Seasonal Dynamics in Coastal Aquifers:
Investigation of Submarine Groundwater Discharge
through Field Measurements and Numerical Models
by
Holly Anne Michael
B.S., Civil Engineering, University of Notre Dame (1998)
Submitted to the Department of Civil and Environmental Engineering in
partial fulfillment of the requirements for the degree of
_
...
MASSACHUSE'ITSINSITWtE
Doctor of Philosophy in the field of Hydrology
OF TECHNOLOGY
FEB 2
at the
MASSACHUSETTS INSTITUTE OF TECHNOLOG
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© 2005 Massachusetts Institute of Technology. All rights reserved
Author ............................................................
.....
.-,.
.
Department of Civil and Environ ental Engineering
October 29, 2004
Certified
by......
..
Charles F. Harvey
Associate Professor, (jivil and Environmental Engineering
Thesis Supervisor
Accepted
by...... ...............................
..
w J. Whittle
Professor of Civil and Environmental Engineering
Chairman, Department Committee on Graduate Studies
Seasonal Dynamics in Coastal Aquifers:
Investigation of Submarine Groundwater Discharge through Field
Measurements and Numerical Models
By
Holly Anne Michael
Submitted to the Department of Civil and Environmental Engineering
on October 29, 2004 in partial fulfillment of the requirements for the Degree of
Doctor of Philosophy in the field of Hydrology
Abstract
The fresh and saline groundwater flowing from coastal aquifers into the ocean
comprise submarine groundwater discharge (SGD). This outflow is an important pathway
for the transport of nutrients and contaminants, and has been shown to adversely affect
coastal ecosystems in many areas of the world. The focus of this work is the
characterization of SGD and the mechanisms that drive it, with a specific emphasis on
seasonal forcing.
Field measurements during five summers in Waquoit Bay, Massachusetts reveal the
pattern and composition of submarine groundwater discharge. Flow is highly variable over
small spatial and temporal scales, and the salinity and radium content of the discharge
demonstrates heterogeneity in groundwater origin. Maximum discharge occurred in two
alongshore bands: brackish outflow nearshore and saline discharge offshore. Most of the
total flow was saline, yet net seawater inflow over a tidal cycle was negligible.
Circulation mechanisms such as tides, waves, and hydrodynamic dispersion cause
significant saline groundwater discharge, and are potentially important for chemical loading
to estuaries. However, these mechanisms can explain only 12-30% of the observed saline
outflow in Waquoit Bay.
A seasonal forcing mechanism is proposed to explain the source of the remaining
observed saline outflow. During periods of high inland recharge, the water table rises,
forcing seaward movement of the freshwater-saltwater interface and outflow of saline
groundwater; the opposite is true during times of low recharge. A series of idealized
simulated systems demonstrates this process for a range of realistic aquifer parameters, and
a time lag between maximum recharge and simulated peak discharge may explain the
observed net discharge during times of low recharge.
Winter hydraulic gradient measurements in Waquoit Bay reveal inflow in the zone
of peak summer saline discharge, confirming seasonal variation in SGD. Investigation of
the subsurface salinity profile and local hydrogeology forms the basis for a hypothesized
groundwater flow pattern that explains the observed discharge. A numerical model of the
system supports the profile and exhibits temporally-lagged inflow and outflow of saltwater
at the sea floor in response to seasonal recharge that may explain the net saline outflow
observed in Waquoit Bay during the summer.
Thesis Supervisor: Charles F. Harvey
Title: Associate Professor of Civil and Environmental Engineering
Acknowledgements
It's been a long road - windy, sometimes rocky, and mostly uphill. In the end,
I've learned that a long hike is easy, and even enjoyable, if the people around you carry
you to where you want to go. This page is a small thank-you to everyone who has
supported me along the way. I will inevitably leave someone out, but you know who you
are, and know that I thank you too.
First, to my committee members, Harry Hemond, Eric Adams, and Ann Mulligan,
I appreciate that you have taken the time to meet with me over the past several years to
discuss this research; your input has been invaluable. Your insights and suggestions have
shaped many parts of this study and have taught me both how to approach scientific
questions and how to critically evaluate the answers. Ann, thank you for the countless
conversations via email and otherwise, your willingness to help, and your positive
feedback; without them the modeling may have taken another year to complete.
I would like to thank my advisor, Charlie Harvey, for six years of patience and
guidance. I admire your intelligence and unending stream of ideas. I have doubted your
optimism as many times as it has proven me wrong, and without it this project would
have ended where it started. Thank you for recognizing what is important in life, for
always having a smile or a joke, and for passing your positive outlook along to your
students, this has been a greater motivator than any deadline.
To group members, past and present: Winston Yu, Brendan Zinn, Kaeo Duarte,
Peter Oates, Ashfaque Khandaker, and Becca Neumann, thanks for being there to help
with research, proofread papers, and most importantly to talk, to laugh, and to celebrate.
The Waquoit Bay National Estuarine Research Reserve is an exceptional facility
that has made the field work in this project possible. I would like to thank the entire
WBNERR staff, especially Chris Weidman and Christine Gault for allowing us to work
there for the past six years. Thanks also to Dr. John Germaine at MIT for the use of his
lab and equipment, and for taking the time to teach me to use them.
To all of the undergraduates who have agreed to work with me: Bridget Brett,
Amber Jaycocks, Connie Yang, and especially Jonathan Lubetsky, thank you. Without
you, much of the field work would have been impossible and certainly less enjoyable.
The Parsons Lab is an amazing place, not only for its science, but also for its
sense of community. I appreciate the support given by everyone in this lab, and I will
mention a few people who have helped me stay sane: Vanja, Fred, Jean, Hanan, Janelle,
Ramahi, Anke, Daniel, Susan, Matt and Emily, thanks. Thanks also to Sheila A., Sheila
F., and Jim, who make everything run smoothly and have always been willing to help.
I would like to thank Cheryl Silva and the MIT Women's Lacrosse teams from
1999 to 2003 for giving me the opportunity to share in their athletic experiences. Thanks
also to the Muddy crew, especially Mike, Monica, Ted and Tom, for your friendship.
To my friends in Boston and far away who I can count on for anything: Jess,
Katie, Erin, Megan, Sara, Danielle, Mario, Alex, Mary, Becky, Tony, Cathy, Sue, and
Big Dave - you're the best.
Finally, I would like to thank my family for their love and support: Mom, Dad,
John, and Heather, and Daryle. Mom and Dad, you have taught me to shoot for the top
and that hard work is the only way to get there. In so many ways I would not be here
without all of you. Thanks for everything.
5
6
Field Assistance
There have been an enormous number of people who have given their time and energy to
help with one or more field expeditions. Clearly this work would have been impossible
without them. I apologize in advance for any omissions.
Kortney Adams
Dror Angel
Roger Beckie
David Bernstein
Bridget Brett
Frederic Chagnon
Jessica Cochrane (USGS)
Kaeo Duarte
Susan Dunne
Freddi-Jo Eisenberg
Rebecca Evans
Giocomo Falorno
David Giehtbrock
Carolyn Gramling
Anke Hildebrandt
Stephanie Hsu
Amber Jaycocks
Ashfaque Khandader
Blake Landry
Jonathan Lubetsky
Bill Lyons
Timothy McCobb (USGS)
Ann Mulligan
Rebecca Neumann
Peter Oates
Daryle Peterson
Theresa Power
Catharine Rockwell
Emily Slaby
Chris Swartz
Connie Yang
Winston Yu
Brendan Zinn
Others:
Tom Chamberlin and GZA Drilling, Inc.
Denis LeBlanc and USGS
Chris Weidman and the WBNERR Staff
Summer 1999
Summer 2003
Summer 2000
Winter 2004
Summer 2003
Summer 2001
Summer 2003
Summer 1999 & 2000
Summer 2000
Summer 1999
Winter 2004
Winter 2004
Summer 2000
Summer 1999, 2000, & 20011
Summer 2001
Summer 2001
Summer 2002
Winter 2004
Winter 2004
Summer 1999 & 2000
Summer 2000
Summer 2003
Summer 2003
Winter 2004
Summer 2001, Winter 2004
November 2002
Summer 2000
Summer 1999
Summer 2001, 2002 & 2003i, Winter 2004
Summer 2000
Summer 2001
Summer 1999
Summer 1999, 2000, 2001, 2002, & 2003
Drilling donation
Field equipment
Research assistance and facilities
7
8
Contents
1. Introduction
21
1.1 The Importance of Groundwater at the Coast
21
1.2 Purpose and Scope of this Work
23
1.3 Significance and Applications
25
References
26
2. Background
27
2.1 Groundwater at the Coast: Underlying Theory
27
2.2 Modeling Submarine Groundwater Discharge
28
2.3 Field Studies
30
References
32
3. Field Investigation in Waquoit Bay
37
3.1 Motivation and Objectives
37
3.2 Study Site Description
38
3.3 Seepage Meters
40
3.4 Discharge Patterns
41
3.4.1 Head of the Bay
41
3.4.1.1 Seepage Meter Grids
41
3.4.1.2 Single Seepage Meter Transects
42
3.4.2 Slug Tests
44
3.4.3 Minimal Freshwater Flow: Island Study
45
3.5 Heterogeneity in Space and Time
3.5.1 Spatial Variability
46
46
3.5.1.1 Head of the Bay Experiments, 50 m Scale
46
3.5.1.2 Cluster Experiments, 1 m Scale
47
9
3.5.1.3 Variability in Discharge Salinity, 5 cm Scale
47
3.5.2 Temporal Variability
49
3.6 Radium Isotope Measurements
51
3.6.1 The Use of Radium as a Tracer
51
3.6.2 Radium Measurements
52
3.6.3 Heterogeneity in Porewater Radium Activity
53
3.7 Discussion
56
3.7.1 How Many Meters are Necessary to Estimate Large56
Scale Discharge?
3.7.2 Discharge Comparison with Freshwater Balance
57
3.7.3 Large-Scale Pattern of Discharge
58
3.8 Summary
60
References
61
4. Circulation of Saline Groundwater
67
4.1 Circulation Mechanisms
67
4.1.1 Tides
68
4.1.2 Waves
71
4.1.3 Dispersion
73
4.2 Saline Circulation in Waquoit Bay
74
4.2.1 Quantification of Saline Discharge Estimates due to
74
Tides and Waves
4.2.2 Mapping Nearshore Saline Circulation Using
Sodium Bromide
75
4.2.3 Discharge Patterns of Saline Circulation
77
4.3 Summary
81
References
82
85
5. Seasonality
85
5.1 Conceptual Model
10
5.2 Idealized Numerical Models
89
5.2.1 FEFLOW
89
5.2.2 Governing Equations
90
5.2.3 Model Properties and Boundary Conditions
91
5.2.4 Simulation Results
94
5.2.4.1 Parameter Effects on SGD
96
5.2.4.2 Parameter Effects on Time Lag
100
5.3 Potential for Seasonality in Actual Aquifers
105
References
106
6. Seasonality at Waquoit Bay
111
6.1 Evidence of Hydrologic Seasonality in the Waquoit Bay
Watershed
111
6.2 Under the Ice: Winter Field Study
115
6.2.1 Methods
115
6.2.2 Results
117
6.2.3 Summary of Saline Circulation in the Unconfined Aquifer
121
6.3 Regional and Local Hydrogeology of Waquoit Bay
123
6.3.1 Regional Geologic Overview
124
6.3.2 Hydrogeology within Waquoit Bay
125
6.3.2.1 Hydraulic Conductivity Estimates
127
6.3.2.2 Hydraulic Head and Salinity Measurements
in Wells
129
6.3.2.3 Geophysical Investigation
133
6.3.2.4 Groundwater Flow Patterns
137
6.4 Numerical Model of Waquoit Bay Cross-Section
138
6.4.1 Model Geometry, Parameters, and Boundary Conditions
138
6.4.2 Results
142
6.5 Summary
146
References
147
11
7. Conclusions, Implications, and Future Directions
151
7.1 Summary and Conclusions
151
7.2 Implications
153
7.3 Future Directions
155
References
160
Appendix A: Field Instrument Construction and Calibration
161
A.1 Submerged Seepage Meter Construction
161
A.2 Intertidal Seepage Meter Construction
163
A.3 Salinity Grid, Porewater Samplers, and Refractometer
Calibration
165
Appendix B: FEFLOW Model Descriptions
169
B. 1 Model Attributes and Parameters
169
B.2 Analysis of Model Output
171
Appendix C: Well Logs
173
Appendix D: Seepage Meter Data
177
D.1 Seepage Meter Flux [m/d]: Head of the Bay Experiments:
177
August 1999 and July 2000
D.2 Seepage Meter Flux [m/d]: Single Transect Experiments:
2002 and 2003
180
D.3 Seepage Meter Flux [m/d]: Washburn Island, 2000
182
D.4 Seepage Meter Flux [m/d]: Multiple Tidal Cycle Experiment,
2001
183
D.5 Seepage Meter Flux [m/d]: Cluster Experiments: 1999
185
References
186
12
List of Figuresand Tables
Chapter Two
Figure 2.1. Schematic of hypothetical coastal groundwater salinity
distribution and flow patterns.
28
Chapter Three
Figure 3.1. Map of Waquoit Bay: experimental design for seepage meter
studies. Five experiments at the head of the bay and one off of
Washburn Island were performed over five summers. Transects E and W
(East and West) are sites of several sets of field measurements. Well 1
refers to a well in the intertidal zone along Transect W used as a
reference point for distance into the bay, and the CCC Well is a deep
multi-level sampling well installed by the Cape Cod Commission.
39
Figure 3.2. (a) and (b) Time-averaged groundwater discharge vs. distance
from approximate mean shoreline for the head of the bay experiments.
Discharge is separated into freshwater, saltwater, and unknown
components. (c) and (d) Corresponding standard deviation in space and
time.
41
Figure 3.3. Temporally and spatially averaged discharge (a) and salinity
(b) vs. distance from shore for the 2002 and 2003 single-transect
experiments. Measurements taken with submerged and intertidal seepage
meters are shown separately.
44
Figure 3.4. (a) Grayscale and contours are time-averaged groundwater
discharge [m/d] for the 2000 head of the bay experiment. (b) Relative
permeability as a function of distance into the bay.
45
Figure 3.5. Groundwater discharge vs. distance from approximate mean
shoreline for the 2000 experiments at the head of the bay and Washburn
Island.
46
Figure 3.6. Daily average groundwater discharge vs. time for selected
seepage meters in the 1999 cluster experiment. Meter 11 is 12 m from
shore. Bars represent ± one standard deviation of the measurements taken
on each day.
48
13
Figure 3.7. Near-surface porewater salinity measurements taken at 5-cm
intervals on a 0.5 m grid in the nearshore and middle discharge zones.
49
Figure 3.8. Spatially-averaged discharge and tidal elevation vs. time for
the 2001 multiple tidal-cycle experiment.
50
Figure 3.9. Excess radium activity vs. distance from shore for all four
radium isotopes. The left axis scale is for 226 Ra, 228Ra, and 223 Ra;the right
axis scale is for
2 24 Ra.
54
Figure 3.10. 226 Ra activity vs. salinity for water samples collected near
Waquoit Bay. Deep porewater samples from CCC-wells (see Figure 3.1
for location), shallow porewater samples near the shore from piezometers,
discharge from seepage meters, and baywater samples are included.
55
Figure 3.11. Spatial distribution with distance from shore, averaged for
each area (row 1, area 1: nearshore, area 2: middle, and area 3: far from
shore) for (a) 226 Ra activity of water collected in seepage meters, (b)
groundwater flux measured in seepage meters, and (c) calculated 226Raflux.
55
Figure 3.12. Short-lived to long-lived (226 Ra) activity ratio vs. distance
from shore. The left axis scale is for
axis scale is
22 8Ra/226 Ra
for 224 Ra/226Ra.Average baywater
and
223 Ra/2 26 Ra;
right
activity ratios are dashed
lines.
56
Figure 3.13. Average percent error in total discharge for subsets of seepage
meters from the 2000 head of the bay experiment. Error bars represent one
standard deviation from the mean.
57
Chapter Four
Figure 4.1. Saline flow patterns induced by circulation mechanisms.
1 - One-dimensional inflow and outflow due to mechanisms such as density
fingering and tidal pumping. 2 - Nearshore circulation due to tides and
waves. 3 - Dispersion-induced
saline circulation.
69
Figure 4.2. Schematic of nearshore saline circulation due to tides and
waves and selected parameters from Li et. al. (1999) equations.
73
Table 4.1. Parameters and calculated values of groundwater circulation
due to tides and waves in Waquoit Bay, MA.
Figure 4.3. Interpretation of NaBr tracer test data. Contours of natural
salinity are shown as grayscale, contours of injected bromide are shown
14
74
as solid lines. Salinity is approximated by electrical conductivity
measurements in mS/cm and bromide concentration is in moles/L. (a)
Experimental set-up and salinity profile. (b)-(h) Approximate subsurface
bromide molarity contours for selected sample times. Dashed contours are
inferred, dashed piezometers indicate screen location and length.
78
Figure 4.4. Data from the 2003 single-transect seepage meter study, as
presented in Section 3.4.1.2. (a) Total discharge vs. distance from the
shoreline. Likely locations of inflow and outflow due to nearshore and
dispersive circulation mechanisms are depicted beneath the x-axis. (b)
Seepage salinity vs. distance from the shoreline.
80
Chapter Five
Figure 5.1. Schematic of interface position in relation to aquifer head level
according to the Ghyben-Herzberg relation (not to scale). Freshwater
discharge at the coast and seasonal saline inflow and outflow at the seafloor
are depicted with arrows.
86
Figure 5.2. Model schematic: flow and transport boundary conditions,
initial concentration profile, and dimensions.
92
Table 5.1. Idealized model simulation parameters.
93
Figure 5.3. Total monthly fresh discharge, saline discharge, and saline
inflow over the sea floor throughout a simulated year. (a)-(f) are models
1, 2, 4, 3, 5, and 6, respectively.
96
Figure 5.4. The effect of model hydraulic conductivity, dispersivity, and
thickness on saline discharge. Total saline circulation and peak saline
discharge as a percentage of peak fresh discharge are plotted against
parameter values.
99
Table 5.2. Total saline circulation as a result of both dispersive entrainment
of saltwater and seasonal interface movement and corresponding percentage
of total fresh discharge over one 365-day cycle for each idealized model.
Peak fresh and saline discharge from monthly estimates and corresponding
percentages reflect the magnitude of the seasonal effect.
100
Table 5.3. Time lag in days between maximums and minimums of system
elements: recharge (R), head (h), freshwater velocity at the origin ((0,0) V),
and saltwater velocity 20 m seaward of the shoreline ((20,0) V).
102
Figure 5.5. The effect of model hydraulic conductivity, dispersivity, and
thickness on time lag. The number of days between peak recharge and peak
15
aquifer head 50 m landward of the shoreline, freshwater velocity at the
shoreline, and saline velocity 20 m offshore are plotted against parameter
values.
103
Figure 5.6. Normalized variation in recharge, aquifer hydraulic head,
interface position, and fresh and saline velocity over one simulation year
for each of the six model runs. Hydraulic head is reported for a point 50 m
landward of the shoreline at sea level. Concentration, or salinity, at a point
20 m landward of the shoreline within the freshwater-saltwater interface
indicates interface movement: highest concentration coincides with the
extent of landward interface motion, and lowest concentration coincides with
the seaward extent. Freshwater velocity at the shoreline and saline velocity on
the seafloor 20 m from the coast indicate discharge variation throughout the
year. Actual values were normalized by dividing its difference from the
minimum by the difference between maximum and minimum values. Seasons
are approximate for a typical yearly recharge cycle within the United States.
Model characteristics are given below each model number: thick (100 m) or
thin (20 m); high K (5x10-4m/s), medium K (1x10 4 m/s), or low K (lx10-5 m/s);
and high dispersivity (D1 = 2 m, Dt = 0.1 m) or low dispersivity (D1 = 0.1 m,
Dt = 0.005 m).
104
Chapter Six
Figure 6.1. Monthly recharge of water to the subsurface estimated from
average monthly rainfall and temperature data (Payne 2004) near Waquoit
Bay using the Thornthwaite (1957) method.
113
Figure 6.2. (a) USGS well head levels above mean sea level (U.S.G.S. 2004a).
(b) Map of well locations (U.S.G.S. 2004b) and depths below mean sea level.
114
Figure 6.3. Hydraulic gradient vs. time over one tidal cycle for the February,
2004 experiment. In the top panel are the three piezometers closest to shore,
and the bottom panel depicts the four piezometers farthest offshore. Tide level
is shown on the right axis. Measurements taken on different days are assigned
a time relative to the tidal height.
118
Figure 6.4. Hydraulic gradient vs. time relative to tidal height over one tidal
cycle for the February, 2004 experiment. Each line represents one piezometer
measured over time.
118
Figure 6.5. Comparison of hydraulic gradient and discharge profiles for
summer and winter investigations along Transect W. (a) Summer discharge
(left axis) and winter hydraulic gradient (right axis). (b) Hydraulic
conductivity estimates from slug tests and interpolated values used to calculate
groundwater discharge from gradient measurements. The conductivity estimate
16
70 m from shore is extrapolated from the measured data. (c) Summer and
winter submarine groundwater discharge. Flow of baywater into the aquifer is
observed where maximum offshore outflow was measured during the summer.
Saline discharge is minimal in the February experiment. A small amount of
freshwater discharges more than 50 m offshore, likely upwelling from a
confined aquifer.
120
Figure 6.6. Porewater salinity and hydraulic gradient vs. distance from Well
1 along Transect W for February 2004 experiment. High upward gradient
offshore corresponds to very low porewater salinity, evidence of a connection
to a confined aquifer.
121
Figure 6.7. Discharge zone summary for saline circulation along Waquoit
Bay Transect W (Figure 3.1). Discharge data from August 2003 and February
2003 is presented in the top panel. Color bars represent approximate extent of
each zone of discharge along the transect. Zone 1 is depicted by cross-hatching
and extends from the shoreline to approximately 28 m into the bay. Zone 2 (red
shading) corresponds to nearshore circulation due to tides and waves and extends
approximately 3 m from the high tide mark. Dispersive circulation discharges in
zone 3 (blue-green shading), along the bayward edge of the fresh discharge.
Seasonal saline outflow occurs in zone 4 (purple shading). It has been measured
between 13 and 35 m from shore, but the zone likely extends to the shoreline,
depicted by the dashed purple bracket, where February measurements were not
possible.
122
Figure 6.8. (a.) Variation in discharge (top panel) and correlation coefficient
(bottom panel) for the August 2003 seepage meter study along Transect W (Figure
3.1). (b.) Variation in hydraulic gradient (top panel) and correlation coefficient
(bottom panel) for the February 2004 piezometer study along Transect W. A
decrease in both the absolute value of the correlation coefficient and the magnitude
of variation indicates a decline in tidal pumping. A correlation coefficient of -1
indicates a perfect inverse correlation between the tide and either discharge or
hydraulic gradient over a tidal cycle, and a correlation coefficient of zero implies
no correlation. The approximate extent of the tidal pumping zone during each
experiment is indicated by the vertical dashed lines.
123
Table 6.1. Hydraulic conductivity estimates from permeameter tests.
129
Figure 6.9. Schematic of measured and estimated parameter values for the four
geologic layers beneath Waquoit Bay. Literature values are listed on the left,
obtained from Masterson et. al. (1997a), Moench et. al. (2001), LeBlanc et. al.
(1991), Garabedian et. al. (1991), and Cambareri and Eichner (1998). Measured
values were obtained in this study through laboratory permeameter experiments,
and slug tests in wells and piezometers. Model values are those used in the
Waquoit Bay cross-section model.
130
17
Table 6.2. Values of porewater conductivity, average hydraulic head over three
typical tidal cycles, and hydraulic conductivity for Waquoit Bay Wells 1-7.
131
Figure 6.10. (a) Hydraulic head measurements over three tidal cycles during a
1-week measurement period from August 27 to September 2, 2003 for all seven
observation wells and the tide. (b) Hydraulic head as in (a.), but only for the
northern well cluster and the offshore Well 5 for clarity. Measurement error in
all wells includes survey error and pressure transducer error, and is
approximately ±2 cm in all wells with the exception of Well 5, which may have
a larger survey error. Well 1 may not be well-sealed, there are inconsistencies
in salinity and hydraulic head magnitude that indicate a possible crack in the
well casing shallower than the screen depth (see note under Table 6.2). (c)
Schematic of well locations and likely geology at the well screens inferred from
well logs, slug tests, and salinity measurements. Scale is approximate.
132
Figure 6.11. (a) Map of all continuous resistivity profile transects obtained by
Marcel Belaval, figure adapted from Belaval, (2003). (b) Schematic of northsouth transects WQ2 and WQ4, west-east transects WQ1 and WQl 1, and
Wells 1, 2, 3, and 4, which were profiled with borehole electromagnetic
induction.
134
Figure 6.12. Four continuous resistivity profiles, adapted from Belaval, (2003).
Higher resistivity (shown as warm colors) indicates lower salinity porewater.
135
Figure 6.13. Borehole (EM) induction and gamma logs, adapted from Belaval,
(2003). EM logging gives the conductivity of the formation surrounding a
borehole, higher conductivity corresponds with higher porewater salinity for
similar geology. Gamma logging gives a measure of clay content, or lithology,
of the formation surrounding the borehole: higher gamma means higher clay
content. (a) EM and gamma for Well 4 (onshore deep well) taken on 1/28/03.
Gamma log indicates that clay content begins to increase at a depth of
approximately 9 m below land surface, while EM shows fresh water to a
depth of 7 m, with a maximum at 9 m, then freshening with depth. (b)
Downhole logs for Well 3 taken on 1/28/03, results are similar to top 6 m of
Well 4. (c) Downhole logs for Well 1 8/26/02: EM suggests saline porewater
from approximately 2-12 m below land surface, with a change in lithology
below 8 m depth. (d) Downhole logs for Well 5 8/26/02: porewater is
consistently saline, but a higher clay content begins at approximately 7 m
below the sea floor.
136
Figure 6.14. Schematic of possible flow pattern in a cross-section of Waquoit
Bay. Dashed line is the position of the freshwater-saltwater interface, with a
Ghyben-Herzberg position in layers 1 and 3 and a nearly horizontal position
along layer 2, a balance between the two layers. Freshwater extends farther
bayward in layer 3 than in layer 1, leading to upward freshwater flow after
layer 2 is breached by the mucky layer. The low hydraulic conductivity muck
18
_ _
may prevent offshore flow of water as the interface moves bayward, resulting
in higher outflow and possibly creating the observed summer banded discharge.
138
Table 6.3. Extent and parameter values of geologic layer representations in
Waquoit Bay model.
139
Figure 6.15. Waquoit Bay model geometry. (a) Model proportions and
coordinates. Colors represent salinity values, where red = 30,000 mg/L and
blue = 0 mg/L. (b) Section from 33 m landward to 140 m bayward of the
shoreline. Colors represent layer locations that correspond with values of
hydraulic conductivity listed in Table 6.3. (c) Closer view of salinity
contours and layer boundaries.
140
Figure 6.16. Example of modeled density fingering. Flow vectors are pictured
in the inset. Colors represent porewater salt concentration: red = 30,000 mg/L
and blue = 0 mg/L. More buoyant fresh water flows upward and denser
saltwater downward, creating complex flow patterns.
141
Figure 6.17. Comparison of modeled salinity and measured conductivity
porewater profiles at Well 4 and Well 1. (a) Well 4 (located 8.8 m landward
of Well 1, Figure 6.10) EM conductivity profile measured by Marcel Belaval
on 1/28/03. (b) Model salinity vs. depth below land surface for low flow
(winter) and high flow (spring/summer) conditions along the line (-9.8, 2),
(-9.8, -15). (c) Well 1 (intertidal zone, Figure 6.10) EM conductivity profile
measured by Marcel Belaval on 8/26/02. (d) Model salinity vs. depth for low
flow (winter) and high flow (spring/summer) conditions along the line (-1, -1),
(-1, -15).
144
Figure 6.18. Total simulated freshwater and saltwater inflow and outflow
across the sea floor (left axis) and recharge (right axis) vs. simulation day for
the Waquoit Bay model. The sum of monthly fluxes calculated from nodal
velocity and salinity values was somewhat lower than the total yearly flux
calculated with the FEFLOW budget analyzer, so monthly values were scaled
to more accurately represent the total flow.
145
Appendices
Figure A.1. Submerged seepage meter. When in use, the seepage meter is
fully submerged, and a bag is attached to the center nozzle.
162
Figure A.2. Intertidal seepage meter schematic.
164
Figure A.3. Intertidal seepage meters in use on 8/14/2003.
165
Figure A.4. Porewater sampler and grid.
166
19
Figure A.5. Refractometer measurement [ppt] vs. conductivity probe
measurement [mS/cm] for a theoretical equation and water samples on
7/21/03 and 8/14/03.
167
Table B.1. Theoretical model mesh size and maximum time step (Chapter 5).
169
Figure C.1. Offshore well drilling system. Outer metal casing is driven into
the sediment by repeatedly dropping it using the rope, pulley, and winch.
Aquifer material was flushed out of the casing and collected for the well logs.
173
Figure D.1. Seepage meter numbering map for 1999 and 2000 head of the
bay experiments. Not to scale.
179
20
Chapter One
Introduction
1.1 The Importance of Groundwater at the Coast
Coastal hydrologic systems are essential to the survival of both human communities and
nearshore marine ecosystems throughout the world. In many coastal areas, aquifers are
the primary source of freshwater to residents, as well as a significant source of nutrients
to coastal habitats. Groundwater accounts for 97 percent of the earth's freshwater
resources (Church 1996) and supplies 1.5 billion people with drinking water throughout
the world (Alley et al. 2002). Despite its abundance, groundwater in many parts of the
world is being threatened by pollution, overuse, and in coastal regions, saltwater
intrusion. In the United States, more than half the population resides in coastal counties,
which constitute less than 20 percent of the land area (Barlow 2003). These populations
are continuing to grow at a rate that may threaten the ability of coastal aquifers to provide
an adequate water supply. The increased extraction of groundwater due to rising demand
causes saltwater to intrude inland, resulting in a deterioration of groundwater quality,
with the potential to undermine the aquifer as a source of freshwater. An analysis of case
studies of groundwater along the Atlantic Coast has shown that saltwater intrusion is
affected by the hydrogeologic setting, saltwater source, groundwater pumping, and
freshwater drainage (Barlow 2003). Therefore, a full understanding of density-dependent
groundwater dynamics and models that incorporate geologic and hydrologic constraints
are necessary to water resource managers in determining sustainable freshwater
extraction rates.
21
Population growth in coastal regions can also adversely affect the chemical composition
of the groundwater by introducing anthropogenic contaminants and elevating nutrient
concentrations from septic systems and agricultural activities. Nutrient transport in
discharging groundwater can result in eutrophication: an increase in primary production
decreases the depth of light penetration, increases the frequency of low oxygen, and
creates a shift in the speciation of flora and fauna in coastal ecosystems (Valiela 1995).
Transport of contaminants into these fragile environments by groundwater also degrades
the health of the ecosystems. Several studies have illustrated the effect of groundwater
discharge on the concentration of constituents in seawater. In Great South Bay, NY,
groundwater discharge accounts for greater than 20% of the estimated nitrogen input
from runoff (Capone and Bautista 1985), and in Flanders Bay, NY, groundwater flow
may contribute up to 58% of the total Cu to the bay (Montlucon and Sanudo-Wilhelmy
2001). A study along the southeast coast of the United States determined that the flux of
several cations as well as nitrate and phosphorous to the coastal ocean is greater from
groundwater seepage than from several major rivers (Simmons 1992). In addition, the
inorganic nutrient concentration in a South Carolina estuary was found to be an order of
magnitude higher in groundwater than in surface water (Whiting and Childers 1989), and
nitrate concentration in the groundwater in Waquoit Bay, MA is up to five orders of
magnitude greater than in the receiving baywater (Valiela et al. 1990). Seawater recycling
may also play an important role in controlling ocean chemistry (Tsunogai et al. 1996). In
each case, the extent to which groundwater seepage affects the coastal ecosystem
depends on the existence of relevant contaminants and nutrients in the groundwater
flowpath, the resulting concentrations in the groundwater, and the amount of seepage that
occurs. Freshwater-saltwater
mixing can also greatly affect chemical concentrations in
groundwater since the ions present in seawater may cause desorption of chemicals
previously fixed to soil particles.
