Methods for estimating groundwater
discharge to streams – summary of
field trials
SUMMARY REPORT

FINAL

January 2012
This project is funded by the Australian Government through
the Water for the Future - Water Smart Australia program
Methods for estimating groundwater
discharge to streams – summary of field
trials

FINAL

January 2012
Methods for estimating groundwater discharge to streams – summary of field trials
Executive Summary
Introduction
This report describes the findings from the field assessments of surface water – groundwater
interaction, which took place in 10 selected catchments in 2008 – 2009. The catchments are being
assessed as part of a project funded by the Australian Government, through the Water for the
Future – Water Smart Australia program aimed at developing methods for quantifying surface
water – groundwater fluxes. This project has been undertaken by Sinclair Knight Merz (SKM) and
the Commonwealth Scientific and Industrial Research Organisation (CSIRO). The project
compared estimates of surface water – groundwater exchange derived using flow differences,
hydraulic gradient analysis, hydrograph baseflow separation and comparison of surface water and
groundwater chemistry. The field studies identified possible strengths and weaknesses in some of
the methods. In one catchment only limited testing was able to be carried out due to low river flow
associated with drought conditions. Assessments of changes in baseflow over time were also made
in each catchment, and compared to volumes of groundwater extraction.
The purpose of the project is to compare methods, and as such, the data and subsequent results for
baseflow estimates in each catchment do not represent agreed absolute estimates of surface water –
groundwater flux volumes. Results are not intended for use as inputs to any rules or regulations
governing water resource extraction in specific catchments. Rather, the results and conclusions of
this study are intended solely as a comparison of scientific methods for measuring surface water –
groundwater exchange.
Appropriateness of methods
The hydraulic gradient method calculates groundwater inflow, based on the local aquifer
properties, and the observed hydraulic gradient between a river gauging station and a nearby
observation bore. This method was useful in determining local groundwater flux conditions close to
gauging stations, although it was shown to not necessarily be representative of inflow rates over
longer reaches.
Sequential river flow measurements were carried out along the river, and the spatial variations in
groundwater inflow or outflow were estimated from the differences between adjacent flow
gaugings. The flow difference method worked best in catchments where the difference between the
downstream flow and the sum of the upstream flow and tributary flow is large relative to the
downstream flow. Overall the flow difference method only produced accurate results in three of the
nine catchments. Results were moderate in a further three catchments, and extremely unreliable in
three catchments.
The longitudinal chemistry method involves measurements of river chemistry along a stream reach
at a point in time. Measurements of groundwater chemistry are also made, and rates of groundwater
inflow are determined from downstream changes in water chemistry using a mass balance
approach. The longitudinal chemistry method produced mostly unreliable results when electrical
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Methods for estimating groundwater discharge to streams – summary of field trials
conductivity (EC) and chloride were used as tracers (accurate results were not obtained in any
catchments, and moderate results were obtained in only one catchment). However the results were
accurate in two catchments and moderate in five catchments when radon tracers were used. The
method was most successful in catchments using tracers with a clear distinction between
groundwater and surface water concentrations, and where the groundwater end-member tracer
concentrations could be accurately estimated. In most catchments radon tracers provided the best
results.
In the current study it was found that by using a method incorporating longitudinal chemistry and
flow gauging results, the errors in inflow calculations could be significantly reduced. When the
model is used the error of the estimates of inflow is likely to be similar to or less than the minimum
estimate of the two individual methods. Thus, groundwater inflows could be accurately estimated
with this combined method in four of the nine catchments, and with moderate accuracy in the
remaining five catchments.
Chemical hydrograph separation involves monitoring changes in tracer concentrations in river
flow over time to determine changes in the relative proportions of surface runoff and groundwater
inflow. At any point in time, the proportion of river flow that is due to groundwater discharge is
estimated by comparing stream water chemistry with assumed values of surface runoff and
groundwater inflow. In this project, the groundwater concentration end-member was assumed to be
equal to the mean EC measured in groundwater, and the surface run-off end-member was assumed
to be equal to the minimum measured EC in river flow. Electrical conductivity sensor probes with
continuous readings were used in this analysis to capture all flow events. Sensors were located at a
total of 13 sites across nine of the catchments. Comparison with fortnightly samples showed major
discrepancies in three of the 13 sites. This is a major issue of concern, and the reliability of EC
sensors in surface water gauging stations should be assessed.
The groundwater inflow calculated from the chemical hydrograph separation method was very
sensitive to the end member values applied. In many catchments the end member values were not
well known and hence groundwater inflow calculations varied significantly depending on the
assumed end member value applied. There were also problems encountered in catchments where
end member values may have changed throughout the study period.
The hydrograph separation method distinguishes streamflow derived from surface runoff and that
derived from groundwater, based on the time-series record of streamflow. Hydrograph separations
were completed for flow data from each catchment using the Lyne and Hollick filter. The results
were sensitive to the operator-controlled parameter (α) selected. Calculated baseflow indices varied
by over 30%, when the applied α value varied between an upper value of 0.98 to a lower value of
0.925. The method produced base flow indices ranging between 0.24 to 0.52 within the ten
catchments using an α value of 0.925.
It should be emphasised that three of the methods trialled in this project enable estimation of
groundwater inflow, but cannot be used for estimating groundwater outflow. Although seven of the
ten studied catchments contained some losing reaches, the study has focussed on estimation of rates
of groundwater inflow. Methods which can only be used for estimating rates of groundwater
outflow have not been specifically trialled in this project.
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Methods for estimating groundwater discharge to streams – summary of field trials
Results of field trials
Mean flow rates within the studied reaches of these rivers ranged from 1.7  107 m3/yr to 2.9  108
m3/yr. The estimates of groundwater inflow as a percentage of streamflow over the study periods
ranged between approximately 15 and 60%. Within the studied stream reaches, groundwater inflow
appeared to be highly variable. Measured at a scale of 1 – 3 km, seven of the ten catchments
contained both losing and gaining reaches at this scale. Maximum estimated rates of gain ranged
over two orders of magnitude – from 0.15 m3/day/m in the Logan River to 17 m3/day/m in the
Cockburn River, with mean rates of gain mostly between 0.1 and 2 m3/day/m. The Belubula and
Cockburn catchments have small sections of river which appear to have distinct spikes in
groundwater inflow. These spikes may suggest groundwater inflow over 10 times that of the nearby
river sections. (Spatial variability of inflow rates would be greater when measured over smaller
distances.) The maximum observed loss rate in any of the catchments was 2 m3/day/m in the
Belubula River.
Baseflow Trends
Prediction of changes in groundwater inflow over time due to groundwater extraction can be made
using numerical groundwater models, or using simple analytical equations. It was not within the
scope of this project to compare these different methods, although estimates of stream depletion
due to pumping were made in each catchment, and these results were compared to historical
changes in baseflow suggested by hydrograph separation of historical streamflow data. Statistically
significant declines in annual baseflow was suggested by baseflow estimates in Cockburn, Elliot
and Tarcutta catchments. In Cockburn and Tarcutta catchments, the estimated reduction in
baseflow appears greater than can be explained by the reduction in recharge due to low rainfall, or
due to groundwater pumping. Other possible reasons for the potential baseflow decline include
reduction in irrigation or improvement in irrigation efficiency (so that less irrigation recharge
occurs), revegetation or development of plantations, however the magnitude of the reduction
appears much greater than can be explained by these processes. It is possible that this apparent
baseflow decline is an artefact of the hydrograph separation method, or that other changes within
the catchment (such as changes to river regulation, or changes in land use that have affected
velocities of surface runoff) have affected the hydrograph separation results. Further work is
required to understand if the apparent reduction in baseflow is real, and if so, what has caused it.
Broader implications
None of the available methods alone measure both spatial and temporal variation in groundwater
inflow. The choice of method for any particular study will therefore depend upon whether
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Methods for estimating groundwater discharge to streams – summary of field trials
information on spatial and temporal variations in groundwater discharge is of interest, or whether
only mean annual values over longer reaches and over time are required. If information on spatial
variations in groundwater inflow rate is needed, then flow difference and longitudinal river
chemistry methods should be considered. However, when information is required on mean annual
rates of groundwater discharge, methods which provide estimates at a single point in time are of
little value. Chemical hydrograph separation is probably the most suitable method for estimating
mean annual groundwater discharge rates, although uncertainty in end-member concentrations can
produce significant uncertainties in estimated groundwater inflow rates. Groundwater inflow
concentrations can be obtained from measurement of groundwater chemistry on bores within the
catchment, however the mean concentration in the sampled bores may not be the mean
concentration of groundwater discharging to the river. Concentrations of groundwater inflow can
also be estimated from modelling of longitudinal chemistry and flow gauging measurements.
Modelling carried out in a number of catchments within this project suggested that the groundwater
inflow concentration was significantly different to the mean groundwater concentration. Also,
changes in groundwater inflow concentration over time are possible, and there was some evidence
that this may be the case in several of the catchments studied in this project. Further work is needed
in this area. Studies to measure the concentration of surface runoff into the river would also be
useful.
The catchments studied in this project represent a range of climatic and geological environments.
Mean streamflow rates within the studied reaches of these rivers ranged from 1.7  107 m3/yr to 2.9
 108 m3/yr, and upstream catchment areas ranged from 150 to 1700 km2. Although the studied
rivers might be considered to be representative of many small to medium Australian rivers, the
project results may not be immediately transferable to much larger river systems. Although several
of the studied rivers were regulated, many of the larger rivers in southern Australia are much more
heavily regulated than those studied. As river regulation increases (and variations in flow are
reduced), hydrograph separation methods become more questionable. In very large catchments,
chemical hydrograph separation methods become more difficult, as spatial variations in endmember concentrations become more pronounced, and so the mean concentration of groundwater
inflow at any point in time becomes more difficult to determine.
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Methods for estimating groundwater discharge to streams – summary of field trials
Contents
1.
2.
3.
4.
5.
6.
Introduction
6
1.1
1.2
6
8
Context and Objectives of the Project
The Importance of Understanding Spatial & Temporal Variability
Spatial & Temporal Patterns of Water Exchange
10
2.1
2.2
2.3
2.4
10
13
14
16
Gaining & Losing Streams
The Water Balance
Influence of Groundwater Pumping on Streams
Consideration of Time Lags
Review of Methods
17
3.1
3.2
3.3
3.4
3.5
3.6
3.7
17
17
20
22
25
28
29
Introduction
Hydrograph Separation
Chemical Hydrograph Separation
Longitudinal River Chemistry
Hydraulic Gradient Analysis
Flow Difference
Suitability of Methods
Field Testing
31
4.1
Estimates of Surface Water – Groundwater Exchange
31
4.1.1
4.1.2
4.1.3
4.1.4
4.1.5
Flow Difference
Longitudinal River Chemistry
Hydraulic Gradient Analysis
Chemical Hydrograph Separation
Hydrograph Separation
32
36
39
41
44
4.2
4.3
Predictions of Streamflow Depletion by Pumping
Analysing Trends in Baseflow
46
50
4.3.1
4.3.2
4.3.3
Statistical Analysis of Hydrograph Separation
SIMHYD Modelling
Summary of Catchment Results
50
51
55
Conclusions - Broader Implications
58
5.1
5.2
5.3
5.4
5.5
58
59
61
62
62
Spatial and Temporal Patterns of Groundwater Discharge
Appropriateness of Methods
Transferability of Results
Effects of Groundwater Extraction
Key Findings
References
64
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Methods for estimating groundwater discharge to streams – summary of field trials
1.
Introduction
1.1
Context and Objectives of the Project
In most instances throughout Australia connected river and aquifer systems are managed as
independent resources and hence, surface water and groundwater are accounted for independently
(i.e. double accounting). This approach often results in the same parcel of water being allocated to
different users (i.e. double allocation). As water resources become over-allocated and streamflow
and groundwater levels decline, the financial (compensatory), legal and environmental implications
of double allocation are beginning to be realised, as is the necessity for integrated management.
To aid in integrated management, the magnitude of double accounting of the water resources must
first be understood. To this end, the Australian Government, through the Water for the Future –
Water Smart Australia program, funded Sinclair Knight Merz (SKM) and the Commonwealth
Scientific and Industrial Research Organisation (CSIRO) to develop a practical and moderately
priced methodology for assessing the range of levels of connection between groundwater and river
systems (i.e. double accounting), both in a spatial and temporal context. The approach was
developed on the basis of existing investigations and refined and tested by conducting field trials in
ten representative catchments (Figure 1). The tools developed will be an integral component of
integrated water resource management within Australia.
Specifically, the five objectives of this project are to:
1)
Develop methods for quantifying the degree of connection between river and groundwater
systems, in both a spatial and temporal context; and
2)
Install the necessary monitoring infrastructure and demonstrate the application of these
methods at ten representative catchments in eastern Australia.
3)
Provide estimates of the level of connection between groundwater and surface water in the ten
trial catchments, and the likely level of double accounting and double allocation of water
resources;
4)
Develop quantitative approaches for assessing surface water – groundwater interaction in
catchments with a range of data availability with consideration of the value of water resources;
and
5)
Communicate the project outcomes to local, State and Commonwealth decision-makers to
enable integrated water resource management.
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Methods for estimating groundwater discharge to streams – summary of field trials
This report describes the methods used for assessing surface water – groundwater connection, and
findings from the ten field studies. Detailed results from each of the ten catchments are described in
separate reports. Since the purpose of the project is to compare methods, the data and baseflow
estimates do not represent agreed absolute estimates of surface water – groundwater flux volumes.
As such they are not intended for use as inputs to any rules or regulations governing water resource
extraction in specific catchments. Rather, the results and conclusions of this report are intended
solely as a comparison of scientific methods for measuring surface water – groundwater exchange.
This report provides a summary of the key findings, focussing on the reliability of the different
methods, and also examines how rates of groundwater inflow are likely to change over time due to
groundwater pumping.
Figure 1. Locations of ten catchments in which methods for assessing surface water –
groundwater interaction are being trialled.

