). The bag can also be weighed and the volume calculated if the density of the water is known. Stream discharge (Q) in L/s is calculated as

Q = V/t Equation 13.3

where V is the contained volume (L) during time t (s). 13.3 Velocity-Area Method In the velocity-area method, stream velocity and water depth measurements are taken along a transect perpendicular to the stream. Total discharge (Q) is calculated by integrating the stream velocities with the cross sectional area of the stream profile defined by the transect. Different types of current meters are available to measure stream velocity. Propeller-type meters have a horizontally aligned vane that rotates in proportion to the stream velocity. The number of rotations can be recorded visually, audibly or digitally. Cup-type meters work on the same principle, but the vane is oriented vertically.


Electromagnetic meters measure the voltage induced when a conducting fluid flows through a magnetic field. These tend to be used in coastal and marine studies due to the high conductivity of saline water. Ultrasonic meters use sound wave propagation in water to measure water velocity. Some versions use the impedance on the time for an ultrasonic wave emitted from one side of the river to reach a receiver on the other side. Other versions use the Doppler Effect by measuring the wavelength of ultrasonic waves reflected off suspended particles in the stream flow. The procedure is to: (i.) Choose a suitable site along the stream with a straight reach, uniform laminar

flow conditions and relatively constant depth and width. Sites with extreme turbulence, protruding obstructions, eddies, stagnant zones or divided channels should be avoided;

(ii.) Set up a tagline consisting of a tape measure perpendicular across the stream to be used for locating the velocity/depth measurements. Measurements are taken along 10-20 verticals across the stream transect. Each vertical should partition stream flow equally, so that verticals should be closer together where water is faster or deeper;

(iii.) At each site, use a current meter to measure stream velocity and a graduated pole to measure stream depth. Typically, flow is measured at a depth considered to reflect average velocity conditions (0.6 of the stream depth measured downward from the surface). Other approaches include two measurements being taken and averaged for each vertical, at 0.2 and 0.8 of the water depth (refer Table 13.1);

(iv.) The stream flow can be calculated using the mid-section method:

2/))((1

1111∑=

+++ +−=n

iiiiii YUYUXXQ Equation 13.4

where the Xi are the distances to successive measurement points along the transect, where stream velocity (Ui) and water depth (Yi) are measured, starting with X1 being the initial point on one bank and Xn being the final measuring point on the opposite bank.

Table 13.1: Different procedures for determining mean velocity at a vertical (after Gordon et al, 2004)

Number of Points in Vertical

Depth of Measurement (from water surface)

Application Mean Velocity Equation

1 0.6D D < 0.5 m or quick measurement required V = V0.6 2 0.2D and 0.8D Preferable where size of meter allows

(D>0.5 m) V = 0.5(V0.2 + V0.8)

3 0.2D, 0.6D and 0.8D Irregularites distort the velocity profile and depth is sufficient

V = 0.25(V0.2 + V0.8 +2V0.6)

1 Just below surface (~0.6m) High or fast-flowing conditions where lowering meter is difficult

V = kVsurf where k typically = 0.85

Many Range of depths including 0.2D, 0.6D and 0.8D

When high precision required or velocity profile is of interest

Integration of the area bounded by the profile divided by D

D = vertical distance between water surface and streambed, measured downwards A simplified version of this method that estimates the water velocity by timing the movement of a float over a set distance can be useful in obtaining quick approximations of stream flow, particularly during floods (Hauer and Lamberti, 1996). This method is to:


(i.) Estimate the cross sectional area (A) of the stream by using a steel tape to measure both the stream width and a few depth measurements;

(ii.) Measure a stream reach of adequate length (L) to allow a travel time of over 20 seconds for the float. Mark the starting and finishing points of this reach with a stake or a string across the stream;

(iii.) Choose a float that is only slightly buoyant to reduce wind effects. An orange, chunk of ice, half-filled fishing float or a waterlogged stick are good options;

(iv.) Place the float upstream of the defined starting point of the reach, so that the float is travelling at the velocity of the stream by the time it reaches the starting marker. Measure the time that the float takes to travel between the upstream and downstream markers using a stopwatch. An average time (t) is obtained by taking multiple readings;

(v.) Use a correction factor (k) to account for surface velocity being faster than the average stream velocity so that:

Q = AkL/t Equation 13.5

The correction factor k generally varies between 0.80 for rough stream beds to 0.90 for smooth, with 0.85 most commonly used.

Instead of a float, a slug of coloured dye can be introduced into the stream and its movement down the stream reach measured (Gordon et al, 2004). This would tend to reflect an average stream velocity rather than the surface velocity that a float is influenced by, so a correction factor (k) is not required. However, due to dispersal during flow it can be difficult to define the centre of the dye slug. 13.4 Slope-Area Method The Slope-Area method indirectly estimates stream flow and can be applied when the use of current meters or floats are not practical. The method is based on Mannings Equation:

Q= 1/n(AR2/3S1/2) Equation 13.6

which involves estimating the hydrological parameters of:

(vii.) The cross sectional area of the stream (A); (viii.) Mannings n which is an index of the roughness of the stream bed, and

increases with turbulence and flow retardance; (ix.) The hydraulic radius (R) which is the ratio of the cross section area (A) of the

stream to its wetted perimeter (which is the cross-sectional distance along the stream bed and banks that is in contact with the water);

(x.) The energy slope (S) which is the change in elevation of the stream over a specified distance. This is assumed to be parallel to the top of the water surface or the base of the stream bed.

A similar approach is used in Chezy’s Equation:

Q=AC(RS)1/2 Equation 13.7

where a channel roughness index (C) is also used.


The procedure is to: (i.) Choose a stream reach that is relatively straight with uniform flows and where

the slope of the water surface and the stream bed are relatively parallel. The stream reach should be greater than six times the mean channel width;

(ii.) Set up a measuring tape across the stream and determine the stream profile by taking a series of depth measurements along intervals along the tape. These depth-distance measurements are used to calculate the cross-sectional area and the wetted perimeter. The method can be applied to particular water levels of interest such as bankfull depth or recent flood heights rather than the existing water level at the time of measurement. These levels can be flagged or pegged along the reach and used in the analysis;

(iii.) Using a staff and a surveyors level, measure the stream bed elevations and water surface elevations for points about 20 m upstream and downstream of the transect. The energy slope (S) can be calculated as the ratio of the difference in water surface elevations to the distance between these points;

(iv.) Estimate the roughness index (n) for the water level considered for the reach. Values for different settings have been suggested (Table 13.2). Also, by walking along the stream reach and observing and categorising the major factors that define channel roughness, these can be combined to estimate Mannings n (Table 13.3).

Table 13.2: Mannings n values for small natural streams (Gordon et al, 2004; after Chow, 1959)

Channel Description Minimum Normal Maximum Lowland Streams

(a) Clean, straight, no deep pools 0.025 0.030 0.033 (b) Same as (a) but more stones and weeds 0.030 0.035 0.040 (c) Clean, winding, some pools and shoals 0.033 0.040 0.045 (d) Same as (c) but some weeds and stones 0.035 0.045 0.050 (e) Same as (c) at lower stages with less effective slopes and sections 0.040 0.048 0.055 (f) Same as (d) but more stones 0.045 0.050 0.060 (g) Sluggish reaches, weedy deep pools 0.050 0.070 0.080 (h) Very weedy reaches, deep pools or floodways with heavy stand of timber and underbrush

0.075 0.100 0.150

Mountain Streams (no vegetation in channel, banks steep, trees and brush on banks submerged at high stages (a) Streambed consists of gravel, cobbles and few boulders 0.030 0.040 0.050 (b) Bed is cobbles with large boulders 0.040 0.050 0.070


Table 13.3: Calculation of Mannings n from field observations (Hauer and Lamberti, 1996) where n = (n0 + n1 + n2 + n3 + n4)m

Channel Condition Value Additive Factors Material Involved n0 Earth 0.020 Rock Cut 0.025 Fine Gravel 0.024 Coarse Gravel 0.028 Cobble 0.030-0.050 Boulder 0.040-0.070 Degree of Surface Irregularity n1 Smooth 0.000 Minor (slightly eroded or scoured) 0.005 Moderate (slumping) 0.010 Severe (badly slumped, eroded banks) 0.020 Variation in channel cross section n2 Gradual 0.000 Alternating occasionally 0.005 Alternating frequently 0.010-0.015 Effect of Obstructions (debris deposits, roots, boulders) n3 Negligible (few scattered) 0.000 Minor (<15% of area) 0.010-0.015 Appreciable (15-50% of area) 0.020-0.030 Severe (>50% of area, turbulent) 0.040-0.060 Vegetation n4 None or no effect 0.000 Low (grass/weeds) 0.005-0.010 Medium (brush, none in stream bed) 0.010-0.025 High (young trees) 0.025-0.050 Very High (brush in streams, full grown trees) 0.050-0.100 Multiplicative Factors m Degree of Meandering Minor (sinuosity 1.0-1.2) 1.000 Appreciable (sinuosity 1.2-1.5) 1.150 Severe (sinuosity >1.5) 1.300