Although formerly thought to be small in comparison to surface water runoff, combined
fresh and saline submarine groundwater discharge (SGD) may contribute up to 40% of
river flow (Moore 1996), with a freshwater component that generally ranges between 6%
and 10% (Taniguchi et al. 2002) of surface water inputs to coastal waters. Consequently,
22
the nutrient and contaminant contribution to nearshore marine ecosystems is potentially
significant. It is therefore necessary to be able to determine quantitatively the amount,
origin, and composition of submarine groundwater discharge in order to fully assess its
chemical contribution to coastal waters and the potential effect of changes in discharge
on the health of nearshore marine ecosystems.
1.2 Purpose and Scope of this Work
Coastal groundwater systems have been studied from many perspectives. Water resource
managers have investigated and modeled coastal aquifers in an effort to maximize
freshwater production while minimizing saltwater intrusion. Mathematicians have
proposed analytical models of idealized coastal systems and furthered understanding of
theoretical density-dependent flow systems. Oceanographers have considered
groundwater discharge to the ocean and studied its effect on the chemical composition of
coastal waters. This work is an attempt to integrate past efforts to better understand the
entire coastal groundwater system and the forces that influence it. The effect of temporal
forcing on both fresh and saline flow patterns are considered. From the seaward side, on a
small time-scale, tides and waves cause saltwater circulation and overtopping of saline
onto fresh water. From the landward side, the longer-scale seasonal recharge cycle causes
fluctuations in freshwater head, resulting in changes in freshwater discharge and
movement of the freshwater-saltwater interface. Through dispersion, the fresh and saline
groundwaters are interconnected: flow in one induces flow in the other.
This work begins with a five-year summer field study designed to estimate the amount,
pattern, and origin (terrestrial or marine) of submarine groundwater discharge in a small
coastal embayment. Small and large-scale spatial variability in discharge is explored, and
temporal variation on a tidal timescale is determined. The radium content of the measured
discharge is investigated to address the potential to estimate total SGD using natural
radium as a tracer. The field work was intended to clarify subsurface flow patterns and
the forcing mechanisms that drive flow, but instead introduced questions regarding the
mass balance and spatial pattern of saline groundwater discharge.
23
A discussion of saline circulation mechanisms beneath the coast in Chapter Four reveals
three potential sources of net saline outflow locally, though inflow must occur elsewhere
in space. These are nearshore circulation due to tides and waves, and deeper circulation
due to dispersion along the freshwater-saltwater interface. Field observations and
calculated estimates of this circulation, however, fail to explain the summer observations
of saline discharge. Thus, a fourth mechanism is proposed that conserves mass in time
rather than space: seasonal forcing of saline groundwater flow.
The theoretical basis for seasonal saline water exchange between aquifers and the coastal
ocean is introduced in Chapter Five. The potential for the existence of this mechanism in
dynamic aquifer systems is investigated through a series of idealized two-dimensional
numerical models. These homogeneous, isotropic aquifers vary in thickness, hydraulic
conductivity, and dispersivity to determine the sensitivity of the system to aquifer
characteristics. The models illustrate the effect of seasonal recharge and motion of the
upland fresh water table on the position of the freshwater-saltwater interface and
submarine groundwater discharge at the sea floor.
In theory, both conceptually and numerically, seasons produce oscillations in saline
groundwater flow, leading to outflow and inflow at different times of year. In Chapter
Six, this theory is investigated in the field through analysis of the regional hydrologic
seasonality and a winter field study. Characterization of the local hydrogeology leads to a
conceptual model of the subsurface flow patterns and salinity profile, which forms the
basis for a two-dimensional numerical model of the field site.
This work is not intended to determine with absolute certainty the groundwater flowpaths
and forcing mechanisms within a complex real aquifer system. Instead, it serves to
introduce seasonal variation as a potentially important forcing mechanism of groundwater
flow in coastal aquifers. This study gives evidence that large-scale saline water exchange
between aquifers and the ocean results from seasonal recharge cycles that drive
oscillations of the position of the subsurface freshwater-saltwater interface. This effort
24
demonstrates a connection between land-based hydrology and density-driven coastal
groundwater systems, with potential implications for chemical loading to coastal
ecosystems.
1.3 Significance and Applications
A more complete understanding of coastal groundwater systems and how they are
affected by temporal changes and aquifer properties will aid in coastal management.
Prediction of the effect of groundwater discharge on coastal ecosystems will also improve
with a more accurate picture of fresh and saline flow patterns. If saltwater is discharging
in much greater amounts and from circulation patterns that extend deeper into the aquifer,
its contribution to the chemical make-up of nearshore seawater may be more significant
than previously estimated. Seasonal inflow and outflow of saline water and large-scale
freshwater-saltwater mixing has been overlooked in the past. Such processes may greatly
affect the subsurface geochemistry, with implications for the global-scale estimation of
the ocean chemical budget.
This research also has important implications for tracer use to estimate submarine
groundwater discharge. Tracers often have different concentrations in fresh and saline
waters due to the water sources as well as effects of competitive sorption. Such
heterogeneity is demonstrated in this work with respect to radium isotopes. If saline
discharge contributes significantly to total SGD, and if tracer concentrations are different
in fresh and saline groundwater, then tracers such as radium cannot be used to estimate
total discharge using one value of endmember concentration.
The results of this work will demonstrate that aquifer parameters in numerical models can
greatly affect the simulated subsurface flow patterns. In particular, high values of
dispersivity, which are often necessary when using a coarse mesh in regional simulations,
may mask seasonal effects and small-scale processes such as density fingering. This can
be applied in future model construction: the question to be addressed by the model output
should be considered when determining input parameters.
25
References
Alley, W. M., R. W. Healy, J. W. LaBaugh, and T. E. Reilly (2002) Hydrology - Flow
and storage in groundwater systems. Science 296(5575): 1985-1990.
Barlow, P. M. (2003) Ground Water in Freshwater-Saltwater Environments of the
Atlantic Coast. U.S. Geological Survey Circular 1262: 113p.
Capone, D. G., and M. F. Bautista (1985) A Groundwater Source of Nitrate in Nearshore
Marine-Sediments. Nature 313(5999): 214-216.
Church, T. M. (1996) An underground route for the water cycle. Nature 380(6575): 579-
580.
Montlucon, D., and S. A. Sanudo-Wilhelmy (2001) Influence of net groundwater
discharge on the chemical composition of a coastal environment: Flanders Bay,
Long Island, New York. Environmental Science & Technology 35(3): 480-486.
Moore, W. S. (1996) Large groundwater inputs to coastal waters revealed by Ra-226
enrichments. Nature 380(6575): 612-614.
Simmons, G. M. (1992) Importance of Submarine Groundwater Discharge (Sgwd) and
Seawater Cycling to Material Flux across Sediment Water Interfaces in Marine
Environments. Marine Ecology-Progress Series 84(2): 173-184.
Taniguchi, M., W. C. Burnett, J. E. Cable, and J. V. Turner (2002) Investigation of
submarine groundwater discharge. Hydrological Processes 16(11): 2115-2129.
Tsunogai, U., J. Ishibashi, et al. (1996) Fresh water seepage and pore water recycling on
the seafloor: Sagami Trough subduction zone, Japan. Earth and Planetary Science
Letters 138(1-4): 157-168.
Valiela, I. (1995) Marine Ecological Processes. New York, Springer-Verlag.
Valiela, I., J. Costa, et al. (1990) Transport of Groundwater-Borne Nutrients from
Watersheds and Their Effects on Coastal Waters. Biogeochemistry 10(3): 177197.
Whiting, G. J., and K. L. Childers (1989) Subtidal advective water flux as a potentially
important nutrient input to southeastern U.S.A. saltmarsh estuaries. Estuarine,
Coastal, and Shelf Science 28: 417-431.
26
Chapter Two
Background
2.1 Groundwater at the Coast: Underlying Theory
Coastal groundwater systems are characterized by a nonlinear density-dependent balance
between freshwater and saltwater that is not fully understood. Fresh groundwater flow to
the sea is driven by regional head gradients and recharge. The freshwater is bounded at
depth by impermeable geology, or the interface with denser saltwater maintained by
hydrodynamic equilibrium. Saltwater flow in the subsurface is driven by mechanisms
such as hydrodynamic dispersion, tidal and wave dynamics, and seasonal fluctuations
(Figure 2.1). Traditionally, conceptual and mathematical models have assumed
freshwater and saltwater to be immiscible, resulting in a sharp freshwater-saltwater
interface. Under this assumption, Ghyben and Herzberg related the depth of the interface
to the elevation of the water table, assuming static conditions, and Muskat and Hubbert
formulated the position of the interface under equilibrium conditions (Reilly and
Goodman 1985). Glover (1959) extended this approximation to include net freshwater
flow to the sea. In reality, we know that saltwater and freshwater do mix, resulting in a
more dispersed interface. Cooper (1959) developed a hypothesis that was quantitatively
validated by Henry (1959). This work was the first to consider hydrodynamic dispersion,
resulting in miscible fluid flow, a mixing zone, and perpetual circulation of saltwater.
When temporal forcing is added to the system, the flow patterns become even more
complicated. Seasonal changes in upland water table height result in variable freshwater
discharge and movement of the freshwater-saltwater interface, while tidal forcing and
wave action cause saltwater circulation and unstable density gradients.
27
I land surface
Figure 2.1. Schematic of hypothetical coastal groundwater salinity distribution and flow
patterns.
2 2 Mdeling Submarine Groundwater Discharge
Despite the complexity of coastal groundwater systems,c m n t understanding of coastal
dynamics is derived primarily from simplified analytical and numerical models (Reilly
and W m a n 1985). Models that incorporate spatial variation and temporal change on an
aquifer scale have in the past been too complex for analytical analysis and too
compurationally expensive for numerical simulation. The majority of numerical models
have simulated coastal systems on a regional scale, and have concentrated only on the
fresh portion of flow in calculating submarine groundwater discharge (SGD) (Oberdorfer
2003). These simplified models predict discharge that is primarily fresh and decreases
monotonically with distance from shore. Recent field studies, however, have shown that
submarine groundwater may be primarily saline and discharge in compiex panem
(Michael et al. 2003; Smith and Zawadzki 2003). Due to improved numerical codes and
computational speed, it i s now possible to sirnulate such cornplexity through
representation of density-driven dynamics.
Regional-scale models that incorporate density effects have been developed to aid in
water resources management, predict saltwater intrusion, and estimate submarine
groundwater discharge to coastal waters. For example, Panday et. al. (1993) conducted a
three-dimensional modeling study considering pumping and non-pumping scenarios in a
model of the Geneva area, Florida (1993). A three-dimensional model was also used to
examine both large and small-scale effects on SGD in a bay in the western Baltic Sea
(Kaleris et al. 2002), and Langevin (2003) estimated the SGD into Biscayne Bay, Florida
using a two-dimensional model calibrated to measured hydraulic heads, known
groundwater fluxes, and the position of the freshwater-saltwater interface.
Numerical studies have also examined the effects of small-scale complexities on coastal
systems that are neglected in simple models. Robinson and Gallagher (1999) and AtaieAshtiani, et. al. (1999) have shown that incorporating tidal dynamics into a model
significantly affects the configuration of salt-concentration and equipotential contours in
the subsurface as well as groundwater flow patterns. As tides are introduced, the interface
becomes more dispersed, the groundwater seepage face widens, and an inverted density
gradient is created. Where a fluid overlies another of lesser density, instability may lead
to density-driven free convection. This has been modeled numerically (Schincariol et al.
1994; Diersch 1998; Ibaraki 1998; Simmons et al. 1999; Eliassi and Glass 2001;
Simmons et al. 2001; Diersch and Kolditz 2002), and can have a significant effect on the
rate of material transport through sediments (Webster et al. 1996). Aquifer heterogeneity
on both large and small spatial scales has also been neglected in the past. However, such
heterogeneity can greatly affect subsurface flow patterns as well as the formation of
density instabilities.
The incorporation of temporal forcing, density, and heterogeneity into numerical models
requires small time steps, a coupled, iterative simulation, and fine grid spacing. These
factors result in a high computational demand. It is therefore difficult to consider all
relevant phenomena in regional coastal hydrogeologic models. Thus, a balance between
the amount of spatial and temporal detail in the model and computational capabilities
must be found.
29
2.3 Field Studies
If models are to be applicable to real systems, they must be corroborated by field data.
Several studies have attempted to characterize coastal groundwater discharge by direct
measurement using seepage meters (Bokuniewicz 1980; Whiting and Childers 1989;
Bokuniewicz and Pavlik 1990; Simmons 1992; Cable et al. 1997a; Burnett 2002;
Taniguchi et al. 2003). The challenges in using point measurements have included a large
variability in discharge as well as a lack of sample density in both space and time.
Automated seepage meters have been developed to obtain continuous measurements of
submarine groundwater discharge. Technologies include the continuous heat-type meter
(Taniguchi and Iwakawa 2001), a heat-pulse type meter (Taniguchi and Fukuo 1993;
Krupa et al. 1998), an ultrasonic-type meter (Paulsen et al. 2001), a dye-dilution seepage
meter (Sholkovitz et al. 2003), and an electromagnetic seepage meter (Rosenberry and
Morin 2004). These seepage meters overcome the temporal limitations of traditional Leetype (Lee 1977) seepage meters, but the expense in design may limit the spatial density of
measurements. Other methods of estimating SGD in the field include head gradient
measurements with Darcy's Law (Tobias et al. 2001; Ullman et al. 2003), water budgets
(Cambareri and Eichner 1998), and hydrograph separation (Zektser 2002). These
methods generally incorporate only the freshwater portion of discharge (Oberdorfer
2003), however, and estimation uncertainty can be high (Burnett et al. 2001).
There is often a significant temperature difference between discharging SGD and the
overlying surface water, allowing for SGD estimation by temperature gradient
measurement (Land and Paull 2001; Taniguchi et al. 2003) and thermal imagery (Banks
et al. 1996). Thermal imagery provides information on spatial flow patterns but flow rates
are difficult to determine, and the flow of surface water into the aquifer cannot be
detected. Temperature gradient measurements allow for point flow rate estimation, but
neither method provides a means to determine seepage salinity.
30
Alternative methods of estimating submarine groundwater discharge include the use of
naturally-occurring tracers such as radium isotopes (Miller et al. 1990; Moore 1996;
Rama and Moore 1996; Krest et al. 2000; Charette et al. 2001; Charette et al. 2003;
Crotwell and Moore 2003; Boehm et al. 2004), radon (Cable et al. 1996; Kim and Hwang
2002; Burnett and Dulaiova 2003; Chanton et al. 2003; Lambert and Burnett 2003),
methane (Cable et al. 1996; Swarzenski et al. 2001) and barium (Moore 1997; Shaw et al.
1998). While such methods may provide a more integrated estimate of total discharge,
spatial and temporal patterns in flux and composition are not likely discernible.
Currently, there is not one completely accurate technique for estimating submarine
groundwater discharge that provides information on small and large-scale spatial and
temporal variability as well as discharge composition. However, incorporating more than
one method can reduce estimation error and more fully characterize SGD. There have
been several intercalibration experiments where multiple researchers have concurrently
tested numerous methods and new technologies for SGD estimation (Burnett et al. 2003;
Taniguchi et al. 2003). Testing at the same site enables the recognition of potential
sources of error as well as more efficient ways to combine measurement techniques and
numerical modeling. Further research attempting to fully characterize coastal
groundwater systems through a multi-faceted and interdisciplinary approach will advance
the overall understanding of coastal dynamics and the importance of submarine
groundwater discharge on a global scale.
31
References
Ataie-Ashtiani, B., R. E. Volker, and D. A. Lockington (1999) Tidal effects on sea water
intrusion in unconfined aquifers. Journal of Hydrology 216(1-2): 17-31.
Banks, W. S. L., R. L. Paylor, and W. B. Hughes (1996) Using thermal-infrared imagery
to delineate ground-water discharge. Ground Water 34(3): 434-443.
Boehm, A. B., G. G. Shellenbarger, and A. Paytan (2004) Groundwater discharge:
Potential association with fecal indicator bacteria in the surf zone. Environmental
Science & Technology 38(13): 3558-3566.
Bokuniewicz, H. (1980) Groundwater Seepage into Great South Bay, New-York.
Estuarineand CoastalMarineScience10(4):437-444.
Bokuniewicz, H., and B. Pavlik (1990) Groundwater Seepage Along a Barrier-Island.
Biogeochemistry 10(3): 257-276.
Burnett, W., J. Chanton, J. Christoff, E. Kontar, S. Krupa, M. lambert, W. Moore, D.
O'Rourke, R. Paulsen, C. Smith, L. Smith, and M. Taniguchi (2002) Assessing
methodologies for measuring groundwater discharge to the ocean. EOS 83(11):
117-123.
Burnett, W. C., H. Bokuniewicz, M. Huettel, W. S. Moore, and M. Taniguchi (2003)
Groundwater and pore water inputs to the coastal zone. Biogeochemistry 66(1-2):
3-33.
Burnett, W. C., and H. Dulaiova (2003) Estimating the dynamics of groundwater input
into the coastal zone via continuous radon-222 measurements. Journal of
Environmental Radioactivity 69(1-2): 21-35.
Burnett, W. C., M. Taniguchi, and J. Oberdorfer (2001) Measurement and significance of
the direct discharge of groundwater into the coastal zone. Journal of Sea Research
46(2): 109-116.
Cable, J. E., G. C. Bugna, W. C. Burnett, and J. P. Chanton (1996) Application of Rn-222
and CH4 for assessment of groundwater discharge to the coastal ocean.
Limnology and Oceanography 41(6): 1347-1353.
Cable, J. E., W. C. Burnett, and J. P. Chanton (1997a) Magnitude and variations of
groundwater seepage along a Florida marine shoreline. Biogeochemistry 38(2):
189-205.
32
Cable, J. E., W. C. Burnett, J. P. Chanton, and G. L. Weatherly (1996) Estimating
groundwater discharge into the northeastern Gulf of Mexico using radon-222.
Earthand PlanetaryScienceLetters 144(3-4):591-604.
Cambareri, T. C., and E. M. Eichner (1998) Watershed delineation and ground water
discharge to a coastal embayment. Ground Water 36(4): 626-634.
Chanton, J. P., W. C. Burnett, H. Dulaiova, D. R. Corbett, and M. Taniguchi (2003)
Seepage rate variability in Florida Bay driven by Atlantic tidal height.
Biogeochemistry 66(1-2): 187-202.
Charette, M. A., K. O. Buesseler, and J. E. Andrews (2001) Utility of radium isotopes for
evaluating the input and transport of groundwater-derived nitrogen to a Cape Cod
estuary. Limnology and Oceanography 46(2): 465-470.
Charette, M. A., R. Splivallo, C. Herbold, M. S. Bollinger, and W. S. Moore (2003) Salt
marsh submarine groundwater discharge as traced by radium isotopes. Marine
Chemistry 84(1-2): 113-121.
Cooper, H. H. (1959) A hypothesis concerning the dynamic balance of fresh water and
salt water in a coastal aquifer. Journal of Geophysical Research 64(4): 461-467.
Crotwell, A. M., and W. S. Moore (2003) Nutrient and radium fluxes from submarine
groundwater discharge to Port Royal Sound, South Carolina. Aquatic
Geochemistry 9(3): 191-208.
Diersch, H. J. G. (1998) FEFLOW finite element subsurface flow and transport
simulation system - user's manual/reference manual/white papers. Release 4.9.
WASY Ltd, Berlin.
Diersch, H. J. G., and 0. Kolditz (2002) Variable-density flow and transport in porous
media: approaches and challenges. Advances in Water Resources 25(8-12): 899944.
Eliassi, M., and R. J. Glass (2001) On the continuum-scale modeling of gravity-driven
fingers in unsaturated porous media: The inadequacy of the Richards equation
with standard monotonic constitutive relations and hysteretic equations of state.
Water Resources Research 37(8): 2019-2035.
Glover, R. E. (1959) The pattern of fresh-water flow in a coastal aquifer. Journal of
Geophysical Research 64(4): 457-459.
Henry, H. R. (1959) Salt intrusion into fresh-water aquifers. Journal of Geophysical
Research 64(11): 1911-1919.
Ibaraki, M. (1998) A robust and efficient numerical model for analyses of densitydependent flow in porous media. Journal of Contaminant Hydrology 34(3): 235246.
33
Kaleris, V., G. Lagas, S. Marczinek, and J. A. Piotrowski (2002) Modelling submarine
groundwater discharge: an example from the western Baltic Sea. Journal of
Hydrology 265(1-4): 76-99.
Kim, G., and D. W. Hwang (2002) Tidal pumping of groundwater into the coastal ocean
revealed from submarine Rn-222 and CH4 monitoring. Geophysical Research
Letters 29(14): 10.1029/2002GL015093.
Krest, J. M., W. S. Moore, L. R. Gardner, and J. T. Morris (2000) Marsh nutrient export
supplied by groundwater discharge: Evidence from radium measurements. Global
Biogeochemical Cycles 14(1): 167-176.
Krupa, S. L., T. V. Belanger, H. H. Heck, J. T. Brok, and B. J. Jones (1998) Krupaseep -
the next generation seepage meter. Journal of Coastal Research 25: 210-213.
Lambert, M. J., and W. C. Burnett (2003) Submarine groundwater discharge estimates at
a Florida coastal site based on continuous radon measurements. Biogeochemistry
66(1-2): 55-73.
Land, L. A., and C. K. Paull (2001) Thermal gradients as a tool for estimating
groundwater advective rates in a coastal estuary: White Oak River, North
Carolina, USA. Journal of Hydrology 248(1-4): 198-215.
Langevin, C. D. (2003) Simulation of submarine ground water discharge to a marine
estuary: Biscayne Bay, Florida. Ground Water 41(6): 758-771.
Lee, D. R. (1977) Device for Measuring Seepage Flux in Lakes and Estuaries. Limnology
and Oceanography 22(1): 140-147.
Michael, H. A., J. S. Lubetsky, and C. F. Harvey (2003) Characterizing submarine
groundwater discharge: a seepage meter study in Waquoit Bay, Massachusetts.
Geophysical Research Letters 30(6): 10.1029/GL016000.
Miller, R. L., T. F. Kraemer, and B. F. Mcpherson (1990) Radium and Radon in Charlotte
Harbor Estuary, Florida. Estuarine Coastal and Shelf Science 31(4): 439-457.
Moore, W. S. (1996) Large groundwater inputs to coastal waters revealed by Ra-226
enrichments. Nature 380(6575): 612-614.
Moore, W. S. (1997) High fluxes of radium and barium from the mouth of the GangesBrahmaputra river during low river discharge suggest a large groundwater source.
Earth and Planetary Science Letters 150(1-2): 141-150.
Oberdorfer, J. A. (2003) Hydrogeologic modeling of submarine groundwater discharge:
comparison to other quantitative methods. Biogeochemistry 66(1-2): 159-169.
Panday, S., P. S. Huyakorn, J. B. Robertson, and B. Mcgurk (1993) A Density-Dependent
Flow and Transport Analysis of the Effects of Groundwater Development in a
34
Fresh-Water Lens of Limited Areal Extent - the Geneva Area (Florida, USA)
Case-Study. Journal of Contaminant Hydrology 12(4): 329-354.
Paulsen, R. J., C. F. Smith, D. O'Rourke, and T. F. Wong (2001) Development and
evaluation of an ultrasonic ground water seepage meter. Ground Water 39(6):
904-911.
Rama, and W. S. Moore (1996) Using the radium quartet for evaluating groundwater
input and water exchange in salt marshes. Geochimica Et Cosmochimica Acta
60(23): 4645-4652.
Reilly, T. E., and A. S. Goodman (1985) Quantitative-Analysis of Saltwater Fresh-Water
Relationships in Groundwater Systems - a Historical-Perspective. Journal of
Hydrology 80(1-2): 125-160.
Robinson, M. A., and D. L. Gallagher (1999) A model of ground water discharge from an
unconfined coastal aquifer. Ground Water 37(1): 80-87.
Rosenberry, D. O., and R. H. Morin (2004) Use of an electromagnetic seepage meter to
investigate temporal variability in lake seepage. Ground Water 42(1): 68-77.
Schincariol, R. A., F. W. Schwartz, and C. A. Mendoza (1994) On the Generation of
Instabilities in Variable-Density Flow. Water Resources Research 30(4): 913-927.
Shaw, T. J., W. S. Moore, J. Kloepfer, and M. A. Sochaski (1998) The flux of barium to
the coastal waters of the southeastern USA: The importance of submarine
groundwater discharge. Geochimica Et Cosmochimica Acta 62(18): 3047-3054.
Sholkovitz, E. R., C. Herbold, and M. A. Charette (2003) An automated dye-dilution
based seepage meter for the time-series measurement of submarine groundwater
discharge. Limnology and Oceanography: Methods 1: 16-28.
Simmons, C. T., T. R. Fenstemaker, and J. M. Sharp (2001) Variable-density
groundwater flow and solute transport in heterogeneous porous media:
approaches, resolutions and future challenges. Journal of Contaminant Hydrology
52(1-4): 245-275.
Simmons, C. T., K. A. Narayan, and R. A. Wooding (1999) On a test case for density-
dependent groundwater flow and solute transport models: The salt lake problem.
Water Resources Research 35(12): 3607-3620.
Simmons, G. M. (1992) Importance of Submarine Groundwater Discharge (Sgwd) and
Seawater Cycling to Material Flux across Sediment Water Interfaces in Marine
Environments. Marine Ecology-Progress Series 84(2): 173-184.
Smith, L., and W. Zawadzki (2003) A hydrogeologic model of submarine groundwater
discharge: Florida intercomparison experiment. Biogeochemistry 66(1-2): 95-110.
35
Swarzenski, P. W., C. D. Reich, R. M. Spechler, J. L. Kindinger, and W. S. Moore (2001)
Using multiple geochemical tracers to characterize the hydrogeology of the
submarine spring off Crescent Beach, Florida. Chemical Geology 179(1-4): 187202.
Taniguchi, M., W. C. Burnett, et al. (2003) Spatial and temporal distributions of
submarine groundwater discharge rates obtained from various types of seepage
meters at a site in the Northeastern Gulf of Mexico. Biogeochemistry 66(1-2): 3553.
Taniguchi, M., and Y. Fukuo (1993) Continuous measurements of ground-water seepage
using an automatic seepage meter. Ground Water 31(4): 675-679.
Taniguchi, M., and H. Iwakawa (2001) Development of continuous heat-type automated
seepage meter and applications in Osaka Bay, Japan. Journal of Groundwater
Hydrology 43(4): 271-277.
Taniguchi, M., J. V. Turner, and A. J. Smith (2003) Evaluations of groundwater
discharge rates from subsurface temperature in Cockburn Sound, Western
Australia. Biogeochemistry 66(1-2): 111-124.
Tobias, C. R., J. W. Harvey, and I. C. Anderson (2001) Quantifying groundwater
discharge through fringing wetlands to estuaries: Seasonal variability, methods
comparison, and implications for wetland-estuary exchange. Limnology and
Oceanography 46(3): 604-615.
Ullman, W. J., B. Chang, D. C. Miller, and J. A. Madsen (2003) Groundwater mixing,
nutrient diagenesis, and discharges across a sandy beachface, Cape Henlopen,
Delaware (USA). Estuarine Coastal and Shelf Science 57(3): 539-552.
Webster, I. T., S. J. Norquay, F. C. Ross, and R. A. Wooding (1996) Solute exchange by
convection within estuarine sediments. Estuarine Coastal and Shelf Science
42(2): 171-183.
Whiting, G. J., and K. L. Childers (1989) Subtidal advective water flux as a potentially
important nutrient input to southeastern U.S.A. saltmarsh estuaries. Estuarine,
Coastal, and Shelf Science 28: 417-431.
Zektser, I. S. (2002) Principles of regional assessment and mapping of natural
groundwater resources. Environmental Geology 42(2-3): 270-274.
36
Chapter Three
Field Investigation in Waquoit Bay
3.1 Motivation and Objectives
Submarine groundwater discharge (SGD) has been investigated using seepage meters,
piezometers, and tracers on both large and small scales. However, a study that captures
small-scale spatial and temporal variations and incorporates them into a large-scale total
discharge estimate had not been attempted prior to this work. There are several reasons
for this. First, the focus of many studies has not been on estimating total discharge but on
the evaluation of measurement methods or obtaining insight into groundwater flux at
specific locations. Another reason is that in many cases the field sites are large and
impossible to cover fully using measurements at specific locations. Lastly, large-scale
estimates of total SGD have been made through the use of tracer measurements, but these
cannot capture the spatial and temporal variability present on smaller scales. Tracer
estimates of SGD can also be highly susceptible to error in coastal systems if tracer
concentrations and chemical reactivity differ in fresh and saline groundwater.
In this study, the field site is a small bay (approximately 3 km 2 in area), where most of
the groundwater discharge is likely confined to the 610 m long head of the bay based on
the upland watershed geometry. It is possible, therefore, to obtain a bay-scale estimate of
total SGD using an instrument field dense enough to minimize the error due to smallscale variability. The objective of this field study is to accurately characterize the rate,
pattern, and variability of submarine groundwater discharge using a dense field of
seepage meters. The temporal variation in seepage on several scales will be addressed. In
37
addition, the variability in discharge over small spatial scales will be used to evaluate the
number of seepage meters necessary to adequately characterize groundwater discharge.
Direct measurement of radium isotopes in water collected in seepage meters will enable
an analysis of the use of radium as a tracer to estimate total SGD. The origin of the
discharge as well as its possible flow pattern in the subsurface will be discussed to further
understanding of the effect of tides and density-dependence on coastal groundwater flow
dynamics.
3.2 Study Site Description
The Waquoit Bay National Estuarine Research Reserve (WBNERR) is located on the
southern shore of Cape Cod, Massachusetts. This coastal embayment (Figure 3.1) has an
average depth of 1 m and a tidal range of approximately 0.5 m. Waquoit Bay has been the
subject of several previous studies (Valiela et al. 1990; Valiela et al. 1992; Barlow and
Hess 1993; Geyer 1997; Cambareri and Eichner 1998; Valiela et al. 2000; Bowen and
Valiela 2001; Charette et al. 2001; Hoefel and Evans 2001; Charette and Sholkovitz
2002; Testa et al. 2002; Abraham et al. 2003; Belaval 2003; Talbot et al. 2003), making it
an ideal site for further investigation of hydrologic processes. The aquifer is 100-120 m
thick in the area along the southern coast, and includes an upper permeable layer
approximately 11 m thick under the head of Waquoit Bay. This is underlain by a less
permeable layer of fine sand, silt, and clay, and bounded below by basal till and bedrock
at a depth of approximately 120 m (Masterson et al. 1997). The head of the bay subwatershed is 0.76 km2 in area, with a recharge rate of approximately 46 cm/year, and an
estimated hydraulic gradient of 0.002 (Cambareri and Eichner 1998). This means that
significant freshwater discharge into the head of Waquoit Bay is expected.
Groundwater discharge into Waquoit Bay is important because of the potential for
introduction of nutrients or contaminants that may greatly affect the estuarine ecosystem.
One major problem is that the nitrogen input to the Waquoit Bay watershed has increased
steadily since the 1930's, due primarily to increased atmospheric deposition, fertilizer
use, and wastewater from a growing population (Valiela et al. 2002). This increased
38
nutrient input enters the estuary primarily through the groundwater and has resulted in
significant eutrophication, causing a shift in the ecosystem. The shift has included a
significant decline in the eelgrass coverage as well as an increase in macroalgae and a
decline in the shellfish population of the bay (Valiela et al. 1990). A better understanding
of the groundwater system in Waquoit Bay is therefore necessary if the mechanisms of
ecological change are to be more fully identified.
Figure 3.1. Map of Waquoit Bay: experimental design for seepage meter studies. Five
experiments at the head of the bay and one off of Washburn Island were performed over
five summers. Transects E and W (East and West) are sites of several sets of field
measurements. Well 1 refers to a well in the intertidal zone along Transect W used as a
39
reference point for distance into the bay, and the CCC Well is a deep multi-level
sampling well installed by the Cape Cod Commission.
3.3 Seepage Meters
Seepage meters have been widely used to measure groundwater discharge into lakes,
rivers and coastal waters (Bokuniewicz 1980; Capone and Bautista 1985; Whiting and
Childers 1989; Simmons 1992; Cable et al. 1997a; Portnoy et al. 1998; Burnett et al.