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Methods for estimating groundwater discharge to streams – summary of field trials
The ten catchments selected were thought to contain generally gaining rivers with small to medium
sized flows with a range of geological, landuse and climatic conditions. Summary statistics for the
ten studied catchments are given in Table 1. Studied river reaches ranged between 20.3 and 40.4
km in length. Mean annual river flow rates within these reaches ranged between 17  106 – 290 
106 m3. Catchment areas of these streams ranged from 153 – 2560 km2. The ratio of mean annual
streamflow to catchment area ranged from 0.03 – 0.9 m/y, and generally increased from south to
north.
Table 1. Catchment summary.

Catchment area is upstream of gauging station located within the study reach, and the flow rate is the mean
annual flow at that gauging station. Mean annual rainfall is for rainfall stations within the catchment upstream
of the studied reach, or adjacent areas. Higher rainfall often occurs within the upland areas of the catchment,
but this is not reflected in these figures.
Catchment
Barron
Belubula
Nambucca
Ourimbah
Tarcutta
Cattle Creek
Cockburn
Elliott
Logan
Hodgson
1
Reach Length
(km)
33.8
40.4
22.5
22.2
29.3
20.4
20.3
32.3
31.3
25.0
Catchment Area
(km2)
228
2560
431
153
1660
326
1130
471
1262
566
Rainfall
(mm)
1367
598
1342
1163
628
1565
686
1018
905
717
Flow
(106 m3/y)
130
140
181
19
150
290
64
17
170
26
Flow value is based on less than 2 years of data and may not be representative of long term average flows.
1.2
The Importance of Understanding Spatial & Temporal Variability
Over the past few decades, there have been numerous studies of surface water – groundwater
interaction. The investigations employ a range of different methods for characterising the level of
groundwater – surface water exchange and are highly valued due to their significant contribution to
our current understanding of the complexities of such processes. Many of these studies involve
intensive instrumentation of relatively short stream reaches. Detailed studies such as these involve
considerable costs and for this reason are unlikely to be routinely applied throughout Australia.
Further to the cost prohibitive nature of some of these studies is the limited consideration that has
been given to the viability and integrity of the range of different methods employed to assess the
process of river – groundwater interaction. That is, the majority of studies focus on a single method
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Methods for estimating groundwater discharge to streams – summary of field trials
for characterising the exchange flux. Studies that have directly compared results from different
methods are rare. Where only a single method is used, the integrity of the results cannot be tested
or verified.
As well as operating over relatively short spatial scales, many of the existing studies have taken
place over only relatively short periods of time. They do not, therefore, allow the long-term nature
of the system to be understood. Studies that consider behaviour of longer term behaviour of large
systems are extremely rare.
For instance, river water and groundwater chemistry have been compared in field trials to quantify
the rate of groundwater discharge to the river system at the time of sampling. The limitation
implicit in this method is that it is not well suited to assessing temporal variations of the
connection. While the analysis of river water hydrographs permits an estimate of changes in
groundwater discharge to river systems over time, the method does not provide information on
spatial variations.
In any one catchment, the level of river – groundwater interaction may be highly variable in both a
spatial and temporal context. For instance, factors controlling the level of interaction are rarely
uniform along a reach of river. The typically heterogeneous nature of the hydrological and
hydrogeological controls may potentially contribute to significant spatial variability in the level of
interaction along a river reach. Likewise, the level of interaction may fluctuate widely in a temporal
context. It is sometimes convenient to express surface water groundwater in terms of the baseflow
index, which is the proportion of river flow that is derived from groundwater. Yet this quantity will
be extremely variable in time. For example, tropical rivers of northern Australia are likely to have a
very small baseflow index when expressed as an annual figure, as the highest river flows occur
within the wet season and are dominated by rainfall runoff. However, in the dry season
groundwater inflows are likely to be the sole source of river flow, and hence, the baseflow index at
this time would be close to 100%.
For these reasons, any representative estimate of the level of interaction between river and
groundwater systems will require a consideration of the potential for both spatial and temporal
variability. Only by considering the range of levels of interaction can the issue and the potential
impacts be truly recognised and appropriate and effective management responses developed.
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Methods for estimating groundwater discharge to streams – summary of field trials
2.
Spatial & Temporal Patterns of Water
Exchange
The process of groundwater and surface water interaction is generally complex and rates of
exchange are highly variable, being dependent upon a range of parameters including geology,
geomorphology and climate. For the purposes of providing a technical underpinning to this project,
an introduction to the physical process of groundwater – surface water interaction is presented
below. Some of these concepts have also been recently summarised in Winter et al. (1998) and
Reid et al. (2009).
2.1
Gaining & Losing Streams
In most instances, groundwater and an adjacent stream form a continuous body of water that
transmits freely between the stream channel and the aquifer. A stream that receives groundwater
discharge is known as a “gaining” stream, whereas a stream that loses water to, or recharges, the
groundwater is known as a “losing” stream.
In gaining streams, the watertable slopes towards the stream, permitting groundwater to discharge
or be added to the stream flow (Figure 2). The component of stream flow that is sourced from
groundwater discharge is known as baseflow. During low flow conditions in a stream, baseflow
constitutes a high proportion of the total stream flow. The rate at which groundwater discharges
into a stream is largely determined by the slope or angle of the watertable (hydraulic gradient) and
the permeability (or hydraulic conductivity) of the adjacent aquifer. Provided all other variables are
equivalent, a steep hydraulic gradient will impart a higher groundwater discharge rate relative to a
site with a subdued groundwater gradient.
A losing connected stream occurs when the water table slopes away from the stream channel,
causing stream flow to seep into (or recharge) the underlying groundwater system. The rate at
which water is recharged to groundwater is controlled in a similar manner as gaining streams, that
is, by the slope of the watertable and the hydraulic conductivity of the aquifer.
Importantly, in some instances, a losing stream becomes disconnected from the underlying
groundwater system (Figure 2). In a disconnected system, an unsaturated zone separates the stream
from the underlying aquifer. Brunner et al. (2009) showed that disconnected streams can only occur
when there is a layer of low permeability beneath the streambed. Importantly, there is still flow of
water from a disconnected stream to the groundwater system, and this flow occurs as drainage
through the unsaturated zone.
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Methods for estimating groundwater discharge to streams – summary of field trials

Figure 2. Classification of Surface Water – Groundwater Exchange
After Winter et al., 1998
For gaining and losing connected streams, the flow rate between the river and the aquifer is
approximately linearly related to the head gradient (Figure 3). However, for losing disconnected
streams, the flow rate from the river to the aquifer is not controlled by the position of the watertable
but by the streambed thickness and hydraulic conductivity alone. Brunner et al. (2009) showed that
a transition stage occurs between losing connected and losing disconnected streams, in which flow
is unsaturated, but the infiltration rate is still related to the watertable position.
Spatial variations in rates of groundwater inflow to streams result from a number of factors. The
most obvious of these is spatial variability in the hydraulic properties of aquifer and river bed
sediments. Hydraulic conductivities of streambed sediments range from less than 10-9 to more than
10-2 m/s (10-4 – 103 m/d) (Calver, 2001). For example, Cey et al. (1998) measured very large
variability in groundwater inflow to a small stream in southern Ontario, with nine of the twenty-one
sites recording zero inflow, and others varying between 0.001 and 0.33 L/s. The authors believed
that the variability was due to extreme heterogeneity in stream bed sediments. Genereux et al.
(2008) measured hydraulic conductivity at 46 sites along a 262.5 m reach of West Bear Creek,
PAGE 11
Methods for estimating groundwater discharge to streams – summary of field trials
North Carolina. Hydraulic conductivity ranged over almost four orders of magnitude, from about
0.01 to 66 m/d. Lamontagne et al. (2005) observed similar variation at much larger scales in the
Murray River, north eastern Victoria. The authors used propagation of rivers flood waves into the
adjacent floodplain aquifer to measured hydraulic diffusivities of 11,000 – 18,000 m2/day for point
bar sand deposits and 10 – 35 m2/day for clay lined river banks. Although flow rates were not
measured in this study, the variation in flow rates would be expected to be similar to the measured
variation in hydraulic diffusivities.
Figure 3. Theoretical Relationship Between Infiltration Rate of a Losing Stream and
Water Table Depth

For connected systems (fully saturated flow), the flow rate is linearly related to the head gradient. However, as
the water table drops, an unsaturated zone can sometimes develop, and then flow rate is no longer linearly
related to head gradient. For a disconnected stream, the flow rate is independent of the watertable position,
and is determined only by the river depth and the properties of the streambed.
Variations in inflow rates can also be associated with river channel morphology. For example,
mountain stream reaches can sometimes be classified into channel units with distinct stream bed
and stream water slopes, which vary on the spatial scale of 1 – 10 channel widths along the stream
(Grant et al., 1990). Groundwater inflow and outflow rates vary spatially in response to variations
in channel units, with areas of river outflow tending to be concentrated toward the end of pools
(with low bed slope), and groundwater inflow concentrated towards to end of steps (with high bed
slope). Harvey and Bencala (1993) observed significant spatial variations in rates of surface water
– groundwater exchange over a 36 metre long reach of a third-order stream in Colorado, which
they attributed to this process. Although the stream was receiving groundwater inflow from the
local aquifer, sections of both inflow and outflow occurred within the studied section of river.
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Methods for estimating groundwater discharge to streams – summary of field trials
Inflow and outflow rates varied over scales of 1-10 metres, and corresponded with changes in the
bed slope of the river.
In lowland streams, regular variations in groundwater inflow rates may be associated with river
meander geometry. Groundwater flow lines will tend to concentrate on the outside of river meander
bends, which should give rise to relatively high rates of groundwater inflow per unit length of river
on the outside of meander bends, and lower rates on the inside of meander bends (Cherkauer and
McKereghan, 1991; Linderfelt and Turner, 2001). Hydraulic principles suggest that groundwater
inflow to streams should be concentrated near the river banks, with lowest rates of discharge in the
middle of the stream (Pfannkuch and Winter, 1984). In practice, of course, the spatial patterns of
groundwater inflow will depend upon the complicated interrelation between channel morphology,
and aquifer and river bed permeability, and will also be influenced by processes within the
catchment, such as spatial variations in aquifer recharge and discharge. It is clear, however, that
spatial variations in exchange rate can be high, and this will have important implications for
measurement approaches.
2.2
The Water Balance
Under stable climatic conditions, groundwater in an aquifer will reach an equilibrium (or steady)
state, where the volume of recharge to the aquifer over a significant time period is equal to the
volume of water discharged. At any point in time, however, recharge may be different from
discharge. Groundwater levels will generally rise in periods when recharge exceeds discharge, and
fall when discharge exceeds recharge. However, under steady state conditions, groundwater levels
will fluctuate seasonally around consistent levels.
If a significant new discharge process occurs (e.g. groundwater pumping) a proportion of
groundwater will initially be removed from the groundwater storage, contributing to a lowering of
groundwater or watertable levels. However, the groundwater levels will reach a new steady state
over time if the pumping discharge continues. Steady state conditions will be achieved by either
reducing some other discharge process (such as base flow) or increasing recharge processes (such
as stream leakage) or possibly both. While the groundwater is moving from one long-term steady
state to another, it is considered to be in a short-term transient state. The time period to reach the
new steady state will be governed by the size of the aquifer and the magnitude of the change in
discharge.
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Methods for estimating groundwater discharge to streams – summary of field trials
2.3
Influence of Groundwater Pumping on Streams
Groundwater levels surrounding a bore will decline upon commencement of pumping. If situated in
the proximity of a gaining stream, the pumping bore and associated declining groundwater levels
will reduce the slope of the watertable, and in turn, the rate of groundwater discharge into the
stream. The lower rate of groundwater discharge (termed “reduced discharge” in Figure 4B)
translates to a decrease in stream flow.
In circumstances where pumping continues for an extended period or the bore is in close proximity
to the stream, the slope of the watertable can reverse and become a losing stream (Figure 4C).
Consequently, groundwater discharge will reduce to zero and the stream will begin to recharge the
groundwater. The loss of stream flow to groundwater is termed “increased recharge” or “induced
recharge”. The decreased flow experienced by a stream that changes from gaining to losing is,
therefore derived from both reduced discharge and increased recharge.
With reference to groundwater pumping in proximity to a losing stream, the stream flow is reduced
only by the process of increased recharge.
A considerable delay may exist between the commencement of pumping and decline in stream
flow. The delay will be governed by a range of variables including the distance between the bore
and stream, the extent and depth of the cone of depression (being dependent on pumping duration,
transmissivity, storage co-efficient and pumping rate). Secondly, the reduction of stream flow will
often be less than the pumped volume on account of there being some increased recharge or
decreased discharge from other sites.
The above discussion has been simplified to demonstrate the concept of groundwater – surface
water interaction and the potential impacts of groundwater pumping on stream flow. The example
provided concerns an unconfined aquifer, which is an aquifer in which there are no confining beds
(layers) between the saturated zone and the surface.
In contrast, a confined aquifer is an aquifer that is overlain by a confining bed (also known as an
aquitard). In the case of a confined aquifer, the hydraulic pressure in the aquifer is above the top of
the aquifer and the level to which water will rise in a bore in a confined aquifer is known as the
potentiometric surface. The confining bed has a lower hydraulic conductivity (i.e. permeability)
than the aquifer and accordingly, the rate of leakage through the aquitard is low relative to the
aquifer.
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Methods for estimating groundwater discharge to streams – summary of field trials

Figure 4. Effect of Groundwater Pumping on a Gaining Stream
Source: Winter et al., 1998.
An aquifer where the confining layer is “leaky” (i.e. transmits a moderate amount of water) is
known as a semi-confined aquifer. The concept of reduced stream flow caused by groundwater
pumping is also relevant for a semi-confined aquifer. As groundwater pressure within a semiconfined aquifer is reduced by groundwater pumping, downward leakage from the overlying
unconfined aquifer through the aquitard is induced, which in effect generates a similar impact on
stream flow as illustrated in Figure 4. The key difference is a longer delay between pumping and
stream flow impact due to the dampening effect of the aquitard. Many aquifers that are considered
to be confined aquifers are in practice semi-confined aquifers, with the degree of confinement
dependent on the hydraulic conductivity, thickness and continuity of the aquitard. The magnitude
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Methods for estimating groundwater discharge to streams – summary of field trials
of the impact on stream flow depends, in turn, on these aquitard properties and the length of the
time lag. Nonetheless, the same fundamental principles apply regardless of whether the aquifer is
unconfined or semi-confined.
2.4
Consideration of Time Lags
A time lag exists between the commencement of groundwater pumping and the impact on stream
flow. In general, as the distance between the pumping bore and stream increases so does the lag or
delay between the commencement of pumping and the impact on stream flow (marked by reduced
discharge and / or induced recharge). A number of methods of varying complexity are available for
quantifying this time lag, and these have been recently reviewed by Rassam and Werner (2008).
When groundwater pumping ceases there is also a time lag between cessation of pumping and the
reduction in stream flow impacts (Figure 5).