13.5 Dilution Gauging Dilution gauging methods measure stream flow on the basis of the rate of dispersal of an introduced tracer with the concentration of the tracer monitored downstream. The method is used in streams with highly turbulent flow conditions where conventional stream velocity measurements are difficult to apply. Chemical tracers such as table salt (NaCl) can be monitored using an electrical conductivity (EC) meter or a specific ion electrode. The mixing of fluorescent dyes such as Rhodamine WT can be monitored using a fluorometer. Alternatively, water samples can be collected at different times for laboratory analysis. The slug-injection method involves introducing a tracer of known volume and concentration into the stream in a single increment or slug. The concentration of the tracer and the time of measurement are monitored at a downstream location, before and after the slug is introduced. The resulting concentration hydrograph maps the dispersed passage of the tracer slug at the measurement point, with stream discharge (Q) represented as the area under the curve and calculated by integration:

∫ −

∀=

2

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t

t

dtcc

cQ Equation 13.8

Where ∀ is the known volume of tracer (L), ct is the tracer concentration of the introduced solution, c0 is the background concentration in the stream, c is the changing tracer concentration as measured downstream, and t1 and t2 are the initial and final times of measurement. As the name suggests, the constant-injection method relies on the tracer solution being introduced into the stream at a constant rate. The tracer concentration at the downstream measurement point should rise and stabilise to a constant level. The stream flow (Q) can be calculated from

tt Q

cccc

Q)()(

100001

1

−−

= Equation 13.9

where c1 is the stabilised tracer concentration at the downstream measurement point, and Qt is the tracer injection rate. 13.6 Thin Plate Weirs Thin-plate weirs are installed within the stream to restrict the channel and to regulate and stabilise the water flow. Stream flow is proportional to the height of water built up on the upstream side of the weir. The geometry and operation of the weir is specified so that flow can be assessed using standard calibration tables. The two types of weir, namely v-notch and rectangular, are based on the geometry of the weir outlet.


13.7 Advantages and Disadvantages A water balance provides an average net seepage flux for the entirety of the surface water feature being considered e.g. for a defined stream reach. This is in contrast to other methods (such as seepage meters), which can only measure seepage at a point at a time. Simple water balances can be estimated relatively quickly and cheaply (using existing stream flow monitoring) to derive an initial rough estimate of the direction and magnitude of seepage. Stream flow measurements at multiple points along a stream can help target hotspots in terms of seepage losses or gains. Box 13.1 gives an example of this approach. The method can also use stream flow monitoring that is commonly collected and publicly available. As stream flow gauging is typically time-series data, temporal changes in seepage can be estimated, highlighting seasonal or longer-term variations. However, measurement errors in stream flow measurements can exceed the magnitude of the seepage flux, particularly for relatively short reaches. The seepage estimates can be misleading if a component of the water balance (eg extraction, evaporation) is not adequately accounted for or effectively measured. The method can be time consuming and expensive depending on how the water balance components are measured. No information is provided on the spatial variability of seepage along the reach being investigated. 13.8 Data Availability The main data requirement for the water budget approach is reliable stream flow gauging data. Brodie et al (2006) outlines the significant surface water monitoring databases in Australia, highlighting the State and Territory agencies involved in water management as the main data custodians. Currently, the Water Resources Station Catalogue http://www.bom.gov.au/hydro/wrsc developed by the Bureau of Meteorology provides a national inventory of the river gauging stations, as well as rainfall and evaporation stations, across the country. Estimating the water budget components of the reach including outputs (such as distributaries, evaporation, extraction) and inputs (such as direct rainfall, runoff, tributaries, irrigation drainage, sewage outfall) typically requires combining data from different sources. Climate data such as rainfall and evaporation are available from the Bureau of Meteorology (http://www.bom.gov.au).

13.9 Relevant Links ANCID Channel Seepage Management Tool http://ancid.org.au/seepage/index.html ANCID Know the Flow http://www.ancid.org.au/ktf/index.html BOM Water Resource Station Catalogue http://www.bom.gov.au/hydro/wrsc Water Resources Applications Software (CGAP) Channel Geometry Analysis Program http://water.usgs.gov/software/cgap.html USGS software for the analysis, interpretation, and quantification of physical properties of an open-channel reach


Water Resources Applications Software (NCALC) http://water.usgs.gov/software/ncalc.html USGS Manning's n value calculation program Water Resources Application Software (SAC) Slope Area Computation http://water.usgs.gov/software/sac.html USGS standardised procedure for computing discharge by the slope-area method

Box 13.1 Stream flow survey on the Alstonville Plateau, NSW Stream flow was measured at specific points on the Alstonville Plateau in November, 2004 (Brodie et al, 2005). Autumn to early summer is typically the driest part of the year with stream flow reflecting baseflow conditions. The measurement sites were selected on the basis of ease of access, with most coinciding with road crossings. Where flow was reasonable, measurements were taken using a propeller-type current meter, with a data logger to store and process the distance, stage and velocity readings. Where flow was too low for the meter, measurements were taken using a bucket and stop watch. The variation of stream flow for the plateau stream is shown in Figure 13.1. Very low flow conditions prevailed in the three eastern subcatchments (Yellow, Gum and Youngmans), with the streams not flowing in their upper reaches. Stream flow in the western subcatchments (Marom and Tucki Tucki) were significantly higher. This is partly due to the rainfall events during the survey of these two catchments. This rainfall appears to have had a greater impact on the Tucki Tucki catchment, with relatively high stream flow in the upper part of the catchment. The urbanised nature of this area around Goonellabah with a larger proportion of pavement and roof area, results in a greater run-off response. The stream flow measurements in the Tucki Tucki subcatchment taken about a week after the rainfall event, are more indicative of baseflow conditions. In general, the larger baseflow regime in the western catchments (Tucki Tucki and Marom) can be attributed to the combination of:

(i.) The larger catchment size increasing access to greater run-off and discharge from the shallow groundwater system;

(ii.) These catchments being to the northwest and down gradient of the flow paths for the deeper groundwater system. With these streams being relatively incised, there is greater opportunity for the deeper aquifers to discharge;

(iii.) The greater level of both surface water and groundwater extraction (as represented by density of water licences in Figure 13.1) in the eastern subcatchments of Youngmans and Gum Creek.

On the plateau itself, the streams are gaining reaches. Youngmans and Marom Creeks in particular show the characteristic downstream increase in stream flow with the aggregation of groundwater discharge. However, the streams appear to lose water where they initially flow onto the coastal floodplain below the plateau escarpment. For example, Marom Creek flows decreased from 28 L/s to 19 L/s along this lower reach (Figure 13.1). The initial flow measurements for this particular reach for the Tucki Tucki were more dramatic, with a reduction from 211 L/s to 11 L/s. These readings may be effected by the fact that the runoff peak due to the rainfall event had not reached the downstream site at the time of measurement. Stream flows at these sites about a week after the rains were measured to be practically identical (46 L/s and 45 L/s).


Figure 13.1: Stream flow measurements taken in November 2004 on the Alstonville Plateau (Brodie et al, 2005). In the Tucki Tucki catchment flow measurements taken during a rainfall event presented in red, about a week later in pink.


14. Acknowledgements The conjunctive water management framework described in this report was produced as part of the Managing Connected Water Resources project. The objective of the project is to progress a more coordinated approach to the management of surface water and groundwater resources. The three main components of the project are the Connected Water website, http://www.connectedwater.gov.au/, a national workshop on managing connected water resources (Fullagar, 2004) and work in two catchments (Border Rivers and Lower Richmond) to test methodologies and approaches This is a collaborative project between the Bureau of Rural Sciences, Australian Bureau of Agriculture and Resources Economics, The Australian National University, NSW Department of Natural Resources and Qld Department of Natural Resources and Water. The project has been partly funded by the Natural Heritage Trust, the National Landcare Programme and the Australian Research Council (ARC-Linkage).



15. References

Alayamani MS, Sen Z, 1993. Determination of hydraulic conductivity from complete grain-size distribution curves. Ground Water 31 (4) 551-555.

Alexander MD and Caissie D, 2003. Variability and comparison of hyporheic water temperatures and seepage fluxes in a small atlantic salmon stream. Ground Water 41: 72-82.

Allen DA 2005. Electrical conductivity imaging of sediment beneath the Border Rivers – NSW/Qld border, Australia. Research Report 046606, Allen Hydrogeophysics, for Bureau of Rural Sciences.

Allen DA and Merrick NP, 2005. Surface water/groundwater interaction investigation using a towed geo-electric array. Conference Proceedings. Irrigation Association of Australia.

Allison, GB, 1988. A review of some of the physical, chemical and isotopic techniques available for estimating groundwater recharge. In Estimation of natural groundwater recharge, ed. I. Simmers, pp 49-72. reidel, Dordrecht.

ANCID, 2000. Open channel seepage and control. Vol 1.1 Literature review of channel seepage identification and measurement. Australian National Committee on Irrigation and Drainage. Prepared by Sinclair Knight Merz.

Atech, 2000. Water monitoring in Australia. Report for Environment Australia and National Land and Water Resources Audit.

Barnes BS, 1940. Discussion of analysis of runoff characteristics. Trans ASCE 105: 106. Baskaran S, Budd KL, Habermehl R and Carter A, 2004. Groundwater-surface water

interactions in upper Lachlan valley: preliminary investigation using hydrochemical and stable isotopes. In: Conference Proceedings 9th Murray Darling Basin Groundwater Workshop Bendigo, Victoria 17-19 February 2004.

Baxter C, Hauer RF and Woessner WW, 2003. Measuring groundwater-stream water exchange: New techniques for installing minipiezometers and estimating hydraulic conductivity. Transactions of the American Fisheries Society 132:493-502.