2003). When used correctly, and where fluxes are large, seepage meters have been shown
to give reproducible flux estimates (Lee 1977; Burnett 2002). Where flows are small,
seepage meters may overestimate flux, perhaps due to effects of currents and waves
(Shinn et al. 2002), although underestimation has also been reported (Belanger and
Montgomery 1992). The waves and currents in Waquoit Bay are minimal, however, and
measured flow rates are up to three orders of magnitude greater than those described in
Shinn et al. (2002). A recent study (Burnett 2002) found that groundwater discharge
measured by seepage meters agreed with total discharge estimated by natural tracers (Ra,
Rn), although the tracer method does not describe the spatial pattern or salinity of
discharge.
Forty conventional submerged seepage meters were constructed from the ends of 55gallon drums similar to those described in Lee (1977) (Appendix A.1). Each meter has a
7.5 cm vent hole, which was left open during meter placement so that pressures quickly
equilibrated with the bay, and then plugged before attaching a thin-walled plastic bag to a
separate quick-connect fitting. Each bag was pre-filled with at least 1 liter of bay water
before placement on the seepage meters to prevent under-filling (Shaw and Prepas 1989)
and so that negative flows may be indicated. During seepage meter experiments, the bags
were left on for two-hour periods before they were replaced. The bags were then weighed
to determine the amount of groundwater seepage and salinity measured with a hand-held
conductivity probe.
40
3.4 Discharge Patterns
3.4.1 Head of the Bay
3.4.1.1. Seepage meter grids. Two sampling campaigns were conducted at the head of
Waquoit Bay during August 1999 and July 2000.Both used 40 seepage meters arrayed in
four rransects perpendicular to the coast (Figure 3.1 ). These were sampled every two
hours over a complete tidal cycle. A distinct band of high discharge parallel to the coast
between 20 and 45 m from the shore (Figure 3.2 (a) and (b), Figure 3.4 (a)) was observed
in all four transects and six time intervals in both campaigns. Inflow was measured in 15
samples in 1999 and 7 in 2000.This inflow occurred primarily in seepage meters located
far from shore, with no apparent correlation to the bay water level over single tidal
cycles.
Figure 3.2. (a) and (b) Time-averaged groundwater discharge vs. distance from
approximate mean shoreline for the head of the bay experiments. Discharge i s separated
into freshwater, saltwater, and unknown (?)components. (c) and (d) Corresponding
standard deviation in space and time.
The proportion of freshwater discharging into the seepage meters was calculated from the
salinity and volume of the discharge and the estimated salinity of the baywater that
recharges the sediments, which may vary over time. An upper bound for this endmember
salinity of inflowing baywater may be estimated as 33.0 ppt, the salinity of the seawater
just outside the mouth of Waquoit Bay in Vineyard Sound, and the lower bound as the
highest salinity among all the seepage meter bags: 29.3 and 30.6 ppt for 1999 and 2000,
respectively. The uncertainty in the salinity of the inflow results in the small 'unknown'
component of outflow depicted in Figures 3.2 (a) and 3.2 (b). However, most of the
discharging water was saline with the exception of some fresher discharge in the row of
seepage meters nearest the shore.
3.4.1.2. Single seepage meter transects. In August of 2002 and 2003, further
investigations were conducted to more accurately quantify nearshore freshwater
discharge and to attempt to measure saline inflow to balance the measured outflow.
Twenty submerged seepage meters were positioned along Transect W in two lines,
m
apart. Eight novel intertidal seepage meters were used in several locations, varying with
the position of the tide (approximate locations are depicted in Figure 3.1). A description
of the intertidal seepage meters and their construction can be found in Appendix A.2. As
in previous experiments, the 28 seepage meters were sampled every two hours over a
tidal cycle. The salinity of the discharge was determined in the submerged seepage
meters by measuring the conductivity in the seepage meter bags before and after
deployment. In 2002, this method was also applied to the intertidal meters, but the mass
balance introduced a large amount of error to the analysis. This was corrected in the 2003
experiment by taking a 2-5 mL porewater sample at a depth of 3 cm as described in
Appendix A.3 and measuring salinity with a refractometer. The refractometer was
calibrated to the conductivity probe so that salinity and conductivity measurements could
be directly compared (Appendix A.3).
A second improvement in the 2003 study was in the amount of initial water in the
submerged seepage meter bags. Analysis of the ability of the seepage meter bags to
42
register inflow has shown that an initial volume of 5 L, rather than 1 L deemed
appropriate by Shaw and Prepas (1989), is necessary to overcome a resistance to
emptying so that the full inflow volume is measured (Emily Slaby, personal
communication). This improvement did not appear to matter greatly in this case,
however, as only 4 submerged seepage meters recorded inflow, at a total of 10 sample
times, 7 of which were within measurement error of zero flux.
The intertidal and submerged seepage meters were in close agreement for both discharge
and salinity, as evidenced by measurements from intertidal and submerged seepage
meters placed at distances of 2.0 m (intertidal) and 3.1 m (submerged) from shore in
2002, and 1.6 m (intertidal), 2.1 m (submerged) and 2.2 m (intertidal) from shore in 2003.
The salinity measurements agree nearly exactly in 2003: the salinity of the porewater at
the 2.1 m intertidal seepage meter location was 3.4 ppt, while the calculated salinity of
the discharge into the submerged seepage meter 2.2 m from shore was 3.8 ppt. The
groundwater discharge at adjacent locations is more variable than the salinity, as
discussed in the next section. Nevertheless, the agreement in discharge measurements
between adjacent submerged and intertidal seepage meters is relatively close. In 2002, the
intertidal and submerged seepage meters registered 0.35 m/d and 0.52 m/d, respectively,
and in 2003, the measurements were 0.39 m/d (intertidal), 0.34 m/d (submerged), and
0.53 m/d (intertidal).
In both 2002 and 2003, the total discharge was mostly saline and the banded discharge
pattern in the middle zone was clearly evident (Figure 3.3). A second band of discharge
very near the shore was measured in the 2003 experiment, comprised of approximately
50% fresh and 50% saline water. Thus the intertidal seepage meters and improved
salinity measurements enabled quantification of nearshore freshwater outflow that was
not captured in previous experiments. Net inflow over a tidal cycle was not measured at
any location in either 2002 or 2003, however. Thus a saline fluid mass balance was not
achieved in any summer seepage meter study presented in this chapter.
43
la tertidal Seepage Metera
0
-S
S
IS
35
25
Diitance tram Well I
45
55
Iml
Figure 3.3. Temporally and spatially averaged discharge (a) and salinity (b) vs. distance
from shore for the 20Q2and 2003 single-transect experiments. Measurements taken with
submerged and intertidal seepage meters are shown separately.
3.4.2
Slug Tests
Slug tests conducted in October 2000 at 6 locations along Transect W (Figure 3.1 )
indicate that the permeability of bay sediments general1y decreases with distance from
shore. This pattern was observed again in August 2002 at 1 8 locations along the same
transect (Figure 3.4 (b)). The pattern in relative permeability (based on the rate of decline
of the water level during a slug test) does not coincide with the high band of discharge
observed in August 2000 (Figure 3.4 (a)), or in other experiments along the head of
Waquoil Bay. Although the trend in permeability is clearly evident, the scatter in the
measurements is an indication of the degree of heterogeneity in the near-surface
sediment. Piezometers screened over the bottom 0.2 m were driven 0,6 m into the
subsurface. A slug of water was added to the clear PVC top section of the piezometers
and the water level was recorded as it dropped. The relative permeability, or the inverse
of the time for the water level to drop 90% of the distance to the bay surface, is plotted in
Figure 3.4 (b).
Shoreline
b
Relative Permeability [irrl]
Figure 3.4. (a) Grayscale and contours are time-averaged groundwater discharge [mld]
far the 2000 head of the bay experiment. (b) Relative permeability as a function of
distance into the bay.
3.43 Minimal Freshwater Flow: Island study
A third seepage meter experiment was conducted on a narrow piece of land jutting off of
Washburn Island (Figure 3.1) in August 2000 to investigate the groundwater discharge
panern where there is little freshwater discharge. Freshwater flow is small because this
narrow spit drains a very small area, thereby virtually eliminating density differences and
a regional gadient while maintaining tidal forces. The twenty seepage meters in this
study were sampled every 2 hours over a tidal cycle with a range of 0.56 m. They reveal
that discharge contains negligible freshwater and ranges from 0.11 to 0.22 m/d when
averaged over time. The discharge pattern is essentially spatially uniform, contrasting
sharply with the pattern at the head of the bay (Figure 3.5). The seepage variation that we
do see likely results from the natural spatial variability observed in all seepage meter
experiments, and the nonzero total discharge may be caused by tidal pumping, or a small
lateral head gradient due to tidal currents around the spit of land.
'II
a..
0
10
20
30
40
50
Distance from Shore [Im
60
70
80
Figure 3.5. Groundwater discharge vs. distance from approximate mean shoreline for the
2000 experiments at the head of the bay and Washburn Island.
3.5 Heterogeneity in Space and Time
3.5.1 Spatial Variability
3.5.1.1 Head of the Bay Experiments, 50 m Scale. The seepage meter grids indicate
large variability in discharge over both time and space (Figures 3.2 (c) and 3.2 (d)). The
spatial standard deviation as a function of distance into the bay was calculated from the
46
four seepage meters in each row after averaging the data from each seepage meter over
time. Similarly, the temporal standard deviation was calculated over the six time periods
after averaging the discharge across the four seepage meters in each row. Comparison of
these plots to the plots of discharge vs. distance from shore (Figures 3.2 (a) and 3.2 (b))
shows that greater discharge correlates with greater variability. Also, the temporal
variability on a tidal cycle scale, while slightly lower, is of the same order of magnitude
as the spatial variability.
3.5.1.2 Cluster Experiments, 1 m Scale. Two further experiments were conducted with
clusters of seepage meters spaced closely together to characterize variability over smaller
areas and longer times. In the 1999 cluster experiment, nine seepage meters were placed
in the nearshore zone (Figure 3.6) and sampled during daylight hours for two-hour
periods on six days over two weeks. The 2001 experiment examined discharge variability
on both a large (50 m) and small (1 m) scale. Eighteen seepage meters arranged in
clusters along Transect E (Figure 3.1) were sampled every two hours over three tidal
cycles, except for those farthest from shore, which were not sampled overnight.
Data from the cluster experiments reveal that differences in discharge over small spatial
scales can be similar in magnitude to differences in discharge over larger spatial scales.
The 1999 cluster experiment (Figure 3.6) indicates that adjacent seepage meters may
differ greatly in discharge. For example, during the same two-hour period, two seepage
meters less than 2m apart registered 0.05 m/d and 0.37 m/d. In the 2001 cluster
experiment, the standard deviation of the time-averaged data is 0.029 m/d for the seepage
meters in the nearshore cluster and 0.053 m/d for the middle zone cluster. All of the data
taken together, spanning nearly 60 m, also has a standard deviation of 0.053 m/d.
3.5.1.3 Variability in Discharge Salinity, 5-cm Scale. In areas where saltwater exists on
top of freshwater, saltwater and freshwater fingers may form due to an unstable density
gradient. Under certain conditions, the instability may lead to upwelling of freshwater
and downwelling of saltwater in lobes of finite dimension. This phenomenon was not
observed in the seepage meter experiments. Although the amount of discharge varies
47
widely between seepage meters, the discharge salinity is relatively consistent as a
function of distance from shore. The seepage meters cover an area of 0.25 m2 , however,
which may average out and obscure smaller-diameter fingers.
---
E 0.25
I'
0.2
a
S 0.15
.
01I
E
0.05
~2
07/15
07/17
07/21 07/23 07/25
Sample Date
07/19
Cluster
Configuration
07/27
07/29
07/31
0
320
26O
4)
2n
Figure 3.6. Daily average groundwater discharge vs. time for selected seepage meters in
the 1999 cluster experiment. Meter 11 is 12 m from shore. Bars represent ± one standard
deviation of the measurements taken on each day.
In order to capture small-scale spatial heterogeneity in discharge salinity, a grid 50 cm
long by 50 cm wide was constructed (see Appendix A) with a spacing of 5 cm. The grid
was placed in two locations, 2.3 m and 15.5 m into the bay from Well 1, along Transect
W (see Figure 3.1 for locations). Salinity was sampled with a syringe as described in
Section 3.4.1.2 at a depth of 3 cm in each grid box, and the results are shown in Figure
3.7. The porewater salinity in the nearshore grid varied from 0-10 ppt and the grid farther
from shore varied from 25-29 ppt, with a baywater salinity of approximately 26 ppt. The
15.5 m location exhibited very high salinities and low variability, so only 35 cm x 50 cm
of the grid was used; freshwater fingers were not encountered in this area. The porewater
at the 2.3 m location was consistently fresh, in agreement with the high freshwater
48
disc-
ohmed there. The salinity variation that was abed was most likely due to
sampling e m (small samples are easily contaminated with baywater) or due to a small
amount of mixing at the sediment-water interface. If the variability were attributable to
actual saltwater fingers, higher salinities would be expected, partlcularly within on1y
centimeten of the saltwater mume. The salinity pattan in Waquoit Bay thus exhibits low
variability on both meter and centimeter scales, and density fingering bas not been
observed
L
23 .mfrom Shore
Figure 3.7. Near-surface porewater salinity measurements taken at 5 c m intervals on a
0.5 m grid in the nearsha and middle disczones.
35.2 Temporal Variabii
The 1999 cluster experiment (Figure 3.6) indicates that discharge varies significantly
from day to day as well as during a tidal cycle. However, the relative temporal discharge
is constant: an area which discharges more than a neighboring area does so steadily, even
as the total discharge increases or decreases. The 2001 duster experiment shows that
temporal variation with the tide changes with distance into the bay (Figure 3.8). The
nearshore (7-16 m from shore) discharge exhibits a clear inverse variation with the tide.
The largest factor of change in this zone occurs during periods of greatest tidal range, and
every seepage meter exhibited a similar inverse variation with the tide. The discharge in
the middle (30-45 m from shore) and far (50-62 m from shore) zones does not vary as
consistently with the tides, although the middle area may exhibit variation due to a
longer-scale driving force such as the spring-neap tidal cycle. Discharge has been shown
to correlate with tidal magnitudes over long time scales (Taniguchi 2002), an observation
that is consistent with both the increased discharge in the middle zone as the tidal
magnitude increased, and the larger discharge in the 2000 head of the bay experiment
relative to 1999.
m
60
E
Q
40 o
20
.t:
ci
W
h-
O
;
W
C
hg
0
60 r
k,
M,
co
A..
W
.40 nte
60
20
06:00
12:00
18:00
00:00
06:00
12:()
18:00
Time
Figure 3.8. Spatially-averaged discharge and tidal elevation vs. time for the 2001
multiple tidal-cycle experiment.
50
PM
.
B
3.6 Radium Isotope Measurements
3.6.1 The Use of Radium as a Tracer
Radium is produced from uranium and thorium, which occur naturally in rocks. As water
moves through an aquifer matrix, it accumulates radium as it is produced, desorbed, or
dissolved, becoming enriched in radium relative to surface and atmospheric water. The
four radium isotopes have identical chemistry but very different half-lives: 1600 yr, 5.75
yr, 11.435 d, and 3.66 d for
226Ra,
228Ra,223Ra,and
22 4Ra,
respectively, which makes
radium potentially useful as a groundwater tracer. Radium has been used extensively by
oceanographers and hydrologists to estimate the amount of SGD to coastal waters (Miller
et al. 1990; Moore 1996; Moore and Shaw 1998; Krest et al. 2000; Kelly and Moran
2002). This has been done by measuring 226Raactivities (disintegration rates) in estuarine
or ocean water, estimating the 226 Ra flux necessary to maintain the distribution
considering the flushing time of the water body, and separating the groundwater
contribution from other sources such as river water, desorption from riverine sediments,
and erosion. Once the groundwater
2 2 6Ra
flux ( 22 6Raex [i.e. dpm/d]) is calculated, an
endmember activity (226 Racw [i.e. dpm/L]) is determined, generally from an average of
well measurements. SGD [d]
is then calculated from the relation (Charette et al. 2001):
SGD
226Raex
226
RaGW
(3.1)
One difficulty in using radium as a tracer in coastal systems, however, is that Ra sorbs
strongly to soil particles, but the distribution coefficient (ion concentration per gram of
sediment / concentration per gram of liquid) differs greatly depending on the porewater
salinity. For example, desorption experiments performed by Li and Chan (1979) give a
distribution coefficient (concentration of exchangeable ion per gram of
sediment/concentration of exchangeable ion per gram of supernate) of 235±20 in
saltwater and 21,000 in freshwater, a difference of nearly two orders of magnitude.
Therefore, it is difficult to estimate one representative concentration in groundwater, or
endmember, for the entire system. Radium activities in fresh and saline groundwater may
differ greatly, and mixing, particularly flow of saline water into a previously fresh region,
51
may result in significant desorption and very high activities in groundwater of
intermediate salinity. This means that in many cases, radium measurements reflect only
the mixed component of SGD, which may be a small fraction of the total flow
(Oberdorfer 2003). Furthermore, if the assumed endmember concentration is calculated
based on measurements in fresh groundwater only, as is often the case, the amount of
total discharge inferred from radium measurements may be much greater than the actual
discharge. It is therefore important to understand the small-scale variability in a coastal
system when applying a tracer method on a large spatial scale.
3.6.2. Radium Measurements
Radium was used as a tracer to estimate SGD into Waquoit Bay by Matthew Charette et.
al. (2001) according to the method described above. In conjunction with that study,
radium measurements were taken during the seepage meter experiments at the head of the
bay in July 2000 (co-investigators M.A. Charette, K.O.Buesseler, and C.F. Harvey) in
order to examine the distribution of the radium isotopes on a smaller scale.
The activity of radium in most groundwater, both fresh and saline, is very low in relation
to the detection limit, so a relatively large volume of water is necessary to obtain an
accurate radium measurement. It was therefore necessary to combine water from seepage
meter bags in order to measure Ra activity in the groundwater discharge. The seepage
meters were separated into four areas based on the relative amount of discharge in each:
row 1 (7 m from shore), area 1 (8 to 17 m from shore), area 2 (20 to 40 m from shore),
and area 3 (45 to 75 m from shore). At each sample time, the bags from each area were
emptied into large polypropylene barrels. The radium was extracted from each sample by
filtering the water through a column of Mn-impregnated fibers. Because row 1 only
consisted of four seepage meters at each sample time rather than 12, as in the other areas,
samples 1 and 2, 3 and 4, and 5 and 6 were combined for a total of 3 measurements rather
than 6 for each of the other areas. The activity of each of the four radium isotopes in each
sample was measured at the Woods Hole Oceanographic Institution as described in
Charette et. al. (2001). The total volume of water was calculated by subtracting the initial
volume in the seepage meter bags from the total volume of flow through the fiber
52
_II
column, thus accounting for the 1 L of pre-filled water in each bag. The seepage meter
bags were pre-filled with 1L of baywater that had been stripped of radium by pumping it
through a Mn-fiber column before filling. The reported activities are therefore accounting
only for radium in the water discharging into the seepage meters.
Samples were taken in wells and piezometers during the summer of 2000 and analyzed
for radium activity. Two sets of data from shallow piezometers in the intertidal zone were
obtained, as well as data from a multi-level sampling well (CCC well, Figure 3.1).
3.6.3. Heterogeneity in Porewater Radium Activity
The excess radium activities of the discharging groundwater (activity of the discharge
minus the average activity of the overlying baywater) are averaged over the six sample
times for the 2000 seepage meter experiment and plotted in Figure 3.9. The radium
content of the row 1 discharge is clearly much greater than the discharge in any other
area. Row 1 was also the only location to exhibit a freshwater component of flow, with an
average salinity of 20.9 ppt, compared to 28.6, 30. 1, and 29.5 ppt (approximately
baywater salinity) in areas 1, 2, and 3, respectively. The porewater 226Raactivities are
plotted against the sample salinity in Figure 3.10. The activity measured in fresh water is
consistently low, with the highest values measured in porewater of intermediate salinity.
This is consistent with measurements in other studies (Moore and Scott 1986; Webster et
al. 1995), and illustrates the large effect that porewater ionic strength has on radium
sorption. This makes it difficult, however, to use one endmember value of 226 Ra activity
to estimate total SGD. The seepage meter data shows that there are three components of
groundwater flow, all of which are significant: fresh, brackish, and saline. The
piezometer and well measurements indicate that radium activity is highly variable in each
of these components. Direct measurement of radium activity in the discharging water
supports this variability. Combining the water flux and activity measurements (Figure
3.11) further illustrates that maximum fluid discharge does not coincide with maximum
radium contribution. It is therefore difficult to accurately estimate total SGD using only
one endmember value in systems where the components of discharge differ in tracer
concentration.
53
It is possible, however, to gain information about subsurface flow through analysis of
radium measurements. The activity ratio (AR) of
to 2 2 6 ~has
a been shown to be
useful in distinguishing groundwater sources to an estuary (Charette and Buesseler 2004).
Ratios of the measured activities of the two short-lived isotopes to
(223~alZZ6Ra
and
224Ra/226~a)
can also give information about the time the water spent in contact with the
aquifer matrix. As recharging water containing little radium comes in contact with
sediment, it wilI pick up the radium as it is produced in proportion to the inverse of the
half-life of the isotope. So very young water will have a high 223~a/2ZbRa
or 2 2 4 ~ a / 2 2 6 ~ a
AR, and water from a longer flowpath will be closer to secular equilibrium, where the
AR of all parents to daughters is 1 . The ratio of 226Ra:228~a:"3~a:Z%
should be
approximately 1: 1 :0.05:1 based on the likely relative abundance of thorium parents in
aquifer solids, (Rarna and Moore 1996).
Figure 3.9. Excess radium activity vs. distance from shore for all four radium isotopes.
The left axis scale is for 226Ra,2 2 8 ~ aand
, 2 2 3 ~ the
a ; right axis scale is for 2 2 4 ~ a .
npore 3.10. 226Raactivity YS. 881inity for water mnplt8 coile~tainear Waquoit ~ a y .
Deep porewater mnplai fmm CCC-web (see F i p e 3,l for l d m } , shallow
~ ~ ~ h m ~ ~ b m E Z W L l p i e ; ~ ~ ~ d i ~ g e f r o t n
baywater samph are included.
Figme 3.11. Spatial &rribution with distance from shore, averaged for each area (row 1,
area 1:ncarshotz,ana2:middle,pndaree3:fsrfranshore)for(a)226Raanivityof
watea C O U E Gin
~ seepage
~
metws, (b) groundwater flux m a d in seepage meters, and
(c) tale 2asRa flux.
The SGD activity ratios are plotted vs. distance in Figure 3.12. All of the ratios, including
baywater, are greater than the expected value based on abundance, demonstrating that the
water is not in secular equilibrium. The 2 2 8 ~ a / 2 2ARs
6 ~ a are very close to the baywater
AR, indicating that if there is another significant source of water to the bay it does not
have a substantially different 2 2 8 ~ a / Z 2ratio.
6 ~ a The ratios of short-lived to long-lived
isotopes decrease with distance, which may indicate that the brackish water has spent less
time in contact with the aquifer matrix than the saline water.
Distance Imj
Figure 3.12. Short-lived to long-lived ( 2 2 6 ~ aactivity
)
ratio vs. distance from shore. The
left axis scale is for Z 2 8 ~ a / 2 and
2 b ~' 'a' ~ a / ~ ~ k right
a ; axis scale is for ' " ~ a i ~ ~ ~ ~ a .
Average baywater activity ratios are dashed lines,
3.7 Discussion
3.7.1 How Many Meters are Necessary to Estimate Large-Scale Discharge?
The large variability in discharge over both time and space raises the question of how
many seepage meters are necessary to adequately estimate the discharge. The grid
spacing in the 1999 head of the bay experiment i s assumed dense enough to accurately
estimate the total discharge. Based on this assumption, we can determine whether a more
sparse spacing will give an adequate estimate. The estimated total groundwater flux from
I000 replicates of random selections of seepage meters in the 1999 head of the bay
experiment was estimated using MATLAB. These selections gtve an average absolute
difference in flux using 10,20, and 30 seepage meters, as compared to all 40 meters, of
348,198 and 1096, respectively. However, if seepage meters
m g e d in transects,
the error is much lower. The flux estimated using one, two, and three tmnsats of 10
seepage meters differs from the flux &mated using all four bansects by 9%, 4%, and
3% respectively (Figure 3.13). Thus, on a 50 m scale, 20 saepqp meters m g e d in
trslnsects appear adequate to estimate total discharge at our site within a reasonable mott
The seepage meter grid in the 2000 head of the bay experiment was more than twice as
large as the grid in the 1999 study as a result of this calcdation.
30 Random
20 Random
10 Random
30
Percent Error
20
Figure 3.13. Average percent cmr in total discharge for subsets of seepage meters from
the 2000 head of the bay experiment. Error bars represent one standard deviation from
the m a ,
3.7.2
Discharge Comparison with Freshwater Balance
Cambareti and Eichner (I W8)estimated the freshwater input to Waquoit Bay from the
head of the bay subwatershed to be 0.012 m3/susing a hydrologic balance based on a
year1y average precipitation of 92 cm.A conservative extraplation of our 1999 and 2000
seepage meter data along the 610 m head of the bay results in total discharge estimates of
0.047 and 0.106 m3/s,respectively. These total discharge estimates an much greater than
the freshwater estimate, indicating that there is significantly more saline than fresh
discharge. Other studies that report the salinity of submarine groundwater discharge have
observed a large amount of saline discharge both in the field (Robinson et al. 1998; Kim
et al. 2003) and with numerical models (Langevin 2003).
The freshwater discharge rate estimates from the 2000 experiment using the upper and
lower bounds for recharging salinity are 0.011 and 0.004 m3 /s, respectively. The values
from 1999 are 0.006 and 0.001 m3 /s. These values are lower than that of the freshwater
balance, but consistent given the likelihood that there is significant freshwater discharge
nearer to shore than the shallowest seepage meter.
The observation that total discharge was a factor of 2.3 larger in the 2000 experiment
than the 1999 experiment may be explained by either the larger tidal magnitude in 2000
or the greater precipitation preceding the 2000 experiment. The tidal range during the
2000 experiment was 40 cm (-1 day prior to spring tide), twice that of the 1999
experiment (midway between spring and neap tide). During the three months and 1 year
prior to the July 2000 experiment 28.7 cm and 129.3 cm (Payne 2004), respectively, of
precipitation fell at Long Pond in Falmouth, MA (-7 km west of Waquoit Bay). This is
significantly more than the 18.9 cm and 93.1 cm that fell during the three months and 1
year prior to the 1999 study.
The 2003 single-transect experiment gives the only direct measurements of nearshore
freshwater discharge. Averaged over a tidal cycle and extrapolated along the 610 m head
of the bay, fresh discharge is 0.025 m3 /s and saline discharge is 0.053 m 3 /s. The
freshwater measurement is greater than the water balance estimate (Cambareri and
Eichner 1998), but consistent because the measurements were taken under a topographic
high, and at only one point in time, whereas the water balance is for an entire average
year over the full head of the bay.
3.7.3
Large-Scale Pattern of Discharge
Despite the small-scale variability, a band of high discharge following the shoreline is
clearly evident. This is an unexpected result since theoretically, discharge should
58
decrease monotonically, approximately exponentially, with distance from the shoreline
(Bokuniewicz 1992). Although the banded pattern has not been previously reported in a
coastal system, there are studies which provide evidence of higher discharge offshore
(Simmons 1992; Cable et al. 1997a; Smith and Zawadzki 2003). It is possible that
diagenetic changes or recent sedimentation on the bottom could create a high
permeability pattern aligned with the shoreline, but slug tests indicate that no such pattern
exists. The uniform discharge observed in the Washburn Island study suggests that
density-dependent flow may play a role in creating the banded discharge pattern.
Washburn Island and the head of the Bay have similar sediments and tides, but the
narrow spit lacks significant freshwater recharge.
Although our results provide a detailed characterization of discharge patterns, they do not
show where the discharging saline water originates. A significant number of seepage
meters indicated inflow, but they do not account for the large net outflow of saline water.
One source of saline water is recharge during a rising tide (Nielsen 1990) that overtops
discharging freshwater and results in an inverse density gradient. A second potential
source is offshore circulation of seawater, which was predicted theoretically by Henry
(1959) and observed along the coast of Florida (Kohout 1960). The theoretical model
relies upon transverse dispersion to mix saline and fresh water at the interface so that
saline water discharges with the freshwater. Such models predict that saltwater flows into
the subsurface far from the shore and circulates toward shore before discharging. Most of
our measured inflow occurred in the zone far from shore, which is consistent with the
possibility of downwelling further offshore, circulation in the subsurface, and discharge
in the middle zone. However, this large amount of inflow could not be confirmed with
seepage meters because the mucky nature of the bay floor beyond 75 m from shore
prevents stable placement of seepage meters. The question remains whether these
mechanisms explain the amount of saltwater discharge and its pattern.
Traditional models of submarine groundwater discharge based on simplifying
assumptions predict a monotonic decrease in primarily fresh discharge with distance from
shore. These simulations do not attempt to represent density-driven free convection in
59
which instabilities, closely related to small-scale heterogeneity, may affect larger-scale
flow (Simmons et al. 2002). Recently published two-dimensional numerical models
(Robinson et al. 1998; Ataie-Ashtiani et al. 1999) show that incorporating tidal dynamics
significantly affects salinity distributions and groundwater flow patterns. The data
presented here demonstrate both tidal effects and high small-scale variability in flow,
raising concerns about models that neglect these factors.
3.8 Summary
This study gives specific information about groundwater flow into Waquoit Bay and also
provides insight into groundwater dynamics in sandy coastal aquifers and the methods
used to investigate discharge. A banded pattern of mostly saline groundwater discharge at
the head of Waquoit Bay suggests that flow follows more complex patterns at the coast
than models have predicted: the interaction between fresh and saltwater, driven by both
tides and freshwater discharge, may create unanticipated circulation cells and flowpaths.
This large-scale pattern is only evident when a sufficient density of measurements is used
to overcome the small-scale variability. The large differences in flow observed over small
spatial scales raise questions about the application of models that assume homogeneity.
In addition, the large proportion of saline discharge may have implications for the use of
geochemical tracers to estimate total submarine groundwater discharge if concentrations
in fresh and saline water differ. This is illustrated by radium isotope measurements in
porewater and submarine groundwater discharge that vary widely with salinity and
location. Lastly, despite efforts to measure inflow of saline water in this study, a fluid
mass balance was not achieved, and questions remain about the flow pattern of the large
amount of saline discharge observed in Waquoit Bay.
60
References
Abraham, D. M., M. A. Charette, M. C. Allen, A. Rago, and K. D. Kroeger (2003)
Radiochemical estimates of submarine groundwater discharge to Waquoit Bay,
Massachusetts. Biological Bulletin 205(2): 246-247.
Ataie-Ashtiani, B., R. E. Volker, and D. A. Lockington (1999) Tidal effects on sea water
intrusion in unconfined aquifers. Journal of Hydrology 216(1-2): 17-31.
Barlow, P. M., and K. M. Hess (1993) Simulated hydrologic responses of the Quashnet
River stream-aquifer system to proposed ground-water withdrawals. U.S.
Geological Survey Water-Resources Investigations Report 93-4074: 52 p.
Belanger, T. V., and M. T. Montgomery (1992) Seepage Meter Errors. Limnology and
Oceanography 37(8): 1787-1795.
Belaval, M. (2003) A geophysical investigation of the subsurface salt/fresh water
interface structure, Waquoit Bay, Cape Cod, Massachusetts. Master of Science
Thesis, Boston College. Department of Geology and Geophysics. Chestnut Hill,
Boston College: 78p.
Bokuniewicz, H. (1980) Groundwater Seepage into Great South Bay, New-York.
Estuarineand CoastalMarine Science10(4):437-444.
Bokuniewicz, H. J. (1992) Analytical Descriptions of Subaqueous Groundwater Seepage.
Estuaries 15(4): 458-464.
Bowen, J. L., and I. Valiela (2001) The ecological effects of urbanization of coastal
watersheds: historical increases in nitrogen loads and eutrophication of Waquoit
Bay estuaries.CanadianJournal of Fisheriesand Aquatic Sciences58(8): 14891500.
Burnett, W., J. Chanton, J. Christoff, E. Kontar, S. Krupa, M. lambert, W. Moore, D.
O'Rourke, R. Paulsen, C. Smith, L. Smith, and M. Taniguchi (2002) Assessing
methodologies for measuring groundwater discharge to the ocean. EOS 83(11):
117-123.
Burnett, W. C., H. Bokuniewicz, M. Huettel, W. S. Moore, and M. Taniguchi (2003)
Groundwater and pore water inputs to the coastal zone. Biogeochemistry 66(1-2):
3-33.