Figure 5. Illustration of the Time Lag Affect
Simulation is for a bore located 500 m from a stream, with aquifer transmissivity T = 100 m/d and specific yield
S = 0.1.
Distance to stream = 500 m, T = 100 m/d , S = 0.1
Streamflow depletion rate (as a % of pumping)
100%
90%
impacts during nonpumping period
80%
impacts during 3 month
pumping period
70%
impact if pumping
continued after 3
months
Summer period
47% impact if pumping were
continuous for 8 months
Maximum impact occurs 3.5 months
after pumping started and 0.5 month
after pumping ceased (ie 0.5 month lag
before impacts begin to decrease)
60%
50%
Pumping ends
after 3 months
40%
Impacts still evident 9 months
after cessation of pumping
30%
Pumping
starts 1
November
20%
10%
No impact for first 2 weeks of
pumping (ie 2 week lag)
0%
Jul
Aug
Sep
Oct
Nov
Dec
Jan
Feb
Mar
Apr
May
Jun
Jul
Aug
Sep
Oct
Nov
Time
I:\wcms\112\GWATER\UTIL\stream_interference(draft)3.xls[jenkinsChartRlinear_date (2 (4)]
PAGE 16
Methods for estimating groundwater discharge to streams – summary of field trials
3.
Review of Methods
3.1
Introduction
Over the past few decades, numerous different methods have been used for studying surface water
– groundwater interaction. In particular, within the last five years, there have been a number of
reviews of these methods (e.g., Kalbus et al., 2006; Brodie et al., 2007; Turner et al., 2009). Many
of the methods that have been used are very labour intensive and hence relatively costly to
implement. In many cases, they provide only local-scale estimates of water flow, which cannot be
directly extrapolated to larger regions. Although valuable for understanding processes at local
scales, these methods are not likely to be widely implemented. For example, seepage meters and
temperature studies are costly and provide only very local-scale estimates of flow. Other methods
provide an indication of groundwater – surface water exchange, but do not allow quantification of
the exchange rate. Ecological indicators, hydrogeological mapping, geophysics and remote sensing
methods, for example, provide only qualitative information on exchange rates and so have not been
considered in this study. Based on these considerations, five methods were identified as potentially
suitable for routine use for quantification of the volume of water flowing between groundwater and
rivers at regional scales. These methods are:
1)
Hydrograph Separation
2)
Chemical Hydrograph Separation
3)
Longitudinal River Chemistry
4)
Hydraulic Gradient Analysis
5)
Flow Difference
The following sections describe the principles of operation of each of these five methods and the
temporal and spatial scales on which they operate.
3.2
Hydrograph Separation
Principle of Method
The hydrograph separation method (sometimes called baseflow separation) aims to distinguish
streamflow derived from surface runoff and that derived from groundwater, based solely on the
time-series record of streamflow. The method is popular because it operates only using readilyavailable streamflow data. Usually, peak flows are attributed to a combination of surface runoff and
groundwater inflow, and the underlying river flow that persists after rainfall has ceased is attributed
PAGE 17
Methods for estimating groundwater discharge to streams – summary of field trials
to groundwater discharge. The method relies on the principle that runoff events are of relatively
short duration, whereas groundwater responds more slowly to rainfall recharge (Figure 6).
Although separation of surface runoff and groundwater during the stormflow peak can be
somewhat subjective, it is usually assumed that (Evans and Neal, 2005):

baseflow recession continues after the rise of the streamflow hydrograph;

baseflow will peak after the streamflow hydrograph because subsurface flows are slower than
surface flows; and

the baseflow hydrograph will rejoin the total hydrograph as direct runoff ceases.
A number of routines have been developed for separating groundwater inflow from surface runoff
based on these principles, many of which are graphical, or use simple data processing or filtering
procedures. Summaries of a number of the more commonly-used routines can be found in Brodie
and Hostetler (2005) and Evans and Neal (2005).
Streamflow data is usually obtained from a gauging station that reads river height at regular
intervals and calculates flow rate from river height using a rating curve. The rating curve is derived
from manual flow gaugings that have been made at different river heights. The gauging station
usually involves a small weir on the river which has been deliberately constructed to increase the
sensitivity of river level to flow rate at low flows.
The baseflow separation method relies on the principle that runoff events are of relatively short
duration whereas groundwater responds more slowly to rainfall recharge. Empirical studies have
determined that the duration of surface water flow following rainfall will be a function of the
catchment area. The most widely used relationship is that of Linsley et al. (1975):
t  0.8278 A 0.2
[1]
Where:
- t is the time (in days) between the storm crest and the end of surface runoff
- A is the catchment area (in square kilometres).
Thus the time for surface runoff to cease is estimated to be approximately 2.1 days for a catchment
area of 100 km2, increasing to 5.2 days for a catchment area of 10 000 km2. A number of baseflow
separation routines use this equation to determine the time after which streamflow is comprised
solely of groundwater inflow. The more difficult issue is to separate surface runoff from
groundwater inflow within the streamflow peak, and the result will depend somewhat on the
routine that is used to do this partitioning.
PAGE 18
Methods for estimating groundwater discharge to streams – summary of field trials

Figure 6. Separation of Streamflow Into Surface Runoff & Groundwater Flow
The flow recession preceeding the streamflow peak is believed to represent groundwater inflow (with no
surface runoff). This baseflow recession continues after the rise of the streamflow hydrograph. The baseflow
peaks after the total streamflow peak and rejoins the streamflow hydrograph as surface runoff ceases.
While a number of papers have compared different methods of hydrograph separation (e.g., Nathan
and McMahon, 1990) there are few studies that have compared hydrograph separation results with
those obtained from other, independent studies. One interesting study, however, is that of
Werner et al. (2006) who estimated groundwater discharge to Sandy Creek in the Pioneer Valley,
Queensland, using a number of different baseflow separation routines. The results of the baseflow
separation were compared with groundwater discharge rates estimated using a coupled surface
water – groundwater model which was calibrated to groundwater levels measured in observation
bores and streamflows measured at a gauging station. Although the model was unable to accurately
reproduce very low river flows (a discrepancy that the authors attributed largely to
evapotranspiration of riparian vegetation and river evaporation) it provided reasonably good
simulation of river flows between 10 and 10,000 ML/day. Werner et al. (2006) found that the
baseflow separation methods of Nathan and McMahon (1990) and Cordery (1993) greatly
overestimated groundwater discharge rates during high flow events, compared to those estimated
from the model. The discrepancy was often more than an order of magnitude.
Hydrograph separation will be most accurate when surface runoff events are well-defined, but
represent a relatively small proportion of the flow to the river. This is likely to be the case in small
catchments, where travel times for surface runoff are short. The method is also most applicable to
PAGE 19
Methods for estimating groundwater discharge to streams – summary of field trials
undeveloped catchments. If river losses occur within the catchment (due to pumping, evaporation,
transpiration of riparian vegetation) then this water will not appear as flows at the gauging station
and so will not be included as groundwater inflow to the river. The method thus estimates net
groundwater inflows within the catchment, rather than total inflows. Evans and Neal (2005) have
also noted that flow releases from upstream reservoirs may also produce a low flow signal that can
be misinterpreted as baseflow.
Scale of Operation
Hydrograph baseflow analysis uses available streamflow data from a gauging station and provides
information on groundwater discharge as a function of time over the period of record. The results
relate to the entire catchment upstream of the gauging station and so the method provides a
spatially integrated assessment and does not provide any information on where within the
catchment the groundwater discharge may be occurring.
3.3
Chemical Hydrograph Separation
Principle of Method
This method examines changes in tracer concentrations in river flow over time and interprets these
in terms of changes in the relative proportions of surface runoff and groundwater inflow. The
method is usually applied as part of an intensive field study, with measurements of river chemistry
made at hourly or more frequent intervals during storm events (Figure 7). At any point in time, the
proportion of river flow that is due to groundwater discharge is calculated using the mass balance
equation:
Qg
Qt

c  c
c g  c
[2]
Where:
- c, cr and cg are the tracer concentrations in the river, in runoff and in groundwater, respectively
- Qt is the measured total river flow
- Qg is the volume of groundwater inflow
While the use of a single tracer allows quantification of the relative proportions of groundwater and
surface runoff in streamflow, use of two or more tracers will allow a greater number of endmembers to be resolved. For example, when two tracers are used, mass balance equations can be
written to determine the relative proportions of three end-members – usually groundwater inflow,
surface runoff and shallow soilwater flow (Uhlenbrook and Hoeg, 2003).
PAGE 20
Methods for estimating groundwater discharge to streams – summary of field trials
The success of the chemical hydrograph separation method ultimately relies on an adequate
chemical differentiation of the source waters (and quantification of the end-member
concentrations). This will often dictate the choice of tracers. Because of their chemical inertness,
2
H, 18O, silica and chloride have usually been preferred. The groundwater concentration is often
determined from sampling river water during baseflow conditions. However, McCallum et al.
(2010) recently pointed out that even during baseflow conditions streamflow can be diluted by
bank storage return flows. It is therefore preferable to directly measure groundwater chemistry to
determine the groundwater end-member concentration. Direct precipitation (runoff) input is
sometimes determined from sampling rainfall or canopy throughfall (Joerin et al., 2002). In threecomponent mixtures, the soil water end-member is determined from sampling soil water using
suction cup lysimeters (Buttle and Peters, 1997; Joerin et al., 2002), throughflow collectors
installed in pits (Buttle and Peters, 1997) or shallow piezometers (Joerin et al., 2002).
The chemical composition of rainfall can change during a storm event, and if this is significant
(which is usually the case with 2H and 18O) then rainfall chemistry needs to be measured over the
duration of the study. The composition of surface runoff can also differ from that of rainfall. For
example, Buttle and Peters (1997) noted that the silica concentration in overland flow was often
significantly greater than that in rainfall, due to interaction between surface runoff and underlying
soil materials. Concentrations in groundwater are likely to be relatively stable over time but can
show significant spatial variability. If end-members are not well-defined, then the estimated endmember contributions will be in error. Bazemore et al. (1994) used a monte-carlo analysis to show
that variability of end-member concentrations in a three-component hydrograph separation applied
to a small forested headwater catchment in Virginia, produced highly uncertain separation of storm
events.
PAGE 21
Methods for estimating groundwater discharge to streams – summary of field trials
Figure 7. Changes in Chloride & Silica Concentration in Mattole River (Northern
California) During a Storm Event

Decreases in concentrations of both tracers is due to dilution with surface runoff, which has lower
concentrations than groundwater. From Kennedy et al. (1986).
Scale of Operation
In principle, this method can be used to determine the relative proportion of groundwater inflow at
any point along a river. The measurement then refers to the entire catchment upstream of the point
of measurement. Usually, however, the method is applied on water chemistry measured at a
permanent gauging station and the availability of river flow data allows a volumetric estimate of
groundwater inflow. Importantly, the method provides no information on where within the
catchment the inflow might have occurred. However, because of the need to accurately define
surface water and groundwater inflow concentrations, the method is best suited to small
catchments, where these parameters are likely to be relatively constant.
3.4
Longitudinal River Chemistry
Principle of Method
Comparison of surface water and groundwater chemistry can also be used to determine spatial
variations in groundwater inflow to a river. In this method, measurements of river chemistry are
made along a stream reach at a particular point in time (usually within a period of 1-2 days).
Measurements of groundwater chemistry are also made and rates of groundwater inflow are
determined from downstream changes in water chemistry using a mass balance approach. Usually,
measurements of river chemistry are performed during baseflow conditions when the only inflows
PAGE 22
Methods for estimating groundwater discharge to streams – summary of field trials
to the river are from groundwater. As the method relies on their being no surface runoff at the time
of measurement it is not necessary to characterise this end-member.
For a conservative tracer (such as the chloride ion), in a gaining river, the equation for
concentration with distance can be expressed:
c I ci  c 