Belanger TV and Montgomery MT, 1992. Seepage meter errors. Limnology and Oceanography 37:1787-1795.

Belanger TV and Walker RB, 1990. Ground water seepage in the Indian River Lagoon, Florida. In: Tropical Hydrology and Caribbean Water Resources: Proceedings of the International Symposium on Tropical Hydrology and Fourth Caribbean Islands Water Resources Congress. American Water Resources Association. 367-375.

Bertin C, Bourg CM, 1994. Radon-222 and chloride as natural tracers of the infiltration of river water into an alluvial aquifer in which there is significant river/groundwater mixing. Environmental Science and Technology. 28, 794-798.

Bierwirth PN, 1997. The use of airborne gamma-emission data for detecting soil properties. Proceedings of the Third International Airborne Remote Sensing Conference and Exhibition. Copenhagen, Denmark.

Blanchfield PJ and Ridgway MS, 1996. Use of seepage meters to measure groundwater flow at brook trout redds. Transactions of the American Fisheries Society 125:613-818.

Bokunieicz H and Pavlik B, 1990. Groundwater seepage along a barrier island. Biogeochemistry 10:437-444.

Boughton WC, 1988. Partitioning streamflow by computer. Inst. Eng. Civ. Eng. Trans. CE30(5), 285-291

Boughton WC, 1995. Baseflow recessions. Australian Civil Engineering Transactions CE37 (1), 9-13


Boussinesq J, 1904. Recherches theoretique sur l’ecoulement des nappes d’eau infiltrees dans le sol et sur le debit des sources. J. Math. Pure Appl. 10 (5): 5-78.

Bouwer 1978 Bouwer, H., Groundwater Hydrology, McGraw-Hill, Inc., New York, NY, 280 p., 1978.

Bouyocous G, 1915. Effects of temperature on some of the most important physical processes in soils. Michigan College of Agriculture Technical Bulletin 24, 63p.

Bowman, R.S., and J.S. Gibbens. 1992. Difluorobenzoates as nonreac-tive tracers in soil and ground water. Ground Water 30:8–14.

Brassington R, 1988. Field Hydrogeology. Geological Society of London Handbook Series. Open University Press.

Bravo HR, Feng J and Hunt RJ, 2002. Using groundwater temperature data to constrain parameter estimation in a groundwater flow model of a wetland system. Water Resources Research 38(8): 1153.

Brodie RS, Baskaran S, Ransley T and Spring J, 2005. The seepage meter: Progressing a simple method of directly measuring water flow between surface water and groundwater systems. International Association of Hydrogeologists Conference, Auckland. NZ

Brodie R, Sundaram B, Tottenham R, Hostetler S and Ransley T, 2007. An adaptive management framework for connected groundwater-surface water resources in Australia. Bureau of Rural Sciences, Canberra.

Brodie RS, Green R, 2002. A hydrogeological assessment of the fractured basalt aquifers on the Alstonville Plateau, NSW. Bureau of Rural Sciences, Canberra.

Browne E, and Firestone, RB 1999. Table of isotopes, 8th Edition. Wiley and Sons, New York.

Busenberg E and Plummer LN, 2000. Dating groundwater with sulphur hexafluoride: Natural and anthropogenic sources of sulphur hexafluoride. Water Resources Research 36:3011-3030.

Byrnes RP and Webster A, 1981. Direct measurement of seepage from earthen channels. Technical Paper No 64. Australian Water Resources Council, Canberra.

Cable JE, Burnett WC, Chanton JP, Corbett DR and Cable PH, 1997. Field evaluation of seepage meters in the coastal marine environment. Estuarine, Coastal and Shelf Science 45:367-375.

Cas RAF, Wright JV, 1995. Volcanic Successions , Modern and Ancient. Chapman & Hall London. 528pp.

Cey EE, Rudolph DL, Parkin GW and Aravena R, 1998. Quantifying groundwater discharge to a small perennial stream in southern Ontario, Canada. Journal of Hydrology 210:21-37.

Chapman T, 1999. A comparison of algorithms for stream flow recession and baseflow separation. Hydrological Processes 13, 701-714.

Cherkauer DS and McBride JM, 1998. A remotely operated seepage meter for use in large lakes and rivers. Ground Water 26:165-171.

Chow VT, 1959. Open-channel hydraulics. McGraw-Hill, New York. Clark I, Fritz P, 1997. Environmental Isotopes in Hydrology. Lewis, Boca Raton, Fl..

http://www.science.uottawa.ca/%7Eeih/ Constanz, J. 1998. Interaction between stream temperature, streamflow, groundwater

exchanges in alpine streams. Water Resources Research. 34, 1609-1616. Constanz J, Stewart AE, Niswonger R, Sarma L, 2002. Analysis of temperature profiles for

investigating stream losses beneath ephemeral channels. Water Resources Research 38(12): 1316.


Cook, PG and Herczeg, AL. 2000. Environmental Tracers in Subsurface Hydrology. Kluwer Academic Publishers.

Cook PG, Favreau G, Dighton JC and Tickell S, 2003. Determining natural groundwater influx to a tropical river using radon, chlorofluorocarbons and ionic environmental tracers. Journal of Hydrology 277:74-88.

Coplen TB, 1993. Uses of environmental isotopes. In Regional groundwater quality, ed. W.M.Alley, pp 227-254. Van Nostrand Reinhold, New York.

Coram JE, Dyson PR, Houlder PA and Evans WR, 2000. Australian Groundwater Flow Systems contributing to dryland salinity. Report by Bureau of Rural Sciences for National Land and Water Resources Audit, Canberra.

Craig H, 1961. Isotopic variations in natural waters. Science 133, 1702-1703. Crandall CA, Katz BG and Hirten, JJ, 1999. Hydrochemical evidence for mixing of river

water and groundwater during high-flow conditions, lower Suwannee River basin, Florida, USA. Hydrogeology Journal 7, 454-467.

Dahm CM and Valett HM, 1996. Hyporheic Zones. In: Hauer FR, Lamberti GA (eds). Methods in Stream Ecology. Academic Press, San Diego: 107-119.

Davis SN, Thompson GM, Bentley HW, Stiles G, 1980. Ground-water tracers – a short review. Ground Water 18, 14-23.

Des Carrieres D, 1883. Etiologies de l’epidimie typhoide quia eclate a Auxerre en septembre 1882. Bull. Mem. Soc. Hosp. Paris, Ser. 2,9:277-284.

Dewandel B, Lachassagne P, Bakalowicz M, Weng Ph & Al-Malki A, 2003. Evaluation of aquifer thickness by analysing recession hydrographs. Application to the Oman ophiolite hard-rock aquifer. Journal of Hydrology, 274:248-269.

Dillon KS, Corbett DR, Chanton JP, Burnett WC, Furbish DJ, 1999. The use of sulfur hexafluoride SF6 as a tracer of septic tank effluent in the Florida Keys. Journal of Hydrology 220:129-140.

Divine CE, McDonnell JJ, 2005. The future of applied tracers in hydrogeology. Hydrogeology Journal 13:255-258.

Drew BP, 1968. A review of the available methods for tracing underground waters. Proc. Br. Speleol. Assoc. 6:1-19.

Drysdale RN, Gale SJ, 1997. The Indarri Falls travertine dam, Lawn Hill Creek, Northwest Queensland, Australia. Earth Surface Processes and Landforms 22, 413-418

Drysdale RN, Taylor MP, Ihlenfeld C, 2002. Factors controlling the chemical evolution of travertine-depositing rivers of the Barkly karst, northern Australia. Hydrological Processes 16, 2941-2962.

Duwelius RF, 1996. Hydraulic conductivity of the streambed, East Branch Grand Calumet River, Northern Lake County, Indiana. US Geological Survey Water-Resources Investigations Report 96-4218.

Eckhardt K, 2005. How to construct recursive digital filters for baseflow separation. Hydrological Processes 19, 507-515.

Ellins KK, Roman-Mas A and Lee R, 1990. Using 222Rn to examine groundwater/surface discharge interaction in the Rio Grande De Manati, Puerto Rico. Journal of Hydrology 155:319-341.

Epstein S, Mayeda TK. 1953. Variations of the 18O/16O ratio in natural waters. Geochimica et Cosmochimica Acta 4: 213.

Fellows CR and Brezonik PL, 1980. Seepage flow into Florida lakes. Water Resources Bulletin 16:635-641

Fetter CW, 1994. Applied Hydrogeology. (3rd ed) Prentice Hall. Flurey M, Wai NN, 2003. Dyes as tracers for vadose zone hydrology. Reviews of Geophysics

41,1:2-27.


Fontes, JC, Edmunds, JN, 1989. The use of environmental isotope techniques in arid zone hydrology. IHP-III project 5.2 UNESCO, Paris.

FREND, 1989. 1: Hydrological Studies. II: Hydrological data. Flow Regimes from Experimental and Network Data. Wallingford, UK

Frohlich K, Frohlich W and Wittenberg H, 1994. Determination of groundwater recharge by baseflow separation: regional analysis in northeast China. FRIEND: Flow Regimes from International Experimental and Network Data, Proceedings of Braunschweig Conference, October 1993. IAHS Publ. No 221

Furey PR and Gupta VK, 2001. A physically based filter for separating base flow from streamflow time series. Water Resources Research 37(11):2709-2722.

Gat, JR, 1996. Oxygen and hydrogen isotopes in the hydrologic cycle. Annu. .Rev. Earth Planet. Sci. 24, 225-262.