61
Cable, J. E., W. C. Burnett, and J. P. Chanton (1997a) Magnitude and variations of
groundwater seepage along a Florida marine shoreline. Biogeochemistry 38(2):
189-205.
Cambareri, T. C., and E. M. Eichner (1998) Watershed delineation and ground water
discharge to a coastal embayment. Ground Water 36(4): 626-634.
Capone, D. G., and M. F. Bautista (1985) A Groundwater Source of Nitrate in Nearshore
Marine-Sediments. Nature 313(5999): 214-216.
Charette, M. A., and K. O. Buesseler (2004) Submarine groundwater discharge of
nutrients and copper to an urban subestuary of Chesapeake Bay (Elizabeth River).
Limnology and Oceanography in press.
Charette, M. A., K. O. Buesseler, and J. E. Andrews (2001) Utility of radium isotopes for
evaluating the input and transport of groundwater-derived nitrogen to a Cape Cod
estuary. Limnology and Oceanography 46(2): 465-470.
Charette, M. A., and E. R. Sholkovitz (2002) Oxidative precipitation of groundwaterderived ferrous iron in the subterranean estuary of a coastal bay. Geophysical
Research Letters 29(10): 10.1029/2001GL014512.
Geyer, W. R. (1997) Influence of wind on dynamics and flushing of shallow estuaries.
Estuarine Coastaland ShelfScience44(6):713-722.
Henry, H. R. (1959) Salt intrusion into fresh-water aquifers. Journal of Geophysical
Research 64(11): 1911-1919.
Hoefel, F. G., and R. L. Evans (2001) Impact of low salinity porewater on seafloor
electromagnetic data: A means of detecting submarine groundwater discharge?
EstuarineCoastaland Shelf Science52(2): 179-189.
Kelly, R. P., and S. B. Moran (2002) Seasonal changes in groundwater input to a well-
mixed estuary estimated using radium isotopes and implications for coastal
nutrient budgets. Limnology and Oceanography 47(6): 1796-1807.
Kim, G., K. K. Lee, K. S. Park, D. W. Hwang, and H. S. Yang (2003) Large submarine
groundwater discharge (SGD) from a volcanic island. Geophysical Research
Letters 30(21): 10.1029/2003GL018378.
Kohout, F. (1960) Cyclic Flow of Salt Water in the Biscayne Aquifer of Southeastern
Florida. Journal of Geophysical Research 65(7): 2133-2141.
Krest, J. M., W. S. Moore, L. R. Gardner, and J. T. Morris (2000) Marsh nutrient export
supplied by groundwater discharge: Evidence from radium measurements. Global
Biogeochemical Cycles 14(1): 167-176.
62
Langevin, C. D. (2003) Simulation of submarine ground water discharge to a marine
estuary: Biscayne Bay, Florida. Ground Water 41(6): 758-771.
Lee, D. R. (1977) Device for Measuring Seepage Flux in Lakes and Estuaries. Limnology
and Oceanography 22(1): 140-147.
Li, Y.-H., and L.-H. Chan (1979) Desorption of Ba and 226 Ra from river-borne sediments
in the Hudson Estuary. Earth and Planetary Science Letters 43: 343-350.
Masterson, J. P., D. A. Walter, and J. Savoie (1997) Use of particle tracking to improve
numerical model calibration and to analyze ground-water flow and contaminant
migration, Massachusetts Military Reservation, western Cape Cod,
Massachusetts. U.S. Geological Survey Water-Supply Paper 2482: 50 p.
Miller, R. L., T. F. Kraemer, and B. F. Mcpherson (1990) Radium and Radon in Charlotte
Harbor Estuary, Florida. Estuarine Coastal and Shelf Science 31(4): 439-457.
Moore, D. G., and M. R. Scott (1986) Behavior of 226Rain the Mississippi River mixing
zone. Journal of GeophysicalResearch91(12): 14317-14329.
Moore, W. S. (1996) Large groundwater inputs to coastal waters revealed by Ra-226
enrichments. Nature 380(6575): 612-614.
Moore, W. S., and T. J. Shaw (1998) Chemical signals from submarine fluid advection
onto the continental shelf. Journal of Geophysical Research-Oceans 103(C10):
21543-21552.
Nielsen, P. (1990) Tidal dynamics of the water table in beaches. Water Resources
Research 26(9): 2127-2134.
Oberdorfer, J. A. (2003) Hydrogeologic modeling of submarine groundwater discharge:
comparison to other quantitative methods. Biogeochemistry 66(1-2): 159-169.
Payne, R. (2004) Falmouth Monthly Climate Reports, Falmouth Water Department,
www.whoi.edu/climate/, Woods Hole Oceanographic Institution. 2004.
Portnoy, J. W., B. L. Nowicki, C. T. Roman, and D. W. Urish (1998) The discharge of
nitrate-contaminated groundwater from developed shoreline to marsh-fringed
estuary. Water Resources Research 34(11): 3095-3104.
Rama, and W. S. Moore (1996) Using the radium quartet for evaluating groundwater
input and water exchange in salt marshes. Geochimica Et Cosmochimica Acta
60(23): 4645-4652.
Robinson, M., D. Gallagher, and W. Reay (1998) Field observations of tidal and seasonal
variations in ground water discharge to tidal estuarine surface water. Ground
Water Monitoring and Remediation 18(1): 83-92.
63
Shaw, R. D., and E. E. Prepas (1989) Anomalous, Short-Term Influx of Water into
Seepage Meters. Limnology and Oceanography 34(7): 1343-1351.
Shinn, E. A., C. D. Reich, and T. D. Hickey (2002) Seepage meters and Bernoulli's
revenge. Estuaries 25(1): 126-132.
Simmons, C. T., M. L. Pierini, and J. L. Hutson (2002) Laboratory investigation of
variable-density flow and solute transport in unsaturated-saturated porous media.
Transport in Porous Media 47(2): 215-244.
Simmons, G. M. (1992) Importance of Submarine Groundwater Discharge (Sgwd) and
Seawater Cycling to Material Flux across Sediment Water Interfaces in Marine
Environments. Marine Ecology-Progress Series 84(2): 173-184.
Smith, L., and W. Zawadzki (2003) A hydrogeologic model of submarine groundwater
discharge: Florida intercomparison experiment. Biogeochemistry 66(1-2): 95-110.
Talbot, J. M., K. D. Kroeger, A. Rago, M. C. Allen, and M. A. Charette (2003) Nitrogen
flux and speciation through the subterranean estuary of Waquoit Bay,
Massachusetts. Biological Bulletin 205(2): 244-245.
Taniguchi, M. (2002) Tidal effects on submarine groundwater discharge into the ocean.
Geophysical Research Letters 29(12): 10.1029/2002GL014987.
Testa, J. M., M. A. Charette, et al. (2002) Dissolved iron cycling in the subterranean
estuary of a coastal bay: Waquoit Bay, Massachusetts. Biological Bulletin 203(2):
255-256.
Valiela, I., J. L. Bowen, and K. D. Kroeger (2002) Assessment of models for estimation
of land-derived nitrogen loads to shallow estuaries. Applied Geochemistry 17(7):
935-953.
Valiela, I., J. Costa, et al. (1990) Transport of Groundwater-Borne Nutrients from
Watersheds and Their Effects on Coastal Waters. Biogeochemistry 10(3): 177197.
Valiela, I., K. Foreman, et al. (1992) Couplings of Watersheds and Coastal Waters Sources and Consequences of Nutrient Enrichment in Waquoit Bay,
Massachusetts. Estuaries 15(4): 443-457.
Valiela, I., M. Geist, J. McClelland, and G. Tomasky (2000) Nitrogen loading from
watersheds to estuaries: Verification of the Waquoit Bay Nitrogen Loading
Model. Biogeochemistry 49(3): 277-293.
Webster, I. T., G. J. Hancock, and A. S. Murray (1995) Modelling the effect of salinity
on radium desorption from sediments. Geochimica Et Cosmochimica Acta 59(12):
2469-2476.
64
-
~
~
__
Whiting, G. J., and K. L. Childers (1989) Subtidal advective water flux as a potentially
important nutrient input to southeastern U.S.A. saltmarsh estuaries. Estuarine,
Coastal, and Shelf Science 28: 417-431.
65
66
-
-
Chapter Four
Circulation of Saline Groundwater
Fresh groundwater flows into coastal waters throughout the year because the upland
hydraulic head is above mean sea level due to recharge from precipitation at the land
surface. Saline groundwater, however, circulates through the subsurface: it is only
recharged to the aquifer from the water body into which it discharges. Thus, a salt mass
balance must be maintained across the sea floor. The measurements of submarine
groundwater flow in Waquoit Bay presented in Chapter Three do not exhibit this saline
mass balance. Far more saline groundwater was observed to discharge into the bay than
flowed into the aquifer from the bay. In this chapter, saline groundwater circulation
mechanisms are considered for two reasons. First, saline water forced into and out of
aquifers affects groundwater flow patterns and has the potential to transport chemicals
from the subsurface into coastal waters, so consideration of the amount of circulation
warrants attention. Secondly, the amount of saline circulation is discussed in an effort to
explain the high saline outflow observed in Waquoit Bay.
4.1 Circulation Mechanisms
Mechanisms that induce fluid motion in the ocean floor have been studied extensively in
relation to sediment transport, ocean biochemistry, and saltwater-freshwater mixing near
the coast. Saline fluid flow can result from density differences due to temperature and
concentration gradients, tides, waves, and dispersion at the freshwater-saltwater interface.
Each of these mechanisms is considered here as a potential means to create the necessary
recharge flux to balance the large amount of saline discharge observed in Waquoit Bay
67
(Chapter Three) and in other locations along the coast (Moore 1996; Moore and Church
1996; Robinson et al. 1998; Taniguchi 2002; Kim et al. 2003; Smith and Zawadzki 2003;
Taniguchi et al. 2003), and as a factor affecting flowpaths.
Density variations due to temperature and salinity gradients in the subsurface may cause
circulation of saline groundwater. Density-driven free convection can result from
unstable stratification of fluids of different density. Although geothermal activity may
cause such groundwater flow, this effect is rare and not a factor on Cape Cod. This
mechanism could occur in Waquoit Bay due to salinity gradients, where saltwater
overtops discharging freshwater at high tide, and potentially where a confined fresh
aquifer underlies a low-permeability layer, separating it from the surficial saline aquifer.
Such instability drives fluid flow that may result in the formation of density fingers, or
lobes of fresh and saline water flowing opposite each other toward a stable configuration.
This convection promotes mixing on a larger spatial scale and shorter time period than
diffusion (Simmons et al. 2001). Density fingering has not been observed in Waquoit
Bay, however, on a meter or centimeter scale (see Section 3.5.1.3), and the overall effect
does not produce a net upward flow of saltwater that would explain the observations of
net saline outflow.
Pressure changes due to tides, waves, and atmospheric pressure fluctuation may induce
flow in and out of the sea floor in one dimension to a degree that depends on the specific
storage and hydraulic conductivity of the aquifer (Figure 4.1, process 1). This mechanism
creates zero net flow over a tidal cycle and cannot explain net saline groundwater
discharge, but could be potentially important to near-surface chemical processes. The
elastic properties of most natural aquifers preclude significant flow in and out of elastic
storage, so this effect occurs near enough to shore that the water table can respond over
the timescale of tides.
There is, however, a two-dimensional effect produced by tides and waves at the coast that
results in net inflow at the beach and net discharge seaward (Figure 4.1, process 2). A
similar circulation, but in the opposite direction, occurs due to dispersion-induced flow
68
along the freshwater-saltwater interface, creating net saline discharge near the coast and
net inflow offshore (Figure 4.1, process 3). None of the mechanisms discussed here
create a net outflow of saltwater averaged spatially and over a tidal cycle, but it is
possible that field techniques have preferentially measured flow in the discharge rather
than recharge areas. Also, these mechanisms have been shown to drive significant
groundwater flow, greatly affecting the amount of saline circulation. A discussion of
these spatially-varying circulation mechanisms is therefore appropriate. h the following
sections,the two-dimensional saline circulation due to tides and waves at the shoreline
and dispersion within the aquifer are discussed in detail, including analytical estimates of
circulation volume. Estimates of the amount of circulation in Waquoit Bay are presented
in Section 4.2, along with a discussion that relates the estimates to direct observations of
submarine groundwater discharge presented in Chapter Three.
Figure 4.1. Saline flow patterns induced by circulation mechanisms.1 - Onedimensional inflow and outflow due to mechanisms such as density fingering and tidal
pumping. 2 - Nearshore circulation due to tides and waves. 3 - Dispersion-induced saline
circulation.
4.1.1 Tides
Tidal rise and fall induces changes in porewater pressure at the beach face and landward.
This action drives movement of the freshwater-saltwater interface near the coast, and can
lead to enhanced mixing and a wider transition zone than would occur in a steady-state
system (Inouchi et al. 1990).
The governing equations given by Nielsen (1990) for shore normal groundwater flow in
an isotropic, homogeneous aquifer with a long, straight beach that is bounded at depth
(D) below mean sea level by an impermeable layer can be written as the Boussinesq
equation under the Dupuit assumption,
h
l
a (hah
a--th~axa ((4.1)
Neglecting waves, the boundary conditions give tidal elevation along the beach slope
with angle ,
h([hide- D] cot f, t) = hid, .
(4.2)
and require that oscillations die out far inland from the beach,
ah
-- >0 ,x -- oo
at
(4.3), (4.4)
These equations can be solved analytically for particular beach geometries and tidal
boundary conditions, but the solutions are limited in representing the actual system. The
formulation fails to explain the temporal skewness in water table fluctuations and the
superelevation of the water table above mean sea level observed at the coast (Turner et al.
1997). These phenomena are due to three mechanisms: the formation of a seepage face
due to decoupling between sea level and the beach water table near low tide, asymmetry
of the boundary condition at a sloping beach face, meaning the beach fills more easily
70
than it drains, and the nonlinearity of the governing equation for horizontal flow below
the water table (Nielsen 1990).
Superelevation of the beach water table due to tides results in saltwater circulation from
the point of infiltration at the coast to a discharge point seaward. Nielsen (1990) presents
an equation for water table elevation in a sloping beach with small tidal amplitude,
neglecting the formation of a seepage face. Integration of this equation by Li et. al.
(1999) leads to an analytical approximation of the discharge associated with this process,
D, =
exp(-2a) cos(x2a) + A'
exp(-a)[cos(a) - sin(a)]+
k~;
SbTt
(4.5)
SbT,
with
r=eo' and a =-,
2KH
Sb
(4.6),(4.7)
where D, is the unit alongshore discharge rate, A, T, and co are the tidal amplitude,
period, and frequency, i1 e is the effective porosity, H is the averaged aquifer thickness,
and Sb is the beach slope.
4.1.2 Waves
As waves near the beach, they shoal and break, leading to changes in the mean water
level known as set-up and set-down (see Figure 4.2) (Raubenheimer et al. 2001). In
general, set-up at the shoreline is approximately 40% of the root-mean-square offshore
wave height above tidal elevation (Turner et al. 1997). Superimposed on the elevated
mean water level due to set-up is the run-up of waves on the beach face, which
potentially increases the area over which infiltration of seawater can occur. Significant
infiltration will only occur, however, when the mean water level is higher than the water
table level within the beach face, which only takes place during high tide and on
particular types of beaches (Turner et al. 1997). Despite this tidal regulation, the overall
effect of wave set-up and run-up is to enhance the superelevation of the water table and
the associated circulation of seawater within the beach face (Li and Barry 2000). Waveinduced pulse forcing due to storm events of significant magnitude and duration has also
71
been shown to significantly affect the position of the freshwater-saltwater interface near
the shore (Cartwright et al. 2004).
A further complication to the tide and wave-driven beach water table dynamics is the
effect of capillarity, which leads to potentially large pressure gradient changes and water
table fluctuations. This effect is negligible for low-frequency oscillations, such as tides,
where mass transport dominates, but is the principal effect for high-frequency wave runup (Li et al. 1997). Analytical solutions cannot currently capture capillarity, but it has
been modeled numerically (Li et al. 1997; Turner and Masselink 1998; Li and Barry
2000; Nielsen and Perrochet 2000; Werner and Lockington 2003). Observations of the
effects of waves in the swash zone and capillary fringe have shown that although there is
a large pressure jump as water infiltrates, which was previously thought to result in a
large upward flow of water, the effect is actually a minute downward infiltration (Turner
and Nielsen 1997; Turner 1998).
An analytical solution for the groundwater circulation due to wave run-up that neglects
the small effects of capillarity is given by Li et. al. (1999):
32
or(sb
-S)
DW = KswL,
with
sW= 3- S
L=
8 +
and
v=
1
1.56 9
Hb
(4.9), (4.10)
- 43.8[1-exp(-19sb)]-Hb
6
l+exp(-19.5s
(4.8)
gTJ2
)
'
(4.11)
where Dwis the discharge rate per unit alongshore distance, K is hydraulic conductivity,
Swis the slope of the wave set-up, L is the distance between the breaker and run-up lines,
a is the breaking index, Hb is the breaking wave height,
Sb is
the beach slope, g is the
magnitude of gravity acceleration, and Twis the wave period.
72
superelevation due to tides
igh Tide Level
lean Tide Level
Figure 4.2. Schematic of nearshore saline circulation due to tides and waves and selected
parameters from Li et. al. (1999) equations.
4.1.3 Dispersion
Sharp-interface models of coastal groundwater systems assume that freshwater and
saltwater are immiscible, simplifying the problem to the coupled flow of two separate
fluids. While such models are capable of representing the general position, shape, and
movement of the interface (Essaid 1986; Larabi and De Smedt 1997; Dagan and Zeitoun
1998; Person et al. 1998; Kooi and Groen 2001), they neglect mixing that can greatly
influence the behavior of the system. Cooper (1959) was first to assert the theory that
diffusion at the freshwater-saltwater interface causes saltwater to circulate from the sea
floor to the zone of diffusion and back to the sea. He noted that while dispersion in
porous media is a result of convection due to velocity variations and molecular diffusion,
it is enhanced by the motion of the saltwater front due to tides and changes in the inland
water table elevation. This theory was supported and quantified by F.A. Kohout (1960)
through field observations, and his calculations suggest that saltwater may amount to
10% or more of the total seaward flow of water. An analytical estimate of the amount of
seawater entrained in freshwater flow to the sea is difficult to obtain, and numerical
modeling of this mechanism gives flow rates that are highly dependent on the value of
dispersivity used in the simulation (Smith 2004).
73
4.2 Saline Circulation in Waquoit Bay
4.2.1 Quantification of Saline Discharge Estimates due to Tides and Waves
It is possible to estimate saline discharge rates due to tides and waves in Waquoit Bay
using the analytical equations 4.5 and 4.8 derived by Li et. al. (1999). Parameter
estimation can be difficult, but upper and lower bounds as well as observed or inferred
values for Waquoit Bay are given in Table 1. The parameters are used to estimate the
potential saline circulation along the head of Waquoit Bay. The numbers given are for
total discharge, but it is important to note that if inflow is included, the net flow is zero.
Table 4.1. Parameters and calculated values of groundwater circulation due to tides and
waves in Waquoit Bay, MA.
Parameter
Minimum Value
Maximum Value
Observed/Estimated
Value*
Units
qle
0.2
0.5
.3
[]
A
0.1
0.8
0.43t
[m]
t
Tt
45000
45000
45000
Sb
0.1
0.04
0.07
H
7
15
[s]
[]1
9.1*
-2
[m]
§
K
2 x 10-5
3 x 10
1.5 x 10-4
[m/s]
Hb
0
0.1
0.0 1 t
[m]
Tw
0.2
5
0.25
[s]
Calculated Values
Dtv
DWY
8.9 x 104
0.64
0.013
[m3/s]
0
0.73
2.5 x 104
[m3/s]
0.014
[m3/s]
8.9 x 10- 4
1.37
Total D
* all values are estimated unless otherwise noted
t Observed on 8/14/2003
· Estimated from well log from shoreline well
BCalculated from slug tests using Hvorslev method (Domenico and Schwartz 1998)
v Calculated groundwater circulation from equations (4.5) and (4.8) for the 610m head of Waquoit Bay
The estimates of total saline discharge due to tides and waves in Waquoit Bay vary by
orders of magnitude depending on the parameter values assumed. Reasonable estimates
and measured parameter values, however, result in a discharge estimate of 0.014 m3/s
along the head of the bay. Measurements of flow and salinity from intertidal and
conventional seepage meters on August 14, 2003 give a total flow (extrapolated along the
74
-------
610 m head of the bay) of 0.078 m3/s, 32% of which is fresh. This flow can be divided
into two bands of discharge, one less and one greater than 15 m from the shoreline. In
order to compare the saline discharge estimates to measurements, it is important to know
where the circulation due to tides and waves is discharging. To do this, the subsurface
flowpath at the beach face is mapped through a sodium bromide tracer test.
4.2.2 Mapping Nearshore Saline Circulation Using Sodium Bromide.
The calculations above give an estimate of the amount of saltwater that infiltrates along
the beach face and discharges bayward. The discharge location, however, is unknown. It
is possible that the infiltrating baywater sinks into the beach and flows vertically
downward due to density differences through the underlying freshwater in a deep
circulation pathway, eventually flowing into the bay and contributing to the band of
saline discharge 25-45 m from shore. Alternatively, the circulation cell may be much
smaller, discharging within a few meters of the point of inflow. Li and Barry (2000) have
observed that circulation due to wave set-up extends to a depth below the beach face that
is comparable to the distance between the breaking point and maximum run-up. This
supports the idea of a small circulation cell in Waquoit Bay because the waves are
generally very small, so this distance to maximum run-up is minimal. It is possible,
however, that tides create a deeper circulation. To test these two hypotheses, a tracer test
was conducted in Waquoit Bay to observe the subsurface saline circulation.
The objective of this tracer test was to qualitatively track the subsurface motion of
baywater infiltrating at the beach face rather than to quantitatively balance inflow and
discharge of fluid and tracer. Thus piezometers were positioned to encompass the plume
(determine its extent), but the number of measurements within the plume was not always
sufficient to obtain the exact position of the concentration contours. Similarly, discharge
of sodium bromide (NaBr) into the bay was not measured, but inferred from the
porewater measurements. The method and results are described below.
On August 27, 2001, a sodium bromide solution was injected into the beach near the high
tide mark, during high tide. The 0.243 M injection solution was designed to have the
75
same density as seawater, 1.025 Kg/L (25 g NaBr + 1L deionized H20), so that it would
track the movement of the infiltrating baywater. The injection solution was diluted with
saline water pumped from both the bay and the subsurface to create 0.1 M, 0.01 M, 0.001
M, and 0.0001 M standards. A multi-meter was connected to a bromide electrode (Cole-
Parmer, Inc.) and the reading in mV translated into molarity through a curve that was
recalibrated at each sample time. Twenty-one 3/4-inchpiezometers were driven into the
sediment at varying depths and distance from shore, as depicted in Figure 4.3.
Piezometers which are not pictured were placed 0.6 m toward the bay from the injection
point: two 0.6 m east of the pictured cluster and two 0.6 m west, at depths of 0.3 and 0.9
m below ground surface at each location. These were included to capture lateral
spreading of the plume, although this was expected to be minimal due to the twodimensional nature of tidal beach face circulation. The 10 piezometers nearest shore were
progressively moved to track the bromide plume. The piezometers were purged after each
installation with a peristaltic pump to fill the pipe with porewater, and subsequent
samples were taken by pumping from the depth of the piezometer screen to minimize the
sampling volume and any resulting flow disturbance. The piezometer pipe volume ranged
from 0.14 - 0.4 L (1.4x10-4 - 4x10 4 m3),depending on the depth, a small volume relative
to the estimated volume of the plume, which was on the order of 0.1 - 1 m3. Thus, initial
purging and subsequent smaller-volume sampling did not likely affect the tracer test
results.
Prior to tracer injection, each piezometer was sampled for initial values of bromide
concentration and porewater conductivity. Although porewater conductivity values were
somewhat variable in time, approximate average contours are depicted in Figure 4.3 (a).
Background NaBr concentrations ranged from x10-5 - 5x10 4 M, increasing with
salinity. The morning high tide on August 27 th occurred at approximately 9:20 AM, 30
cm above the subsequent low tide, and injection began at 9:36 AM at a depth of 0.4 m
and lasted 34 minutes. Samples were taken from the piezometers every 1-2 hours during
daylight until 6:50 PM on August 30, resulting in 32 total sample times. The approximate
0.1 M, 0.01 M, and 0.001 M contours for five time periods are shown in Figures 4.3 (b) -
76
4.3 (h). Only concentrations greater than or equal to lx10- 3 M are considered part of the
plume and contoured in Figure 4.3.
Figure 4.3 (a) depicts a clear inverse density gradient: highest porewater salinity at the
top of the beach face where saltwater infiltrates at high tide, and decreasing salinity with
depth. Similar salinity profiles have been observed at Waquoit Bay (Talbot et al. 2003)
and at Nauset Marsh, Cape Cod (Urish 2001). The salinity profile alone implies that
circulation due to tides and waves is contained within a few meters from shore, and the
NaBr tracer confirms this assumption. The plume moved from the injection point and
spread both horizontally and vertically over time. Lateral spreading was detected in both
0.3 m deep piezometers placed on either side of the transect, but not in the piezometers
screened at a depth of 0.9 m. The plume traveled downward initially and then circulated
upward, the center moving roughly 1 m/d. The bayward edge of the plume appears to
begin to flow into the bay approximately 40 hours after injection, discharging between 2
and 3 m from the position of high tide and the injection point. Bromide was never
detected in piezometers driven to depths greater than 1.2 m, which supports the assertion
that saline circulation due to tides and waves in Waquoit Bay is confined to the first few
meters into the bay, closer to shore than most of the discharging fresh water.
4.2.3 Discharge Patterns of Saline Circulation
The discharge profile from the August 14, 2003 seepage meter study is chosen for
comparison to circulation estimates because the flow rates are similar to the other head of
the bay experiments, and it is the only set of data that accurately characterizes both
discharge and salinity in the nearshore zone. From the data in Figure 4.4, there appear to
be three separate zones of saline discharge: less than 4 m, between 4 and 16 m, and
greater than 16 m. From the tracer test, it is likely that any baywater infiltrating from
tides and waves under normal conditions discharges within the first 4 m from shore. This
is supported by the discontinuous pattern of saline discharge 3.8 m from shore. The
amount of saline discharge less than 4 m from shore calculated from the 2003 seepage
meter data is 0.004 m3/s over the head of the bay. This is much less than the combined
value of 0.014 m3 /s estimated in Section 4.2.1, but greater than the minimum estimated
77
0
-0s
-1
-1.5
.1
0
1
2
3
4
0
1
Mstance from Injection Point [m]
2
3
4
Figure 4.3. Interpretation of NaBr tracer test data. Contours of natural salinity are shown
as grayscale, contours of injected bromide are shown as solid lines. Salinity is
approximated by electrical conductivity measurements in mS/cm and bromide
concentration is in molesll. (a) Experimental set-up and salinity profile. (b)-(h)
Approximate subsurface bromide molarity contours for selected sample times. Dashed
contours are inferred, dashed piezometers indicate screen location and length.
value of 8.9 x 10-4 m3 /s. One explanation for the analytical overprediction of circulation
due to tides and waves is the large amount of fresh discharge into Waquoit Bay. The
water table at the beach due to the upland hydraulic gradient is relatively high, limiting
the amount of saline water that can infiltrate at the beach face, and thus the amount of
saltwater discharge estimated by integration of the water table. It has been noted by
Ataie-Ashtiani et. al. (2001) that increasing the regional hydraulic gradient can overcome
the effect of tidal overheight in numerical simulations. This overprediction of beach face
circulation may also occur because Nielsen's (1990) equation did not account for
decoupling of the water table and tide due to a seepage face, which occurs in nearly all
beaches except those that are very steep and coarse-grained (Turner et al. 1997). The
seepage face width has been shown to be sensitive to the inland hydraulic gradient
(Raubenheimer et al. 1999), which, again, is significant in Waquoit Bay. Thus, Li et. al.'s
(1999) integration of this equation to obtain the estimates above also relies on a boundary
condition that may artificially raise the level of the water table at high tide, leading to an
overestimate of saline circulation due to tides.
Theoretically, saltwater circulating due to dispersion will discharge along the bayward
edge of the freshwater discharge. This discharge is clearly demonstrated by the measured
discharge in Figure 4.4 approximately 4 m from shore, where the freshwater flux begins
to decrease and saline flux increases. The measured freshwater discharge less than 16 m
from shore is 0.024 m3 /s, and saltwater discharge between 4 and 16 m from shore is
0.023 m3 /s. Kohout (1960) estimated saline discharge due to dispersion to be 10% or
more of the total discharge at a field site in Florida. There, the freshwater-saltwater
interface was quite dispersed: the distance from the 5% to 95% salinity contours was well
over 100 m at the narrowest part of the interface, resulting in a large amount of saline
circulation. In Waquoit Bay, however, the aquifer dispersivity is estimated to be small (D1
= 0.96 m; Dt = 0.018 m) (Garabedian et al. 1991), and the interface is less than 5 m thick
along Transect E (Figure 3.1) (Talbot et al. 2003). It is therefore likely that saline
discharge due to dispersion is less than 10% of the total flow in Waquoit Bay, and
consequently unlikely that dispersive saline circulation is equal in magnitude to the
79
freshwater discharge. Thus the total amount of saline outflow observed at Waquoit Bay
cannot be explained by dispersive circulation.
r
FW: 0.024 m3/s
A
SW: 0.023 m 3/s
a.
0.8
-
0.7
V
0.6
,
-
0.5
0.4
W
C
Ia
0.3
0.2
0.1
0
SW: 0.004
5o
5
25
35
45
55
45
55
T'ides and
Waves
Dispersion
b.
!
a
M
2003
It
en
go
v
-5
5
15
25
35
Distance from Shore Iml
Figure 4.4. Data from the 2003 single-transect seepage meter study, as presented in
Section 3.4.1.2. (a) Total discharge vs. distance from the shoreline. Likely locations of
inflow and outflow due to nearshore and dispersive circulation mechanisms are depicted
beneath the x-axis. (b) Seepage salinity vs. distance from the shoreline.
80
I_
4.3 Summary
This section has addressed three mechanisms for two-dimensional saline circulation and
discharge into coastal waters: nearshore circulation due to tides and waves, and
dispersive circulation along the saltwater-freshwater interface. In Waquoit Bay, the wave
action is usually very small, so most of the nearshore circulation results from tidal action.
Baywater flows into the beach face at high tide, circulates to a depth of approximately 1.2
m, and discharges within the first few meters from shore. Analytical estimates of
circulation due to tides and waves at the coast can explain at most 25% of the total saline
discharge (0.053 m3 /s) observed in August 2003 and extrapolated along the head of the
bay, but observations indicate that this number is closer to 7%. Dispersion along the
freshwater-saltwater interface entrains saline groundwater in the freshwater flowpath.
This mechanism draws baywater into the aquifer farther from shore. The inflow could
occur approximately 15 m from the shoreline, where observed outflow was discontinuous
in 2002 and 2003 (Figure 3.3), although net inflow was not observed. Circulation due to
dispersion, if assumed to be 10% of total freshwater flow, only accounts for another 5%
of the saline discharge. This means that 70-88% of the observed saline outflow in
Waquoit Bay cannot be explained by known forcing. The circulation mechanisms
discussed in this section clarify the subsurface flow patterns of saline groundwater
beneath Waquoit Bay, but they do not explain the magnitude of outflow observed during
the summer field studies.
81
References
Ataie-Ashtiani, B., R. E. Volker, and D. A. Lockington (2001) Tidal effects on
groundwater dynamics in unconfined aquifers. Hydrological Processes 15(4):
655-669.
Cartwright, N., L. Li, and P. Nielsen (2004) Response of the salt-freshwater interface in a
coastal aquifer to a wave-induced groundwater pulse: field observations and
modelling. Advances in Water Resources 27(3): 297-303.
Cooper, H. H. (1959) A hypothesis concerning the dynamic balance of fresh water and
salt water in a coastal aquifer. Journal of Geophysical Research 64(4): 461-467.
Dagan, G., and D. G. Zeitoun (1998) Seawater-freshwater interface in a stratified aquifer
of random permeability distribution. Journal of Contaminant Hydrology 29(3):
185-203.
Domenico, P. A., and F. W. Schwartz (1998) Physical and Chemical Hydrogeology. New
York, N.Y., John Wiley & Sons, Inc.