x
Q
[3]
Where:
- c is concentration of the tracer in the river
- ci is its concentration in groundwater inflow
- I is the groundwater inflow rate
- Q is the river flow rate
- x is distance downstream
In principle, a number of different tracers may be used to identify groundwater inflow, although in
practice tracers that have distinct surface water and groundwater concentrations are usually chosen.
The ideal tracer will be one which has a relatively uniform concentration in groundwater and
whose groundwater concentration differs markedly from its concentration in the river. The simplest
tracer is electrical conductivity, as rivers will usually have lower electrical conductivity than
groundwater. Where the electrical conductivity of groundwater is very high, then this can be a
sensitive indicator of groundwater inflow. In some areas, however, electrical conductivity is not a
particularly useful tracer because surface water and groundwater concentrations are not sufficiently
distinct. However, particular ions may still provide the required discrimination. For example, Cook
et al. (2003) used chloride and magnesium concentrations to assist identification of locations of
groundwater inflow to the Daly River, Northern Territory, while Genereux et al. (1993) used
calcium to help quantify groundwater discharge to a small creek in Tennessee, USA. It should be
noted, however, that in situations in which groundwater provides a significant contribution to river
flow the surface water will become more similar to groundwater as the proportion of groundwater
in the river increases with distance downstream. In these situations, the ability of ionic chemistry or
electrical conductivity to identify groundwater inflow will diminish.
Another tracer that is beginning to be more widely used for identifying locations of groundwater
discharge to rivers is radon. Radon (222Rn) is a radioactive gas with a half-life of 3.8 days. It is
produced by the radioactive decay of uranium-series isotopes (its immediate parent is 226Ra).
Within the saturated zone radon produced by decay of uranium isotopes attached to the aquifer
matrix is immediately dissolved in the groundwater. After groundwater containing radon
discharges to surface water bodies, radon concentrations decrease due to gas exchange with the
atmosphere (which is low in radon) and radioactive decay. High radon concentrations are therefore
present in surface waters only in the immediate vicinity of points of groundwater inflow and for
PAGE 23
Methods for estimating groundwater discharge to streams – summary of field trials
relatively short distances downstream of such locations. It has therefore been used as a tracer of
groundwater discharge to rivers and there are documented studies from Puerto Rico (Ellins et al.,
1990), USA (Lee et al., 1993; Genereux et al., 1994) and Japan (Yoneda et al., 1991). In Australia
it has been used in the Daly River, Northern Territory (Cook et al., 2003), Burdekin River,
Queensland (Cook et al., 2004) and Cockburn River, New South Wales (Cook et al., 2006).
Because radon is continually lost from the river by the combined processes of radioactive decay
and gas exchange, radon activities in the river will always be less than those in the groundwater.
Often, radon activities in the river will be less than those in the groundwater by several orders of
magnitude, making radon a particularly sensitive tracer and able to identify very low rates of
groundwater inflow.
The longitudinal river chemistry method relies on the existence of a clear differentiation between
river and groundwater chemistry and also requires accurate determination of the groundwater endmember. The choice of tracer for the analysis will thus usually be determined based on an
examination of river and groundwater chemical composition. It should be noted that some ions are
not conservative and so changes in concentrations within rivers might be due to chemical reactions
rather than groundwater inflow. In the case of radon quantification of the groundwater inflow rate
requires estimation of the gas transfer velocity, which describes the rate at which radon moves from
the surface water into the atmosphere. Another difficulty with radon is distinguishing radon input
to the river from regional groundwater inflow from that contributed by water that may move in and
out of the hyporheic zone (Cook et al., 2006). Water that moves into the hyporheic zone will
accumulate radon from decay of 226Ra contained in these sediments, which will be transported to
the river as water moves back from the hyporheic zone into the river. Furthermore, radon may not
be able to distinguish between regional groundwater discharge and bank storage return flow. As
with hyporheic exchange, rivers waters that are temporarily stored in the river banks will acquire
radon from these sediments when water in bank storage is transported back into the river as the
river level drops. Chloride is perhaps the most reliable tracer of regional groundwater discharge
provided that there is sufficient differentiation between surface water and groundwater
concentrations. Chloride is conservative and will not be affected by bank storage return flow or
water exchange with the hyporheic zone. However, in most cases, the use of more than one tracer
is preferred.
Importantly, longitudinal tracer studies do not directly allow measurement of rates of groundwater
outflow. In fact, river losses do not cause any downstream change in concentration. If areas of both
groundwater inflow and outflow occur in a river then longitudinal tracer methods will generally
estimate the total groundwater inflow rate and will be relatively insensitive to river losses (Cook et
al., 2006). However, combining river chemistry with flow gauging can permit regions of both loss
and gain to be identified and the rates of exchange quantified.
PAGE 24
Methods for estimating groundwater discharge to streams – summary of field trials
Scale of Operation
This method provides information on groundwater inflows only on the day(s) on which sampling
takes place. The method is most useful for identifying locations of groundwater inflow and for
quantifying the spatial variations in inflow rates. It can operate over reasonable large scales and has
been used to study river reaches up to 120 kilometres in length (Cook et al., 1993). The intensity of
sampling (distance between sampling points) that will produce the greatest resolution of
groundwater inflow can be estimated from river characteristics (particularly flow velocity) using an
equation developed by Cook et al. (2006).
3.5
Hydraulic Gradient Analysis
Principle of Method
The direction of flow between the groundwater and a river can be determined by comparing the
hydraulic heads within the groundwater with the water level in the river. If the river level is higher
than the groundwater there will be a potential for the river to leak water into the groundwater.
Conversely, if the river level is lower than the groundwater level adjacent to the river then there is a
potential for groundwater to flow into the river. It is possible to estimate the magnitude of the water
exchange using Darcy’s Law, which calculates flow as the product of the hydraulic gradient and
transmissivity:
q T
h
x
[4]
Where:
- Q is the flow rate (per unit length of river)
- T is the transmissivity
- h is the hydraulic head
- x is distance
The method can be applied either using a bore adjacent to the river (h is the difference between
the water level in the bore and the level in the river, and x is the perpendicular distance of the bore
from the river); or using a mini-piezometer installed beneath the stream bed specifically for this
purpose (h is the difference between the water level in the mini-piezometer and the level in the
river, and x is the vertical distance between the river bed and the top of the piezometer screen;
Figure 8). In each case, accurate estimation of the transmissivity between the river and the
piezometer is critical.
PAGE 25
Methods for estimating groundwater discharge to streams – summary of field trials
Figure 8. Estimation of Groundwater Inflow or Outflow from Difference Between the
Water Level in a Mini-Piezometer

The water level in the river is often measured using a stilling well, which is screened within the river. This
diagram depicts river loss, because the water level in the mini-piezometer is lower than the river level. From
Brodie et al. (2007).
Scale of Operation
This method can operate on a range of different spatial and temporal scales, depending on the
distance of the bores from the river. In the case of mini piezometers installed beneath the river bed,
the method provides a very local estimate of the flux with a high temporal resolution. If the bore is
further from the river, then more regional and temporally averaged flow rates are obtained (see
below).
As discussed above, groundwater inflow and outflow rates show relatively high variations over
small spatial scales. Therefore, pieozometers need to be positioned to so that measured hydraulic
heads reflect those in the regional aquifer, and not local-scale processes. If the bore is placed very
close to the river, then the head gradient can be difficult to measure and the estimated flow rate
may not be representative of the larger region. However, as the distance between the bore and the
river increases, other problems are created.
The use of mini-piezometers, or piezometers placed very close to the river, will measure very local
exchange rates which may not be representative of the river reach. For example, Cey et al. (1998)
calculated rates of groundwater inflow and outflow for a small stream in southern Ontario from
vertical hydraulic gradients measured beneath the stream and 21 mini-piezometers installed one
metre beneath the stream bed, together with estimates of hydraulic conductivity estimated from
slug tests performed on these mini-piezometers. The authors measured very large variability in the
PAGE 26
Methods for estimating groundwater discharge to streams – summary of field trials
hydraulic gradients, with values ranging from zero to 0.55. The mean estimated rate of
groundwater inflow was only one quarter of the net inflow estimated using the flow difference
method. The authors believe that the mini-piezometers failed to provide an accurate estimate of
groundwater discharge due to the extreme heterogeneity, with most groundwater discharge
assumed to be occurring through high conductivity sediments that were not sampled. The authors
also noted that the use of mini-piezometers installed beneath the streambed only allows estimation
of vertical upward flow into the river through the base of the stream. Bank seepage, and shallow,
horizontal flow to the stream are neglected, even though they may contribute a substantial
proportion of the groundwater inflow. The presence of horizontal, low conductivity layers beneath
the streambed would enhance the importance of horizontal flow into the river, which would not be
measured using mini-piezometers. Similarly, calculation of the exchange from using a bore
adjacent to the river assumes that flow is entirely horizontal. It may thus underestimate the true
exchange rate if the aquifer is thick and the bore is very close to the river.
If the piezometer is placed a significant distance from the river, estimation of the appropriate value
of hydraulic conductivity becomes more difficult. The magnitude of flow between stream and
aquifer may be controlled by the hydraulic conductivity of streambed sediments rather than the
aquifer hydraulic conductivity (especially if the hydraulic conductivity of the river bed sediments is
lower than that of the aquifer). In this case, use of the aquifer transmissivity in Equation [4] will
overestimate the exchange rate. Instead, the mean transmissivity between the bore and the river is
required, a parameter which may be difficult to determine.
Also, as the distance of the bore from the river increases it becomes less likely that the hydraulic
gradient measured between the bore and the river is representative of the hydraulic gradient
immediately adjacent to the river. For example, where the river rises following a major runoff
event and flow from the river to the groundwater occurs, the water level is a bore located some
distance from the river may still be above river level, suggesting flow towards the river. Thus, as
the distance of the bore from the river increases, the method will progressively fail to measure the
short-term dynamics of the relationship between the groundwater and the river.
Use of Equation [4] assumes a linear relationship between head gradient and flow rate. If this is the
case, then temporal changes in hydraulic gradient can be used to estimate temporal changes in flow
rate, even if this rate is not accurately quantified. However, it is not clear that this linearity will
hold in heterogeneous systems. In particular, consider a system in which the river bed has a low
permeability but the river banks have higher permeability. We might envisage now that the
exchange rate will depend upon the depth of water in the stream (or the height of the bank where
the river is connected to the aquifer). Further examination of the extent of this linearity is required.
Finally, it should be noted that this technique cannot be applied to losing streams, when the point at
which disconnection of the river from the groundwater is approached. Equation [4] implies a linear
PAGE 27
Methods for estimating groundwater discharge to streams – summary of field trials
relationship between head gradient and flow rate. While this approach may be reasonable for a
saturated system, it does not predict the transition from a connected to a disconnected system
(Brunner et al., 2009).
3.6
Flow Difference
Principle of Method
This method is perhaps the most straightforward of all of the methods for characterising surface
water – groundwater interactions. At the simplest level, the flow rate in the stream at a time when
surface runoff is likely to be negligible is assumed to be equal to the groundwater discharge rate in
the catchment upstream of the gauging station. Usually measurements of stream flow rate can be
carried out at a number of locations along the river on the same day and the groundwater inflow (or
outflow) is estimated from the differences between adjacent flow gaugings.
The method assumes that the changes in river flow are due to groundwater inflows and outflows
and so will only be accurate if all other inflows and outflows are negligible or have been quantified.
The method is usually applied at sufficient time after rainfall so that it can be assumed that
downstream increases in river flow are due to groundwater inflow. However, groundwater inflow
will be underestimated if river losses occur over that same reach. Although evaporation will usually
be small if pumping from the river (or from groundwater adjacent to the river) is significant then
these losses may dominate the observed variations in river flow, making this technique of limited
value. The method only determines net inflow (or outflow) between gauging locations. If areas of
inflow and outflow occur between gauging stations, then these will not be identified.
The error on individual flow gaugings also limits the ability of this technique to accurately quantify
low rates of groundwater inflow. Errors in manual flow gaugings will usually be between 5 and
15%, depending largely upon the number of flow measurements that are made across the width of
the stream (Carter and Anderson , 1963; Cey et al., 1998; Langhoff et al., 2006). However, errors
may be larger than this if a significant fraction of the river flow occurs beneath the stream bed.
Errors in flow rate estimated at permanent gauging stations will depend upon the data used to
construct the rating curve, and the stability of the rating curve over time.
Scale of Operation
This method provides information on groundwater inflows only on the day(s) on which sampling
takes place. The method is most useful for identifying locations of groundwater inflow and for
quantifying the spatial variations in inflow rates. Conceptually, if more than one gauging station is
located along a reach of river (and there is no tributary inflow between the stations), then
PAGE 28
Methods for estimating groundwater discharge to streams – summary of field trials
differences between these station records can yield some information on temporal variations in
groundwater inflow. The existence of more than one gauging station along a river reach, however,
would be rare.
3.7
Suitability of Methods
The five different methods for estimating groundwater inflow described in this report, operate over
different spatial and temporal scales (Table 2). While the hydrograph separation and chemical
hydrograph separation methods provide information at the whole-of-catchment scale, hydraulic
gradient analysis provides a more local estimate of the exchange rate. Only the longitudinal river
chemistry and flow difference methods provide data on the spatial pattern of exchange. Conversely,
longitudinal river chemistry and flow difference methods provide data only at a single point in
time, whereas hydrograph separation, chemical hydrograph separation and hydraulic gradient
methods provide information on temporal variations in the exchange rate.
Although it is not the focus of this project, some of the methods are also able to be used to estimate
groundwater outflow as well as groundwater inflow (Table 3). Quantifying rates of outflow can be
important if both inflow and outflow occur along a river reach, and net rates of exchange are
required.

Table 2. Spatial & Temporal Scales
Spatial Scale
Method
Point
Integrated
Temporal Scale
Spatial
Point
Temporal
Hydrograph separation


Chemical hydrograph separation



Longitudinal river chemistry
Hydraulic gradient analysis



Flow difference



Table 3. Inflow Versus Outflow
Method
Inflow
Outflow
Hydrograph separation

Chemical hydrograph separation

Longitudinal river chemistry

Hydraulic gradient analysis


Flow difference


PAGE 29
Methods for estimating groundwater discharge to streams – summary of field trials
Table 4 ranks the relative costs of the methods, both in terms of infrastructure costs and on-going
operational costs. In many river systems, the infrastructure costs have already been incurred (e.g.
setting up gauging stations), and so only the on-going costs should be considered.
Table 2 to Table 4 however, do not provide any information on the accuracy of the different
methods. Some methods are only suitable for application in relatively small, undeveloped
catchments, whereas others are more generally applicable. The five methods all have their own
advantages and limitations. However, there are few studies that have compared groundwater inflow
estimates obtained using different methods. In part this is probably due to differences between the
spatial and temporal scales of the different methods that limit the value of comparisons.