Gaspar E. (ed) 1987. Modern trends in tracer hydrology (Vol II). CRC Press, Boca Raton. Fl. Gordon ND, McMahon TA, Finlayson BL, Gippel GJ, Nathan RJ, 2004. Stream Hydrology -

An introduction for ecologists. 2nd Edition, Wiley. Grayson, RB, Argent, RM, Nathan, RJ, McMahon, TA and Mein, RG. 1996. Hydrological

recipes. Estimation techniques in Australian hydrology. CRC for Catchment Hydrology, Dept. Civil Engineering, Monash University.

Griffiths GA and Clausen B, 1997. Streamflow recession in basins with multiple water storages. Journal of Hydrology 190, 60-74.

Grimm NB and Fisher SG, 1984. Exchange between surface and interstitial water: implications for stream metabolism and nutrient cycling. Hydrobiologia 111: 219-228.

Hagley R, 1996. The Tuckean Project. Proceedings 2nd National Conference on Acid Sulfate Soils. Coffs Harbour, 5-6 September 1996.

Halevy E, Moser H, Zellhofer O, Zuber A, 1967. Borehole dilution techniques: a critical review. Proceedings, 1966 Symposium of the International Atomic Energy Agency, 531-564.

Hall FR, 1968. Base flow recession - a review. Water Resources Research 4(5): 973-983. Hamilton S, 2004. River-groundwater interactions in the Cudgegong Valley, Mudgee NSW.

In: Conference Proceedings, 9th Murray Darling Basin Groundwater Workshop Bendigo, Victoria 17-19 February 2004.

Harvey JW, Wagner BJ, Bencala KE, 1996. Evaluating the reliability of the stream tracer approach to characterise stream-subsurface water exchange. Water Resources Research 32(8):2441-2451.

Harvey FE and Lee DL, 2000. Discussion on "The effects of bag type and meter size on seepage meter measurements" by SA Isiorho and JH Meyer, May-June 1999 issue, v 37, no 3:411-413. Ground Water 38(3):326-327.

Hauer FR and Lamberti GA, 1996. Methods in stream ecology. Academic Press. Hazen A, 1893. Some physical properties of sands and gravels. Massachusetts State Board of

Health, 24th Annual Report. Healy RW and Ronan AD, 1996. Documentation of the computer program VS2DH for

simulation of energy transport in variably saturated porous media – Modification of the US Geological Survey’s computer program VS2DT. US Geological Survey Water Resources Investigation Report 96-4230, 36p.

Hem JD. 1989. Study and interpretation of the chemical characteristics of natural water. 3rd edition. U.S.Geological Survey, Water-Supply Paper 2254, 263 pp.

Herczeg A, Lamontagne S, Pritchard J, Leaney F and Dighton J, 2001. Groundwater-surface water interactions: testing conceptual models with environmental tracers. 8th Murray Darling Basin Groundwater Workshop, Victor Harbor, South Australia. P. 6B.3.


Hewlett JD and Hibbert AR, 1963. Moisture and energy considerations within a sloping soil mass during drainage. J. Geophys. Res. 64, 1081-1087

Hino M and Hasebe M, 1984. Identification and prediction of nonlinear hydrological systems by the filter-separation autoregressive (AR) method: extension to hourly hydrologic data. Journal of Hydrology 68: 181-210

Horton RE, 1933. The role of infiltration in the hydrological cycle. Transactions of the American Geophysical Union 14:446-460.

Howard KW, Lloyd JV. 1983. Major ion characterisation of coastal saline groundwaters. Ground Water. 21, 429-437.

Institute of Hydrology, 1980. Low flow studies. Res. Rep. 1. Institute of Hydrology, Wallingford, UK.

Isiorho S and Meyer JH, 1999. The effects of bag type and meter size on seepage meter measurements. Ground Water 37:411-413.

Israelson OW and Reeve RC, 1944. Canal lining experiments in the Delta Area, Utah. Utah Agricultural Experimental Station, Bulletin 313: 15-35.

James LD and Thompson WO, 1970. Least squares estimation of constants in a linear recession model. Water Resources Research, 6, 1062-1069.

Jaynes DB, 1994. Evaluation of fluorobenzoate tracers in surface soils. Ground Water 32, 532-538.

John PH and Lock MA, 1977. The spatial distribution of groundwater discharge into the littoral zone of a New Zealand lake. Journal of Hydrology 33: 391-395.

Johnson EA and Dils RE, 1956. Outline for compiling precipitation, runoff and groundwater data from small watersheds. Southeastern Forest Expt. Sta. Paper 68

Kaleris V, 1998. Quantifying the exchange rate between groundwater and small streams. Journal of Hydraulic Research 36(6): 913-932.

Kass W, 1998. Tracing techniques in geohydrology. AA Balkema, Broofield Vt. Kendall C, Coplen TB. 1985. Multisample conversion of water to hydrogen by zinc for stable

isotope determination. Analytical Chemistry 57: 1437-1440. Klijn, F and Witte, JP. 1999. Eco-hydrology: groundwater flow and site factors in plant

ecology. Hydrogeology Journal. 7, 65-77. Kraemer, T.F. and Genereux, D.P. 1998. Applications of uranium- and thorium-series

radionuclides in catchment hydrology studies. In: C. Kendall and J.J. McDonnell (Eds.), Isotope tracers in catchment hydrology, Elsevier, Amsterdam, pp. 679-722.

Krupa SL, Belanger TV, Heck HH, Brok JT, Jones BJ, 1998. Krupaseep- the next generation seepage meter. Journal Coastal Research 25, 210-213.

LaFont MA, Durbec A and Ille C, 1992. Oligochaete worms as biological describers of the interactions between surface and groundwaters: a first synthesis. Regulated Rivers: Research and Management 7, 65-73.

Lamontagne S, Dighton J, Ullman W, 2002. Estimation of groundwater velocity in riparian zones using point dilution tests. Technical Report 14/02. CSIRO Land and Water.

Landon MK, Rus DL and Harvey FE, 2001. Comparison of instream methods for measuring hydraulic conductivity in sandy streambeds. Ground Water 39:870-885.

Langbein WB, 1938. Some channel-storage studies and their application to the determination of infiltration. Trans. Am. Geophys. Union, 19, 435-445

Lapham WW, 1989. Use of temperature profiles beneath streams to determine rates of vertical ground-water flow and vertical hydraulic conductivity. US Geological Survey Water Supply Paper 2337

Larkin RG and Sharp JM Jr, 1992. On the relationship between river-basin geomorphology, aquifer hydraulics and groundwater flow direction in alluvial aquifers. Geological Society of America Bulletin 104(12): 1608-1620.


Lawrence AR, Lloyd JW, Marsh JM. 1976. Hydrochemistry and groundwater mixing in part of the Lincolnshire Limestone aquifer, England. Groundwater 14, 320-327.

Lee DR, 1977. A device for measuring seepage flux in lakes and estuaries. Limnology and Oceanography 22(1):140-147.

Lee DR, 1985. Method for locating sediment anomalies in lakebeds that can be caused by groundwater inflow. Journal of Hydrology 79:187-193.

Lee DR and Cherry JA, 1978. A field exercise on groundwater flow using seepage meters and mini-piezometers. Journal of Geological Education 27:6-10.

Lee DR and Hynes HBN, 1978. Identification of groundwater discharge zones in a reach of Hillman Creek in southern Ontario. Water Pollution Research Canada 13:121-133.

Lee RW and Hollyday EF, 1993. Use of radon measurements in Carters Creek, Maury County, Tennessee, to determine location and magnitude of ground-water seepage. In: Gunderson LCS and Wanty RB (eds), Field studies of radon in rocks, soils and water. CK Smoley, 237-242.

Lerner DN, Issar AS and Simmers I, 1990. Groundwater recharge, a guide to understanding and estimating natural recharge. International Association of Hydrogeologists, Kenilworth.

Lewis JB, 1987. Measurements of groundwater seepage-flux into a coral reef: Spatial and temporal variation. Limnology and Oceanography 32:1165-1169.

Libelo EL and McIntyre WG, 1994. Effects of surface-water movement on seepage meter measurements of flow through the sediment-water interface. Hydrogeology Journal 2:49-54.

Lindgren RJ, Landon MK, 2000. Effects of ground-water withdrawals on the Rock River and associated valley aquifer, Eastern Rock County, Minnesota. US Geological Survey Water-Resources Investigation Report 99-4157.

Linsley RK, Kohler MA, Paulhus JLH, Wallace JS, 1958. Hydrology for engineers. McGraw Hill, New York

Loaiciga HA and Zekster IS, 2002. Estimation of submarine groundwater discharge. Water Resources 30(5): 473-479.

Lodge DM, Krabbenhoft DP and Striegl RG, 1989. A positive relationship between groundwater velocity and submerged macrophyte biomass in Sparling Lake, Wisconsin. Limnology and Oceanography 34: 235-239.

Lusczynski NJ, 1961. Head and flow of ground water of variable density, Journal of Geophysical Research, 66(12) 4247 - 4256.

LWBC, 2003. Minimum Construction requirements for water bores in Australia. Land and Water Biodiversity Committee. http://www.iah.asn.au/pdfs/mcrwba.pdf

McBride MS and Pfannkuch HO, 1975. The distribution of seepage within lakebeds. Journal of Research, US Geological Survey, 3(5):505-512.