Essaid, H. I. (1986) A Comparison of the Coupled Fresh-Water Salt-Water Flow and the
Ghyben-Herzberg Sharp Interface Approaches to Modeling of Transient-Behavior
in Coastal Aquifer Systems. Journal of Hydrology 86(1-2): 169-193.
Garabedian, S. P., D. R. Leblanc, L. W. Gelhar, and M. A. Celia (1991) Large-Scale
Natural Gradient Tracer Test in Sand and Gravel, Cape-Cod, Massachusetts .2.
Analysis of Spatial Moments for a Nonreactive Tracer. Water Resources
Research 27(5): 911-924.
Inouchi, K., Y. Kishi, and T. Kakinuma (1990) The Motion of Coastal Groundwater in
Response to the Tide. Journal of Hydrology 115(1-4): 165-191.
Kim, G., K. K. Lee, K. S. Park, D. W. Hwang, and H. S. Yang (2003) Large submarine
groundwater discharge (SGD) from a volcanic island. Geophysical Research
Letters 30(21): 10.1029/2003GL018378.
Kohout, F. (1960) Cyclic Flow of Salt Water in the Biscayne Aquifer of Southeastern
Florida. Journal of Geophysical Research 65(7): 2133-2141.
Kooi, H., and J. Groen (2001) Offshore continuation of coastal groundwater systems;
predictions using sharp-interface approximations and variable-density flow
modelling. Journal of Hydrology 246(1-4): 19-35.
82
· ____
Larabi, A., and F. De Smedt (1997) Numerical solution of 3-D groundwater flow
involving free boundaries by a fixed finite element method. Journal of Hydrology
201(1-4): 161-182.
Li, L., and D. A. Barry (2000) Wave-induced beach groundwater flow. Advances in
Water Resources 23(4): 325-337.
Li, L., D. A. Barry, J. Y. Parlange, and C. B. Pattiaratchi (1997) Beach water table
fluctuations due to wave run-up: Capillarity effects. Water Resources Research
33(5): 935-945.
Li, L., D. A. Barry, F. Stagnitti, and J. Y. Parlange (1999) Submarine groundwater
discharge and associated chemical input to a coastal sea. Water Resources
Research 35(11): 3253-3259.
Moore, W. S. (1996) Large groundwater inputs to coastal waters revealed by Ra-226
enrichments. Nature 380(6575): 612-614.
Moore, W. S., and T. M. Church (1996) Submarine groundwater discharge - Reply.
Nature 382(6587): 122-122.
Nielsen, P. (1990) Tidal dynamics of the water table in beaches. Water Resources
Research 26(9): 2127-2134.
Nielsen, P., and P. Perrochet (2000) Watertable dynamics under capillary fringes:
experiments and modelling. Advances in Water Resources 23(5): 503-515.
Person, M., J. Z. Taylor, and S. L. Dingman (1998) Sharp interface models of salt water
intrusion and wellhead delineation on Nantucket Island, Massachusetts. Ground
Water 36(5): 731-742.
Raubenheimer, B., R. T. Guza, and S. Elgar (1999) Tidal water table fluctuations in a
sandy ocean beach. Water Resources Research 35(8): 2313-2320.
Raubenheimer, B., R. T. Guza, and S. Elgar (2001) Field observations of wave-driven
setdown and setup. Journal of Geophysical Research-Oceans 106(C3): 46294638.
Robinson, M., D. Gallagher, and W. Reay (1998) Field observations of tidal and seasonal
variations in ground water discharge to tidal estuarine surface water. Ground
Water Monitoring and Remediation 18(1): 83-92.
Simmons, C. T., T. R. Fenstemaker, and J. M. Sharp (2001) Variable-density
groundwater flow and solute transport in heterogeneous porous media:
approaches, resolutions and future challenges. Journal of Contaminant Hydrology
52(1-4): 245-275.
83
Smith, A. J. (2004) Mixed convection and density-dependent seawater circulation in
coastal aquifers. Water Resources Research 40(8): W08309
doi: 10. 1029/2003WR002977.
Smith, L., and W. Zawadzki (2003) A hydrogeologic model of submarine groundwater
discharge: Florida intercomparison experiment. Biogeochemistry 66(1-2): 95-110.
Talbot, J. M., K. D. Kroeger, A. Rago, M. C. Allen, and M. A. Charette (2003) Nitrogen
flux and speciation through the subterranean estuary of Waquoit Bay,
Massachusetts. Biological Bulletin 205(2): 244-245.
Taniguchi, M. (2002) Tidal effects on submarine groundwater discharge into the ocean.
Geophysical Research Letters 29(12): 10.1029/2002GL014987.
Taniguchi, M., J. V. Turner, and A. J. Smith (2003) Evaluations of groundwater
discharge rates from subsurface temperature in Cockburn Sound, Western
Australia. Biogeochemistry 66(1-2): 111-124.
Turner, I. L. (1998) Monitoring groundwater dynamics in the littoral zone at seasonal,
storm, tide and swash frequencies. Coastal Engineering 35(1-2): 1-16.
Turner, I. L., B. P. Coates, and R. I. Acworth (1997) Tides, waves and the super-
elevation of groundwater at the coast. Journal of Coastal Research 13(1): 46-60.
Turner, I. L., and G. Masselink (1998) Swash infiltration-exfiltration and sediment
transport. Journal of Geophysical Research-Oceans 103(C13): 30813-30824.
Turner, I. L., and P. Nielsen (1997) Rapid water table fluctuations within the beach face:
Implications for swash zone sediment mobility? Coastal Engineering 32(1): 4559.
Urish, D. W., and Gomez, A.L. (2001). The temporal and spatial distribution of coastal
groundwater seepage. First International Conference on Saltwater Intrusion and
Coastal Aquifers, Essaouira, Morocco.
Werner, A. D., and D. A. Lockington (2003) Influence of hysteresis on tidal capillary
fringe dynamics in a well-sorted sand. Advances in Water Resources 26(11):
1199-1204.
84
Chapter Five
Seasonality
The saline circulation mechanisms discussed in Chapter Four cannot explain the amount
of saline discharge observed in Waquoit Bay. In this chapter, a new mechanism is
proposed in which saltwater is forced in and out of coastal aquifers in response to upland
seasonal variation in recharge and hydraulic head. This process is presented conceptually
and modeled numerically in order to determine its potential effect on saline submarine
groundwater discharge (SGD) in actual aquifers.
5.1 Conceptual Model
Under static conditions, the Ghyben-Herzberg relation (Hubbert 1940) predicts a
freshwater-saltwater interface that is a distance below mean sea level proportional (40
times larger for a freshwater density of 1.000 Kg/L and a typical ocean saltwater density
of 1.025 Kg/L) to the head level above mean sea level at any location upland of a
saltwater body,
z
P-Pf
=ah,
(5.1)
where pf is the freshwater density, p, is the saltwater density, z is the depth below mean
sea level (MSL) to a point on the interface, and h is the water table elevation at that point
(Figure 5.1). This relationship is only approximate for an actual system with moving
groundwater and a dispersed interface, but it remains true that motion of the water table
at a timescale long enough to induce head changes at depth will result in interface
movement. It follows that seasonal changes in recharge will generate seasonal changes in
85
tbe interface that are potentially magnified by 40,or the freshwater density divided by the
density difference (equation 5.1 ), over the change in height of the water table for
umnfined aquifers. During a time of high recharge, freshwater will flow toward the
interface as it moves to establish dynamic equilibrium, thereby lowering the water table
and decreasing the 40-fold magnification over the original water table height. The overall
effect, however, will be an increase in the depth of the interface, effectively a seaward
shift. Fluid mass must be conserved in all physical systems, so as the interface moves
seaward. freshwater must move down to replace saltwater, and saltwater must move out
ofthe system; the opposite is true for landward motion.
Head
,
T
Figure 5.1. Schematic of interface position in relation to aquifer head level according to
the Ghyben-He&g
relation (not to scale). Freshwater discharge at the mast and
seasonal saline inflow and outflow at the seaflmr are depicted with arrows.
Theoretically, this seasonal movement of the freshwater-saltwater interface could induce
seasonal inflow and outflow of saline water at the sea floor. The vast majority of field
exprhents have measured SGD during the summer. Of these, studies reporting
discharge salinity have found that SGD is predominantly saline, but do not report flow of
seawater into the aquifer in an amount sufficient to balance the outflow (Robinson et al.
1998; Taniguchi 2002; Kim et al. 2003; Michael et al. 2003; Smith and Zawadzki 2003;
Taniguchi et al. 2003). Moore (1996) estimated that SGD was -100 m3 /d per m length of
shoreline in July 1994, equivalent to 40% of river discharge and, from consideration of
the regional freshwater balance, concludes that most of this discharge is seawater
circulation (Moore and Church 1996). Studies that have directly measured discharge
throughout the year report a total discharge that is consistently greatest during the
summer and lowest in the winter months along the Atlantic coast of the United States
(Simmons 1992; Cable et al. 1997a; Cable et al. 1997b). Local radium fluxes measured
over several years along the South Atlantic Bight indicate that SGD is larger in the
summer than the winter and spring (Moore 1987; Bollinger and Moore 1993; Moore
1996), and monthly groundwater discharge estimated from radium fluxes in Rhode Island
show a distinct pattern that peaks in the summer (Kelly and Moran 2002). Along the
Ganges Delta, radium fluxes are also out of phase with the water table elevation. Radium
fluxes are largest in the winter (Moore 1997), but the water table elevation is maximum
in the summer because, unlike the Atlantic coast of the US, recharge from the summer
monsoon dominates evapotranspiration. Thus, observations of SGD reveal highly saline
outflow that is unbalanced by inflow, and a seasonal pattern of discharge that is highest
during summer in the eastern United States. Freshwater discharge is expected to occur
year-round, but saline inflow could explain the decrease in total SGD measured during
winter. Seasonal movement of the freshwater-saltwater interface could account for the
large amount of observed saline discharge in the summer, balanced by an inflow of
saltwater during winter months.
Aquifers are recharged by the net infiltration of precipitation after evapotranspiration.
Although precipitation may vary seasonally, in temperate climates the seasonal variation
in recharge is forced primarily by the incoming solar radiation, creating a consistent
seasonal pattern in evaporation, soil moisture content, river flow, and groundwater levels
(Eltahir and Yeh 1999). In monsoonal climates, seasonal recharge patterns are caused by
precipitation, but have a similar effect on the regional hydrology. This seasonal change in
groundwater level may induce seasonal movement of the freshwater-saltwater interface in
coastal aquifers. Increasing hydraulic head levels will cause a seaward shift and saline
87
groundwater discharge, while decreasing head will allow the interface to intrude inland,
leading to inflow of saline water at the sea floor. Throughout the United States, analysis
of precipitation (P) and evapotranspiration (ET) indicates that highest recharge (P-ET)
generally occurs during late winter, and minimum recharge occurs during the summer
when ET is highest (Thornthwaite 1948; Evans and Jakeman 1998). This would seem to
imply that saline water should flow into the aquifer in the summer and out during the
winter, the opposite of what has been observed in the summer field studies. However,
there is evidence that water levels in unconfined aquifers do not change instantaneously
in response to recharge, aquifer hydraulic head may lag recharge by several months.
Statistical analysis of precipitation and shallow groundwater well levels in Illinois by
Changnon et. al. (1988) indicates that aquifer head lags precipitation by 0-3 months, most
commonly 1-2 months, depending primarily on soil type and the depth from the land
surface to the water table. Another study of the regional hydrologic cycle in Illinois
indicates that peak solar radiation, or minimum recharge, leads the minimum
groundwater level by approximately 3 months (Eltahir and Yeh 1999). These studies
were conducted in unconfined aquifers of Illinois, where the soil is not as sandy as in the
Cape Cod aquifer, but it seems likely that the lag experienced in Illinois may occur in
other regions with different soil types. The freshwater-saltwater interface has also been
shown to lag fluctuations of the water table (Essaid 1986) and flow rate in an unconfined
aquifer (Isaacs and Hunt 1986) in numerical simulations. If the seasonal cycle in
hydraulic head in a coastal aquifer lags the seasonal recharge, and if that head induces a
lagged change in the position of the interface that then translates to offshore saline inflow
and discharge, it is conceivable that recharge peaking in late winter and early spring
could induce saline discharge during the summer. The timescale and magnitude of the
translation of recharge-induced water table movement into saline inflow and outflow
along the sea floor is evidently complicated. Exploration with numerical simulations of
simplified aquifers and field measurements in actual systems can further our
understanding of these processes and the potential effect of seasonality on coastal
ecosystems.
88
5.2 Idealized Numerical Models
5.2.1 FEFLOW
The numerical models presented here have been run using the finite-element simulation
system, FEFLOW (Finite Element subsurface FLOW system) (Diersch 1998). FEFLOW
was developed to model variable density flow and transport in porous media. It is capable
of solving coupled flow and transport equations in two and three-dimensions to predict
groundwater flow patterns that are affected by density-dependence and temporal forcing.
The ability of numerical models to accurately represent variable density flow and
transport are often tested by comparison to benchmark examples such as the Henry,
Elder, and salt dome problems. FEFLOW was found to be in good agreement with
Henry's semi-analytic solution and other numerical results for the advance of a saltwater
front in a confined aquifer (Kolditz et al. 1998). Kolditz et al. (1998) also demonstrated
that FEFLOW simulations for the Elder fingering problem and the salt dome problem
agree well with prior results, although the mesh discretization may affect flow paths and
salt contours. FEFLOW has been used successfully in extending the Elder and salt dome
problems to include thermohaline convection processes in both two and three dimensions
(Diersch and Kolditz 1998).
FEFLOW has been used in several studies to simulate flow and transport in physical
systems. The flow pattern beneath the Swan-Canning Estuary in Western Australia has
been modeled by Smith and Turner (2001) to reveal density-driven free convection.
These convection cells result from density contrasts between the brackish river and fresh
groundwater and transport high levels of nutrients to the estuary. Contaminant transport
in non-saline environments has also been simulated using FEFLOW. For example,
Christoph and Dermietzel (2000) modeled the movement of DCE (trans-1,2dichloroethene) through a layered aquifer system to assess the potential effect of this
contaminant on groundwater quality. FEFLOW has also been used by Smith and
Zawadzki (2003) to model submarine groundwater discharge in the northeast Gulf of
Mexico, a site very similar to Waquoit Bay. Density-dependent circulation of saltwater
89
due to mixed convection and hydrodynamic dispersion was simulated by Smith (2004)
using both FEFLOW and the widely-used SUTRA code, with consistent results.
Numerical models have been used to simulate motion of the freshwater-saltwater
interface in relation to water management and saltwater intrusion. Some examples
include a sharp-interface approach to model saltwater intrusion in response to
groundwater pumping developed by Essaid (1990). Emekli et. al. (1996) simulated the
transient movement of the interface due to seasonal irrigation pumping to investigate
water resources in a coastal aquifer of Turkey. The effect of monsoonal rainfall on the
reversal of saltwater intrusion was simulated by Mahesha and Nagaraja (1996) as a
potential means of reclaiming brackish aquifers in India. These models demonstrate the
relationship between upland hydraulic head and the position of the freshwater-saltwater
interface, but the literature does not address the effect of interface movement on saline
submarine groundwater discharge.
5.2.2 Governing Equations
Density-dependent groundwater flow is governed by equations of conservation of fluid
and solute mass and conservation of momentum, or Darcy's Law. These equations must
take into account advective and dispersive solute transport as well as mass transfer
between the fluid and solid phases, although in the simulations that follow, only salt
transport is considered and it is assumed to exist only in the fluid phase. First, there is a
fluid phase mass balance which incorporates equations of state:
S-+
. (v) = Q -q,-
- Uv.(
vc)
(5.2)
S = 8h + ah
(5.3)
The Boussinesq approximation, which neglects density variations in all terms other than
the momentum conservation equation, is often introduced to reduce computational effort.
This approximation is appropriate where density variations are small in comparison to the
reference density (as in the natural coastal systems modeled below), but becomes
insufficient for large density gradients such as those found in high-concentration brines or
90
-
~~~ ~ ~ ~
fluids with high temperature gradients (Diersch 1998; Diersch and Kolditz 2002). This
approximation results in a simplified mass balance equation:
S- + v.(V)= Qp
at
(5.4)
The equation of solute mass conservation in terms of mass concentrations can be written
as:
-- + Iv VC - V(D.
VC)+ CQp = Qc
(5.5)
The third governing equation is the momentum equation, the generalized form of Darcy's
Law:
q = v =--
k
(VP- pg)
(5.6)
S = specific storativity of the porous medium with respect to hydraulic head changes [Pa']
h = hydraulic head related to the mass density of water [m]
C = mass concentration of the solute component [kg/m3]
r = porosity
v = macroscopic velocity [m/s]
fic = coefficient of expansivity resulting from the change of mass concentration of solute at constant
pressure [m 3 /Kg]
= coefficient of compressibility of the fluid resulting from the change of the hydraulic head at constant
mass fraction of the solute [m- ']
ah = coefficient of compressibility of the porous medium due to hydraulic head variations [m- ']
fIh
D = tensor of hydrodynamic dispersion [m2/s]
k = tensor of permeability of a porous medium [m2]
pQ,, = source term for fluid mass [kg/m3/s]
Qc = source term of solute component in terms of mass concentration [kg/m3/s]
Equation notation from Kolditz et al. (1998).
5.2.3 Model Properties and Boundary Conditions
A series of two-dimensional variable-density models has been constructed using
FEFLOW to examine the validity of the conceptual model of seasonality in coastal
aquifers. These idealized models are designed to illustrate the potential effect of a
seasonal recharge pattern on submarine groundwater discharge. Each model extends 500
m landward and 200 m seaward from the shoreline. The aquifers are unconfined and
homogeneous, with zero flow and zero mass transport boundary conditions along the
base and sides (Figure 5.2). The recharge boundary condition along the landward model
91
top varies sinusoidatly in time, with an average value of 0.002 mtd, an amplitude of
0.0025 d d , and a 365 day period. The average value of recharge results in a net yearly
freshwater inflow to the aquifer of 73 an,which is a reasonable value for a coastal
watershed in the eastern United States (and like1y many temperate climates elsewhere).
Near Waquoit Bay, recharge has been estimated as 46 cm (Carnbareri and Eichner 1998),
but the watershed extends more than 500 m landward and is wider than the 610 m head of
the bay. Thus the modeled recharge is reasonable for a Iarger watershed with less
recharge or a watershed of similar s i z (500 m landward per meter length of shoreline)
with more recharge than occurs on Cape Cod. The amplitude was chosen so that negative
recharge, or net evapotranspiration,occurs for approximately 80 days during the year,
which is reasonable for a climate with a 3-month summer season. The sea floor flow
boundary condition is constant head, and the transport condition allows for fluid flow
outward across the boundary that has a lower concentration than the seawater value by
assigning a constant concentration (30,000 m a ) where flow is inward, and zero
concentration gradient where flow is outward.
H=Om;C=30,000
mgll for irrflow, zsro
-re
5.2. Model schematic: flow and transport boundary conditions, initial
concentration profile, and dimensions.
Three aquifer parameters have been varied in the simulations to determine the effect of
each on the seasonal motion of the freshwater-saltwater interface and the resulting pattern
of SGD. These are aquifer thickness (b), hydraulic conductivity (K),and longitudind and
transverse dispersivity (DJand Dl),
all others are constant (see Appendix B). The
thickness of an aquifer has an effect on the length of the fkshwater-sdtwater interface. A
longer contact area increases the effect of dispersion and interface movement on SGD.
Two values of thickness were chosen: b=100 m allows the interface to extend landward
the full length of the model, while b=20 m causes the bottom boundary to intersect the
interface, cutting it short. The hydraulic conductivity of the subsurface affects the
freshwater head and the resulting depth of the interface. Three values of K are used in the
simulations: 5x10-4 , 1x10-4, and 5x10-5 m/s, within the wide range of values for finecoarse sand, typical material for unconfined coastal aquifers. According to Domenico and
Schwartz (1998), the hydraulic conductivity of sand ranges from 2x 10- 7 to 6x 10-3 m/s,
while Todd (1980) lists values from 3x10-5 to 5x10-4 m/s. Dispersivity affects the extent
of interaction between the fresh and saline water at the interface. In the past, models have
been assigned larger values than are likely to occur in natural systems in order to
maintain numerical stability. Here we use two sets of values: D 1=2 m, Dt=0. m, and
D=0. 1 m, Dt=0.005 m (D1/ Dt = 20). Analysis of 88 dispersivity estimates in porous
media by Gelhar et. al. (1992) indicates that on a 100 to 1000 m scale, the range of
longitudinal dispersivity from the most reliable estimates is 0.2 to 3 m. Thus our selected
values are average to high and a lower bound, respectively, and within the range of
longitudinal to transverse dispersivity ratios. Six simulations (Table 5.1) have been
analyzed to assess the effect of each parameter on seasonal submarine groundwater
discharge.
Table 5.1. Idealized model simulation parameters.
Model
Thickness
[m]
Hydraulic
Conductivity [m/s]
Longitudinal
Dispersivity [m]
Transverse
Dispersivity [m]
1
100
Ix104
0.1
0.005
2
20
Ix10-4
0.1
0.005
20
4
2
0.1
-4
2
0.1
3
Ix10-
4
100
Ix10
5
100
5x10 -4
2
0.1
100
-5
2
0.1
6
5x10
The mesh for each model is triangular, initially with a low nodal density, generated
automatically by FEFLOW, and then refined to balance numerical accuracy and stability
93
with computational efficiency. The mesh discretization and time step length can have a
large effect on simulation error. A mesh that is too coarse or a large time step can result
in growing numerical dispersion and instability, while a very fine mesh and small time
steps can lead to round-off error (Woods et al. 2003). Some guidelines have been
established based on non-dimensional numbers to help maintain this stability, particularly
where the density contrast is high. The Peclet number (Peg) can be defined as:
Pe-IVmAL
Peg, = D
g Do+aml
(5.7)
where vma is the maximum velocity parallel to AL, AL is the element length parallel to
flow, Do is the molecular diffusivity, and a is the dispersivity. In general, a value of Peg
less than 4 is sufficient for stability (Weatherill 2004), and a value less than 2 prevents
numerical dispersion (Boufadel 2000). The Courant number (Cr) is the ratio of the
maximum movement of the advection front in one time step to the element length in the
direction of flow.
Cr = IVaIAt
AL
(5.8)
Where Cr < 1, the fluid does not move through an entire element during a time step, so
numerical stability is generally maintained (Weatherill 2004). The number of elements in
the six models ranges from 93,532-597,638, a higher nodal density in models with lower
dispersivity values. This small grid spacing is necessary in coastal models; larger spacing
has been shown to underestimate brackish and saline water flux (Langevin 2003). Within
each model, the nodal density is higher in areas where concentration varies and near the
coast where velocity is highest in order to maintain a low Peclet number. The time steps
are constrained to a maximum value 0.5-1.5 d, depending on the model, to prevent
violation of the Courant criterion.
5.2.4 Simulation Results
Each of the six models was allowed to run to dynamic equilibrium, that is, the
concentration and velocity profiles were unchanging from year to year within a very
small error (less than -0.1% for concentration and -0.2% for velocity along the seafloor
boundary), although the nodal values changed seasonally within each year. The velocity
94
values in areas of the model that could potentially violate the Peclet and Courant criteria,
near the coastline for example, were monitored and checked for violations.
Unfortunately, a balance between the Pe and Cr numbers was difficult to maintain at all
times and locations within the models. For example, if grid spacing were increased to
conform to the Courant criterion, the Peclet number would be too large. The Peclet
criterion was never violated near the coastline, but far from it, where grid spacing is
larger, some violations did occur. Similarly, the Courant criterion was violated only very
near the coastline where flow is focused and velocities are large. Despite these occasional
violations, in all cases the models maintained numerical stability and eventually reached
dynamic equilibrium within a very small amount of error. Also, any numerical dispersion
that was introduced through a large Peclet number occurred far from the coast, where
velocities were very small. The model output was analyzed both within FEFLOW
through the budget and fluid flux analyzers and by exporting values of velocity,
concentration, and head at specified times and locations within the model (see Appendix
B).
Analysis of the simulations indicates that seasonal recharge leads to seasonal changes in
the water table, causing motion of the interface that induces inward and outward flow of
saltwater along the seafloor. Figure 5.3 illustrates the total flux of fresh and saline water
across the sea floor boundary in a day during each month for the six idealized models.
Saline inflow and outflow are plotted individually to separate the dispersion-induced
circulation from seasonal flow. Dispersive circulation occurs throughout the year,
flowing into the aquifer away from the shore, the magnitude decreasing monotonically
with distance, and out of the aquifer on the seaward edge of the freshwater discharge.
Thus inflow and outflow occur simultaneously but in different places along the boundary.
The seasonal component of saline flow also decreases monotonically with distance, but
the direction is either in or out, depending on the time of year rather than the position
along the boundary.
95
-
I.
I-
I
OD
'T
o
e:
a
-0.
-
I
Dispersive
Recharge
.
,Saltwater Out
Circulation
FreshwaterOut
-
Saltwater In
Figure 5.3. Total monthly fresh discharge, saline discharge, and saline inflow over the
sea floor throughout a simulated year. (a)-(f) are models 1, 2, 4, 3, 5, and 6, respectively.
5.2.4.1 Sensitivity of SGD to Hydrogeologic Parameters. An analysis of the sensitivity
of the modeled SGD to parameter values reveals the relative effects of aquifer thickness,
dispersivity, and hydraulic conductivity on the simulated system as well as the potential
magnitude of the seasonal discharge (Figure 5.4). The thickness of the aquifer has an
effect on the surface area of the interface. Greater contact between fresh and saline water
results in more dispersion-induced circulation and an enhanced seasonal effect in which
interface movement causes saltwater to move in and out of the aquifer. The relative
amount of dispersive circulation in each model can be examined by considering the
amount of saline outflow in days 0-120 of the model year, as indicated in Figure 5.3 on
model 5. If dispersion did not induce circulation, there would be no saline outflow during
96
I
_
this time. The effect of the length of the interface (or model thickness) on the amount of
dispersive circulation is small but significant. For the same values of hydraulic
conductivity and dispersivity, total saline outflow during days 0-120 in the low
dispersivity thick aquifer (b=100 m, model 1) is 6.9 m3 , while the corresponding thin
aquifer (b=20 m, model 2) exhibits 6.2m3 discharge. Similarly, the high dispersivity thick
aquifer model 4 (b=100 m) and corresponding thin model 3 (b=20 m) have dispersioninduced saline outflow during this time of 16.7 and 14.8 m3 , respectively. A more evident
effect of aquifer thickness is its influence on seasonal discharge. Peak saline discharge for
equal values of hydraulic conductivity and dispersivity in the thick aquifer (b=100 m,
model 1) is 51% of peak freshwater discharge, while the thin aquifer (b=20 m, model 2)
maximum percentage is 26%. The total seasonal discharge does not scale consistently
with the interface length because seasonal motion is greatest where the interface is
shallower, near the coast. The amplification of the motion of the interface in response to
motion of the water table is highly sensitive to the vertical distance between them. Thus
the closer spatial proximity from the water table to the interface near the coast appears to
outweigh the dampened change in head near the shoreline, so although model 1 has an
interface several times longer than the model 2 interface, it has only about twice as much
saline discharge.
The dispersivity of the aquifer also has an effect on saline circulation. As dispersivity
increases, more saltwater is entrained in the freshwater flow along the interface, resulting
in a greater amount of saltwater circulation. This dispersion-induced circulation cell
creates of inflow far from shore and nearshore saline discharge consistently in time. In
thick aquifers with equal hydraulic conductivity, total saline circulation is clearly greater
where there is a higher dispersivity (model 4) than in the low dispersivity model (model
1), as illustrated in Figure 5.3 (a) and (d) by comparing the saline outflow during days 0120 in each model. Increasing dispersivity appears to slightly decrease the impact of
seasonal recharge, as total saltwater circulation is 24% and 21 % of total freshwater
discharge, yet the peak saline discharge is 45% and 51% of peak freshwater discharge for
the high and low dispersivity models, respectively. This means that the low dispersivity
model 1 exhibits greater seasonal variability than the identical but high dispersivity
97
model 4. This effect also exists in models 2 and 3, identical thin aquifers but with low
and high dispersivity, respectively. This may occur because the Ghyben-Herzberg
relation holds only when the interface is sharp. Increasing the dispersivity of the aquifer
creates a thicker, more diffuse interface. This lessens the effect of the sharp density
contrast that causes the factor of 40 amplification from water table elevation to interface
depth in hypothetical sharp-interface, static aquifers. Therefore, the seasonal change in
aquifer head will result in a smaller motion of the interface in more dispersed aquifers. It
is possible that this seasonal effect has been overlooked in the past by numerical
modelers because dispersivity values are often much higher in modeled aquifers than in
actual ones.
Lastly, the aquifer hydraulic conductivity has a significant effect on both dispersive and
seasonal saltwater circulation as well as the position of the freshwater-saltwater interface.
In higher conductivity aquifers, the recharge is more easily discharged at the coast, so
hydraulic head does not build up as high and the interface is shallower than in an
identical but lower K aquifer. Because of this, seasonal variation in recharge produces a
more pronounced head variation in low conductivity aquifers. For example, at point (50,0), the head varies 0.29 m in model 6 (K = 5x10O5 m/s), 0.16m in model 4, (K = 1x10-4
m/s), and 0.05 m in model 5, (K = 5x10-4 m/s), with respective head maximums of 0.82,
0.56, and 0.23 m. Since the change in head drives the interface movement, one might
expect an increase in the seasonality of offshore saline flow with decreasing K. In fact,
the opposite is the case for the combination of parameters in our models for two reasons.
First, a lower K results in a lower groundwater velocity, thus increasing the time to reach
dynamic equilibrium and decreasing the distance the interface can move in one season.
Secondly, decreasing hydraulic conductivity increases the average height of the water
table in addition to its rate of change. The greater average water table height results in a
much lower average interface location, which increases the physical distance between the
recharge and the interface. This distance reduces the effect that a changing head has on
the position of the interface, likely because a longer timescale of change is required to
affect a deeper interface. A lower hydraulic conductivity also reduces the amount of
saline circulation due to dispersion, a result of a decrease in both the amount of saltwater
98
entrained in the fteshwater flow and the velocity of circulation. Both the total yearly
circukion and the p a k d i n e discharge as a m t a g e of freshwater discharge increase
with increasing K Total saline circul6ltioa is 19%, 24%, and 52% of -water
discharge, and peak saline disharge is 37% 47% and 10046 of p a k f r e s h w e
discharge for models 6,4, and 5 0< = 5x10*~.1x10~.and 5 x 1 0 ~mts), icspstively. Table
5.2 lists the total yearly saltwater
onas a percentage of W freshwater flow (365
m31yrin every madel) and the peak saline outflow as a percentage of the cornsponding
freshwater discharge during the peak flow month in, slll six thmtical models.
Figure 5.4. The effect of model hydraulic conductivity, dispersivity, and thickness on
saline discharge. Total saline chdation and peak &e discharge as a pcentage of
pe4lkfteshdischargeamplotted~paramebervdrres.
Table 5.2. Total saline circulation as a result of both dispersive entrainment of saltwater
and seasonal interface movement and corresponding percentage of total fresh discharge
over one 365-day cycle for each idealized model. Peak fresh and saline discharge from
monthly estimates and corresponding percentages reflect the magnitude of the seasonal
effect.
Model
Total Saline
Discharge
[m3/yr]
Total Saline
Discharge as % of
Fresh Discharge*
Peak Fresh
Discharge
[m 3/d]t
Peak Saline
Discharge
[m3 /d] t
Peak Saline
Discharge as %
of Peak Fresh
1
76.6
21%
35.1
18.8
51%
2
48.5
13%
33.0
10.4
32%
3
49.5
14%
33.1
6.3
19%
4
87.8
24%
35.6
16.7
47%
5
188.9
52%
31.6
31.4
100%
6
68.2
19%
36.2
13.3
37%
3
*Total fresh discharge is 365m /yr in all models.
tThe method used to determine monthly discharge may underestimate the total fresh and saline discharge
over the year as determined by the FEFLOW budget analyzer, a more accurate calculation. The budget
analyzer can only be used over an entire year due to an error in the buoyancy term calculation. This
problem has been fixed in the newest version of FEFLOW. In the version (4.9) used in this study, however,
monthly velocities and concentrations must be exported and fresh and saline flux calculated separately,
introducing error. Monthly discharge is therefore scaled uniformly to sum to the known yearly value to
correct this discrepancy.