Table 4. Relative Costs
Method
Infrastructure Cost
Ongoing Cost
Hydrograph separation


Chemical hydrograph separation


Longitudinal river chemistry
o
/1
Hydraulic gradient analysis


Flow difference
o

1
The cost is small if electrical conductivity sensors are used, but expensive for other tracers that require manual sampling.
PAGE 30
Methods for estimating groundwater discharge to streams – summary of field trials
4.
Field Testing
4.1
Estimates of Surface Water – Groundwater Exchange
This project used five different methods to estimate surface water – groundwater exchange in ten
catchments in eastern Australia over a period of approximately twelve months. River reaches of
approximately 20 – 30 km in length were selected in each catchment. These reaches were chosen
so that they included at least one existing gauging station for river heights and flows. In each
catchment, three or four sites were established at which river heads and adjacent groundwater
levels could be measured so that hydraulic gradients towards the stream could be directly
determined. Surface water samples were collected at these sites at approximately two-week
intervals and analysed for electrical conductivity, chloride concentration and radon activity. In most
catchments, a data logger with EC sensor was used to collect finer resolution water quality data
from at least one of the sites.
In nine of the ten catchments a detailed longitudinal stream sampling survey took place at one point
in time, in which samples were collected at approximately 1 km intervals over the entire studied
reach of the river and from the major tributaries. These samples were analysed for electrical
conductivity, chloride concentration and radon activity. At the same time, a number of manual flow
gaugings were also performed along the length of the river, using a manual electromagnetic
flowmeter. (In Hodgson catchment sequential flow gaugings and longitudinal stream chemistry
sampling were not able to be carried out due to low river flow associated with drought conditions.)
Where possible, estimates of groundwater inflow were derived using hydrograph separation,
chemical hydrograph separation, longitudinal river chemistry, hydraulic gradient analysis and flow
difference. The first two methods provide an integrated estimate of discharge over the entire
upstream catchment (including tributaries) (in m3/day or ML/day). Longitudinal river chemistry
and flow difference methods allow estimation of variations in the groundwater inflow along the
stream (m3/m/day), and these are summed to give a volumetric inflow within the study reach
(m3/day). Strictly speaking, hydraulic gradient analysis multiplies the observed gradient by an
estimate of transmissivity to give discharge per unit length of stream (m2/day). However, rather
than estimating hydraulic conductivity independently, we equate the hydraulic gradient measured at
a particular point in time with the groundwater discharge rate at that time. The ratio of discharge
rate to hydraulic gradient is assumed to be constant, and this constant (actually equal to the product
of transmissivity and reach length) is used to estimate groundwater discharge at other times based
on the head gradient. The implicit assumption with this approach is that the measured hydraulic
gradients are representative of the exchange flux along the entire river reach, and that the
transmissivity of the aquifer does not change with time.
PAGE 31
Methods for estimating groundwater discharge to streams – summary of field trials
4.1.1
Flow Difference
This method is perhaps the most straightforward of all of the methods for characterising surface
water – groundwater interactions. Measurements of stream flow rate are carried out along the river
at a time when surface runoff is thought to be negligible (at least several days since the last rain
event) and the spatial variations in groundwater inflow or outflow are estimated from the
differences between adjacent flow gaugings. The method was applied to nine of the ten studied
catchments. In Hodgson catchment sequential flow gaugings were not able to be carried out due to
low river flow associated with drought conditions. Studied river reaches were between 12.3 and
40.4 km in length, and between 8 and 12 gauging were performed along this river length (an
average spacing of 2.6 km between gaugings). Where possible, tributaries were also gauged.
Manual flow gaugings were also carried out adjacent to several of the automatical gauging stations.
Discrepancies between manual gaugings and gauging station values ranged between zero and 46%,
but with a mean difference of 18%. This difference reflects both the uncertainty of the manual
measurement and errors in the gauging station rating curve.
The flow difference method was most useful for estimating groundwater inflow in catchments
where the inflow was a large fraction of the streamflow. The Barron River was one such system.
Figure 9 shows the flow gaugings made along the Barron River between 28 and 29 October 2008.

Figure 9. Longitudinal profiles of river flow measured on the Barron River.
Closed circles denote manual flow measurements on the river, open triangles represent gauging station data,
and open squares are manual flow measurements on tributaries.
1.20
Flow (m3/s)
1.00
0.80
0.60
0.40
0.20
0.00
0
5
10
15
20
25
30
35
Distance (km)
PAGE 32
Methods for estimating groundwater discharge to streams – summary of field trials
River flow increased from 0.045 m3/s at the start of the river reach to 0.96 m3/s at the end (33.8 km
downstream), an increase of approximately 0.91 m3/s. However, a number of tributaries flow into
the Barron River. Three tributaries were gauged and contributed 0.34 m3/s. Evaporation is
estimated to be approximately 0.03 m3/s. It is believed that the remaining 0.55 m3/s is due to
groundwater inflow. This represents a mean inflow rate of approximately 1.5 m3/day per metre
length of river.
It is relatively straightforward to estimate the error associated with this estimate of groundwater
inflow. If groundwater pumping is negligible, then errors in the net groundwater inflow rate depend
largely on the uncertainty of the flow gauging measurements. The inflow rate is given by:
I  QD  QU  QT  ELw  P
[5]
where I is the groundwater inflow rate (m3/s), QU and QD are the river flow rates at the upstream
and downstream ends of the reach being considered (m3/s), QT is the tributary inflow rate (m3/s), E
is the evaporation rate from the river (m/s), L is the length of the river reach (m), w is the mean
river width (m), and P is the rate of pumping from the river (m3/s). (The pumping term also
includes loss rates from the river to the aquifer, which may be induced by pumping.) To calculate
the uncertainty of the sum of different terms, we simply add the squares of the individual errors,
and then take the square root. Thus, the relative uncertainty of the estimated groundwater inflow
rate is equal to
( I )
I

 QD 2   QU 2   QT 2   ELw2   ( P )2
QD  QU  QT  ELw  P
[6]
Table 5shows estimated uncertainties for groundwater inflow rate derived from flow difference in
each of the catchments, based on an assumed relative uncertainty of measurement for flow gauging
of 10%, and uncertainty of river evaporation of 20%. (A relative uncertainty of 10%, or 0.1, means
that the estimated value is within 10% of the mean value two thirds of the time, and within 20% of
the estimated value 90% of the time.) In each case we have assumed that pumping from the river is
known to be zero, although this unlikely to be correct in the Cockburn catchment. (Where river
losses due to pumping are unknown, uncertainties will therefore be larger than those calculated.) In
the Ourimbah catchment, we have used a relative uncertainty of 50% for tributary inflows, since
these were visually estimated and were not gauged. In the Elliot catchment, we have used a relative
uncertainty of 50% for the evaporation term (ELw). This river contains a number of very large
pools, and so river width is highly variable and difficult to estimate. In each catchment we have
used an evaporation rate of E = 7 mm/d, although the sensitivity to this term is small.
PAGE 33

Table 5. Estimated groundwater inflow rates calculated from flow difference, and estimated uncertainties.
Negative values of I reflect a net outflow along the study reach. (Flows were gauged in Logan catchment on two occasions during this project, and data
shown here is from June 2008.)
Catchment
Barron
Belubula
Nambucca
Ourimbah
Tarcutta
Cattle Creek
Cockburn
Elliott
Logan
Reach Length
(km)
33.8
40.4
22.5
12.3
29.3
20.4
20.3
32.3
24.0
QD
(m3/s)
0.96
0.50
4.4
0.31
0.52
1.65
0.048
0.24
2.13
QU
(m3/s)
0.045
0.49
1.5
0.15
0.49
1.1
0.0025
0.014
0.32
QT
(m3/s)
0.34
0.26
2.1
0.11
0
0.6
0.045
0.10
1.44
ELw
(m3/s)
0.03
0.03
0.03
0.01
0.03
0.04
0.02
0.02
0.02
I
(m3/s)
0.55
-0.28
0.77
0.05
0.00
-0.09
-0.02
0.10
0.35
PAGE 34
(I)
(m3/s)
0.10
0.07
0.51
0.06
0.07
0.21
0.01
0.03
0.26
(I)/I
(-)
0.19
0.27
0.66
1.38
14.74
2.31
0.40
0.28
0.73
The estimate of groundwater inflow derived from flow difference is considered to be reliable only
in the Barron, Elliott and Belubula catchments, where the relative uncertainty is less than 0.3. In
these catchments, the difference between the downstream flow and the sum of the upstream flow
and tributary flow is large relative to the downstream flow. Estimates are poorest in Tarcutta,
where error in flow gauging was much greater than the difference between upstream and
downstream gaugings; and in Ourimbah and Cattle Creek, where there are a large number of
tributaries.
Of course, if pumping is occurring from the river at the time of the survey, then errors in
groundwater inflow are likely to be much larger, unless the pumping rates are very accurately
known. Also, the groundwater inflows derived from this method represent net inflows over
between the flow gaugings. If both inflow and outflow occur within a reach, then only the
difference between inflow and outflow rates will be determined.
Figure 10 depicts the variation in groundwater inflow rate along the Barron River. Inflow rates
have been estimated between each of the flow gaugings. The total groundwater inflow along the
entire reach can be estimated much more accurately than the smaller scale variations. In particular,
while the total groundwater inflow is estimated with an error of approximately 19%, the error is
greater than 20% between each flow gauging, and is greater than 100% for four of the nine
sections. Thus, the flow difference method is only able to estimate groundwater inflow rates
reliably over relatively long river reaches.

Figure 10. Groundwater inflow along the Barron River from flow difference.
The red line depicts the estimated inflow rate, and the broken lines show the uncertainty.
PAGE 35
Methods for estimating groundwater discharge to streams – summary of field trials
4.1.2
Longitudinal River Chemistry
Comparison of surface water and groundwater chemistry can also be used to determined spatial
variations in groundwater inflow to a river. There are a number of different tracers that can be used,
each with their own benefits and limitations. Ultimately, however, the accuracy of this method is
highly dependent on the accuracy with which the groundwater end-member concentration can be
estimated, and the relative difference between the groundwater end-member concentration and the
surface water concentration.
In this project, estimates of groundwater inflow were made using a mass balance model of stream
chemistry that was calibrated to flow gaugings and measurements of electrical conductivity,
chloride concentration and radon activity made along the stream. End-member concentrations of
each of the tracers were obtained by sampling bores within the catchment, and in most cases the
mean value measured on the bores was used as the groundwater inflow value for the entire stream
reach. Net inflow rates along the sampled stream reach are presented in Table 6.
Table 6. Estimated groundwater inflow and outflow rates obtained from modelling of
longitudinal stream chemistry and flow gauging data.

Catchment
Barron
Belubula
Nambucca
Ourimbah
Tarcutta
Cattle Creek
Cockburn
Elliott
Logan
Reach Length
(km)
33.8
40.4
22.5
22.2
29.3
20.2
11.7
32.3
26.4
Inflow
(m3/s)
0.71
0.069
0.78
0.15
0.033
0.33
0.19
0.16
0.022
Outflow
(m3/s)
0
0.33
0
0
0.058
0.4
0.17
0.02
0
Difference
(m3/s)
0.71
-0.26
0.78
0.15
-0.025
-0.07
0.02
0.14
0.022
Table 7 shows the variation in concentration of different tracers in surface water (c) and
groundwater (ci) within the various catchments. As can be seen, surface water and groundwater
values of electrical conductivity and chloride concentration overlap each other in most of the
catchments, whereas surface water and groundwater radon activities overlap only in the Barron and
Cockburn catchments. Because surface water and groundwater values tend to be distinct for radon,
this makes this tracer particularly useful for identifying groundwater inflow. The variability of the
groundwater value (ci) is also usually less for radon than for chloride and electrical conductivity,
and so the end-member concentration is more easily defined. The uncertainty in groundwater
PAGE 36
Methods for estimating groundwater discharge to streams – summary of field trials
inflow rate is proportional to the uncertainty in (ci – c). Since the uncertainty in c is small relative to
the uncertainty in ci, the uncertainty in (ci – c) is approximately equal to the uncertainty in ci. The
relative uncertainty in (ci – c) is then equal to:
 ( ci  c )
ci  c

 ( ci )
ci  c

 ( ci )
E( ci )  E( c )
[7]
where E is the expected (or mean) value.
Table 8 shows the estimated uncertainty in groundwater inflow from longitudinal chemistry
sampling using EC, chloride and radon, based on Equation 7. The uncertainty in groundwater
inflow concentration has been assumed to be equal to the coefficient of variation of the
groundwater concentration measured at sampled bores. (The coefficient of variation is equal to the
standard deviation divided by the mean.) In general, relative uncertainties approaching 1 might be
considered too large for the technique to be useful. Columns 2 – 4 of Table 8 represent the
uncertainties if particular tracers are used independently. Values are greater than 1 for chloride in
all catchments and greater than 0.9 for electrical conductivity in all catchments except Logan.
Radon values are greater than 0.75 in only two of the nine catchments. Of course, Table 8 only
shows the effects of uncertainty in (ci – c) and not other parameters involved in the calculation.
Although estimation of groundwater inflow using radon will also be affected by errors in gas
exchange rate, the uncertainty in this parameter is usually less than 20-30% and so the uncertainty
in (ci – c) still dominates the overall uncertainty in I.
When several tracers are used together, or when flow gauging data is used in association with water
chemistry, errors are much lower than suggested by the above analysis and this approach is always
to be recommended. In the current study, a mass balance model was calibrated to all three tracers
and to river flow data. The uncertainties of the derived estimates of inflow are difficult to
determine but are likely to be similar to or less than the minimum uncertainties of the individual
methods (Table 8). The relative uncertainty of estimated groundwater inflow rate is thus likely to
be less than 0.75 in all catchments, and less than 0.4 in five of the nine catchments.
PAGE 37
Table 7. Variation in measured concentration of different tracers within surface water (c) and groundwater (ci) within the
various catchments. CV is the coefficient of variation, which is the standard deviation divided by the mean.