McCarthy KA, McFarland WD, Wilkinson JM and White LD, 1992. The dynamic relationship between ground water and the Columbia River: using deuterium and oxygen-18 as tracers. Journal of Hydrology 135, 1-12.

McMahon PB, Tindall JA, Collins JA, Hull KJ, Nuttle JR, 1995. Hydrologic and geochemical effects on oxygen uptake in bottom sediments of an effluent-dominated river. Water Resources Research 31(10) 2561-2569.

Maillet E, 1905. Essais d’hydraulique souterraine et fluviale. Librairie Sci. Hermann Paris, 218pp.

Malcolm RL, Aiken GR, Thurman EM, Avery PA, 1980. Hydrophilic organic solutes as tracers in groundwater recharge studies. In: Baker RA (ed) Contaminants and Sediments, Butterworth-Heinemann, Woburn, Mass. 71-88.


Mazor, E, 1997. Chemical and isotopic groundwater hydrology: The applied approach. M.Dekker, New York, 413 pp.

Meyboom P, 1961. Estimating ground-water recharge from stream hydrographs. Journal of Geophysical Research, 66(4), 1203-1214

Moore RD, 1997. Storage-outflow modelling of streamflow recessions, with application to a shallow-soil forested catchment. Journal of Hydrology 198:260-270

Mull DS, Liebermann TD, Smoot JL, Woosley LH, 1988. Application of dye-tracing techniques for determining solute transport characteristics of groundwater in karst terranes. Report EPA904/688-001, US Environmental Protection Agency.

Murdoch LC and Kelly SE, 2003. Factors affecting the performance of conventional seepage meters. Water Resources Research 39(6), 1163.

Nathan RJ and McMahan TA, 1990. Evaluation of automated techniques for baseflow and recession analysis. Water Resources Research. 26(7):1465-1473.

NCDWQ, 2004. Stream Identification Method Version 3 (Draft). North Carolina Division of Water Quality.

Neal BP, Nathan RJ, Evans R, 2004. Survey of baseflows in unregulated streams of the Murray-Darling Basin. Proceedings 9th Murray-Darling Basin Groundwater Workshop, Bendigo

Nield SP, Townley LR and Barr AD, 1994. A framework for quantitative analysis of surface water – groundwater interaction: Flow geometry in a vertical section. Water Resources Research 30(8): 2461-2475.

Otz MH, Otz HZ, Otz I and Siegel DI, 2003. Surface water/groundwater interaction in the Piora aquifer, Switzerland: evidence from dye tracing tests. Hydrogeology Journal 11: 228-239.

Oxtobee JPA and Novakowski K, 2002. A field investigation of groundwater/surface water interaction in a fractured bedrock environment. Journal of Hydrology 269:169-193.

Paulsen RJ, Smith CF, O'Rourke D and Wong TF, 2001. Development and evaluation of an ultrasonic ground water seepage meter. Ground Water 39:904-911.

Petalas CP, Diamantis IB, 1999. Origin and distribution of saline groundwater in the upper Miocene aquifer system, coastal Rhodope area, northeastern Greece. Hydrogeology 7:305-316.

Pettyjohn WA and Henning R, 1979. Preliminary estimate of ground-water recharge rates, related streamflow and water quality in Ohio. Ohio State University Water Resources Centre Project Completion Report No 552, 323pp.

Pritchard J, Herczeg A and Lamontagne S, 2002. The use of environmental tracers for estimating seasonal contributions of groundwater to stream flow. In: Proceedings of International Groundwater Conference, May, International Association of Hydrogeologists, Darwin.

Robinson M, Gallagher D and Reay W, 1998. Field observations of tidal and seasonal variations in ground water discharge to tidal estuarine surface water. Ground Water Monitoring and Remediation 18(1):83-92

Ronan AD, Prudic DE, Thodal CE, Constanz J, 1998. Field study and simulation of diurnal temperature effects on infiltration and variably saturated flow beneath an ephemeral stream. Water Resources Research 34:2137-2153.

Rorabaugh MI, 1964. Estimating changes in bank storage and groundwater contribution to streamflow. Int. Ass. Scientific Hydrology. Publ 63, 432-441

Rose PE, Benoit WR, Adams MC, 1998. Tracer testing at Dixie Valley, Nevada using Pyrene Tetrasulfonate, Amino G and Fluoroscein. Transactions Geothermal Resource Council 22:583-587.


Rosenberry DO, 2000. Unsaturated-zone wedge beneath a large natural lake. Water Resources Research 36(12) 3401-3409.

Rosenberry DO and Morin RH, 2004. Use of an electromagnetic seepage meter to investigate temporal variability in lake seepage. Ground Water 42(1):68-77.

Rutledge AT, 1998. Computer programs for describing the recession of ground-water discharge and for estimating mean ground-water recharge and discharge from streamflow records-Update. US Geological Survey, Water Resources Investigations Report 98-4148

Rutledge AJ, Daniel III CC, 1994. Testing an automated method to estimate ground-water recharge from stream flow record. Ground Water 32(2):180-189

Schincariol RA and McNeil JD, 2002. Errors with small volume elastic seepage meter bags. Ground Water 40(6):649-651.

Sebestyen SD and Schneider RL, 2004. Seepage patterns, pore water and aquatic plants: hydrological and biogeochemical relationships in lakes. Biogeochemistry 68: 383-409.

Shaw RD and Prepas EE, 1989. Anomalous, short term influx of water into seepage meters. Limnology and Oceanography 34:1343-1351.

Shaw RD and Prepas EE, 1990. Groundwater-lake interactions: I. Accuracy of seepage meter estimations of lake seepage. Journal of Hydrology 119:105-120.

Shinn EA, Reich CD and Hickey TD, 2002. Seepage meters and Bernoulli's revenge. Estuaries 25:126-132.

Sholkovitz E, Herbold C, Charette M, 2003. An automated dye-dilution based seepage meter for the time-series measurement of submarine groundwater discharge. Limnology and Oceanography: Methods 1, 16-28.

Silliman SE and Booth DF, 1993. Analysis of time-series measurements of sediment temperature for identification of gaining vs. losing portions of Juday Creek, Indiana. Journal of Hydrology 146: 131-148.

Simmons Jr GM and Love FG, 1986. Water quality of newly discovered submarine ground water discharge into a deep coral reef habitat. Final Report to Sanctuary Program Division, Office of Ocean and Coastal Resource Management. National Oceanic and Atmospheric Administration, Washington.

Singh VP, 1989. Hydrologic Systems Vol 11: Watershed Modeling. Prentice Hall. Sinokrat BA and Stefan HG (1993). Stream temperature dynamics: Measuring and modelling.

Water Resources Research 29(7): 2299-2312. Sloto RA and Crouse MY, 1996. HYSEP: A computer program for streamflow hydrograph

separation and analysis. US Geological Survey, Water Resources Investigations Report 96-4040

Smakhtin VY, 2001. Low flow hydrology: a review. Journal of Hydrology 240:147-186. Sophocleous M, 2002. Interaction between groundwater and surface water: the state of the

science. Hydrogeology Journal 10: 52-67. Springer AE, Petroutson WD, Semmens BA; 1999. Spatial and temporal variability of

hydraulic conductivity in active reattachment bars of the Colorado River, Grand Canyon. Ground Water 37(3) 338-344.

S S Papadopulos and Associates, 2004. Evaluation of tracer tests conducted at the Northwest Wellfield, Miami-Dade County, Florida.

Stonestrom DA and Constanz J (eds), 2003. Heat as a tool for studying the movement of ground water near streams. US Geological Survey Circular 1260.

Sujono J, Shikasho S and Hiramatsu K, 2004. A comparison of techniques for hydrographic recession analysis. Hydrological Processes 18, 403-413


Suzuki S, 1960. Percolation measurements based on heat flow through soil with special reference to paddy fields. Journal of Geophysical Research 65(9): 2883-2885.

Swanson GA, Adomaitis VA, Lee FB, Serie JR and Shoessmith JA, 1984. Limnological conditions influencing duckling use of saline lakes in south-central North Dakota. Journal of Wildlife Management 48: 340-349.

Szilagyi J, Parlange MB, 1998. Baseflow separation based on analytical solutions of the Boussinesq equation. Journal of Hydrology 204:251-260

Tallaksen LM, 1995. A review of baseflow recession analysis. Journal of Hydrology 165, 349-370.

Taniguchi M, Burnett WC, Smith CR, Paulsen RJ, O'Rourke D, Krupa SL and Christoff JL, 2003. Spatial and temporal distributions of submarine discharge rates obtained from various types of seepage meters at a site in the Northeastern Gulf of Mexico. Biogeochemistry 66:35-53.

Taniguchi M, Fukuo, Y, 1993. Continuous measurements of ground-water seepage using an automatic seepage meter. Ground Water 31, 675-679.

Taniguchi M, Iwakawa H, 2001. Measurement of submarine groundwater discharge rates by a continuous heat-type automated seepage meter in Osaka Bay, Japan. Journal of Groundwater Hydrology 43, 271-277.

Toebes C and Strang DD, 1964. On recession curves 1: Recession equations. J. Hydrology, New Zealand 3(2), 2-15

Toth J, 1963. A theoretical analysis of groundwater flow in small drainage basins. Journal of Geophysical Research 68, 4785-4812.