5.2.4.2 Sensitivity of Time Lag to Hydrogeologic Parameters. A notable effect in the
seasonal models is a time lag between recharge, water table change, interface movement,
and fresh and saline discharge. One reason for this lag is the time for recharge at the
ground surface to percolate through the unsaturated portion of the aquifer and change the
level of the water table. A second, potentially more important reason for a lag between
peak recharge and peak aquifer head level is that any landward recharge greater than the
amount of fresh water discharging at the coast will raise the level of the water table
regardless of whether the recharge is increasing or decreasing with time. The water table
will therefore continue to rise after the peak in recharge until the amount of discharge
induced by the head gradient is large enough to balance the recharge. The flux of
freshwater to the ocean is directly proportional to the spatial gradient in hydraulic head.
The head at the coast is constant at mean sea level if tides are excluded, so the maximum
gradient coincides with maximum aquifer head, resulting in zero time lag between
freshwater discharge and hydraulic head. The relationship between seasonal saline
100
-----
discharge and aquifer head is a bit more complicated. Saline discharge is driven by
seaward movement of the freshwater-saltwater interface, which is forced by the aquifer
head. Unlike the relationship between recharge and head (head continues to increase past
the peak in recharge until freshwater outflow equals the recharge), seaward motion of the
interface will halt when the peak hydraulic head is reached (or when the temporallylagged effect of maximum head reaches the interface) since the position of the interface
relates directly to the magnitude of hydraulic head. Also, maximum saline discharge
velocity is caused by the maximum velocity of the interface rather than its most seaward
extent. So the highest saline outflow may occur before the interface reaches its deepest
point. The timescales of forcing between the peaks in aquifer recharge, hydraulic head
increase, seaward interface rate of motion, and saline discharge are unclear due to the
complicated nature of the relationship, but likely result in a saline discharge time lag that
differs from the fresh discharge time lag with respect to peak recharge.
The time lag between each aspect of the simulated coastal system is affected by the set of
model parameters. Only the parameters that varied between simulations are discussed
here: hydraulic conductivity, dispersivity, and aquifer thickness. The time between
maximum recharge and maximum hydraulic head decreases with increasing hydraulic
conductivity or aquifer thickness, as shown in Figure 5.5. The same is true for the time
between peak recharge and peak fresh velocity, which is expected since freshwater
discharge is directly affected by aquifer head. Increasing hydraulic conductivity also
decreases the lag between maximum recharge and saline velocity, but this lag is
unaffected by aquifer thickness. The dispersivity of the aquifer has no effect on time lag
for the values considered here. Table 5.3 lists the time lag in days for maximum and
minimum values of several components of the system. In general, the same pattern is
true: time lag is inversely proportional to hydraulic conductivity and thickness, but is
unaffected by the dispersivity. Increasing K allows the freshwater to flow out of the
aquifer faster, allowing the head to equilibrate with the high recharge more quickly,
which results in a lower time lag. The relationship to aquifer thickness is slightly more
complicated, but is likely related to the significantly greater change in head from
minimum to maximum in thinner aquifers (0.25 m in model 2 as opposed to 0.16 m in
101
model 1 at point (-50,0)). This occurs because the much shorter interface allows a
reduced aquifer volume for freshwater to flow into as the interface moves seaward, so the
same amount of freshwater (recharge fluxes at the top boundary in thin and thick aquifers
are equal) must move through a smaller space. Thus the change in water table elevation is
greater in thinner aquifers, taking a longer time to build up, and resulting in a longer time
lag.
Table 5.3. Time lag in days between maximums and minimums of system elements:
recharge (R), head (h), freshwater velocity at the origin ((0,0) V), and saltwater velocity
20 m seaward of the shoreline ((20,0) V).
Model 1
Model 2
Model 3
Model 4
Model 5
Model 6
100
125
125
100
60
110
200
210
200
200
170
210
Max h & Max (0,0) V*
Min h & Min (0,0) V*
-0
-0
-0
-0
-0
-0
-0
-0
-0
-0
-0
-0
Max R & Max (0,0) V
Min R & Min (0,0) V
Max Ah & Max (20,0) V
90
120
120
90
60
120
180
210
210
180
150
210
40
40
40
40
20
60
Min Ah & Min (20,0) V
170
160
145
180
195
135
Max R & Max (20,0) V
Min R & Min (20,0) V
90
90
90
90
30
120
180
180
180
180
150
210
Days Between:
Max R & Max h
Min R & Min h
* Velocity is reported every 30 d, and head is reported every 10d, so a lag of less than or equal to 20 d is
given as approximately zero.
The yearly oscillation in land-based recharge, aquifer head, position of the interface, and
offshore freshwater and saltwater velocity illustrates the relationship between each
element and the dependence of this relationship on the set of aquifer parameters. Figure
5.6 depicts the normalized variation (maximum=l, minimum=0) over one year of
simulation for each of the six theoretical models. Overall, the pattern and timing of
aquifer head, fresh discharge velocity, and saline discharge velocity oscillations are very
similar, particularly in thick aquifers with lower hydraulic conductivities. The lack of
dependence on dispersivity is clear in the nearly identical patterns in models 1 and 4 and
models 2 and 3. In the thin and high K models, saline velocity does not track head and
fresh velocity exactly; instead it exhibits a slightly lower lag from the maximum
102
recharge, likely because it is a result of the interface velocity, which is affected by the
rate of change of aquifer head rather than its magnitude. The salt concentration within the
interface20 m landward of the shoreline indicates that the interface begins to move
seaward 0-30d after the aquifer head begins to rise.
Figure 5 5 . The effect of model hydraulic conductivity, dispersivity, and thickness on
time lag. The number of days between peak recharge and peak aquifer head 50 m
landward of the shoreline, freshwater velocity at the shoreline, and saline velocity 20 m
offshore are plotted against parameter vaiues.
Winter
Spring
Summer
Recharge
Fall
Winter
-- -- -- Couceatratlon (-20, interface)
----- Head (-50,O)
Spring
--- ---
Summer
Fall
tseasonl
Fresh Velaci ty (0,O)
Saline Velocity (20,O)
Figure 5.6. Normalized variation in recharge, aquifer hydraulic head, interface p s i tion,
and fresh and saline velocity over one simulation year for each of the six model mns.
Hydraulic head is reported for a point 50 m landward of the shoreline at sea level.
Concentration, or salinity, at a point 20 m landward of the shoreline within the
freshwater-saltwater interface indicates interface movement: highest concentration
coincides with the extent of landward interface motion, and lowest concentration
coincides with the seaward extent. Freshwater velocity at the shoreline and saline
velocity on the seafloor 20 m from the coast indicate discharge variation throughout the
year. Actual values were normalized by dividing its difference from the minimum by the
difference between maximum and minimum values. Seasons are approximate for a
typical yearly recharge cycle within the United States. Model characteristics are given
below each model number: thick (100 m) or thin (20 rn); high K ( 5 x l 0 - mls),
~
medium K
( 1 x 1 0 ~m/s), or low K ( 1 x 1 0 ~d ~s ) ; and high dispersivity (Dl = 2 rn, D,= 0.1 rn) or low
dispersivity (Dl= 0.1 m, D, = 0.005 m).
5.3 Potential for Seasonality in Actual Aquifers
In theory, seasonally varying land-based recharge to an unconfined coastal aquifer will
induce changes in hydraulic head and the position of the freshwater-saltwater interface
that will drive inflow and outflow of saltwater at the sea floor. This was investigated
through a series of idealized numerical simulations of two-dimensional homogeneous and
isotropic aquifers. The simulations support the plausibility of the seasonal theory,
confirming that for a range of realistic hydrogeologic characteristics, seasonal inflow and
outflow is induced at the sea floor on a yearly timescale. Moreover, the peak saline
discharge over the year may be as large as the peak fresh discharge, an occurrence that
has been widely observed (Simmons 1992; Moore and Church 1996; Kim et al. 2003;
Michael et al. 2003; Smith and Zawadzki 2003), but has only been simulated numerically
without temporal forcing using dispersivity values much higher than those estimated in
real aquifers (Smith 2004). The numerical simulations also clarify the relationships
between the observable aquifer characteristics such as recharge, head, and discharge,
providing evidence of a significant time lag between the temporal forcing and the
observed changes. Thus a simple, idealized depiction of seasonality in coastal aquifers
explains two phenomena that have been previously unexplained: why peak submarine
groundwater discharge occurs during a period of low recharge, and why a high proportion
of this discharge is saline. Naturally, real aquifers are not homogeneous, isotropic, or
idealistic. However, the seasonal changes are clear for every set of parameters in our
simulations, indicating that while the complexity of true aquifers makes it difficult to
predict the magnitude or spatial regularity of the seasonality, it is likely that such effects
exist to a smaller or larger extent in a wide range of coastal systems.
105
References
Bollinger, M. S., and W. S. Moore (1993) Evaluation of Salt-Marsh Hydrology Using
Radium as a Tracer. Geochimica Et Cosmochimica Acta 57(10): 2203-2212.
Boufadel, M. C. (2000) A mechanistic study of nonlinear solute transport in a
groundwater-surface water system under steady state and transient hydraulic
conditions. Water Resources Research 36(9): 2549-2565.
Cable, J. E., W. C. Burnett, and J. P. Chanton (1997a) Magnitude and variations of
groundwater seepage along a Florida marine shoreline. Biogeochemistry 38(2):
189-205.
Cable, J. E., W. C. Burnett, J. P. Chanton, D. R. Corbett, and P. H. Cable (1997b) Field
evaluation of seepage meters in the coastal marine environment. Estuarine
Coastaland Shelf Science45(3): 367-375.
Cambareri, T. C., and E. M. Eichner (1998) Watershed delineation and ground water
discharge to a coastal embayment. Ground Water 36(4): 626-634.
Changnon, S. A., F. A. Huff, and C.-F. Hsu (1988) Relations between precipitation and
shallow groundwater in Illinois. Journal of Climate 1(12): 1239-1250.
Christoph, G. a. D., J (2000) The impact of a contaminated lignite seam on groundwater
quality in the aquifer system of the Bitterfield region - modeling of groundwater
contamination. Water, Air, and Soil Pollution 122: 421-431.
Diersch, H. J. G. (1998) FEFLOW finite element subsurface flow and transport
simulation system - user's manual/reference manual/white papers. Release 4.9.
WASY Ltd, Berlin.
Diersch, H. J. G., and 0. Kolditz (1998) Coupled groundwater flow and transport: 2.
Thermohaline and 3D convection systems. Advances in Water Resources 21(5):
401-425.
Diersch, H. J. G., and 0. Kolditz (2002) Variable-density flow and transport in porous
media: approaches and challenges. Advances in Water Resources 25(8-12): 899944.
Domenico, P. A., and F. W. Schwartz (1998) Physical and Chemical Hydrogeology. New
York, N.Y., John Wiley & Sons, Inc.
106
Eltahir, E. A. B., and P. A. J. F. Yeh (1999) On the asymmetric response of aquifer water
level to floods and droughts in Illinois. Water Resources Research 35(4): 11991217.
Emekli, N., N. Karahanoglu, H. Yazicigil, and V. Doyuran (1996) Numerical simulation
of saltwater intrusion in a groundwater basin. Water Environment Research 68(5):
855-866.
Essaid, H. I. (1986) A Comparison of the Coupled Fresh-Water Salt-Water Flow and the
Ghyben-Herzberg Sharp Interface Approaches to Modeling of Transient-Behavior
in Coastal Aquifer Systems. Journal of Hydrology 86(1-2): 169-193.
Essaid, H. I. (1990) A Multilayered Sharp Interface Model of Coupled Fresh-Water and
Saltwater Flow in Coastal Systems - Model Development and Application. Water
Resources Research 26(7): 1431-1454.
Evans, J. P., and A. J. Jakeman (1998) Development of a simple, catchment-scale,
rainfall-evapotranspiration-runoff model. Environmental Modelling & Software
13(3-4): 385-393.
Gelhar, L. W., C. Welty, and K. R. Rehfeldt (1992) A Critical-Review of Data on FieldScale Dispersion in Aquifers. Water Resources Research 28(7): 1955-1974.
Hubbert, M. K. (1940) The theory of ground-water motion. Journal of Geology 48(8):
785-944.
Isaacs, L. T., and B. Hunt (1986) A Simple Approximation for a Moving Interface in a
Coastal Aquifer. Journal of Hydrology 83(1-2): 29-43.
Kelly, R. P., and S. B. Moran (2002) Seasonal changes in groundwater input to a wellmixed estuary estimated using radium isotopes and implications for coastal
nutrient budgets. Limnology and Oceanography 47(6): 1796-1807.
Kim, G., K. K. Lee, K. S. Park, D. W. Hwang, and H. S. Yang (2003) Large submarine
groundwater discharge (SGD) from a volcanic island. Geophysical Research
Letters 30(21): 10.1029/2003GL018378.
Kolditz, O., R. Ratke, H. J. G. Diersch, and W. Zielke (1998) Coupled groundwater flow
and transport .1. Verification of variable density flow and transport models.
Advances in Water Resources 21(1): 27-46.
Langevin, C. D. (2003) Simulation of submarine ground water discharge to a marine
estuary: Biscayne Bay, Florida. Ground Water 41(6): 758-771.
Mahesha, A., and S. H. Nagaraja (1996) Effect of natural recharge on sea water intrusion
in coastal aquifers. Journal of Hydrology 174(3-4): 211-220.
107
Michael, H. A., J. S. Lubetsky, and C. F. Harvey (2003) Characterizing submarine
groundwater discharge: a seepage meter study in Waquoit Bay, Massachusetts.
Geophysical Research Letters 30(6): 10.1029/GL016000.
Moore, W. S. (1987) Radium 228 in the South Atlantic Bight. Journal of Geophysical
Research 92(C5): 5177-5190.
Moore, W. S. (1996) Large groundwater inputs to coastal waters revealed by Ra-226
enrichments. Nature 380(6575): 612-614.
Moore, W. S. (1997) High fluxes of radium and barium from the mouth of the GangesBrahmaputra river during low river discharge suggest a large groundwater source.
Earth and Planetary Science Letters 150(1-2): 141-150.
Moore, W. S., and T. M. Church (1996) Submarine groundwater discharge - Reply.
Nature 382(6587): 122-122.
Robinson, M., D. Gallagher, and W. Reay (1998) Field observations of tidal and seasonal
variations in ground water discharge to tidal estuarine surface water. Ground
Water Monitoring and Remediation 18(1): 83-92.
Simmons, G. M. (1992) Importance of Submarine Groundwater Discharge (Sgwd) and
Seawater Cycling to Material Flux across Sediment Water Interfaces in Marine
Environments. Marine Ecology-Progress Series 84(2): 173-184.
Smith, A. J. (2004) Mixed convection and density-dependent seawater circulation in
coastal aquifers. Water Resources Research 40(8): W08309
doi: 10. 1029/2003WR002977.
Smith, A. J., and J. V. Turner (2001) Density-dependent surface water-groundwater
interaction and nutrient discharge in the Swan-Canning Estuary. Hydrological
Processes 15(13): 2595-2616.
Smith, L., and W. Zawadzki (2003) A hydrogeologic model of submarine groundwater
discharge: Florida intercomparison experiment. Biogeochemistry 66(1-2): 95-110.
Taniguchi, M. (2002) Tidal effects on submarine groundwater discharge into the ocean.
GeophysicalResearchLetters 29(12): 10.1029/2002GL014987.
Taniguchi, M., J. V. Turner, and A. J. Smith (2003) Evaluations of groundwater
discharge rates from subsurface temperature in Cockburn Sound, Western
Australia. Biogeochemistry
66(1-2): 111-124.
Thornthwaite, C. W. (1948) An approach toward a rational classification of climate.
Geographical Review 38(1): 55-94.
Todd, D. K. (1980) Groundwater Hydrology. New York, John Wiley & Sons, Inc.
108
Weatherill, D., Simmons, C.T., Voss, C.I., and Robinson, N.I. (2004) Testing densitydependent groundwater models: two-dimensional steady state unstable convection
in infinite, finite, and inclined porous layers. Advances in Water Resources 27:
547-562.
Woods, J. A., M. D. Teubner, C. T. Simmons, and K. A. Narayan (2003) Numerical error
in groundwater flow and solute transport simulation. Water Resources Research
39(6): 1158.
109
110
Chapter Six
Seasonality at Waquoit Bay
The field studies presented in Chapter Three provide substantial evidence that a large
amount of saline groundwater discharges into Waquoit Bay during July and August, with
little or no inflow to balance the outflow. The mechanisms of saline circulation discussed
in Chapter Four fail to explain the observed outflow. In Chapter Five, the concept of
seasonal inflow and outflow of seawater at the bay floor due to the motion of the
freshwater-saltwater interface is presented, along with a set of idealized numerical
models that exhibit this behavior for a range of aquifer parameters. This chapter
investigates seasonal changes in recharge, aquifer head, and submarine groundwater
discharge at the bay floor in the Waquoit Bay watershed. The subsurface hydrogeology is
discussed in relation to observations, and a pattern of salinity and groundwater flow
beneath the bay is proposed. Finally, a numerical model of Waquoit Bay exhibits a
salinity profile and seasonal discharge variation similar the observations presented
throughout this study.
6.1 Evidence of Hydrologic Seasonality in the Waquoit Bay Watershed
A climate cannot be classified as moist or dry based on the amount of precipitation alone,
the potential for evapotranspiration (ET) in the system must also be considered
(Thornthwaite 1948). Evapotranspiration is very difficult to measure directly, but its
magnitude has been tied to atmospheric elements such as solar radiation, air temperature,
wind speed, and humidity. According to Thornthwaite (1948), the potential
evapotranspiration (PET) (actual ET depends on both PET and factors such as the amount
111
of precipitation and soil moisture storage) is highest in the southern United States and
lowest in the north, everywhere varying from winter to summer in a uniform pattern,
generally reaching a maximum value in July. Data have shown that air temperature is
most closely related to ET, and an empirical relation between PET, latitude, and
temperature can be used to produce a reasonable PET estimate. The water balance for a
specific location can be analyzed using monthly or daily air temperature and precipitation
data, soil information, and published conversion tables (Thornthwaite 1957). Although
this method is empirical and does not incorporate factors such as wind speed and
humidity, it is a good first approximation and has been shown to agree with more recent
numerical models (Evans and Jakeman 1998).
Using monthly average temperature and precipitation data measured at Long Pond in
Falmouth by the Falmouth Water Department and published by Richard Payne (2004), an
approximate monthly water balance for the Waquoit Bay watershed was calculated. The
method for these calculations is detailed in Thornthwaite (1957). Several assumptions are
required for this method. First, the amount of soil moisture storage available was
assumed to be 150 mm, appropriate for fine sandy loam with moderately to deep-rooted
vegetation, which includes most crops, pastures, and shrubs. A second assumption in the
calculations is that runoff is equal to approximately 10% of precipitation, which differs
from the the Thornthwaite (1957) recommendation that runoff is 50% of moisture
surplus. Runoff as a proportion of precipitation rather than as a proportion of surplus
(which may result in months with zero runoff) is potentially more realistic since 100% of
the precipitation cannot immediately infiltrate into the subsurface up to the water holding
capacity, particularly in inhabited areas such as Cape Cod where a fraction of the ground
is paved. Other assumptions include the validity of the tabulated values such as daily
potential evapotranspiration as a function of mean monthly temperature and soil moisture
retention as a function of potential evapotranspiration. Despite these assumptions, this
method can be used to approximate the yearly fluctuation in recharge on Cape Cod. The
precipitation data does not exhibit a strong seasonal trend, but the recharge of water to the
soil (precipitation - runoff - actual ET) is extremely variable from winter to summer. The
monthly precipitation and recharge over the time period of the Waquoit Bay field study
112
(January 1999 to December 2003) is plotted in Figure 6.1.Clearly recharge is lowest
during the summer and early fall and greatest in the winter and spring throughout the
study petid.
figwe 6.1. Monthly recharge of water to the subsurfe estimated from average monthly
rainfall and temperaoure data payne 2004) near Waquoit Bay using the Thornthwaite
(1957) method.
Water level data from wells installed by the United States Geological Survey (USGS) in
the vicinity of Waquoit Bay provide evidence that the d!rend in recharge is
translated into seasonal variation in the position of the water table, The water level in
many of these wells has been monitored over several years and is published as red-time
data (U.S.G.S, 2004a), Figure 6,2(a) gments water lever measurements for 7 wells
nearest the h a d of the bay, and Figure 6.2 (b) gives the depths and lucations of the wells.
All of the wells are screened in the upper, unconfined aquifer, except for MIW-26, which
is 89 m deep. However, the head measurements for this deep well are newly identical to
MIW-29, which extends to only 7.3 m below mean sea level in the same location,
indicating that either the d q well is fully screened or m e d within the upper aquifer,
or that the confining layer in this area is not present. Geologic maps of the subsurface
(Mastenon et d, 1997a) depict a come-grained formation extending to a depth of
approximately 100 m to the northwest of Waquoit Bay, beginning just east of MIW-26,
confirming that dl seven wells are likely in the unconfined upper aquifer. The data
Figure 6.2. (a) USGS well head levels above mean sea level (U.S.G.S. 2004a). @) Map
of well locations (U.S.G.S. 2004b)and depths below mean sea level.
indicate that the water table is highest from March to June and lowest from October to
December. This lag of approximately 3 months from peak recharge to peak head in the
upper aquifer is in very close agreement to the average lag of 100 days from maximum
recharge to maximum head in the theoretical models (Table 5.3).
6.2 Under the Ice: Winter Field Study
6.2.1 Methods
A winter pattern of submarine groundwater discharge at the head of Waquoit Bay that
differs significantly from the consistent summer saline discharge would indicate seasonal
variability of SGD in response to upland changes in recharge and hydraulic head. Field
work was therefore attempted during February 2004 to investigate the winter patterns of
SGD in Waquoit Bay.
Waquoit Bay is shallow and protected from the open ocean, communicating with
Vineyard Sound through a small opening at its mouth. The bay is therefore susceptible to
freezing over and can be covered with a layer of ice for several weeks or months during
very cold winters. January 2004 was a particularly cold month, resulting in a floating
layer of ice nearly 0.5 m thick in some places. The ice precluded the use of seepage
meters to measure winter groundwater discharge, but enabled instrumentation of the bay
from the top of the ice sheet. Between February 3 and 20, 2004, piezometers were
installed in several locations along Transect W (Figure 3.1), where seepage meters were
arranged in August of 2002 and 2003. The piezometers consisted of either 6 or 12-inch
screens attached to 3/4-inchsteel pipe of varying length, driven to depths of between 0.6
and 0.9 m. The piezometers were used to measure the vertical hydraulic gradient between
the bay and the screen location beneath the bay floor. Where the hydraulic head was
greater in the piezometer than in the bay, groundwater must be discharging, and where
the head was greater in the bay (a negative gradient by our convention), baywater must be
flowing into the aquifer.
115
After installation and periodically throughout sampling, each piezometer was purged to
remove fines, fill the pipe with porewater, and to ensure that water was flowing freely
between the aquifer and the piezometer. Salinity was measured both in the piezometer
and in the bay at several depths to account for salinity stratification resulting from
melting of the fresher ice sheet. The hydraulic gradient was initially measured using
manometers, but problems such as freezing within the tubing and exacerbation of density
effects caused measurement error. It was consequently determined that an electronic
water-level meter gives more accurate and consistent data.
Measurements using the electronic water-level meter were taken approximately every
hour during daylight in seven piezometers installed between 14 and 70 m from shore
from February 10 and February 13. The hydraulic gradient between the piezometer tip
and the bay was calculated from these measurements at each piezometer over a tidal
cycle. On February 20, four piezometers were driven into the sediment between 11 and
32 m from shore in order to obtain additional data where the band of high saline
discharge was observed during the summer. Again, measurements were taken
approximately once per hour over an 1 -hour period.
The salinity of the porewater was different than that of the baywater, so density
differences in the water columns had to be taken into account when calculating the head
gradient. First, conductivity (c), measured with a hand-held conductivity probe, was
translated into density according to the relation,
p = p0 exp(0.6923cm),
(6.1)
and
cm[kg/kg] = 0.69778 104c[S
cm],
(6.2)
where po is the freshwater density, taken to be 998.23 kg/m 3 , and Cmis the mass fraction
of NaCl (Holzbecher 1998).
The water from the bay and piezometers to be analyzed for conductivity was collected in
small sample vials. These vials were allowed to equilibrate with the air temperature
(approximately 1-3 C) before measurements were taken so that corrections for
116
measurement temperature as described in Perkin and Lewis (1980) could be avoided.
Also, density differences due to temperature variation between the porewater and
baywater were neglected based on the following calculation. The temperature at which
water is most dense, 4 °C,can be assumed for the baywater as an upper bound on density.
The porewater is likely warmer, but still cold considering that the piezometers are
screened less than a meter below the bay floor and the piezometer shaft is in contact with
the baywater. As an upper bound, 10 °C is chosen as the porewater (and piezometer)
temperature. Using Holzbecher's (1998) equations (2.2) for the density of freshwater at
different temperatures in the range below 40 °C,the density values calculated for 4 °C
and 10 °C differ by 0.027%. This is significantly less than the density difference due to
the salinity variation in Waquoit Bay, and well within measurement error incurred in the
conductivity and water level measurements.
The hydraulic head (h) measurements were converted to freshwater head (hf) using the
relation,
hf = P h-z,
Po
(6.3)
where po is the freshwater density, p is the density of the water column, and z is the
height above a datum. The hydraulic gradients were calculated from the converted head
values, and then averaged over a tidal cycle by weighting each measurement according to
the time increment between them.
6.2.2 Results
The vertical hydraulic gradients, calculated from the converted head values, for each
piezometer are plotted over one tidal cycle in Figures 6.3 and 6.4. The piezometers
closest to shore exhibit the strongest variation with the tide, with the largest gradient
during low tide and the smallest, or most negative, gradient during high tide. This inverse
variation becomes less apparent with distance from shore, which is consistent with the
seepage meter measurements from the summer experiments.
117
Figure 63. Hydraulic gradient vs. time over one tidal cycle for the February 2004
experiment. In the top panel are the three piemmeters closest to shore, and the bottom
pane1 depicts the four piemeters fatthest offshore. Tide level is shown on the right axis.
Measurements taken on d i f f a n t days are assigned a time relative to the tidal height.
D b n c e from Well 1 (mi
Figure 6.4. H y h u l i c gradient vs. time relative to tidal height over one tidal cycle for the
February, 2004 experiment. Each line qmsents one piezometer measured over time.
The values of hydraulic gradient over a tidal cycle were weighted according to the time
increment between them and averaged in order to obtain one representative gradient
value for each location (Figure 6.5 (a)). A net negative gradient indicating inflow was
measured in five locations corresponding to the position of the band of high saline
outflow observed during the summers of 1999 to 2003. Slug test data from October 20,
2000 and August 28, 2002 were converted to estimates of hydraulic conductivity using
the Hvorslev method (Domenico and Schwartz 1998) (see Section 6.3.2.1). The estimated
hydraulic conductivity (K) decreases with distance from shore, as described in Section
3.4.2, and is plotted in Figure 6.5 (b). Interpolated K values corresponding with
piezometer locations were used to estimate flow (q) according to Darcy's Law:
dh
q = -K dh
dx
(6.4)
The magnitude of discharge is plotted vs. distance from shore in Figure 6.5 (c), with
negative values representing flow from the bay into the aquifer. Along this transect,
extrapolated along the 610 m head of the bay for comparison to other discharge
estimates, the net inward flux is estimated as 0.01 m3 /s, although this number depends
greatly on the method used to estimate K.
The porewater salinity measurements exhibit a trend that was not observed in prior
seepage meter studies due to the extremely low flow rates in meters more than 50 m from
shore. The porewater was essentially the salinity of the baywater between 11 and 37 m
from shore, but became increasingly fresh as distance increases. This pattern coincides
inversely with the hydraulic gradient, which changes from negative to very strongly
positive with distance (Figure 6.6). This large upward gradient corresponding to fresh
porewater measurements indicates upwelling from a confined aquifer, which is possible
considering the geology of Waquoit Bay, which is discussed in the following sections.
119
-
-
Summer'03
Summer '02
Winter Gradient
+ Interpolated or
Extmpohted Values
Measured Values
-!I.&
7
-
-10
0
10
20
30
40
50
60
70
80
Distance from Well 1 Iml
Figure 6.5. Comparison of hydraulic gradient and discharge profiles for summer and
winter investigations dong Ttansect W.(a) Summer discharge (left axis) and winter
hydraulic gradient (right axis). (b) Hydraulic conductivity estimates from slug tests and
interpolated values used to calculate groundwater discharge from gradient measurements.
The conductivity estimate 70 m from shore is extrapolated from the measured data. (c)
Summer and winter submarine groundwater discharge. Flow of baywater into the aquifer
is observed where maximum offshore outflow was measured during the summer. Saline
discharge is minimal in the February experiment. A small amount of freshwater
discharges more than 50 m offshore, likely upwelIing from a confined aquifer.
F
w
i 6.6. Porewater salinity a d hydraulic gradient vs. distance from Well 1 along
T m t W for February 2004 experiment. Higb upward gradient offshore m s p o n d s to
very low porewater salinity, evidence of a connection to a confined aquifer.
6.23 Summary of W n e Chulation in the Unconfined Aquifer
In Chapter Four, twdmensional mechanisms of saline circulation in Waquoit Bay were
discussed: nearshore circulation due to tides and waves, and dispersive entrainment dong
the freshwater-saltwater interface. Tidal pumping can also be important in one dimension
near the coast, as evidenced by the tidally-correlated variation in discharge and hydraulic
gradient ohmed in both summer and winter experiments. In this section, we have
presentedevidence for a reversal of saline flow between summer and winter in Waquoit
Bay: net saline outflow during the summer balanced by net saline M o w in the winter.
Figure 6.7 sumnarkm the approximate discharge zones for each saline circulation
mechanism. Zone 1 corresponds to the area of obsewed tidal pumping in both August
2003 and February 2004. The extent of this zone offshore is determined by a denrease in
both the magnitude of variation over a tidal cycle in either discharge (August) or
hydraulic gradient (February) and tbe absolute value of the correlation coefficient
between the tide and discharge or gradient, The approximate extent of tidal pumping and
m s p o n d h g data are presented in Figure 6.8. Zone 2 in Figure 6.7 corresponds to
circulation of baywater due to tides and waves at the coastline. In Chapter Four,evidence
was presented that confines this circulation to the first few meters from the high tide
mark, nearer to shore than the discharging freshwater, Just offshore of the fresh discharge
is outflow of dispersion-inducedcirculating seawater (zone 3 in Figure 6.7). This outflow
is brackish to saline, likely extending no farther than the low discharge zone 12-15 m
from the shoreline. finally, seasonal exchange of saltwater between the aquifer and the
bay is represented by zone 4 in Figure 6.7.February measurements begin 13 m from
shore, but seasonal effects likely exist shoreward of this measurement, as depicted by the
dashed bracket in the diagram.
Figure 6.7. Discharge zone summary for saline circulation along Waquoit Bay Transect
W (Figure 3.1 ). Discharge data from August 2003 and February 2003 is presented in the
top panel. Color bars represent approximate extent of each zone of discharge along the
transect. Zone 1 is depicted by cross-hatching and extends from the shoreline to
approximately 28 m into the bay. Zone 2 (red shading) corresponds to nearshore
circulation due to tides and waves and extends approximately 3 rn from the high tide
mark. Dispersive cMation discharges in zone 3 (blue-gmn shading), along the
bayward edge of the fresh discharge. Seasonal saline outflow occurs in zone 4 (purple
shading). It has been measured between 13 and 35 m From shore, but the zone likely
extends to the shoreline, depicted by the dashed purple bracket, where February
measurements were not possible.
Figum 63. (a)Variation in discharge (top panel) and correlation coefficient (bottom
panel) for the August 2003 seepage meter study along Tmsect W (Figure 3.11, (b.)
Variation in hydraulic gradient (top panel) and correlation coefficient (bottom panel) for
the February 2004 piezometer study along Transect W.A damaw in both the absolute
value of the correlation coefficient and the magnitude of variation indicates a decline in
tidal pumping. A m l a t i o n d c i e n t of - 1 indicates a perfect inverse correlation
m
e
e
n the tide and either discharge or hydraulic gradient over a tidal cycle, and a
correlation coefficient of zero implies no m l a t i o n . The approximate extent of the tidal
pumping zone during each experiment is indicated by the verticd dashed lines.