Chloride (mg/L)
EC (S/cm)
c
Barron
Belubula
Nambucca
Ourimbah
Tarcutta
Cattle Creek
Cockburn
Elliott
Logan
MEAN
Range
54-113
550-623
73-98
145-234
208-222
112-162
281-558
329-629
195-436
ci
Range
102-369
620-2200
90-446
88-4890
140-1690
192-2260
295-1950
195-1530
912-2660
c
CV
54%
48%
72%
188%
81%
108%
59%
68%
39%
80%
Range
6.4-8.2
34-42
12-14
18-49
20-24
11-16
8-32
83-174
26-69
Radon (Bq/L)
ci
Range
7.2-16
29-590
11-83
13-1700
25-407
12-630
12-230
39-307
78-984
c
CV
29%
104%
103%
242%
93%
186%
129%
68%
91%
116%
PAGE 38
Range
0.5-6.4
0.1-2.8
1.1-2.7
0.2-1.2
0.1-0.4
0.5-2.2
0.6-63
0.1-1.7
0.1-0.4
ci
Range
6.2-93
13-68
3.6-203
8.2-37
11-55
23-190
4-580
2.4-32
16-43
CV
69%
31%
143%
55%
54%
71%
142%
52%
30%
72%
Methods for estimating groundwater discharge to streams – summary of field trials

Table 8. Relative uncertainty in groundwater inflow estimated using different methods.
Uncertainties in inflow estimated using individual tracers are based on Equation 7. Uncertainties in inflows
derived from flow difference are from Table 5. The model used in this project combines longitudinal chemistry
and flow gauging data, and so has an uncertainty which is equal to or less than the minimum uncertainty of
the individual methods.
Barron
Belubula
Nambucca
Ourimbah
Tarcutta
Cattle Creek
Cockburn
Elliott
Logan
4.1.3
Longitudinal Chemistry Alone
EC
Chloride
Radon
1.00
1.24
0.75
0.95
1.28
0.31
1.41
2.29
1.47
2.55
2.97
0.56
1.11
1.06
0.55
1.40
2.15
0.72
1.06
1.78
1.58
3.11
6.97
0.53
0.46
1.01
0.30
Flow
Difference
0.19
0.27
0.66
1.38
14.74
2.31
0.40
0.28
0.73
Combination
<0.19
<0.27
<0.66
<0.56
<0.55
<0.72
<0.40
<0.28
<0.30
Hydraulic Gradient Analysis
The flow rate between the groundwater and the river can also be estimated using Darcy’s Law,
which assumes that flow is proportional to the head difference between the groundwater and the
surface water. The difficulty lies in estimation of the constant of proportionality - the aquifer
transmissivity. One approach is to use a measurement of the exchange flux at a particular point in
time to calibrate the method.
In this study, we equate the hydraulic gradient measured at the time of the manual flow gauging
and river chemistry survey, with the groundwater discharge rate determined from these methods.
The ratio of discharge rate to hydraulic gradient is assumed to be constant and this constant
(actually equal to the product of transmissivity and reach length) is used to estimate groundwater
discharge at other times based on the head gradient.
For example, river flow rates were measured along the length of the Elliott River on 19 - 20 June
2008, at a time when it can be assumed that flow was entirely derived from groundwater discharge.
At this time, the increase in flow along the river that is not accounted for by tributary inflows is
assumed to be groundwater inflow and was estimated to be 0.14 m3/s (Table 6). The observed head
gradient between the groundwater and the river at this time was 0.014 and was directed towards the
river. The flow rate between the groundwater and the river can be written as
Q  2 LT
i
x
[8]
PAGE 39
Methods for estimating groundwater discharge to streams – summary of field trials
where Q is flow rate (m3/s), L is the river length (m), T is the transmissivity (m2s) and i/x is the
head gradient. (The value of two assumes that similar flows occur from each side of the river.)
Equating the observed head gradient of 0.014 measured on 20 June and the observed inflow rate of
0.14 m3/s gives 2LT = 10 m3/s. Since measurements were made over a river length of L = 32.3 km,
then this gives T = 13.2 m2/day. Groundwater inflow rates can then be estimated at other times
based on the measured head gradient, and this estimated value of transmissivity. Results from nine
catchments (data was not available for Hodgson River) are presented in Table 9.

Table 9. Groundwater inflow rates derived from hydraulic gradient analysis.
Negative values reflect a net outflow along the study reach.
Catchment
Barron
Belubula
Nambucca
Ourimbah
Tarcutta
Cattle Creek
Cockburn
Elliott
Logan
Reach Length
(km)
33.8
40.4
22.5
22.2
29.3
20.2
11.7
32.3
26.4
No. Days
(-)
396
350
215
373
390
346
400
422
379
I
(106
m3)
22.2
-6.4
7.2
7.5
18.3
18.3
5.5
3.3
4.4
An important question, though, relates to the representativeness of the head gradient data. In this
project, paired groundwater – surface water monitoring stations were set up on each river stretch to
monitor hydraulic gradients and changes in hydraulic gradients over time. Between two and four
paired sites were located along each river, although in some cases these were not operational at the
time that the groundwater inflow rate was measured. Thus, the distance of river over which the
observed head gradients were assumed to apply ranged from 4.0 to 32.3 km.
The results of these studies indicated that hydraulic gradients are highly spatially variable. In the
Barron River catchment, for example, on 29 October 2008 the hydraulic gradients were measured
to be 0.0054, 0.082 and 0.102, at the three monitoring sites (at 0, 17.3 and 33.8 km, respectively).
In the Nambucca catchment, hydraulic gradients measured on 9 July 2009 were 0.0027, 0.0021,
and -0.0023. Large differences between hydraulic gradients at different sites suggest that these are
strongly affected by local conditions and may not be a useful indicator of regional flow conditions.
The difficulty in using hydraulic gradient to extrapolate groundwater inflow rates measured at one
point in time over a longer period is further exemplified by the temporal variations in hydraulic
gradient measured at sites at 0 and 17.3 km in the Barron River catchment (Figure 11). At the
PAGE 40
Methods for estimating groundwater discharge to streams – summary of field trials
upstream site, hydraulic gradients show little variation throughout the year, while at the site further
downstream, the gradient is three times greater in February 2009 than in December 2008. Clearly
the temporal changes in hydraulic gradient will vary spatially, but the lack of representativeness of
a small number of paired sites is cause for concern.
Figure 11. Temporal variations in hydraulic gradient at two sites within the Barron River
catchment. In both cases, flow is directed towards the river.
Head Gradient
Head Gradient

4.1.4
0.2
0.18
0.16
0.14
0.12
0.1
0.08
0.06
0.04
0.02
0
20-Aug-08 19-Oct-08 18-Dec-08 16-Feb-09 17-Apr-09 16-Jun-09 15-Aug-09
0.02
0.018
0.016
0.014
0.012
0.01
0.008
0.006
0.004
0.002
0
20-Aug-08 19-Oct-08 18-Dec-08 16-Feb-09 17-Apr-09 16-Jun-09 15-Aug-09
Chemical Hydrograph Separation
Where groundwater and surface water have different chemical signatures, then changes in tracer
concentrations in river flow over time can be used to infer changes in the relative proportions of
surface runoff and groundwater inflow. Because the relative proportions of surface runoff and
groundwater inflow will change significantly during rainfall events, measurements of river
chemistry need to be made at frequent intervals. At any point in time, the proportion of river flow
PAGE 41
Methods for estimating groundwater discharge to streams – summary of field trials
that is due to groundwater discharge is estimated by comparing stream water chemistry with
assumed values of surface runoff and groundwater inflow.
Usually, the method is used to provide estimates of changes in groundwater inflow over time
through a rainfall event (hours to days), with samples collected at approximately hourly intervals or
more frequently. In the present study, fortnightly water samples were collected from each of the
river reaches studied (and analysed for radon, chloride and electrical conductivity). Electrical
conductivity data was also obtained using sensor probes. Although the fortnightly EC samples
proved valuable for verifying the data obtained from the sensor probes, the fortnightly data alone
was not considered useful for estimation of groundwater inflow because the occasional sampling
failed to capture many of the flow events. The more continuous data obtained using EC sensor
probes was considered much more amenable to this form of analysis.
Electrical conductivity sensors were located at a total of thirteen sites across nine of the
catchments. Comparison with fortnightly samples showed major discrepancies in three of the 13
sites. This is a major issue of concern, and the reliability of EC sensors in surface water gauging
stations should be assessed.
The success of the chemical hydrograph separation method ultimately relies on an adequate
chemical differentiation of the source waters and accurate quantification of the end-member
concentrations. In this project, the groundwater concentration end-member was assumed to be
equal to the mean EC measured in groundwater and the surface run-off end-member was assumed
to be equal to the minimum measured EC in river flow. Results from the nine catchments where the
method was used are presented in Table 10.
Table 10. Estimated groundwater inflow rates obtained from chemical hydrograph
separation.

Catchment
Barron
Nambucca
Ourimbah
Tarcutta
Cattle Creek
Cockburn
Elliott
Logan
Hodgson
No. Days
(-)
396
215
373
390
346
400
422
379
71
I
(106
m3)
23.6
11.6
5.9
6.4
51.1
4.2
3.1
4.8
4.5
I/Q
(-)
0.17
0.39
0.42
0.34
0.27
0.27
0.48
0.16
0.13
PAGE 42
Methods for estimating groundwater discharge to streams – summary of field trials
Figure 12 shows how the estimated groundwater discharge depends on the adopted values for
groundwater inflow and surface runoff end-members. Between 1 August 2008 and 30 September
2009, the electrical conductivity of the Barron River (measured at the Picnic Crossing gauging
station) varied between 46 and 105 S/cm. Based on this, the concentration of surface runoff has
been assumed to be 46 S/cm. The mean concentration of groundwater inflow to the river has been
assumed to be 191 S/cm, based on measured groundwater concentrations in bores close to the
river. The baseflow index (ratio of baseflow to total streamflow) for the Barron River upstream of
Picnic Crossing is thus estimated to be 17%. However, the end member concentrations are not well
known. If the electrical conductivity of surface runoff is 30 S/cm water rather than 46 S/cm, then
the baseflow index is estimated to be 26%, whereas if it is 60 S/cm, the baseflow index is
estimated to be 10%. Similarly, if the mean concentration of groundwater inflow is 250 S/cm
rather than 191 S/cm, then the baseflow index is estimated to be 12%, whereas if it is 105 S/cm,
the baseflow index is estimated to be 45%.
Of course, the use of a constant value for these end-members may also cause problems if endmembers change throughout the year. This appears to be the case in several of the catchments. For
example, if groundwater chemistry changes downstream, then the mean concentration of inflow
can change through the year as the headwaters migrate upstream and downstream in response to
changes in the watertable. Further studies should examine methods to calculate end-member
concentrations and the variability of end-members in space and time.
PAGE 43
Methods for estimating groundwater discharge to streams – summary of field trials
Figure 12. Effect of end-member values on estimated groundwater inflow derived from
chemical hydrograph separation using electrical conductivity (Barron catchment).

Values in legend refer to assumed values of groundwater and surface water end-members (in
S/cm), respectively.
100
River Flow
191/46
10
250/46
Flow (m3/s)
105/30
191/30
191/60
1
0.1
0.01
1/8/08
4.1.5
20/9/08
9/11/08
29/12/08
17/2/09
8/4/09
28/5/09
17/7/09
5/9/09
25/10/09
Hydrograph Separation
Although there are a number of concerns with its accuracy (e.g., Halford and Maher, 2000; Werner
et al., 2006), hydrograph separation based on streamflow data is one of the most widely used
methods for quantifying rates of estimating surface water – groundwater interaction. It is therefore
useful to calculate flow rates using this method to compare with those obtained from other
methods. In Australia, the Lyne and Hollick filter method (Nathan and McMahon, 1990) is most
widely used, and so was chosen for the current project. The method is automated, but involves an
operator-controlled parameter, α. The estimated baseflow will depend upon the chosen value for α,
although a value of α = 0.925 was recommended by Nathan and McMahon (1990) on the basis that
it commonly provides a reasonable match to the baseflow index (BFI) values achieved from
manual hydrograph separation methods. However, choice of the  parameter remains somewhat
subjective, and depending on the catchment, values between 0.925 and 0.99 have been found to
provide the closest match to the BFI values achieved from manual baseflow separations (Murphy et
al., 2008).
Using the parameter value of α = 0.925, the estimated baseflow index for nine of the ten
catchments ranges from 0.36 to 0.62, with the lowest estimated value measured in the Cockburn
catchment and the highest estimated value measured in the Tarcutta catchment (Table 11). The
PAGE 44
Methods for estimating groundwater discharge to streams – summary of field trials
baseflow index for Hodgson River is much lower at 0.01. A long-term drought in this catchment
meant that groundwater levels mostly below river level, and the river was not flowing for much of
the observation period. The only significant flow events occurred during high rainfall events when
surface runoff dominated the hydrograph.
Table 11. Baseflow indices (%) derived from hydrograph separation using the Lyne and
Hollick filter method (=0.925).

Catchment
Barron
Belubula
Nambucca
Ourimbah
Tarcutta
Cattle Creek
Cockburn
Elliott
Logan
Hodgson
No. Days
(-)
396
350
215
373
390
346
400
422
379
71
I
(106
I/Q
m3)
86.3
6.0
15.9
8.1
10.5
88.5
6.3
3.4
11.8
0.43
0.62
0.39
0.53
0.46
0.56
0.48
0.36
0.52
0.42
0.01
Because hydrograph separation is not a physically-based method, it is not possible to accurately
estimate the uncertainty of the method. However, the difference between the baseflow index
obtained using different α-parameters gives some indication of the uncertainty. Differences
between baseflow indices using α = 0.925 and α = 0.98 parameters, were relatively large, with
differences exceeding 30% in all catchments (Figure 13).
It is also instructive to compare the results of hydrograph separation and chemical hydrograph
separation. Figure 13 indicates that generally the hydrograph separation method overestimates
baseflow relative to the chemical hydrograph separation method. The difference is largest in the
Barron, Cattle Creek and Logan catchments. The Barron and Logan catchments have intense wet
seasons causing the combination of river hydrograph peaks. These effects dampen the hydrograph
spikes and may cause the hydrograph separation method to overestimate groundwater inflow. The
Logan catchment is partially regulated, which also creates dampened hydrograph spikes and
overestimated groundwater inflow compared with estimates from the hydrograph separation
method. However, when results are considered over the same period of time, in only one of the
eight catchments are the results of the two methods within 30% of each other and they are within
50% in only three of seven catchments (=0.925). Results are somewhat closer if the =0.98 filter
parameter is used. Of course, differences between results obtained from the different methods does
not in itself provide any information on which method is more likely to be correct. Nevertheless,
PAGE 45
Methods for estimating groundwater discharge to streams – summary of field trials
for the hydrograph separation method to be reliable, the  filter parameter must be known a priori.
The fact that it is not, and that results are sensitivity to the parameter choice, is a cause for concern
when absolute rates of inflow are required.
Figure 13. Comparison between estimates of baseflow percentage estimated using
hydrograph separation (=0.925 and =0.98) and chemical hydrograph separation.