Valiela I, Costa J, Foreman K, Teal JM, Howes B and Aubrey D, 1990. Transport of groundwater-borne nutrients from watersheds and their effects on coastal waters. Biogeochemistry 10:177-197

Valett HM, Fisher SG, Grimm NP, Camill P, 1994. Vertical hydrologic exchange and ecological stability of a desert stream ecosystem. Ecology 75(2): 548-560.

Van de Griend AA, De Vries JJ, Seyhan E, 2002. Groundwater discharge from areas with a variable specific drainage resistance. Journal of Hydrology 259:203-220

Vaux WG, 1968. Intragravel flow in interchange of water in a streambed. US Fish Wildlife Service Fishery Bulletin 66:479-489.

Voss CI, 1984. A finite-element simulation model for saturated-unsaturated, fluid density-dependent ground-water flow with energy transport or chemically-reactive single species solute transport. US Geological Survey Water Resource Investigations Report 84-4369, 409p.

Vukovic M, Soro A, 1992. Determination of hydraulic conductivity of porous media from grain-size composition. Water Resources Publications, Littleton, Colorado

Ward JV, 1989. The four-dimensional nature of lotic ecosystems. Journal of the North American Benthological Society 8, 2-8.

Ward JV and Stanford JA, 1989. Groundwater animals of alluvial river systems: a potential management tool. In: Grigg NS (ed) Proceedings of the Colorado Water Engineering and Management Conference, Colorado Water Resources Research Institute, Fort Collins, Colorado.

Winter TC, 1999. Relation of streams, lakes and wetlands to groundwater flow systems, Hydrogeology Journal 7, 28-45.

Winter, TC, Judson, WH, Franke, OL and Alley WM. 1998. Groundwater and surface water a single resource. Circular 1139, U.S. Geological Survey, Denver.

Woessner WW, 2000. Stream and fluvial plain ground water interactions: Rescaling hydrogeologic thought. Ground Water 38(3):423-429.


Woessner WW and Sullivan KE, 1984. Results of seepage meter and mini-piezometer study, Lake Mead, Nevada. Ground Water 22(5): 561-568.



16. Appendices Appendix 16.1: Inventory of published hydrogeological maps in Australia (Brodie et al, 2004)

Region Author(s) Date Title Scale Organisation ACT Evans W.R. 1984 Hydrogeology of Canberra ACT and environs 100000 BMR AUS AWRC 1975 Groundwater resources of fractured rocks in Australia 5000000 AWRC AUS AWRC 1975 Groundwater resources of sedimentary basins in Australia 5000000 AWRC AUS AWRC 1975 Groundwater resources of shallow unconsolidated sediments in Australia 5000000 AWRC

AUS AWRC 1975 Principal groundwater resources of Australia 5000000 AWRC AUS Coram J.E.,Dyson P.R.,Houlder P.A. and Evans

W.R. 1999 Australian groundwater flow systems 5000000 BRS

AUS Jacobson G. and Lau J.E. 1987 Hydrogeology of Australia 5000000 BMR

AUS NLWRA 2001 Australian dryland salinity assessment 2500000 NLWRA AUS NLWRA 2000 Australian groundwater management units, unincorporated areas and provinces 5000000 NLWRA

GAB Habermehl M.A. and Lau J.E. 1997 Hydrogeology of the Great Artesian Basin Australia 2500000 AGSO

MDB ? ? Murray Darling Basin groundwater flow systems ? MDBC MDB Evans W.R., Kellett J.R., Williams R.M., Gates G.,

Lawson S., van der Lelij A., Barnett S.R., Lakey R.C., Lawrence C.R., Ife D., Thorne R. and Macumber P.G.

1987 Preliminary shallow groundwater salinity map of the Murray Basin 1000000 BMR

MDB Evans W.R., Hillier J. and Woolley D.R. 1995 Hydrogeology of the Darling River drainage basin 1000000 AGSO

NSW ? 1997 COFFS HARBOUR groundwater vulnerability map ? NSW DLWC NSW ? ? HUNTER RIVER groundwater vulnerability map ? NSW DLWC NSW ? ? SNOWY RIVER groundwater vulnerability map ? NSW DLWC NSW Bradd J.M. 1995 NSW dryland salinity hazard map 4,500,000 NSW DLWC NSW Brodie R.S. 1992 ANA BRANCH Hydrogeological Map 250,000 AGSO

NSW Brodie R.S. 1994 MENINDEE Hydrogeological Map 250,000 AGSO

NSW Broughton A.K. 1994 Hydrogeological map of the Liverpool Plains Catchment 250,000 NSW DWR NSW Hamilton S., Ross J.B. and Williams R.M. 1988 MOREE hydrogeological map 250,000 NSW DWR

NSW Hamilton S., Ross J.B. and Williams R.M. 1988 NARRABRI hydrogeological map 250,000 NSW DWR NSW Hamilton S., Williams R.M. and Woolley D 1987 Groundwater in New South Wales : Assessment of Pollution Risk. 2,000,000 NSW DWR

NSW Kellett J. R. 1994 IVANHOE Hydrogeological Map 250,000 AGSO

NSW Kellett J. R. 1994 MANARA Hydrogeological Map 250,000 AGSO

NSW Kellett J. R. 1989 POONCARIE Hydrogeological Map 250,000 BMR

NSW Kellett J.R. 1994 BALRANALD Hydrogeological Map 250,000 AGSO

NSW Krumins I., Bradd J.M. and McKibbon, D. 1997 Hawksbury-Nepean catchment availability map ? NSW DLWC

NSW Piscopo et al 2001 2nd order catchments in NSW availability map various NSW DLWC NSW Piscopo et al 2001 State wide 2nd order catchments in NSW availability map 2,000,000 NSW DLWC

NSW Piscopo G. 2001 Castlereagh catchment groundwater vulnerability map 250000 NSW DLWC

NSW Piscopo G. 2001 Lachlan catchment groundwater vulnerability map 250000 NSW DLWC

NSW Piscopo G. 2001 MacIntyre catchment groundwater vulnerability map 250000 NSW DLWC

NSW Piscopo G. 2001 Macquarie catchment groundwater vulnerability map 250000 NSW DLWC

NSW Piscopo G. and Please P. ? 2nd order catchments in NSW availability map various NSW DLWC

NSW Piscopo G. and Please P. 1997 Murrumbidgee Catchment groundwater vulnerability map 1000000 NSW DLWC

NSW Piscopo G. and Please P. ? State wide 2nd order catchments in NSW availability map 2,000,000 NSW DLWC

NSW Ross J.B., Williams R.M. and Woolley D.R. 1986 FORBES hydrogeological map 250,000 NSW WRC NSW Sinclair P. 1997 Wagga Wagga groundwater vulnerability map 100,000 NSW DLWC

NSW Williams R.M. and Woolley D. 1992 DENILIQUIN Hydrogeological Map 250,000 AGSO

NSW Woolley D. 1994 CARGELLIGO Hydrogeological Map 250,000 AGSO NSW Woolley D. 1991 NARRANDERA Hydrogeological Map 250,000 BMR

NSW Woolley D.R. 1992 HAY Hydrogeological Map 250,000 AGSO

NSW Woolley D.R. and Williams R.M. 1994 BOOLIGAL Hydrogeological Map 250,000 AGSO

NSW NSW DLWC 2000 Predicted Salinity Extent 2000 in New South Wales 250,000 NSW DLWC



NT Domahidy G. 1990 Tanami Mining region hydrogeological map 250000 NT PAWA NT McGowan R.J. 1989 Pine Creek hydrogeological map 250000 NT PAWA NT Verma M.N. 2002 Hydrogeology Map Darwin 1:250000 250000 NT DLPE

NT Haigh T. 2002 Water Resources map of Tiwi Islands 250000 NT DLPE

NT Verma M.N. 1992 Helen Springs hydrogeological map 250000 NT PAWA NT Tickell S.J. & Qureshi H. 1998 Groundwater map of Darwin 50000 NT DLPE

NT Tickell S.J. and Rajaratnam L.R. 1996 Water resource development map of Hooker Creek Station and Lajamanu Community

100000 NT DLPE

NT Yin Foo D. 2000 Water resource development map Avago, Birdum Creek, Maryfield, Middle Creek, Sunday Creek, Tarlee, Vermelha and Western Creek Stations

250000 NT DLPE


NT Yin Foo D. 2000 Water resource development map Elsey station and Wubalawun Aboriginal Land Trust

250000 NT DLPE

NT Yin Foo D. 2000 Water resource development map Kalala and Hidden Valley Stations 250000 NT DLPE

NT Yin Foo D. 1999 Water resource development map Bloodwood Downs, Cow Creek, Lakefield, Larrizona, Gilnockie, Margaret Downs, Nenen, Gorrie, Dry River and Wyworrie Stations

250000 NT DLPE

NT Tickell S.J. and Rajaratnam L.R. 1996 Water resource development map of Mount Sanford Station, part of Victoria River Downs Station

100000 NT DLPE

NT Tickell S.J. and Rajaratnam L.R. 1996 Water resource development map of Fitzroy Station and NT Portion 3122 100000 NT DLPE

NT Tickell S.J. and Rajaratnam L.R. 1995 Water resource development map of Inverway Station 100000 NT DLPE

NT Tickell S.J. and Rajaratnam L.R. 1995 Water resource development map of Kirkimbi Station and NT Portion 3540 100000 NT DLPE

NT Tickell S.J. and Rajaratnam L.R. 1996 Water resource development map of Limbunya Station and part of Riveren Station