6.3 Regional and Local Hydrogeology of Waquoit Bay
The groundwater discharge and hydraulic gradients observed in Waquoit Bay are the
product of a complex subsurface hydraulic head distribution and groundwater flow
pattern. This complexity is a result of the interaction between forcing, such as the
regional hydraulic gradient and saline circulation mechanisms, and the physical structure
of the subsurface. In this section, a description of the regional geology will lead into a
discussion of the subsurface hydrogeology local to Waquoit Bay. A combination of field
observations, a geophysical investigation, and published gsologic information are the
background for a proposed subsurface salinity distribution and flow pattern that forms the
basis for the numerical model presented in Section 6.4.
6.3.1 Regional Geologic Overview
The great ice age of the Pleistocene Epoch began two to three million years ago. During
this time, a worldwide lowering of sea level occurred, and glaciers advanced in the area
of Cape Cod, Massachusetts at least four times (Oldale 1981), each time removing
previous geologic deposits and re-depositing them. The most recent ice sheet advance
reached its maximum extent near the present-day Martha's Vineyard and Nantucket
Islands between 18,000 and 25,000 years ago. Two ice lobes contributed to the formation
of western Cape Cod: the Buzzard's Bay lobe to the west over Buzzard's Bay and
Vineyard Sound, and the Cape Cod Bay lobe to the east over present-day Nantucket
Sound. During their advance, the ice sheets deposited a layer of compact basal till onto
the surface of older bedrock. Approximately 15,500 years ago, the ice began to recede,
damming a large proglacial lake in the area of Nantucket Sound. Sediment from the ice
sheet was transported into the lake by meltwater streams (Masterson et al. 1997a). As the
lobes continued to retreat, chunks of ice broke off and were left behind, and the lake
expanded north. Initially, ice marginal deposits of unsorted clay, sand, and gravel were
deposited atop basal till, followed by finer deltaic and lacustrine deposits. The ice
recession halted at the Buzzard's Bay moraine on the western shore of Cape Cod.
Meltwater streams deposited very coarse deltaic material into the lake just to the east of
the moraine. At the same time, the Cape Cod lobe retreated to the present day north shore
of Cape Cod to form the Sandwich moraine. According to Masterson (1997a), meltwater
streams from the interlobate area built extensive progradational deltaic deposits of the
Mashpee pitted plain: coarser foreset beds as streams entered the lake, and finer
bottomset beds beneath it. The ice blocks buried during this depositional regime melted
over time, forming the numerous kettle-hole lakes currently existing on Cape Cod.
Continued melting eventually led to a sea level rise of approximately 300 ft (Oldale 1981;
Oldale and Barlow 1986; Masterson et al. 1997a).
124
Waquoit Bay is located along the southern shore of present-day Cape Cod, within the
Mashpee pitted plain deposits depicted by Oldale and Barlow (1986). The stratigraphy in
this area is uncertain because there is very limited subsurface information along the coast.
Regionally, however, the subsurface is characterized by approximately three depositional
layers. Mulligan and Uchupi (2003) describe these as (1) glacial outwash consisting of
primarily sand and gravel with lenses of silt; (2) very fine sand, silt, and clay deposited in
the proglacial lake; (3) till composed of poorly sorted gravel, sand, and silt atop bedrock.
Hydrogeologic sections based on USGS well logs and geologic analyses are mapped by
Masterson et al. (1997a). The transect, J-J' includes well FSW 183, which is located
approximately 1 km north of the head of Waquoit Bay. Four layers are depicted beneath
FSW 183, corresponding to those described by Mulligan and Uchupi. The top layer of
coarse-grained Mashpee pitted plain deposits exists to approximately 10 m below sea
level, and is composed of delta glaciolacustrine foreset beds consisting of medium to fine
sand with some silt. Below this to a depth of about 30 m is a layer of delta
glaciolacustrine bottomset beds of silt and clay with a small amount of very fine sand that
is moderately sorted and horizontally laminated. Lacustrine lake bottom beds of silt and
clay are depicted between 30 and 100 m below sea level. Beneath this thick layer is a
relatively thin (-10 m thick) layer of till composed of compact, unsorted sand, silt, clay
and scattered gravel, which overlies bedrock. The sediment distribution in the deltaic and
lacustrine meltwater deposits generally coarsens upward and fines from north to south
(Masterson et al. 1997a).
6.3.2 Hydrogeology within Waquoit Bay
The watershed contributing to groundwater discharge at the head of Waquoit Bay extends
approximately 2 km north, with a maximum width of about 1 km between the Childs and
Quashnet Rivers (Masterson and Walter 2000), and a total area of approximately 0.76
km2 (Cambareri and Eichner 1998). The hydraulic gradient near the head of the bay is
approximately 0.0015-0.0030, with a flow direction from north to south (Barlow and
Hess 1993; Masterson et al. 1997b).
125
The lithology described in 6.3.1 is strongly correlated to the hydraulic conductivity of the
sediments. Masterson et al. (1997a) give values of hydraulic conductivity and anisotropic
ratios that correspond to the sedimentary facies of the Cape Cod subsurface. Hydrologic
parameters such as hydraulic conductivity and anisotropy are difficult to estimate
accurately and often differ greatly over small distances. As a result, there are a wide
range of estimates of these parameters and others, such as specific yield, based on
methods such as aquifer and slug tests, permeameter measurements, and grain-size
analysis given in the literature (Masterson et al. 1997b; Moench et al. 2001). Some of the
parameter estimates for each layer underlying Waquoit Bay are depicted in Figure 6.9.
Because hydraulic conductivity generally decreases with grain size, and the deposits fine
from north to south, the K values near Waquoit Bay are likely on the low side of the
overall estimates.
The hydrogeology local to Waquoit Bay may differ from the regional description. The
slug tests, well logs, and well head measurements described below give some insight into
the local structure of the subsurface and corresponding parameter estimates.
In order to investigate the hydrogeology beneath the bay, five 2-inch diameter PVC wells
were installed at the head of the bay along transect W (Well 1 and transect W are shown
in Figure 3.1). Each well is screened over a 0.3 m interval to a depth of between 3 and 14
m below the surface. Wells 2, 3, and 4 are clustered 8.8 m north of Well 1, and screened
at 3.3, 6.4, and 12.6 m below the surface, respectively. Well 1 is located between the
average extent of low and high tide and is screened to a depth of 13.3 m. Well 5 is
located 18 m into the bay from Well 1, and is screened at a depth of 1 1.0 m. A sixth well
(referred to as Well 8 in Appendix C) was installed 36 m bayward of Well 1 to a depth of
9 m, but was destroyed a few days later by a boat. The wells were installed by GZA
Drilling, Inc. by driving a steel casing into the subsurface with a weight hung from a
tripod and pulley system (Appendix C, Figure C. 1). The casing was periodically flushed
with water and the cuttings collected at approximately 1 m intervals. The well logs
(Appendix C) indicate that a top layer of primarily medium sand exists in all five wells
from 0 to 7-9 m below the surface. A dramatic change in lithology from brown medium
126
sand to gray silt with some very fine sand occurs at depths of 11 m and 13 m in Wells 1
and 4, respectively. Although Well 4 is screened at 13.6 m, the outer casing was driven to
a depth of approximately 18 m during installation, where the lithology changed again
from the gray silt to light reddish brown fine sand. The well log depths are approximate,
however, due to compaction and the method by which the cuttings were flushed from the
holes.
The top two layers described above may be the continuous layers 1 and 2 described by
Mulligan and Uchupi (2003), but there is not sufficient evidence to connect them to the
regional system. The deep deposit of fine sand could be a localized formation, but it is
consistent with, though shallower than, the layer of gray fine sand at a depth of 40 m
along Transect E (Figure 3.1) depicted by Cambareri and Eichner (1998). Inspection of
the well log recorded during installation of the 46 m deep CCC Well (pictured in Figure
3.1 along Transect E and in Cambareri and Eichner (1998)) obtained from Desmond Well
Drilling, Inc. gives further detail for comparison to the well logs in Appendix C. The log
depicts sand and fine gravel to a depth of approximately 10 m below land surface,
followed by very fine gray silty sand to a depth of at least 19 m. Below this, the sediment
appears to coarsen, with very fine gray sand at 27 m, then very fine brown sand at 37 m,
and finally very fine white and gray sand at 46 m, the bottom of the borehole. This
appears to be consistent with the top sandy layer, second silty layer, and brown sand at
depth recorded in the well logs along Transect W. The depth of the fine sand in the CCC
Well appears to indicate that this may be a deep aquifer confined by the silty layer above
it, continuous at least locally beneath Waquoit Bay. The top layer of medium sand,
deeper layer of very fine silty sand, and deepest logged layer of very fine sand will be
referred to as layers 1, 2, and 3, respectively throughout this chapter.
6.3.2.1. Hydraulic Conductivity Estimates. The hydraulic conductivity of the nearsurface sediments of the bay floor has been estimated from slug tests as described in
Sections 3.4.2 and 6.2.2, exhibiting a pattern of decreasing conductivity with distance
from shore. The hydraulic conductivity of the deeper material surrounding the screens of
each of the five wells was also estimated with slug tests. The tests were performed by
127
withdrawing a fixed volume of water with a 1.5-cm diameter bailer, then measuring the
water level every 0.5 second as it recovered with a pressure logger (Leveloggers, Solinst
Canada, Ltd). The hydraulic conductivity values were calculated according to the
Hvorslev method (Domenico and Schwartz 1998):
K
r2 ln(L/r)
(6.5)
2LTo
where K is hydraulic conductivity, r is the radius of the borehole, L is the length of the
intake area, and To is the time from the start of the test (when head, h=ho) to the time
when h=0.37ho. The results of the slug tests are listed in Table 6.2.
A laboratory permeameter test was performed on samples collected during well
installation. The material from layers 2 and 3, however, was too fine for the permeameter
apparatus, so only the hydraulic conductivity of layer 1 could be estimated using this
method. The two samples tested were from Well 5: Sample 1 is from the top 4.5 m, and
Sample 2 is sediment from 4.5 to 7.5 m deep. The material was placed in a column 7.5
cm in diameter, with screens on the bottom and top to contain the sample, and a spring at
the top to compress it. Small-diameter tubing attached to two ports spaced 10.3 cm apart
along the column enabled measurement of the change in hydraulic head over the flow
distance. Water flowed in from a constant-head reservoir through tubing connected to the
bottom opening, upward through the sample, and then out the top. Air bubbles were
removed from the column and tubing before measurement. The flow rate at 30-second
intervals over a period of 7 minutes was measured and used to calculate the hydraulic
conductivity according to Darcy's Law (equation 6.4). This method was repeated for
several head gradient values for each sample, the results are given in Table 6.1. These
hydraulic conductivity estimates are only approximate because the method results in at
least two potential sources of error. First, the samples were collected by catching the
sediment as it was flushed out of the hole by a water jet during drilling, so it is possible
that some of the fine material was lost during sample collection. Secondly, the samples
were flushed out and repacked into the permeameter, which introduces a potentially
significant error due to the sensitivity of hydraulic conductivity to the sorting and packing
128
of aquifer material. However, the permeameter estimates are a useful supplement to the
slug test estimates, which are also prone to error.
The hydraulic conductivity estimates reported in this section as well as some of the
parameter estimates given in the literature are compiled and pictured in Figure 6.9. The
parameter estimates and lithology given by USGS sources (Garabedian et al. 1991;
LeBlanc et al. 1991; Masterson et al. 1997b; Moench et al. 2001) refer to the regional
geology, while hydraulic conductivity estimates and subsurface structure obtained from
this study and Cambareri and Eichner (1998) are local to Waquoit Bay. The modeled
values are discussed in Section 6.4.
Table 6.1. Hydraulic conductivity estimates from permeameter tests.
Sample 2 (4.5-7.5 m depth)
Sample 1 (0-4.5 m depth)
Ah [cm]
K [m/s]
0.3
-4
6.7 x 10
0.3
3.2 x 10 4
0.5
6.1 x 10 -4
0.9
5.7 x 10 -4
0.7
6.3 x 10-4
1.15
5.1 x 10
1.1
6.0 x 10-4
1.55
5.3 x 10 4
1.2
6.5 x 10 4
2.95
5.6 x 10 4
1.3
6 .4
Average
5.0 x 10 4
Ah [cm]
Average
K [m/s]
4
x 104
6.3 x 104
6.3.2.2. Hydraulic Head and Salinity Measurements in Wells. The hydraulic head in
each well has been measured and recorded over several days, and in come cases several
months, with pressure transducers (Leveloggers, Solinst Canada, Ltd). From August 27 to
September 2, 2003, Wells 1-5 and two additional shallow wells were instrumented with
leveloggers that recorded pressure simultaneously every 15 minutes. The additional
wells, 6 and 7 (Figure 6.10 (c)), are 1-inch diameter PVC with a 0.3 m screen, installed
using a hand auger and PVC casing to depths of 1.8 and 2.6 m, respectively (Ann
Mulligan, pers. comm.). The conductivity of the fluid in each of the wells was measured
by a combination pressure, temperature, and conductivity Levelogger. The conductivity,
129
average hydraulic head over three tidd cycles, and hydraulic conductivity at the screen
determined by slug tests are listed in Table 6.2. The head variation in each well over three
typical tidd cycles during that week is shown in Figure 6.10 (a) md (b), and a diagram of
well locations (Figure 6.10 (c))depicts the inferred geology at the depth of the well
screens, also listed in Table 6.2. The results indicate that there is a close hydraulic
; Work!
; Values
I "k:
I
I
I
w
r .I
K IW.0
V M * 0.f9.33
VM ~ 0 0 1 0 . 1
SIB n d day
K -03s
V M = ODf
v
I
Ylrddm an4 mom dlt
'1 1-36
1
014-S.l
:
I
I
I
K = 5.6 3
slua lerrs K
I Slug t.rWK
I
I
= 11-18
g 0.6
.
I
Fhe u n d
1I KWM-0.06
-Of
I
I
I
I
1
I
I w-aa
Near-surlace slug tosts, K = 0 4-5 B
I K = 3
Well logs: med-frne sand, some fines I VtH = 1: 0 3. 0
Well logl:wry R M mmnd. .I&. and day: N.
K 0.w~
1 VIH=O.I:O.PE
I
I
, K = 0.1
Appmxlmm 8mb
I
1
I
1
$
1
1
I
L
Figure 6.9. Schematic of measured and estimated parameter values for the four geologic
layers beneath Waquoit Bay. Literature values are listed on the left, obtained from
Masterson et. al. (1 997a), Moench et. al. (2001), LeBlanc et. al. (1 99 l), Garabedian et. al.
(1 9911, and Cambareri and Eichner (1 998). Measured values were obtained in this study
through laboratory permemeter experiments, and slug tests in wells and piezometers.
Model values are those used in the Waquoit Bay cross-section model.
connection between the bay and the subsurface, since all seven wells show significant
variation with tidal height. Well 1, although screened in layer 2, is likely damaged and
communicating with the bay or shallow subsurface and will not be included in this
discussion. Well 4 is clearly screened in layer 2 due to its low salinity and low hydraulic
conductivity as determined from slug tests, and Well 5 appears to be screened near the
interface of layers I and 2. Both wells exhibit head levels that are much greater than the
layer 1 wells and the tide, indicating that it is acting as a confining unit to layer 3, which
has a higher hydraulic head. The wells in layer 1 have much lower heads, some of which
are lower than mean tidal height, although this may be within measurement error. The 4well cluster 8.8 m landward of Well 1 indicates that flow is converging from the top and
bottom on the middle of layer 1: Well 3 has the lowest head (-3 cm), with a very high
head (17 cm) below it in layer 2, and somewhat higher heads (2 cm and -1 cm) in Wells 2
and 6 above it. Well 7, a shallow well near the location of average tide, has the lowest
head, which supports the expectation that groundwater flow converges and discharges at
the shoreline. These results are similar to those reported by Cambareri and Eichner
(1998), who measured hydraulic heads in a transect of multi-level sampling wells (CCCwells) along Transect E, less than 100 m east of transect W (Figure 3.1).
Table 6.2. Values of porewater conductivity, average hydraulic head over three typical
tidal c cles, and hydraulic conductivity for Waquoit Bay Wells 1-7.
Screen (likely)
Dist. into
Bay from
Well 1 [ml
Average PW
Conductivity
Avg. Head
[cm above
t
Estimated K
from Slug
13.3
Layer 2
--
14.4
-2 §
2.5 x 10-3
2
3.3
Layer 1
-8.8
0.13
2
1.2 x 10 -3
3
6.4
Layer 1
-8.8
20.4
-3
1.5 x 10-3
4
12.6
Layer 2
-8.8
0.13
17
5.0 x 10 5
5
11.0
Layer 1,2 (border?)
18
32.9
12*
N/A
6
1.8
Layer 1
-8.8
N/A (fresh)
-
N/A
7
2.6
Layer 1
0
1.95
-6
N/A
46.4
0
--
Well
Depth
[m]
1
tide
Geoloy at
- ---
[mS/cm]
tide]
Tests [m/s] *
Values are within approximately I2 cm.
§ Well I may not be well-sealed. Borehole electromagnetic induction results show that the porewater at the
screen should be fresh (Belaval 2003), but the well water is brackish, indicating a possible leak into the
well shallower than the screen. Slug test K is also much higher than expected in layer 2 (nearly 2 orders of
magnitude greater than Well 4).
* Measurement error may be significantly greater than 2 cm.
· Using Hvorslev method for K estimation.
131
Figure 6.10. (a) Hydraulic head measurements over three tidal cycles during a 1-week
measurement period from August 27 to September 2,2003 for all seven observation wells
and the tide. (b)Hydraulic head as in (a), but only for the northern well cluster and the
offshore Well 5 for clarity. Measurement aror in all wells includes survey emor and
pressure transducer error, and is approximately S cm in all wells with the axceptron of
Well 5, which may have a larger survey emr. Well I may not be well-sealed, there are
inconsistencies in salinity and hydraulic head magnitude that indicate a possible crack in
the well casing shallower than the screen depth (see note under Table 6.2). (c) Schematic
of well locations and likeIy geology at the well screens inferred from well logs, slug tests,
and salinity measurements. Scale is approximate.
6.3.2.3. Geophysical Investigation. The well logs, hydraulic conductivity estimates, and
porewater conductivity measurements reinforce the hypothesis that there are three locally
continuous geologic layers beneath the head of Waquoit Bay. Layer 1 is the most
permeable, the unconfined aquifer connecting the watershed to the bay. The permeability
of layer 2 drops dramatically, impeding any connection between the more permeable
layer 3 and layer 1. Because layer 3 is unable to discharge freshwater from its recharge
area into the bay, it maintains a higher hydraulic head and remains fresh under the bay
where layer I and part of layer 2 are saline. This is supported by offshore continuous
resistivity profiling by Marcel Belaval in June, 2002. Profiles of resistivity along several
transects both parallel and perpendicular to the shoreline at the head of the bay indicate
the presence of brackish water deeper than 5 m beneath the surface, extending at least
450 m into the bay (Belaval 2003). The inverted resistivity profiles along four transects
mapped in Figure 6.11 are shown in Figure 6.12. This approximately horizontal interface
that appears freshest toward the head of the bay is likely a balance of saltwater from the
surficial aquifer (layer 1) and freshwater from the confined aquifer (layer 3) below.
Borehole electromagnetic (EM) induction and gamma logging conducted by Marcel
Belaval (2003) in Wells 1, 3, 4, and 5 (Figure 6.13) provide additional information about
the salinity distribution and lithology of the subsurface in Waquoit Bay. Layer 2 appears
as an increase in the gamma logs at depths below land surface of 9 m, 8 m, and 7 m in
Wells 4, 1, and 5, respectively. Because the land surface slopes down from north to south,
this is a relatively consistent or slightly upsloping depth to a datum. The salinity profiles
with depth are evidence of a Ghyben-Herzberg-type interface in layer 1: salinity increases
at a depth of 7 m in Well 4, 2 m in Well 1, and Well 5 is completely saline. Porewater
salinity decreases again in Wells 4 and 1 at depths of 11 m and 12 m below land surface,
below the depth of the start of layer 2, giving further evidence of a somewhat horizontal
saltwater-over-freshwater interface occurring in layer 2 as a result of the fresh permeable
layer 3 beneath it.
133
1
F
w
i
. 6.11, (a) Map of all continuous resistivity profile mansects obtained by Marcel
Belaval, figure adapted from Belaval, (2003). (b) Schematic of north-south transects
WQ2 and WQ4, west-east msects WQI and WQI I , and Wells 1,2, 3, and 4, which
were profiled with borehole electromagnetic induction.
Hgtm &12. Fwr continuous resistivity profiles, adapted from BeLstvd, (2003). Higher
resistivity (shown as w m oolors) i n d i m lower salinity pewaterater
ConductivityImS/ml
(onductivityImSi/ml
a-
a.
D.
Well 3
Well 4
B
E
d.
Well 5
Well A.
I
I
.2
C-
0
20
40
60
o #
t
1o11
12141
20
(amma Icpsl
441
60
NO
IOU
(;Gma IepSl
Figure 6.13. Borehole (EM) induction and gamma logs, adapted from Belaval, (2003).
EM logging gives the conductivity of the formation surrounding a borehole, higher
conductivity corresponds with higher porewater salinity for similar geology. Gamma
logging gives a measure of clay content, or lithology, of the formation surrounding the
borehole: higher gamma means higher clay content. (a) EM and gamma for Well 4
(onshore deep well) taken on 1/28/03. Gamma log indicates that clay content begins to
increase at a depth of approximately 9 m below land surface, while EM shows fresh
water to a depth of 7 m, with a maximum at 9 m, then freshening with depth. (b)
Downhole logs for Well 3 taken on 1/28/03, results are similar to top 6 m of Well 4. (c)
Downhole logs for Well 1 8/26/02: EM suggests saline porewater from approximately 212 m below land surface, with a change in lithology below 8 m depth. (d) Downhole logs
for Well 5 8/26/02: porewater is consistently saline, but a higher clay content begins at
approximately 7 m below the sea floor.
136
I_
6.3.2.4. Groundwater Flow Patterns. Beginning approximately 50 m from shore, a
layer of thick, low permeability, organic muck begins to develop. This layer thickens
with distance from the shoreline, filling in what may be a kettle-hole in the center of the
bay. This kettle-hole likely cuts through the thinning top permeable layer of the bay floor,
into and potentially through the confining layer 2. A 10 m core taken -200 m from shore
exhibits coarse organic silt to a depth of 9.3 m, changing abruptly to sand below that
depth (Rosen 2004). The depth of sand is deeper than the 7 m deep layer 2 at Well 5,
indicating that the sand at 9.3 m is likely the top of layer 3. This would expose a
connection between the bay and the fresh layer 3, which would explain both the fresh
porewater and high upward gradient observed in the mucky sediment greater than 50 m
from shore in February 2004. Although the gradient is very high, upward flow through
the muck is impeded by its low permeability, thereby maintaining the high hydraulic head
observed in layer 3.
The existence of a confining layer beneath layer 1 indicates that fresh discharge observed
nearer than 50 m from shore is likely a result of flow in the unconfined layer only. A
large part of the summer saline discharge and winter inflow can be explained by seasonal
motion of the freshwater-saltwater interface within that layer. A portion of the saline
discharge may also be explained by the motion of the horizontal interface within the
confining layer, although the seasonal head variation within layer 3 has not been
observed directly in this study. The saline outflow as the interface rises may also
contribute to the observed banded pattern of saline discharge since flow will be diverted
around the low permeability muck. This hypothesized subsurface flow pattern is depicted
in Figure 6.14 and tested with a numerical model.
137
muck creates band?
winter inflow
1 /
'
HIghK:Iaywl
c-#
t-
res
Saline
!J-
All IPK
&5,
M i u r n K:
higbr W ~FOW
##
Prxrrxr u p h d
e 0# #
1
,#LLmwthrowonlrrlm*
/'
*Led
,u*----------
.-A1
.
Figure 6.14. Schematic of possible flow pattern in a cross-section of Waquoit Bay.
Dashed line is the position of the freshwater-saltwaterinterface, with a Ghyben-Henberg
position in layers I and 3 and a nearly horizontal position along layer 2, a balance
between the two layers. Freshwater extends farther bayward in layer 3 than in layer I ,
leading to upward freshwater flow after layer 2 is breached by the mucky layer. The low
hydraulic conductivity muck may prevent offshore flow of water as the interface moves
bayward, resulting in higher outflow and possibly creating the observed summer banded
discharge.
6.4 Numerical Model of Waquoit Bay Cross-Section
6.4.1 Model Geometry, Parameters, and Boundary Conditions
The Waquoit Bay model geometry is based on the regional and local geology discussed
in Section 6.3. The model is two-dimensional and extends 400 m landward and 250 m
bayward of the shoreline. Three layers with distinct sets of parameter values are
represented, similar to the top three layers depicted in Figure 6.9. The depths and
parameter values of each layer are listed in Table 6.3, Each layer is homogeneous with
respect to horizontal hydraulic conductivity, but the anisotropy factor and dispersivity
values may differ within single layers. A general trend of increasing anisotropy with
depth is given by isotropic conditions (V/H = 1) in approximately the top 3 m across the
top of the model, then a higher value of V/H, 0.3, to a depth of - 8 m below sea level,
with a much lower anisotropy, 0.05, below, consistent with published values for deeper
layers (Masterson et al. 1997a). The mucky layer differs from the other three layers in its
lateral extent. Observations within Waquoit Bay indicate that the muck begins as a thin
layer approximately 50 m from the shoreline, progressively deepening with distance. This
is represented in the model geometry: the muck cuts into layer 2 approximately 70 m
from shore, until it reaches the top of layer 3 (Figure 6.15).
Table 6.3. Extent and parameter values of geologic layer representations in Waquoit Bay
model.
Depth from sea
level (top/bottom)
Lateral Extent
(distance from
Hydraulic
Conductivity
Anisotropy
Factor
[m]
shoreline) [m]
[m/s]
(V/H)
I
Surface / -11
-400 - 64
3x10-4
0.3, -3 > y > -8
0.05, y < -8
2
-11/-18
-400 - 70; 100- 250
0.01x104
0.05
3
-18 / -80
-400 - 250
0.2x104
0.05
Muck
Surface/ -15
50 - 250
0. lx10
Layer
1.0, top 3 m
4
i, top 4 m
0.3, y < -5
The longitudinal and transverse dispersivity values were set as 0.5 m and 0.05 m,
respectively, throughout the model domain. These values are similar to the estimates
from Garabedian et. al. (1991), small enough to maintain the sharp freshwater-saltwater
interface observed in the shoreline wells (Figure 6.13). These small values of dispersivity
combined with the high value of hydraulic conductivity in layer 1 can lead to instabilities
where denser saltwater overlies freshwater, as observed in Waquoit Bay. After multiple
years of simulation with the above parameters, as the model advances toward steadystate, density fingers form within the model layer I (Figure 6.16). However, small and
large-scale porewater salinity measurements in Waquoit Bay indicate that density
fingering does not occur within the first 50 m from shore (Section 3.5.1.3). Thus, either
the parameter estimates are incorrect, or coastal processes not accounted for in the model
139
Layer 1
Layer 2,
Layer 3b
I
FJgure 6.15. Waquoit Bay model geometry. (a) Model proportions and coordinates.
Colors represent salinity values, where red = 30,000 mglL and blue = 0 mgk.(b) Section
from 33 m landward to 140 m baywacd of the shoreline. Colors represent layer locations
that correspond with values of hydraulic conductivity listed in Table 6.3. (c) Closer view
of salinity corttours and layer boundaries.
reduce the existence of density fingers.The buadary conditions (discussed blow)
represent seasund variations with smooth functions of time, and smaller-waletemporal
forcing from tides, waves, and individual atom events is not included. These high-
frequency fluctuations of hydraulic hesad at the sea floor have an o v d l effect of
increasing mixing within the fitst few meters of the aquifer. Thus,the dispersivity values
to a depth of
approximately 7 m beneath the sea floor have been set as 5 m and 0.5 m for
longitudinal and transverse dispmivity, respectively, to account for small-timescaIe
mixing. These higher dispersivity values were also assigned far from the shoreline, in
areas where concentration is constant, in order to reduce the simulation time.
Figmre 6.16. Example of modeled density fmgering. Row vectors are pictured in the
inset. Colm repsent porewater salt concenimion: red = 30,000 q L and blue = 0
m g L More buoyant freshwater Wws upward and denser SEiltwakr downward, creating
complex flow pattam.
The boundary conditions in the model depict seasonal variations similar to those in the
idealized models presented in Chapter Five. Across the top boundary, the freshwater flux
is specified as a sinusoidal function with a mean value of 0.0015 m/d, an amplitude of
0.0025 m/d, and a period of 365 d. This results in a total yearly recharge to the top layer
of 0.54 m, consistent with the 0.46 m of recharge estimated by Cambareri and Eichner
(1998), but somewhat larger considering the smaller landward extent of the model
compared to the actual watershed. Layer 3 is considered to be hydraulically separated
from layer 1 by a laterally continuous confining layer (layer 2), and recharged at a
hypothetical landward location. This assumption is supported by the salinity profile
beneath the bay: the lower aquifer must be confined to transport freshwater farther
seaward than the discharging freshwater at the coast. This separation requires an
independent boundary condition, which is represented by a hydraulic head condition on
the landward boundary, adjacent to layer 3, along the line (-400, -15), (-400, -80). This
boundary is also given a seasonal variation, with a mean value of 0.8 m, an amplitude of
0.7 m, and the same 365 d period as the flux boundary condition. The seafloor boundary,
(0, -1), (250, -1), is assigned a constant head of 1 m (the bay is simulated 1 m deep), with
constant concentration (30,000 mg/L) where flow is inward, and zero concentration
gradient where flow is outward. The constraint condition to obtain the zero gradient was
removed from x = 5 to x = 30 m from shore, where flow is primarily inward, to prevent
artificially low salinity in the top layer at dynamic equilibrium. All other boundaries are
maintained with zero fluid flux and zero mass transport. The parameter values and
boundary conditions described in this section were calibrated by trial-and-error to obtain
the approximate salinity profile observed in Waquoit Bay.
6.4.2 Results
Similar to the models presented in Chapter Five, the Waquoit Bay model was allowed to
run until it reached dynamic equilibrium. The change in velocity at nodes along the
simulated sea floor from year to year was less than 0.5%, and the change in salinity less
than 0.6%. The salinity profile (Figure 6.15), obtained by varying model parameters
within the estimated ranges, is very similar to that depicted in Figure 6.14. Freshwater
142
exists within layer 3 beneath the bay and the saline layer 1 due to a higher hydraulic head
in the confined layer translated from upland. Where the confining layer (layer 2) is
breached by the mucky sediment within a possible submerged kettle-hole approximately
70 m from shore, freshwater upwells into layer 1. This is consistent with the porewater
salinities measured in February 2004 (Figure 6.6). The simulated salinity profiles along
line segments that correspond to the approximate locations of the borehole EM profiles
taken by Marcel Belaval are also compared in Figure 6.17. The measured and simulated
profiles correspond well to each other, in the magnitude of salinity as well as in the width
of the transition zone along the interface, which is represented by the slope of the profile
as it changes from low to high salinity or high to low salinity. Both low and high flow
conditions are shown for the simulated profiles. During low flow times (fall or winter),
there is slightly more saline porewater than during times of high flow (spring or summer).
The total net freshwater input to the model over one year of simulation was 233 m3 ,
which included 219 m3 along the top flux boundary (inflow to layer 1) and 14 m3 that
flowed into the model along the head boundary on the landward side of layer 3. Although
the hydraulic head in layer 3 was high enough to drive freshwater beneath the bay, the
hydraulic conductivity of layers 2 and 3 was low enough that only a small amount of
freshwater flowed through this layer over one year. This is similar to what is expected in
Waquoit Bay since the salinity distribution within the bay (Charette et al. 2001) does not
indicate a large flux of freshwater into the bay far from the shoreline. The total saline
circulation in and out of the sea floor, which includes both seasonal exchange and
dispersive entrainment, over a yearly cycle is -103 m3 , or 44% of the freshwater flow.
The fresh outflow and saline inflow and outflow are plotted every month against the
simulation day, or recharge cycle, in Figure 6.18. The result is similar to the Chapter Five
simulation results: a yearly cycle in both fresh and saline flow is clearly evident. The
peak saline discharge is 63% of the peak fresh discharge, a significant proportion
considering the relatively thin geometry of the layer 1 aquifer. Also, the peak saline
outflow lags the highest recharge by 90 days, and the peak freshwater outflow lags peak
recharge by 120 days.