Results shown are slightly different from those in Table 10, because the period of time is slightly
different.
4.2
Predictions of Streamflow Depletion by Pumping
The impact of groundwater pumping within each catchment has been estimated using a simple
model that calculates the rate of streamflow depletion due to pumping from individual bores, based
on their distance from the river and the aquifer transmissivity. In this project, the model of Jenkins
(1968) is used. Streamflow depletion is thus calculated using:
Q
Sl 2
 erfc(
)
Qw
4Tt
[9]
where:
PAGE 46
Methods for estimating groundwater discharge to streams – summary of field trials
T = aquifer Transmissivity (m2/day)
S = aquifer storage co-efficient (-)
l = shortest distance between the pumping bore and the stream (m)
t = pumping duration (days)
Qw = pumping rate in bore (m3/day)
ΔQ = stream depletion flow rate (m3/day)
ΔQ/Qw = rate of stream flow depletion as a proportion of the pumping rate (%)
The calculation is performed for each bores, and summed to give the total impact on the river.
A key point to note is that the model assumes that there is only one site (i.e. a single stream) that
can be intercepted by the pumping bore. In reality there may be several sources of groundwater
discharge (e.g several streams, wetlands and phreatophytic vegetation) that could be intercepted by
the drawdown cone. As a result, stream flow depletion calculated by the model should be
considered as a maximum value. Lower rates of depletion will occur if these others forms of
groundwater discharge are significant.
The model also makes a number of other assumptions. These include that the stream penetrates to
the bottom of the aquifer and that the streambed hydraulic conductivity is that same as that of the
aquifer. The effects of these assumptions have been considered by Sophocleous et al. (1995), and
are also discussed by Rassam and Werner (2008). Although the model applied is simplistic, lack of
information about aquifer thickness and streambed hydraulic conductivity mean that application of
more realistic models is difficult. It should be noted, though, that uncertainty of the estimated
stream depletion rates is relatively high.
Figure 14 presents an example of the streamflow depletion estimates for one catchment – the
Elliott River based on continuation of the current pumping (using 2008/2009 data). The annual
total streamflow depletion in 2009 is estimated to be approximately 3350 ML. Thus only about one
half of the total estimated impact of groundwater pumping within the catchment has already
impacted the river. Based on the current extraction rates, an estimated annual stream depletion of
an additional 1300 ML is expected by 2050. Eventually, the annual streamflow depletion is 7340
ML, equal to the total assumed extraction.
PAGE 47
Methods for estimating groundwater discharge to streams – summary of field trials
Figure 14. Modelled streamflow depletion resulting from groundwater pumping within
the Elliot River catchment.

Table 12 depicts estimated streamflow depletion rates for all catchments in 2009, in 2050 and in
steady state conditions - the latter is equivalent to the assumed total groundwater pumping rate.
Figure 15 presents the current streamflow depletion and the total groundwater extraction in each
catchment. It is important to note, however, that the streamflow depletion analysis was only carried
out for the studied stream reach and so pumping in areas further upstream has not been factored
into these calculations. The only exception to this is the Elliot catchment, where the headwaters of
the stream were included within the studied reach. Figure 15 and Table 12 show that the
Cockburn, Elliot and Belubula catchments have the largest volume of groundwater usage and
estimated streamflow depletion. The extent to which the extracted volume of groundwater impacts
on stream flow varies between the catchments. In the Barron and Elliot catchment most of the
groundwater extraction is thought to have impacted on the river flows, while in the Tarcutta
catchment the streamflow depletion is small relative to the extracted volume of groundwater. The
degree of streamflow impact from groundwater extraction largely depends on the distance from the
river of the extraction bores and the properties of the aquifer. The Barron and Elliot catchments
have high transmissivity aquifers in close connection with the river, while the Tarcutta catchment
groundwater extraction occurs from a semi-confined aquifer separated from the river by an
PAGE 48
Methods for estimating groundwater discharge to streams – summary of field trials
aquitard. In many cases the full impact of groundwater pumping on streamflow may be delayed
several years.
Table 12. Estimated streamflow depletion by pumping.

Catchment
Streamflow Depletion by Pumping (ML)
Current
782
2590
1716
3350
714
553
2959
390
104
Barron
Belubula
Cockburn
Elliott
Hodgson
Logan
Nambucca
Ourimbah
Tarcutta
2050
842
2619
1947
4661
879
680
3041
415
198
Steady State
892
2652
2196
7430
1056
830
3149
437
465
Figure 15. Estimated groundwater extraction volumes and current streamflow depletion
due to groundwater pumping

8000
stream
depletion
7000
6000
Groundwater
usage
4000
3000
2000
1000
Hodgson
Nambucca
Cattle
Belubula
Ourimbah
Elliott
Barron
Cockburn
Logan
0
Tarcutta
Volume (ML)
5000
PAGE 49
Methods for estimating groundwater discharge to streams – summary of field trials
4.3
Analysing Trends in Baseflow
Although flow-based hydraulic separation is the most subjective of the methods used in this report
to estimate groundwater discharge, since it only uses streamflow data, is it able to be applied over
long time periods. For each catchment, flow-based hydrograph separation is used to estimate
groundwater discharge over the period of streamflow record and statistical methods are used to
estimate trends in groundwater discharge over time. To examine the possible causes of any trends
in groundwater discharge two additional techniques are applied. SIMHYD is a simple rainfallrunoff-infiltration model which is used to predict how variations in rainfall and evapotranspiration
may have resulted in changes in groundwater discharge to the river. Stream depletion analysis
calculates the timelag between groundwater pumping and reduction in streamflow, and is hence
used to examine the likely impact of groundwater pumping on the river. Climate variability and
groundwater pumping are two of the processes which are likely to have caused long-term changes
in groundwater discharge. Other factors, such as land-use change are not specifically examined.
4.3.1
Statistical Analysis of Hydrograph Separation
Analysis of changes in baseflow over time have been made using estimates derived from
hydrograph separation. Although this method for calculation of baseflow may be inferior to some
of the other methods described above, it offers the benefit of requiring only readily available river
flow data and so is amenable to trend analysis. Trends in baseflow over time have been examined
by summing daily baseflow figures to produce annual values, and then carrying out a linear
regression on the annual data. The equation:
Baseflow = a0 + a1[Rainfall]+ a2[time]
[10]
was fit to the annual baseflow data, where a0 , a1 and a2 are coefficients determined in the regression
process. The model assumes that annual baseflow is linearly related to annual rainfall, and analyses
for linear trends in baseflow over time.
Figure 16 shows an example of this analysis for the Elliot stream gauge. The blue circles denote
annual baseflow values estimated from hydrograph separation (using baseflow parameter α =
0.925). The black line represents the expected baseflow volumes calculated from rainfall data and
the red line indicates a trend in baseflow over time. The analysis shows a downward trend in annual
baseflow, equivalent to a reduction of 8000 ML over the 50 year period. The the effects of reduced
PAGE 50
Methods for estimating groundwater discharge to streams – summary of field trials
rainfall are subtracted from the trend, the analysis suggests a reduction of 3300 ML over the same
period.
Figure 16. Elliot results of linear regression between annual baseflow, annual rainfall
and year, using baseflow parameter α = 0.925.

Baseflow parameter = 0.925
Aggregation period = Jan to Dec
Regression results
30000
25000
Observed Baseflow
Flow (ML)
20000
15000
Estimated Baseflow
10000
Estimated Trend
5000
0
1960
4.3.2
1965
1970
1975
1980
1985
1990
1995
2000
2005
SIMHYD Modelling
One of the limitations of the above trend analysis is the simplicity of the model, particularly the
assumption of linear dependence between baseflows and rainfalls. To verify these results, baseflow
trends were also assessed using a method that does not assume such linear dependence. An
appropriate method in this context is to use a lumped conceptual rainfall-runoff model. The
SIMHYD model developed by Chiew et al. (2002) estimates daily streamflow from daily rainfall
and areal potential evapotranspiration data. The SIMHYD model represents interception,
infiltration, groundwater recharge and baseflow. A schematic of the model is given in Figure 17.
The parameters of the SIMHYD model are selected to fit the observed streamflows in the
catchment. As the observed streamflows may be affected by diversions (both groundwater and
PAGE 51
Methods for estimating groundwater discharge to streams – summary of field trials
surface water) the modelled streamflows may not match the observed streamflows over time.
However, the model parameters were selected to avoid an overall bias in the estimated flows. Once
the SIMHYD model is calibrated for the catchment the baseflow contribution can be extracted from
the model which accounts for climatic variation and any residual trends with time will thus reflect
the influence of exogenous influence.

Figure 17. SIMHYD rainfall-runoff model
Once fitted, the baseflow contribution to streamflow were extracted from the model, and
aggregated to annual values. The annual SIMHYD baseflow estimates were compared with the
‘observed’ baseflow obtained from the hydrograph separation using the digital filter. Where the fit
between the observed and measured data changes over time, this might be due to factors that are
affecting baseflow which are not included in the SIMHYD analysis, such as, for example,
groundwater pumping. Figure 18 shows this comparison for the Elliott catchment. A reasonably
good match is observed at early times, but modelled baseflow overestimates observed baseflow at
later times. Figure 19 then shows the difference between these two curves. The difference between
the two curves changes by approximately 3700 ML over the 50 year period.
PAGE 52
Methods for estimating groundwater discharge to streams – summary of field trials
Figure 18. Comparison between ‘observed baseflow’ (obtained from baseflow separation
analysis of streamflows) and SIMHYD-derived modelled estimates of baseflow for Elliot
River.
30,000
Annual baseflow (ML)
25,000
20,000
15,000
10,000
5,000
0
1950
1960
SimHyd
1970
1980
Lynne Hollick 0.925 filter
1990
2000
2010
2020
Year
Figure 19. Time series of difference between observed annual baseflows and modelled
data for Elliot River.

Annual difference in baseflow (ML)
6,000
4,000
2,000
0
-2,000
-4,000
-6,000
1950
1960
1970
1980
SIMHYD - Lynne-Hollick 0.925 filter
1990
2000
2010
2020
Year
PAGE 53

Table 13. Apparent changes in baseflow over time.
Catchment
Barron
Belubula
Cattle Creek
Cockburn
Elliott
Logan
Ourimbah
Tarcutta
Time Period
1972-2008
1958-2009
1967-2009
1978-2009
1958-2008
1965-2008
1980-2008
1980-2008
Trend Analysis (=0.925)
Total Baseflow
Baseflow Not Due
(ML)
To Rainfall (ML)
-27000
-17000
-16,000
-20,000
-8000
-63000
-4600
-95000
-14000
2000
9000
-19000
-3300
-28000
-450
-60000
SIMHYD Analysis
Time Period
 (Modelled-Observed)
(ML)
1972-2008
-8500
1967-2009
1978-2009
1958-2008
1965-2008
1980-2008
1980-2008
16000
-18000
-3700
-23000
-1200
-44000
PAGE 54
4.3.3
Summary of Catchment Results
Table 13 presents the results of baseflow trends analysis for all ten catchments. Estimated trends in
annual baseflow, after the effects of rainfall variation have been removed, range from +9000 ML to
-60,000 ML over the period of record. Estimates obtained from this statistical analysis are similar
to those obtained from comparison of the SIMHYD data and the hydrograph separation.
Figure 20 presents an overview of estimated long term trends in baseflow for all the catchments
analysed, including the estimates of total change in baseflow and baseflow independent of climate
change using the two different methods. The Hodgsons Creek and Nambucca River catchments
have been omitted from the time trend analysis due to lack of data in these catchments.
Figure 20 shows that estimated baseflow declines in most catchments (red bars). The two green
bars indicate the estimated change in baseflow that cannot be attributed to changes in rainfall,
calculated from analysis of hydrograph separation and from SIMHYD modelling. In the Tarcutta,
Logan, Barron, and Elliot catchments, approximately one third to a half of the total estimated
baseflow depletion is due to exogenous factors, with the remaining half to two thirds due to
changing climate. In the Cockburn catchment most of the estimated baseflow depletion appears to
be due to exogenous factors. In contrast, in the Ourimbah, Belubula, and Cattle Creek catchments
the reductions in estimated baseflow appear almost entirely due to changes in climate. However,
although six of the eight catchments show reductions in baseflow estimates that appear not to have
been caused simply by variations in rainfall (and hence might be due to groundwater pumping), the
trends are significant in only three catchments: Tarcutta, Cockburn and Elliot. The other
catchments show trends, but these are within the range of annual variation.
PAGE 55
Methods for estimating groundwater discharge to streams – summary of field trials

Figure 20. Comparison of change in baseflow estimates.
The labels on the x-axis give the catchment name and the number of years over which the analysis
was conducted. The y-axis shows the change in annual baseflow. The red bars indicate the total
change in annual baseflow, based on statistical analysis of hydrograph separation results. The two
green bars indicate the change in baseflow that cannot be attributed to changes in rainfall. This has
been calculated both from analysis of hydrograph separation and from SIMHYD modelling.
Figure 21 presents baseflow reduction estimates relative to the average streamflow in each
catchment. Negative values of baseflow reduction (Belubula River and Cattle Creek) represent an
increase in estimated baseflow. The Tarcutta and Logan catchments have the highest proportion of
estimated baseflow reduction relative to average streamflow, while the Belubula, Barron and Cattle
creek catchments have the lowest estimated baseflow reductions relative to average streamflow.
PAGE 56
Methods for estimating groundwater discharge to streams – summary of field trials
Figure 21. Change in baseflow estimates relative to average streamflow.
Decrease in Baseflow Over Time Period (ML) or
Average Streamflow (ML/year)