250000 NT DLPE

NT Tickell S.J. and Rajaratnam L.R. 1994 Water resource development map of Legune Station 100000 NT DLPE

NT Tickell S.J. and Rajaratnam L.R. 1996 Water resource development map of Mistake Creek Station 100000 NT DLPE

NT Tickell S.J. and Rajaratnam L.R. 1994 Water resource development map of Newry Station and Keep River National Park

100000 NT DLPE

NT Sanders R. and Rajaratnam L.R. 1994 Water resource development map of Rosewood Station 100000 NT DLPE

NT Tickell S.J. and Rajaratnam L.R. 1995 Water resource development map of Riveren Station 100000 NT DLPE

NT Tickell S.J. and Rajaratnam L.R. 1994 Water resource development map of Spirit Hills Station 100000 NT DLPE

NT Tickell S.J. and Rajaratnam L.R. 1994 Water resource development map of Waterloo Station 100000 NT DLPE

NT Tickell S.J. and Rajaratnam L.R. 1994 Water resource development map of Bullo River Station 100000 NT DLPE

NT Sanders R. and Rajaratnam L.R. 1994 Water resource development map of Auvergne Station 100000 NT DLPE

NT Tickell S.J. and Rajaratnam L.R. 1995 Water resource development map of Buda Station 100000 NT DLPE

NT Tickell S.J. and Rajaratnam L.R. 1995 Water resource development map of Birrindudu Station 100000 NT DLPE

NT Tickell S.J. and Rajaratnam L.R. 1994 Water resource development map of Bradshaw 250000 NT DLPE

NT Tickell S.J. and Rajaratnam L.R. 1997 Water resource development map of Daguragu 100000 NT DLPE

NT Tickell S.J. and Rajaratnam L.R. 1995 Water resource development map of Wallamunga Station 100000 NT DLPE

NT Sanders R. and Rajaratnam L.R. 1994 Water resource development map of Amanbidji Station 100000 NT DLPE

NT Mathews I. 1997 DELAMERE hydrogeological map 250000 NT DLPE

NT Karp D. 1997 VICTORIA RIVER DOWNS hydrogeological map 250000 NT DLPE

NT Karp D. 1997 WAVE HILL hydrogeological map 250000 NT DLPE

NT Verma M.N. 1997 Berry Springs-Noonamah Area hydrogeological map 50000 NT DLPE

NT Ride G. 1997 Water resource development map of Alambi Station 100000 NT DLPE

NT Ride G. 1997 Water resource development map of Todd River Station 100000 NT DLPE

NT Ride G. 1997 Water resource development map of Ltyentye Apurte 100000 NT DLPE

NT Ride G. 1997 Water resource development map of The Gardens Station 100000 NT DLPE

NT Ride G. 1998 Water resource development map of UndoolyaStation 100000 NT DLPE

NT Ride G. 1998 Water resource development map of Maryvale Station 100000 NT DLPE

NT Ride G. 1998 Water resource development map of Henbury Station, (two sheets) 100000 NT DLPE

NT Ride G. 2000 Water resource development map of Ltalatuma, Ntaria, Rodna, Roulmaulpa and Uruna

100000 NT DLPE

NT Ride G. 1997 Water resource development map of Orange creek Station 100000 NT DLPE

NT Ride G. 1998 Water resource development map of Hamilton Downs Station 100000 NT DLPE

NT Ride G. 1997 Water resource development map of Deep Well Station 100000 NT DLPE

NT Ride G. 1998 Water resource development map of Amoonguna Station 15000 NT DLPE

NT Ride G. Water resource development map of Mount Allen Station 100000 NT DLPE

NT Ride G. 1998 Water resource development map of Iwupataka and Standley Chasm 100000 NT DLPE

NT Ride G. 1998 Water resource development map of Owen Springs Station 100000 NT DLPE

NT Zaar U. 2002 Water Resources West Central Arnhem Land 250000 NT DLPE

NT Zaar U. 2002 Water Resources West Arnhem Land 250000 NT DLPE

NT Prowse G., Zaar U. & Matthews I. 1999 Water Resources South East Arnhem Land 250000 NT DLPE

NT Read R. and Seidel G. 1999 Stuart Highway and Railway Corridor, Water Prospects 1000000 NT DLPE

NT Wischusen J. 1998 Hydrogeology of the Papunya-Yuendumu-Kintore region NT 500000 AGSO

NT Yin Foo D. 2001 Water resource development map Sturt Plateau Region 250000 NT DLPE

NT George D. 2001 Water Resources of the Katherine Region and South West Arnhem land 250000 NT DLPE

NT Tickell S.J., Rajaratnam L.R. and Sanders R. 1998 Water Resources of the Victoria River District 500000 NT DLPE

NT Rooke E. 1997 Water Resources map of Barkly/Gulf Study Region 1 500000 NT DLPE

NT Rooke E. 1998 Water Resources map of Barkly/Gulf Study Region 2 500000 NT DLPE

NT D.Chin & Paiva J. 1998 Water resource availability map, within 150km. of Katherine 500000 NT DLPE

NT D.Chin 1995 Douglas Station water availability map 100000 NT DLPE

NT D.Chin 1995 Claravale and Jindare water availability map 100000 NT DLPE

NT Mathews I. 1997 Great Artesian Basin Hydrogeological map 1000000 NT DLPE

NT Verma M.N. 2000 Groundwater supply prospects of the Litchfield Shire 100000 NT DLPE

NT Prowse G., Zaar U., Tickell S.J. & Matthews I. 1999 Water Resources North East Arnhem Land 250000 NT DLPE


NT Prowse G., Zaar U. & Matthews I. 1999 Water Resources East Central Arnhem Land 250000 NT DLPE

NT Verma M.N. 1998 Groundwater supply prospects of the Coomalie Shire 100000 NT DLPE

QLD Laycock J.W. and Wecker H.R.B. 1971 Groundwater resources of Queensland 2500000 GSQ, QIWSC

QLD Stenson M.P. and Hansen A 1998 - 2002

Groundwater vulnerability mapping in the Upper Condamine River catchment 250 000 - 2 500 000

QNR&M

QLD ? 2000 Queensland statewide salinity hazard map ? QNR&M

QLD Schmiede D.L. 1987 Groundwater resources of Queensland 2500000 QWRC, QDM

SA Shepherd R.G. (SA Department of Mines and Energy)

1982 Groundwater resource map South Australia 2000000 GSSA

SA Barnett S.R. (SA Department for Water Resources) 1994 ADELAIDE-BARKER Hydrogeological Map 250000 AGSO

SA Barnett S.R. (SA Department for Water Resources) 1991 PINNAROO Hydrogeological Map 250000 AGSO

SA Barnett S.R. (SA Department for Water Resources) 1994 BURRA-CHOWILLA-OLARY Hydrogeological Map 250000 AGSO

SA Cobb M.A. and Barnett S.R. (SA Department for Water Resources)

1994 NARACOORTE Hydrogeological Map 250000 AGSO

SA Barnett S.R. (SA Department for Water Resources) 1991 RENMARK Hydrogeological Map 250000 BMR

SA SA Department for Water Resources 1997 Depth to shallow aquifer watertable map 2000000 SA DWR

SA SA Department for Water Resources 1997 Salinity of shallowest groundwater aquifer map 2000000 SA DWR

SA SA Department for Water Resources 1997 Shallow aquifer yield map 2000000 SA DWR

SA Barnett S.R. and Henschke, C.J. (SA Department for Primary Industries and Resources)

2001 Dryland salinity extent and risk map South Australia 50000 DPIRSA

SA Henschke, C.J. and Barnett S.R. (SA Department for Primary Industries and Resources)

2001 Groundwater flow systems map South Australia 50000 DPIRSA

TAS P.B. Nye 1921 The Underground Water Resources of the Midlands 160000 TDM

TAS P.B. Nye 1922 The Underground Water Resources of the Jericho - Richmond -Bridgewater Area 160000 TDM

TAS P.B. Nye 1924 The Underground Water Resources of the Richmond -Brdgewater - Sandford District

89000 TDM

TAS P.B. Nye 1926 The Campbell Town - Conara - St Marys District 89000 TDM

TAS A.B.Gilline 1959 The Underground Water Resources of the Smithton District 123000 TDM

TAS Leaman D.E. 1967 The Groundwater Resources of the Cygnet District 40000 TDM

TAS Leaman D.E. 1971 The Geology and Groundwater Resources of the Coal River Basin 63360 TDM

TAS Matthews W.L. 1979 Geology and Groundwater Resources of the Longford Tertiary Basin 100000 TDM

TAS W.C.Cromer 1979 Greens Beach groundwater resources 25000 TDM

TAS Moore W.R. 1992 Hydrogeology of the Scottsdale sedimentary basin (Map 1 Geology - 1993 and Map 2 Hydrogeology-1990)

60000 TDM

TAS W.C.Cromer 1993 Geology and Groundwater resources of the Devonport-Port Sorell -Sassafras Tertiary Basin

50000 TDM

TAS W.L.Mathews R.C.Donaldson 1999 1:500,000 digital groundwater prospectivity Tasmania 500000 MRT

TAS K.Taylor 2000 Groundwater Resources of the Northern Midlands and Fingal Regions (Map1 Study Area Borehole locations, Map 2 Hydrogeology of the Macquarie River Drainage Basin, Map3 Hydrogeology of the South Esk River Drainage Basin and Map 4 Groundwater Prospectivity of Macquarie River Drainage Basin

100000 MRT

TAS M.Latinovic 2001 Sorell Groundwater Project (Map 1 Hydrogeology and Map 2 Geology) 50000 MRT

TAS M.Latinovic 2001 Groundwater Prospectivity - Great Forester Catchment Map 100000 MRT

TAS M.Latinovic 2002 Tasmanian Groundwater Flow Systems for Dryland Salinity Planning 500000 TDPIWE, MRT

VIC ? ? Avoca dryland 25000 ?