280,000
average streamflow
baseflow depletion
Lynne-Holick 0.925
Simhyd
180,000
80,000
Cattle - 41
Belubula - 41
Ourimbah - 28
Elliott - 50
Barron - 36
Cockburn - 31
Logan - 44
Tarcutta - 29
-20,000
Comparison of Table 13 and Figure 12, however, suggests that the estimated decline in baseflow
that cannot be attributed simply to a reduction in rainfall in recent years is much greater than can be
explained by groundwater pumping. For example, in the Tarcutta catchment, the analysis suggests
a decline in annual baseflow not attributed to rainfall reduction of between 44,000 and 60,000 ML,
but groundwater pumping is only 465 ML/yr, and of this only 104 ML/yr is estimated to have
currently impacted on streamflow. The exception is in the Elliot River catchment, where the
estimated decline in annual baseflow not due to rainfall reduction is between 3300 and 3700 ML,
and the streamflow depletion from pumping is estimated to be 3100 ML. Thus, in this catchment
the observed decline in baseflow can be explained by groundwater pumping.
PAGE 57
Methods for estimating groundwater discharge to streams – summary of field trials
5.
Conclusions - Broader Implications
Accurate, regional scale estimates of groundwater discharge to streams are very difficult to obtain.
This report describes the findings from the field assessments of surface water – groundwater
interaction which took place in 10 selected catchments in 2008 – 2009. Five different methods for
estimating groundwater discharge were trialled. All of these methods were considered suitable for
providing regional scale estimates of discharge, but the methods had not previously been
systematically compared and evaluated.
5.1
Spatial and Temporal Patterns of Groundwater Discharge
In this project, estimates of spatial patterns of groundwater inflow were obtained from
measurements of longitudinal river chemistry and flow gauging. Water chemistry was measured at
1 – 2 km intervals, and river flow was measured at 2 – 3 km intervals. These methods therefore
provide estimates of spatial variations in groundwater inflow with a resolution of 1 – 3 km. The
relatively coarse resolution of these methods avoids many of the issues associated with techniques
that measure at finer spatial scales (e.g. seepage meters).
Previous studies have shown extremely high spatial variability in streambed hydraulic properties
and groundwater inflow rates at scales of metres to tens of metres (Harvey and Bencala, 1993;
Calver, 2001) which limit the usefulness of small scale measurements for regional studies. Even at
the scale of a few kilometres, this study found groundwater inflow to be highly variable. For
example, seven of the ten catchments appeared to contain both losing and gaining reaches at this
scale, even at relatively low river flows. Maximum estimated rates of gain ranged over two orders
of magnitude – from 0.15 m3/day/m in the Logan River to 17 m3/day/m in the Cockburn River.
Mean estimated rates of gain were mostly between 0.1 and 2 m3/day/m. The maximum estimated
loss rate in any of the catchments was 2 m3/day/m in the Belubula River. Variations in inflow and
outflow rates at scales of kilometres may be due to changes in the geometry of the river relative to
prevailing groundwater flow direction, variations in geology (and hence in hydraulic properties of
the aquifer system), and changes in the potentiometric surface (particularly in areas affected by
groundwater pumping).
Temporal variations in groundwater inflow rate also occur and have important implications for
measurements made at a point in time. Over short timescales large temporal variations in
groundwater inflow rates occur due to changes in river height. In many cases, the flow direction of
predominantly gaining streams will reverse and river water will flow into the aquifer when the river
PAGE 58
Methods for estimating groundwater discharge to streams – summary of field trials
level rises following a rainfall event in the catchment. At one of the Nambucca River sites, for
example, the gradient towards the stream averaged approximately 0.002 m/m, but a gradient
of 0.05 m/m away from the stream was observed during a large flow event in February 2009. Over
longer time periods, variations in groundwater inflow (or outflow) would be expected to occur in
response to seasonal variations in the water table, as well as seasonal variations in river height.
Where groundwater pumping occurs close to the river, the timing of pumping can also impact on
the temporal variability of groundwater discharge to the river. If groundwater extraction bores are
further from the river, groundwater pumping will still affect river flows, but temporal variations in
the pumping rate should be damped and so should not result in large temporal variations in
discharge to the river.
5.2
Appropriateness of Methods
The five methods used in this report operate over different spatial and temporal scales. The choice
of method for any particular study will therefore depend upon whether information on spatial and
temporal variations in groundwater discharge is of interest or whether only mean annual values
over longer reaches are required.
Hydrograph separation offers a simple method which usually can be implemented with existing
data obtained from a surface water gauging station. It provides estimates of groundwater discharge
over time for the entire catchment upstream of the gauging station. However, the method may be
highly inaccurate and based on a number of questionable assumptions. Often there is a large
element of subjectivity, such as in the choice of the  parameter for the Lyne and Hollick filter.
Furthermore, the method will never produce a negative discharge rate (flow of water to the aquifer)
even though this occurs during many flow events. The method is better suited to examining
changes in time or for broad scales comparisons of different river catchments. The baseflow values
produced by hydrograph separation analyses are likely to be correlated with groundwater inflow
rates but should not be considered to be an accurate measure of groundwater inflow.
In this study, the hydrograph separation method generally provided baseflow estimates that were
closer to those of chemical hydrograph separation method when the 0.98 filter parameter was used,
rather than 0.925. The method worked poorly in catchments where there was either regulation of
river flows or intense wet season rainfalls (e.g. more than 70% of annual rainfall within 3 months).
Chemical hydrograph separation, which used water chemistry at the gauging station, as well as
flow), provides an alternative means for estimating changes in groundwater discharge over time for
the upstream catchment. Electrical conductivity data is available at many gauging stations.
Although the data is under-utilised, this project has revealed potential problems with a number of
PAGE 59
Methods for estimating groundwater discharge to streams – summary of field trials
the gauging stations, with significant discrepancies between automated and manual measurements.
However, if the data obtained from these stations is first verified, then this method can provide an
additional way to improve the accuracy of the determination. The accuracy of this method,
however, is ultimately related to the accuracy with which end-member concentrations can be
determined. The method will be most accurate when groundwater electrical conductivity is high (so
that the difference between surface runoff and groundwater inflow is large).
Flow gauging can be used to estimate groundwater inflow rates at a point in time in catchments
where the inflow rate is high (so that errors in individual gaugings are small). However, for water
management purposes, interest is usually on regional groundwater inflow and it should be
recognised that not all groundwater inflow to a stream is regional groundwater. During river flow
events, river water can enter the aquifer, and this can discharge back to the river after the river level
has subsided. Flow gaugings therefore need to be carried out when these flows are unlikely to be
significant. Longitudinal river chemistry measurements can improve the accuracy of flow gauging,
and can also assist differentiation between bank discharge and regional groundwater, provided that
the groundwater end-member concentrations are well known. Electrical conductivity is relatively
cheap to measure, and provides a large amount of additional information. Other tracers (such as
radon, for example) are more expensive, but can further improve the accuracy of the results.
Hydraulic gradient measurements are perhaps the most direct measurement of groundwater
discharge rate and one with which hydrogeologists are likely to be most comfortable. However, this
project has suggested that the information obtained from this data may be highly site specific, and
not representative of the larger river system. The instrumentation, however, is valuable, and will
assist calibration of groundwater models, and clearly reveal changes in trends over time. However,
we caution against reliance on groundwater inflow rates calculated from gradients observed at a
single location.
None of the available methods alone measure both spatial and temporal variation in groundwater
inflow. If information on spatial variations in groundwater inflow rate is needed, then flow gauging
and longitudinal river chemistry methods should be considered. However, when information is
required on mean annual rates of groundwater discharge, methods which provide estimates at a
single point in time are of little value. Extrapolating these estimates in time using hydraulic
gradient measurements from a small number of bores is questionable because of concerns on the
representativeness of the bore data. Chemical hydrograph separation is probably the most suitable
method for estimating mean annual groundwater discharge rates, although uncertainty in endmember concentrations can produce significant uncertainties in estimated groundwater inflow
rates. Further studies are required to determine the best method for estimating groundwater inflow
concentrations. Although groundwater inflow concentrations can be obtained from measurement of
groundwater chemistry on bores within the catchment, the mean concentration in the sampled bores
may not be the mean concentration of groundwater discharging to the river. Concentrations of
PAGE 60
Methods for estimating groundwater discharge to streams – summary of field trials
groundwater inflow can also be estimated from modelling of longitudinal chemistry and flow
gauging measurements. Modelling carried out in a number of catchments within this project
suggested that the groundwater inflow concentration was significantly different to the mean
groundwater concentration. Also, changes in groundwater inflow concentration over time are
possible and there was some evidence that this may be the case in several of the catchments studied
in this project. Further work is needed in this area. Studies to measure the concentration of surface
runoff into the river would also be useful.
5.3
Transferability of Results
The catchments studied in this project represent a range of climatic and geological environments.
Mean flow rates within the studied reaches of these rivers ranged from 1.5  107 m3/yr to
1.5  108 m3/yr and upstream catchment areas ranged from 150 to 1700 km2. The project included
both regulated and unregulated rivers. Although the studied rivers might be considered to be
representative of many small to medium Australian rivers the project results may not be
immediately transferable to much larger river systems. Although several of the studied rivers were
regulated, many of the larger rivers in southern Australia are much more heavily regulated than
those studied. As river regulation increases (and variations in flow are reduced), hydrograph
separation methods become more questionable. In very large catchments, chemical hydrograph
separation methods become more difficult, as spatial variations in end-member concentrations
become more pronounced, and so the mean concentration of groundwater inflow at any point in
time becomes more difficult to determine. Where rivers have wide floodplains discharge to the
river floodplain may occur which either does not ultimately enter the river or which enters the river
as surface flow (e.g., Lamontagne et al., 2005; Langhoff et al. 2006). Although this is unlikely to be
the case for any of the streams studied in this project, as the floodplains were usually relatively
narrow or non-existent, it is likely to pose additional difficulties in large rivers with deeply incised
floodplains.
It should also be emphasised that three of the methods trialled in this project enable estimation of
groundwater inflow, but cannot be used for estimating groundwater outflow. Although seven of the
ten studied catchments contained some losing reaches, the study has focussed on estimation of rates
of groundwater inflow. Methods which can only be used for estimating rates of groundwater
outflow have not been specifically trialled in this project.
PAGE 61
Methods for estimating groundwater discharge to streams – summary of field trials
5.4
Effects of Groundwater Extraction
Prediction of changes in groundwater inflow over time due to groundwater extraction can be made
using numerical groundwater models or using simple analytical equations. It was not within the
scope of this project to compare these different methods. However, estimates of stream depletion
due to pumping were made in each catchment using one of the simple analytical methods and these
results were compared to historical changes in baseflow suggested by hydrograph separation of
historical streamflow data.
Trend analysis of these baseflow estimates has suggested declines in annual estimated baseflow of
between 4600 and 95000 ML over the last 28 to 50 years. Once the effects of reduced rainfall are
taken into account, then the reduction in baseflow beyond that attributed to reduced rainfall is
estimated to be between -9000 and 60000 ML. In two catchments, the observed decline in baseflow
is less than predicted to occur due to rainfall reduction, suggesting an increase in annual baseflow
of up to 9000 ML, although this was not statistically significant. Results are statistically significant
in three of the catchments, and in two of these (Tarcutta and Cockburn), the reduction cannot be
explained by groundwater pumping.
Thus, the observed estimated reductions in baseflow are much greater than can be explained by
either reduction in rainfall or groundwater pumping. Possible reasons for this might include
reduction in irrigation or improvement in irrigation efficiency (so that less irrigation recharge
occurs), revegetation or development of plantations, however the magnitude of the effect is much
greater than would appear to be able to be explained by these processes. It is also possible that this
is an artefact of the hydrograph separation method or that other changes within the catchment (such
as changes to river regulation, or changes in land use that have affected velocities of surface runoff)
have affected the hydrograph separation results. Further work is required to understand if the
apparent reduction in baseflow is real, and if so, what has caused it.
5.5

Key Findings
Estimated groundwater inflow was found to be highly variable along the river reaches. For
example, seven of the ten catchments contained both losing and gaining reaches at a scale of a
few kilometres.
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Methods for estimating groundwater discharge to streams – summary of field trials

Estimated groundwater inflow rates measured by river flow gauging were reliable in only three
of the nine catchments where this method was used.

The longitudinal river chemistry method should be used in conjunction with river flow data.
When this data is used together, the estimates of inflow are more accurate than when either
method is used separately. The combined method provided reasonable estimates of
groundwater inflow in all studied catchments, although the estimates relate only to the point in
time at which the data was obtained.

The hydraulic gradient method may be highly site specific, and not representative of a larger
river reach

The hydrograph separation method is easy to use and requires only available river flow data.
However, since the method is not physically based, the errors in the calculated inflow rates
cannot be directly determined. It is recommended that it be used for comparative analyses,
rather than to obtain quantitative estimates of groundwater inflow. In the current project, the
Lynne Hollick filter parameter of 0.98 provides a better match than 0.925 to the results of the
chemical hydrograph separation method.

Neither the chemical hydrograph separation method nor the hydrograph separation method
account for losing river reaches and therefore may overestimate groundwater inflow. A
quantitative model that incorporates sequential flow gaugings and longitudinal river chemistry
can account for river losses in the groundwater inflow estimates.

The groundwater inflow estimates produced by the chemical hydrograph separation method
are very sensitive to groundwater and surface water runoff end members applied. Temporal
and spatial variations in groundwater chemistry can make groundwater end members difficult
to define. Although further work is needed, the method offers perhaps the best potential for
providing catchment scale estimates of groundwater inflow over time.
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Methods for estimating groundwater discharge to streams – summary of field trials
6.
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