VIC ? 2000 Broken Plains 25000 ?

VIC ? ? Campaspe 25000 ?

VIC Nahm G.Y. 1982 Groundwater resources of Victoria 1000000 GSV

VIC Australian Groundwater Consultants 1986 Hydrogeological Map of BALLARAT 250000 BMR

VIC Tickell S. and Humphreys W.G. 1985 Hydrogeological map of BENDIGO and part of DENILIQUIN 250000 VDITR

VIC Lakey R. and Tickell S.J. 1980 Hydrogeological map of Western Port Basin 100000 GSV

VIC ? ? Loddon dryland 25000 ?

VIC ? 2000 MDB-Victoria dryland salinity???? 100000 ?

VIC Bradley J., Stanley D., Mann B., Chaplin H. and Foley G. (Vic. Rural Water Corporation)

1994 BALLARAT Hydrogeological Map 250000 AGSO

VIC Dimos A., Chaplin H., Potts I., Reid M. and Barnewall (Vic. Rural Water Corporation)

1994 BENDIGO Hydrogeological Map 250000 AGSO

VIC Mann B., Chaplin H. and Stanley D. (Vic Rural Water Corporation)

1994 HAMILTON Hydrogeological Map 250000 AGSO

VIC McAuley C., Evans C., Robinson M., Chaplin H. and Thorne R. (Vic Rural Water Corporation)

1992 HORSHAM Hydrogeological Map 250000 AGSO

VIC Woolley D., Lee J., Yoo E.K. and Williams R.M. (NSW Department of Water Resources)

1992 JERILDERIE Hydrogeological Map 250000 AGSO

VIC Thorne R., Pratt M., Hoxley G., McCauley C. and Chaplin H. (Vic. Rural Water Commission)

1991 MILDURA Hydrogeological Map 250000 BMR

VIC Robinson M., Thorne R., Basocak C. and Chaplin H. (Vic Rural Water Corporation)

1992 OUYEN Hydrogeological Map 250000 BMR

VIC Dudding M., Chaplin H. and O'Rorke M. (Vic. Rural Water Corporation)

1993 ST ARNAUD Hydrogeological Map 250000 AGSO

VIC O'Rorke M.E., Bolger P., Thorne R. and Chaplin H. (Vic. Rural Water Commission)

1992 SWAN HILL Hydrogeological Map 250000 BMR

VIC Hennessy J., Reid M. and Chaplin H. (Vic. Rural Water Corporation)

1994 WANGARATTA Hydrogeological Map 250000 AGSO

VIC Sinclair Knight Merz 1996 Beneficial Use of Lower Tertiary Layer of Victoria 500000 VDNRE

VIC Sinclair Knight Merz 1996 Beneficial Use of Middle Tertiary Layer of Victoria 500000 VDNRE

VIC Sinclair Knight Merz 1996 Beneficial Use of Upper Tertiary Layer of Victoria 500000 VDNRE


VIC Sinclair Knight Merz 2000 Victorian Dryland Salinity Assessment 2000 - Current Depth to Water table 1998 250000 VDNRE

VIC Sinclair Knight Merz 2000 Victorian Dryland Salinity Assessment 2000 - Predicted Depth to Water Table Surface in 2050 under Best Case Trend Scenario

250000 VDNRE

VIC Sinclair Knight Merz 2000 Victorian Dryland Salinity Assessment 2000 - Predicted Depth to Water Table Surface in 2050 under Worst Case Trend Scenario

250000 VDNRE

VIC Sinclair Knight Merz 2000 Victorian Dryland Salinity Assessment 2000 - Predicted Depth to Water Table Surface in 2020 under Worst Case Trend Scenario

250000 VDNRE

VIC Sinclair Knight Merz 2000 Victorian Dryland Salinity Assessment 2000 - Predicted Depth to Water Table Surface in 2020 under Best Case Trend Scenario

250000 VDNRE

VIC Sinclair Knight Merz 2000 Victorian Dryland Salinity Assessment 2000 - Salinity Risk Classification under Best Case Trend Scenario

250000 VDNRE

VIC Sinclair Knight Merz 2000 Victorian Dryland Salinity Assessment 2000 - Salinity Risk Classification under Worst Case Trend Scenario

250000 VDNRE

VIC Sinclair Knight Merz 2000 Victorian Dryland Salinity Assessment 2000 - Best Case Trends 250000 VDNRE

VIC DCNR 1995 Eastern Victoria regional aquifer systems ? VDCNR

VIC DCNR 1995 Eastern Victoria water table aquifers ? VDCNR

VIC DCNR 1995 North Western Victoria regional aquifer systems ? VDCNR

VIC DCNR 1995 North Western Victoria water table aquifers ? VDCNR

VIC DCNR 1995 South Western Victoria regional aquifer systems ? VDCNR

VIC DCNR 1995 South Western Victoria water table aquifers ? VDCNR

VIC Sinclair Knight Merz 1996 Watertable Beneficial Use of Victoria 1000000 VDCNR

VIC Foley G., Walker G. and Chaplin H. 1995 WARRAGUL/SALE Hydrogeological Map 250000 SKM

WA WA Water and Rivers Commission 1997 Perth Groundwater Atlas 15000 WA WRC

WA De Silva, J. Smith R.A. Rutherford J. Ye L. 1995 Hydrogeological map of Blackwood Catchment Area 500000 WA WRC

WA Davidson W.A. 1995 Hydrogeology and groundwater resources of the Perth Region, WA 600000 GSWA

WA Rutherford J.L. 2000 COLLIE hydrogeological map 250000 WA WRC

WA Leonard E.L. 2000 DUMBLEYUNG hydrogeological map 250000 WA WRC

WA Johnson S.L. 1999 LAVERTON hydrogeological map 250000 WA WRC

WA Johnson S.L. 1999 LEONORA hydrogeological map 250000 WA WRC

WA Johnson S.L. 1999 SIR SAMUEL hydrogeological map 250000 WA WRC

WA Dodson W. and Brereton S. 1998 NEWDEGATE hydrogeological map 250000 WA WRC

WA De Silva, J. 2000 PEMBERTON-IRWIN INLET hydrogeological map 250000 WA WRC

WA Kern A.M. 1995 BOORABBIN hydrogeological map 250000 GSWA

WA Dodson W. 1996 BREMER BAY hydrogeological map 250000 GSWA, WA WRC

WA Laws A.T. 1991 BROOME hydrogeological map 250000 GSWA

WA Smith R.A. 1992 DERBY hydrogeological map 250000 GSWA

WA Baddock L.J 1996 ESPERANCE-MONDRAIN ISLAND hydrogeological map 250000 GSWA,WA WRC

WA Kern A.M. 1995 KALGOORLIE hydrogeological map 250000 GSWA

WA Kern A.M. 1996 KURNALPI hydrogeological map 250000 GSWA

WA Smith R.A. 1995 MT BARKER-ALBANY hydrogeological map 250000 GSWA

WA McGowan R.J. 1987 PERENJORI hydrogeological map 250000 GSWA

WA Johnson S.L. 1996 RAVENSTHORPE hydrogeological map 250000 GSWA, WA WRC

WA Kern A.M. 1996 WIDGIEMOOLTHA hydrogeological map 250000 GSWA

WA Appleyard S.J., Commander D.P. and Allen A.D. 1993 Groundwater vulnerability to contamination of the Perth Basin 500000 GSWA

AGSO (Australian Geological Survey Organisation) AWRC (Australian Water Resources Council) BMR (Bureau of Mineral Resources) BRS (Bureau of Rural Sciences) DPIRSA (Department of Primary Industries and Resources SA) GSQ (Geological Survey of Qld) GSSA (Geological Survey of SA) GSV (Geological Survey of Victoria) GSWA (Geological Survey of WA) MDBC (Murray Darling Basin Commission) MRT (Mineral Resources Tas) NLWRA (National Land and Water Resources Audit) NSW DLWC (NSW Department of Land and Water Conservation) NSW DWR (NSW Department of Water Resources) NSW WRC (NSW Water Resources Commission) NT PAWA (NT Power and Water Authority) NT DLPE (NT Department of Lands Planning and Environment) QDM (Qld Department of Mines) QIWSC (Qld Irrigation and Water Supply Commission) QNR&M (Qld Department of Natural Resources and Mines) QWRC (Qld Water Resources Commission) SA DWR (SA Department of Water Resources) SKM (Sinclair Knight Mertz) TDM (Tas Department of Mines) TDPIWE (Tas Department of Primary Industries, Water and Environment) VDCNR (Vic Department of Conservation and Natural Resources) VDITR (Vic Department of Industry, Technology and Resources) VDNRE (Vic Department of Natural Resources and Environment) WA WRC (WA Water and Rivers Commission

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An Overview of Tools for Assessing Groundwater-Surface ...data.daff.gov.au/.../data/assessinggroundwatersurfacewaterconnecti… · resource management. Assessing groundwater-surface