143
C*
Well I
-8
8
e"
u
a
A
i
9
a"
i nl
I
Measured Conductivity
Modeled Salinity
Figure 6.17. Comparison of modeled salinity and measured conductivity porew ater
profiles at Well 4 and Well 1 . (a) Well 4 (located 8.8 m landward of Well 1, Figure 6.10
fc)) EM conductivity profile measured by Marcel Belaval on 1/28/03. (b) Model salinity
vs. depth below land surface for low flow (winter) and high flow (springlsumer)
conditions along the line (-9.8,2), (-9.8,-15). (c) Well 1 (intertidal zone, Figure 6.10 (c))
EM conductivity profile measured by Marcel Belaval on 8/26/02. (d) Model salinity vs.
depth for low flow (winter) and high flow (spring/summer) conditions along the line (-1 ,
-11, (-1. -15).
1.0
0.004
0.8
0.003
0.6
4
0.002 Z
a
0.001
0.2
O 0.0
L.
U
oo
0.000
-0.2
-0.001
-0.4
-0.002
0
50
100
150
Recharge
-|--
200
250
-
Saltwater Out
*
FreshwaterOut
300
3.50
Saltwater In
Figure 6.18. Total simulated freshwater and saltwater inflow and outflow across the sea
floor (left axis) and recharge (right axis) vs. simulation day for the Waquoit Bay model.
The sum of monthly fluxes calculated from nodal velocity and salinity values was
somewhat lower than the total yearly flux calculated with the FEFLOW budget analyzer,
so monthly values were scaled to more accurately represent the total flow.
The results of the numerical model suggest that the hypothesized flow system represented
in Figure 6.14 can be simulated approximately using parameter estimates and boundary
conditions that are reasonable for Waquoit Bay. The simulated salinity profile very nearly
matches measurements in wells and along resistivity profile transects (Figure 6.12). Also,
the model exhibits a net inward flow of seawater in the winter and outward flow in the
summer that corresponds to estimates from field data. The amount of total and peak
saline outflow with respect to fresh outflow, however, is less than measured in Waquoit
Bay. There may be several reasons for this. First, the combination of values of hydraulic
conductivity, anisotropy, and dispersivity, as well as other parameters that have not been
varied in the modeling in this study, such as porosity and diffusion coefficient, may have
145
a considerable effect on simulated flow rates. Since all possible combinations were not
explored in this work, it is likely that significantly more saltwater flow can be achieved
through further variation of aquifer parameters. Secondly, the geometry chosen in the
model is only approximate based on a small number of observations, so additional
complexity in the actual aquifer may also increase flow. Thirdly, the hydraulic
conductivity was decreased and anisotropy increased near the horizontal freshwatersaltwater interface at the top of layer 2 to prevent fingering and maintain the observed
salinity profile in the model. However, if natural processes not present in the model, such
as tidal fluctuation, prevent fingering in Waquoit Bay, the hydraulic conductivity near
this interface may actually be larger than modeled. This would allow greater vertical
movement of the interface over a seasonal cycle, creating larger flows of saltwater into
and out of the aquifer. This hypothesis could be tested in the future by applying a tidal
boundary condition to the sea floor of the model and varying parameter values.
6.5 Summary
Clear seasonal variations in recharge, hydraulic head, and saline groundwater discharge
have been observed in the Waquoit Bay watershed, indicating that the seasonal effect
modeled in Chapter Five may exist in this real system. Analysis of the local
hydrogeology reveals a layered aquifer system with a salinity profile that differs
dramatically from the Ghyben-Herzberg interface that exists in theory and in idealized
models. The combination of data and numerical modeling results presented in this
chapter are evidence for a postulated groundwater flow pattern that could explain both
the spatial and temporal pattern of submarine groundwater discharge observed at the head
of Waquoit Bay.
146
__
References
Barlow, P. M., and K. M. Hess (1993) Simulated hydrologic responses of the Quashnet
River stream-aquifer system to proposed ground-water withdrawals. U.S.
Geological Survey Water-Resources Investigations Report 93-4074: 52 p.
Belaval, M. (2003) A geophysical investigation of the subsurface salt/fresh water
interface structure, Waquoit Bay, Cape Cod, Massachusetts. Master of Science
Thesis, Boston College. Department of Geology and Geophysics. Chestnut Hill,
Boston College: 78p.
Cambareri, T. C., and E. M. Eichner (1998) Watershed delineation and ground water
discharge to a coastal embayment. Ground Water 36(4): 626-634.
Charette, M. A., K. O. Buesseler, and J. E. Andrews (2001) Utility of radium isotopes for
evaluating the input and transport of groundwater-derived nitrogen to a Cape Cod
estuary. Limnology and Oceanography 46(2): 465-470.
Domenico, P. A., and F. W. Schwartz (1998) Physical and Chemical Hydrogeology. New
York, N.Y., John Wiley & Sons, Inc.
Evans, J. P., and A. J. Jakeman (1998) Development of a simple, catchment-scale,
rainfall-evapotranspiration-runoff model. Environmental Modelling & Software
13(3-4): 385-393.
Garabedian, S. P., D. R. Leblanc, L. W. Gelhar, and M. A. Celia (1991) Large-Scale
Natural Gradient Tracer Test in Sand and Gravel, Cape-Cod, Massachusetts .2.
Analysis of Spatial Moments for a Nonreactive Tracer. Water Resources
Research 27(5): 911-924.
Holzbecher, E. (1998) Modeling Density-Driven Flow in Porous Media: Principles,
Numerics, Software. Berlin, Springer-Verlag.
LeBlanc, D. R., S. P. Garabedian, et al. (1991) Large-Scale Natural Gradient Tracer Test
in Sand and Gravel, Cape-Cod, Massachusetts .1. Experimental-Design and
Observed Tracer Movement. Water Resources Research 27(5): 895-910.
Masterson, J. P., B. D. Stone, D. A. Walter, and J. Savoie (1997a) Hydrogeologic
framework of western Cape Cod, Massachusetts. U.S. Geological Survey
Hydrologic Investigations Atlas HA-741: 1 pl.
147
Masterson, J. P., and D. A. Walter (2000) Delineation of Ground-Water Recharge Areas,
Western Cape Cod, Massachusetts. Water-Resources Investigations Report 004000. Reston, VA, U.S. Department of the Interior, Geological Survey.
Masterson, J. P., D. A. Walter, and J. Savoie (1997b) Use of particle tracking to improve
numerical model calibration and to analyze ground-water flow and contaminant
migration, Massachusetts Military Reservation, western Cape Cod,
Massachusetts. U.S. Geological Survey Water-Supply Paper 2482: 50 p.
Moench, A. F., S. P. Garabedian, and D. R. LeBlanc (2001) Estimation of Hydraulic
Parameters from an Unconfined Aquifer Test Conducted in a Glacial Outwash
Deposit, Cape Cod, Massachusetts. Professional Paper 1629, U.S. Department of
the Interior, Geological Survey: 51p.
Mulligan, A., and E. Uchupi (2003) New Interpretation of Glacial History of Cape Cod
May Have Important Implications for Groundwater Contaminant Transport. EOS
84(19): 177, 182-183.
Oldale, R. N. (1981) Geologic history of Cape Cod, Massachusetts. Washington, D.C,
U.S. Dept. of the Interior, Geological Survey.
Oldale, R. N., and R. A. Barlow (1986) Geologic Map of Cape Cod and the Islands,
Massachusetts. Miscellaneous Investigations Series, I-1763. Reston, VA, U.S.
Geological Survey: Miscellaneous Investigations Series, 1-1763.
Payne, R. (2004) Falmouth Monthly Climate Reports, Falmouth Water Department,
www.whoi.edu/climate/", Woods Hole Oceanographic Institution. 2004.
Perkin, R. G., and Lewis, E.L. (1980) The Practical Salinity Scale 1978: Fitting the Data.
IEEEJournal of OceanicEngineering5(1):9-16.
Rosen, G. P. (2004). Department of Geological Sciences. Gainesville, University of
Florida: 58p.
Thornthwaite, C. W. (1948) An approach toward a rational classification of climate.
Geographical Review 38(1): 55-94.
Thornthwaite, C. W. (1957) Instructions and tables for computing potential
evapotranspiration and the water balance. Publications in Climatology X(3): 185-
243.
U.S.G.S. (2004a) U.S. Geological Survey National Water Information System
(NWISWeb) data available on the World Wide Web, accessed [July 30, 2004], at
URL http://nwis.waterdata.usgs.gov/usa/nwis/gwlevels.
148
U.S.G.S. (2004b) U.S. Geological Survey, The National Map, Seamless Data Distribution
System data available on the World Wide Web, accessed [July 30, 2004], at URL
http://seamless.usgs.gov/viewer.htm.
149
150
Chapter Seven
Conclusions, Implications, and Future Directions
7.1 Summary and Conclusions
This study is an investigation of coastal groundwater systems, focusing on the
mechanisms controlling submarine groundwater discharge, in an effort to more fully
understand the complicated flow patterns of fresh and saline groundwater and the
resulting contribution to nearshore surface waters. A detailed characterization of the
subsurface hydrogeology and groundwater discharge into an estuary, Waquoit Bay,
Massachusetts, has answered questions regarding the amount, pattern, and salinity of
submarine groundwater discharge, its spatial and temporal variability, and the potential
use of natural tracers to estimate groundwater flux. With these answers are new questions
concerning the origin of the discharge: how can it be explained using our knowledge of
subsurface flow patterns within the system? Theoretical and numerical examination of
saline circulation due to small-scale forcing of tides and waves as well as large-scale
regional forcing of upland recharge leads to the conclusion that seasonal recharge has an
important impact on saline groundwater discharge. Waquoit Bay is a case study for what
is likely occurring along the coast in many parts of the world. Where inland aquifer
recharge varies seasonally, due to cyclic incoming solar radiation in temperate climates
and precipitation in monsoonal climates, saline water may be forced in and out of the sea
floor in a seasonal cycle.
In summary, we have found significant submarine groundwater discharge into Waquoit
Bay during the summers of 1999-2003. Fresh discharge measurements are slightly greater
151
than, but consistent with, a water balance estimate, but saline discharge is much larger
than expected. More than twice as much saltwater than freshwater discharges into
Waquoit Bay during July and August, most of it occurring in a band 20-45 m from shore.
Only 12-30% of this saline discharge can be explained by circulation mechanisms due to
waves, tides, and dispersion along the freshwater-saltwater interface, although direct
measurements of inflow are negligible. A hydraulic gradient indicating net flow of water
from the bay into the aquifer over a tidal cycle was measured in February 2004,
coinciding with the location of the summer band of high saline discharge.
Idealized aquifer models show that for realistic values of aquifer parameters and
boundary conditions, seasonally varying recharge results in cyclic motion of the water
table, or aquifer head, which induces motion of the freshwater-saltwater
interface. The
interface movement forces groundwater to discharge along the sea floor as it moves
seaward (due to rising head), and draws seawater into the aquifer as it retreats landward.
The models reveal a time lag between maximum recharge and maximum saline discharge
at the coast that can result in peak outflow during the summer if peak recharge is in the
late winter and early spring. Moreover, the maximum saline discharge can equal the
maximum fresh outflow for a typical value of hydraulic conductivity. A similar model
that incorporates the complex local hydrogeology of Waquoit Bay reveals a subsurface
salinity profile and temporal discharge pattern that correspond well to observations. This
further supports the proposed mechanism of seasonal saline circulation in coastal
groundwater systems.
The seasonal forcing of saline submarine groundwater inflow and discharge thus explains
the observations of high saline SGD in this and other studies as well as seasonal variation
that gives rise to maximum outflow during the summer along the Atlantic Coast of the
United States.
152
7.2 Implications
The direct measurements of submarine groundwater discharge in this study provide
insight into temporal and spatial variability that is essential for the design of similar field
work in the future. We have shown that high spatial variability in groundwater discharge
requires a dense field of measurements for an estimate of total flow, and that
measurement over a complete tidal cycle is necessary, particularly near the coast. This
work also demonstrates the significance of seasonal variability and the importance of
measurement throughout a yearly cycle if results are to be extrapolated in time. The
spatial variability in the radium activity of groundwater also highlights the complications
involved in the use of natural tracers to estimate groundwater flow, particularly at the
coast where porewater ionic strength varies dramatically. In terms of field equipment, we
have introduced a new intertidal seepage meter with the ability to capture the previously
elusive fresher discharge at the shoreline.
Improvements in submarine groundwater discharge measurement techniques are only
necessary if SGD is important to coastal ecosystems: if the constituents in groundwater
have the potential to significantly affect the composition of coastal waters. The total
groundwater discharge measured during the summer in Waquoit Bay is approximately
10% of the Quashnet and Childs River discharge estimated by Cambareri and Eichner
(1998), a significant proportion of flow. The fresh groundwater carries nutrients from
septic systems and fertilizer that has greatly impacted the Waquoit Bay ecosystem over
the past several decades by increasing microalgal growth, thereby greatly altering
vegetation and increasing mortality of benthic fauna (Valiela et al. 1990). Submarine
groundwater discharge influences the productivity, biomass, species distribution, and
zonation of estuarine systems (Johannes 1980). This is particularly important in shallow
estuaries where the seagrass biomass is a habitat for a variety of organisms and a nursery
to others, but its dependence on light makes seagrass an easy victim of nutrient loading
that results in shading from phytoplankton and seaweed (Valiela 1995). These effects are
widespread in shallow coastal ecosystems and can be anticipated throughout the world
153
where nutrient levels are increasing due to the anthropogenic effects of human population
growth.
It has been argued that saline groundwater discharge is an insignificant contributor to
surface water constituents since its source is the surface water itself (Younger 1996).
However, the ammonium concentration in brackish and saline porewater beneath the
intertidal region of Waquoit Bay is up to 14 times greater than that of the baywater and
exhibits an increasing trend with salinity (Talbot 2003). High levels of nutrients and
contaminants have been measured in other areas throughout the world as well. Whiting
and Childers (1989) found that porewater advecting into a South Carolina saltmarsh is of
the same salinity as the overlying surface water, but with an order of magnitude higher
concentration of ammonium and four times more phosphate. This saline advection, likely
due to tidal pumping, contributes three times as much ammonium and an equal amount of
phosphate to the water column as low tide runoff from the marsh surface, making it a
major source of nutrients to the marsh. Another study in South Carolina finds that high
discharge of saline groundwater containing elevated concentrations of DOC contributes
significantly to the overall carbon budget in the North Inlet estuary (Goni and Gardner
2003). Simmons (1992) measured SGD and its composition using seepage meters along
the Florida Keys. Measured levels of Ca, Na, K, and Mg ions were nearly identical to the
overlying seawater, but levels of Zn, Cd, Pb, and Ni ions as well as nitrate and total
phosphate were significantly higher. Comparison of concentrations of the same ions and
nutrients in seepage meter discharge and water from the James, Savannah, and Altamaha
Rivers reveals that discharge levels are consistently much higher than river levels, often
by several orders of magnitude. Along the coast of Jeju Island, Korea, the fluxes of
inorganic nitrogen, inorganic phosphorous, and silicate into the sea from groundwater
were measured to be 22, 530, and 46 times greater, respectively, than the flux from fresh
groundwater in Eastern Jeju, and 2.2, 5.0, and 4.5 times greater in Western Jeju (Kim et
al. 2003). Concentrations of these nutrients in fresh and saline groundwater are similar,
the higher nutrient flux from saline groundwater is a result of its higher volume flux, but
seawater concentrations are significantly lower, indicating that SGD has a large effect on
the supply of nutrients to this portion of the South Sea of Korea. Thus, saline
154
groundwater flowing in and out of a coastal aquifer in a seasonal pattern can carry
harmfully elevated levels of nutrients and contaminants. This, combined with the
widespread observation of significantly greater saline than fresh discharge, indicates that
saline submarine groundwater discharge can have an important effect on receiving
waters.
If saline groundwater contributes significantly to coastal waters, high seasonal discharge
may increase its impact. Along the eastern United States, highest discharge occurs during
the summer, when microbial activity is greatest and organisms are most active and
reproducing, thus having a larger potential effect on the ecosystem than a lower level of
discharge occurring throughout the year. If unchecked, the contribution of nutrients and
contaminants from submarine groundwater discharge may lead to eutrophication and
degradation of estuarine ecosystems that are an essential and unique habitat for an
abundance of marine species.
This work makes a connection between land-based aquifer recharge and submarine saline
groundwater flow. Changes in the upland freshwater system induce saltwater motion with
the potential to affect the chemical composition of the adjoining seawater. Thus the
freshwater and saltwater systems are coupled and should be treated as one in studies of
coastal dynamics and total submarine groundwater discharge.
7.3. Future Directions
Several aspects of this work can be explored further to better understand coastal systems
and the effects of seasonal variability on the chemistry of the coastal ocean. First, the
seasonal motion of the freshwater-saltwater interface presented in this work has not been
well-studied previously. Hydrogeological and geophysical investigations that map the
two or three-dimensional position of the interface and the water table in time would
provide a means to analyze the interaction between system components in an actual,
rather than idealized, aquifer. Concurrent monitoring of precipitation, evapotranspiration,
and submarine groundwater discharge would further our understanding of the full
155
hydrologic system and the primary factors that force it. Monitoring of submarine
groundwater discharge with arrays of automatic seepage meters that continuously
measure discharge, inflow, and salinity over yearly cycles in order to obtain a salt and
fluid mass balance would further close the system and confirm the seasonal modeling and
field results presented here.
Continuous monitoring of SGD would also clarify the temporal pattern of discharge as
well as the time lag between recharge and discharge in an actual system. When
considering the effect of SGD on coastal waters, it is important to know the period of
time that discharge is occurring, whether the amount of time is equal to or much different
than the amount of time that seawater is flowing into the aquifer, and whether there are
times of year that only dispersive saline circulation occurs, without the seasonal
component. It is also useful to know when saline discharge is occurring relative to the
recharge pattern, and how the discharge pattern in real systems responds to non-uniform
recharge. For example, is the time lag in real systems longer than the 1-4 months in the
simulated aquifers? Does peak discharge occur in the middle of the discharge season or
closer to the beginning or end, or does it depend on the recharge pattern? Also, how do
the fresh and saline components of SGD respond to individual rain events, is there an
immediate response or is average recharge the major factor affecting flow? Is it possible
to predict the SGD and consequent chemical loading to a coastal system based on rainfall
and ET? These questions can be answered through continuous monitoring in the field as
well as further numerical modeling of idealized systems.
The modeling presented in this work, both idealized (Chapter Five) and specific to
Waquoit Bay (Chapter Six), can be expanded. The parameters used in the idealized
models can be explored further to encompass other aquifer types. In particular, the
dispersivity values should be increased to determine the effect of dispersion on seasonal
discharge. Anisotropy, heterogeneity, and three dimensions can also be introduced to the
theoretical cross-section. Similarly, the Waquoit Bay model can be expanded to three
dimensions and the aquifer characterized in more detail. Simulation of temporal forcing
on smaller timescales, such as tides, waves, and individual precipitation events can also
156
be added to more accurately depict actual coastal systems. This modeling work will
improve our general understanding of temporal effects on subsurface flow patterns at the
coast.
This study emphasizes the spatial and temporal variability of submarine groundwater
discharge. The composition of this discharge is likely similarly variable, as evidenced in
the radium analysis presented in Chapter 3. This has implications for both the use of
tracer techniques to track SGD and the prediction of chemical loading to coastal waters.
First, the heterogeneity in tracer content between the components of discharge is
important when calculating total SGD, and potentially useful for identifying the relative
amount of each component of discharge. For example, if the radium content and activity
ratios are significantly different in fresh, brackish, and saline porewater, and if these are
further variable with location, an accurate estimate of total SGD using one value of
radium activity is impossible. But if sufficient measurements are taken to understand this
variability, separation into the fresh, nearshore brackish, nearshore saline, and offshore
saline groundwater discharge is a possibility if several isotopes and activity ratios are
measured. This variability also likely exists for tracers other than radium, such as radon
and barium, and should be considered in any tracer study.
Chemical analysis of submarine groundwater discharge is essential for estimating and
predicting its effect on coastal ecosystems. This has been done in the past, but rarely in
conjunction with an in-depth field and modeling effort which fully characterizes the
system in both space and time. The chemical composition of each component of SGD
and an estimate of the amount of discharge is necessary to determine the flux of nutrients
and contaminants at the coast. In areas where seasonal variability is a factor,
measurements should be taken throughout the year. This is particularly important when
considering the seasonal saline discharge because the first water to discharge is likely that
which entered last, and the last water to discharge before the inflow season was likely the
first to flow in the previous year, and so had the longest residence time. Thus, the
chemical composition of discharge likely varies not only from season to season but
within the same discharge season as well. The temporal variation in the chemical content
157
of discharge is further variable by the seasonal changes in the activity of microbes and
the oxidation state of the top portion of the sea floor. Investigation of chemical and
microbial reactions on a small scale over time may be important for understanding the
full effects of seasonal chemical loading on coastal ecosystems.
The idealized numerical models in this study are generically applicable to a wide range of
coastal systems, but the field investigation is limited to one estuary with specific
characteristics. Extending the concepts presented here to sites in different climates, and
with diverse qualities such as the magnitude of tides and waves, aquifer properties, and
human influence will enable a more accurate assessment of the global implications of
chemical loading via submarine groundwater discharge. Climate has an effect on the
seasonal variability of the system as well as factors such as the type and activity of
organisms in coastal ecosystems and their susceptibility to the effects of nutrient and
contaminant loading. Tides and waves affect the dispersivity of the system, the
magnitude of mixing between fresh and saline groundwater, and the amount of saline
circulation. Comparison of high and low dispersivity systems will clarify the effect of
such mixing on the chemistry and amount of SGD. Aquifer properties have a very large
effect on the amount of SGD, chemical transport, and the pattern of discharge. For
example, in highly stratified aquifers, SGD may occur very far offshore where confined
aquifers meet the ocean. This has implications for effects on deeper marine species and
makes it difficult to predict and detect the discharge patterns. Aquifer hydraulic
conductivity and the proximity of land-based chemical loading to the coast affect the
amount and state of contamination that reaches coastal waters. A long transport time as a
result of low conductivity or long distances may allow the breakdown of contaminants by
chemical or microbial reaction before they are released into the ocean. Assessment of
areas with high and low levels of anthropogenic influence will reveal the effects of
human activities, with potential implications for the regulation of releases such as
fertilizer and septic into coastal aquifers.
Submarine groundwater discharge can significantly affect the chemical composition of
coastal waters. Seasonally-varying discharge is a compounding factor that may greatly
158
increase the effects on estuarine ecosystems. Further investigation is therefore essential to
ensure the viability of important coastal habitats throughout the world.
159
References
Cambareri, T. C., and E. M. Eichner (1998) Watershed delineation and ground water
discharge to a coastal embayment. Ground Water 36(4): 626-634.
Goni, M. A., and L. R. Gardner (2003) Seasonal dynamics in dissolved organic carbon
concentrations in a coastal water-table aquifer at the forest-marsh interface.
Aquatic Geochemistry 9(3): 209-232.
Johannes, R. E. (1980) The Ecological Significance of the Submarine Discharge of
Groundwater. Marine Ecology-Progress Series 3(4): 365-373.
Kim, G., K. K. Lee, K. S. Park, D. W. Hwang, and H. S. Yang (2003) Large submarine
groundwater discharge (SGD) from a volcanic island. Geophysical Research
Letters 30(21): 10.1029/2003GL018378.
Simmons, G. M. (1992) Importance of Submarine Groundwater Discharge (Sgwd) and
Seawater Cycling to Material Flux across Sediment Water Interfaces in Marine
Environments. Marine Ecology-Progress Series 84(2): 173-184.
Talbot, J. M., Kroeger, K.D., Rago, A., Allen, M.C., and Charette, M.A. (2003) Nitrogen
flux and speciation through the subterranean estuary of Waquoit Bay,
Massachusetts. Biological Bulletin 205: 244-245.
Valiela, I. (1995) Marine Ecological Processes. New York, Springer-Verlag.
Valiela, I., J. Costa, et al. (1990) Transport of Groundwater-Borne Nutrients from
Watersheds and Their Effects on Coastal Waters. Biogeochemistry 10(3): 177197.
Whiting, G. J., and K. L. Childers (1989) Subtidal advective water flux as a potentially
important nutrient input to southeastern U.S.A. saltmarsh estuaries. Estuarine,
Coastal, and Shelf Science 28: 417-431.
Younger, P. L. (1996) Submarine groundwater discharge. Nature 382(6587): 121-122.
160
Appendix A
Field Instrument Constructionand Calibration
A.1 Submerged Seepage Meter Construction
The forty submerged seepage meters (Figure A. 1) were constructed by cutting off the top
and bottom 20 cm of 55-gallon steel drums (approximately 60 cm in diameter) with an
angle-grinder. A 3/4-inchhole was drilled into the center of the end piece and fitted with a
brass barbed 3 /8 -inch nozzle. The nozzle was tightly connected to a large nut inside the
drum, with washers on either side. All parts were sealed with caulk and marine sealant to
prevent leaks. A 3-inch hole was then cut into the top of the meter to allow for
equilibration during placement. This was sealed during seepage meter use with a 3-inch
expandable well plug. Around the rim of the seepage meter, three /4-inch holes were
drilled 2 cm below the top, spaced equally apart. Two holes were fitted with a nut, 3-inch
bolt, and two washers, and the third was fitted with a 3-inch circular hook, nut, and bolt.
Each was sealed with caulk and marine sealant. The protrusions allow the seepage meters
to be placed an equal distance above the sea floor, and prevent them from sinking into the
bottom sediment. The hook allows for attachment of ropes, either between adjacent
seepage meters to maintain consistent separation distances or attached to a buoy. Safety
precautions were taken with buoy ropes to prevent diver entanglement: a strong Velcro
connection on the rope near the sea floor allows for a quick release. The seepage meters
were entirely coated with acrylic primer and a top paint coat to prevent rusting and
radium adsorption to the metal. Each was tested thoroughly for leaks.
161
The brass barbed fitting on each seepage meter is connected to a short length of 3/einch
cleat. rubber tubing with a hose clamp. The female end of a quickdsconnect fitting
(ColeParmer bswment Company) on the end of the tubing attaches to the male on the
end of the bag attachment. The seepage meter bags are 24x30-inch, 2 mil polyprqylene
autoclave bags (VWR International). They are attached to %-inch tubing with a plastic
hose clamp and rubber band around a plastic insert that keeps the tubing from collapsing.
The tubing is then connected to a 2-way barbed stop valve, and finally to the male quickdisconnect fitting. AII tubing junctions are sealed with hose damps. This configuration
allows for easy and Id-free bag auachment and detachment. Eighty bag attachments
were assembled so that all forty bags can be quickly exchanged during large-scale
seepage meter experiments. The air in each bag was moved before attachment to
prevent artificial flow measwments due to floating.
FSgure A.1. Submerged seepage meter. When in use, the seepage meter is fully
submerged, and a bag is attached to the center nozzle.
A.2 Intertidal Seepage Meter Construction
The intertidal seepage meters were designed to measure flow in and out of the sediment
in water depths too shallow for the submerged seepage meters described in Section A. 1.
An essential element in the design of submerged seepage meters is that zero hydraulic
gradient exists across the seepage meter wall. This is true if the seepage meter bag is fully
submerged, translating the hydraulic pressure in the seawater to the water within the
seepage meter, and if the head loss across the bag connection is negligible. If zero
gradient exists across the wall, artificial flow will not be induced, and only natural flow
will be measured. In order to maintain zero gradient across a seepage meter that is not
fully submerged, the pressure changes due to tidal rise and fall must be maintained within
the meter. This is accomplished by attaching a very large plastic bag to the side of an
open seepage meter just above the sediment. If water is able to flow freely between the
bag and the seepage meter, and if the seepage meter is open to the atmosphere,
equilibrium is maintained across the wall. The water therefore needs only to be deep
enough to reach the top of the opening and allow full submersion of the bag. If the tidal
rise and fall is measured while the seepage meter bag is connected, the inflow or
discharge can be calculated as the difference between the total change in volume of the
bag and the change in volume due to tides:
=aVmeasured
AVseepage
where
Ameter
-Ahtide
Ameter ,
(A.
1)
is the area of the seepage meter at the location of the water level. This is
illustrated in Figure A.2.
The intertidal seepage meters were constructed from aluminum trash barrels in two sizes.
Seepage meters 1-5 are 58 cm high, with a top diameter of 44 cm and a bottom diameter
of 38 cm. Meters 6-8 are 67 cm high, designed for slightly deeper water, with a top
diameter of 52 cm and a bottom diameter of 43 cm. The bottom of each barrel was
removed with metal cutters and the sharp edges smoothed. A 3/4-inchhole was drilled 7
cm from the bottom of each barrel and fitted with a /2-inch brass barbed nozzle. The
nozzle was secured to the drum with a large nut and two washers, which were then coated
with marine sealant to prevent leaks. A very short length of /2-inch inner diameter rubber
163
tubing was attached to the nozzle, with a %-inchfemale polycarbonate quick-disconnect
fitting (Cole Parmer Instrument Company). The maIe end was connected to a plastic
insert with rubber tubing, and the bag secured to that with a large hose clamp. The
seepage meter bags must hold a large volume of water due to changing tide, so Husky 3
mil, 83.8 cm x 1.22 m Contractor Clean-up bags were used.
Figure A.2. Intertidal seepage meter schematic.
The inside and outside of each seepage meter was marked vertically every centimeter for
easy tide level reading. A clear length of tubing with cotton in the bottom end was
secured next to the outside markings to dampen any wave effects and enable more
accurate water level readings. During use, each meter should be checked often to verify
that the inside and outside water levels match, ensuring zero hydraulic gradient across the
seepage meter wall. This equilibrium is essential for accurate measurement of seepage,
but is easily disturbed. Waves are a major source of this error, so intertidal seepage
meters of this design must be used in very calm waters. In Waquoit Bay, it is important
that the wind come from the north during sampling so that the land protects the water
from wind. A second source of disequilibrium across the seepage meter wall is mis-
estimation of the initid volume of water in the bags. If tide is rising, water will flow from
the bag into the meter to maintain equilibrium. If there is not enough water available,
however, the water inside the drum will be tca low, creating an artificial gradient. If the
tide is falling, water will flow from the drum into the bag, so the initial volume in the bag
should be smdl to prevent it from filling to capacity. The intertidal seepage meters are
pictured in Figure A.3, during the 2003 seepage meter transect experiment.
Figure A 3 . Intertidal seepage meters in use on 8/14/2003.
A.3 Salinity Grid, Porewater Samplers, and Refractometer Calibration
A grid was constructed by stapling wire fencing to a wooden frame in order to assess
small-scale salinity variation in near-surface porewater. Every other square was labeled
with a letter and number for reference (Figure A.4). A small volume of porewater was
removed from each labeled square with the plastic syringe sampler pictured in the center
of the grid. A washer has been glued to each syringe to prevent baywater From entering
the subsurface during sampling and to ensure consistent sampling depths. The porewater
samplers were also used during the 2003 transect experiment to measure near-surface
salinity next: to the nozzle of each intertidal seepage meter.
AA. Porewater sampler and grid.
The porewater samplers can only remove a small volume (3-5 mL)of water if baywater
is not to be drawn into the sampler. A refractometer is used to measure salinity because it
requires only a few drops of water. The salinity of the water in the seepage meter bags
was measured using a conductivity probe (Cole-Pmer Instrument Company), however.
Thus, intercalibration of the ~fractometerand conductivity probe is necessary if the
measurements are to be compared d i m 1y . A series of water samples over the full
salinity range was measured with both instruments during the grid porewater sampling on
July 21,2003 and the seepage meter transect experiment on August 14,2003. The
measurements on each date fall very nearly along a straight line (Figure AS), although
the lines differ slightly on each date due to re-calibration of the conductivity probe. The
equations for each line are estimated as:
c[ppt] = 0.7 179 +0.7556 c[mSlcm] ,
with an It2 value of 0.9953 for the July 2 1 calibration, and
c[ppt] = 0.6586 c[mSlcm] - I -1773 ,
with an R
' value of 0.9908 for the August 14 experiment. A theoretical equation
(Holzbecher 1998) for the =lationship between salinity and conductivity is given by:
c[ppt] = -3.83
+ 0.699 .c[mS/cm] ,
tA.4)
The slope of these lines is more important than the intercept since it is the salinity
difference, or gradient, that drives flow. The slopes of the calibrated lines differ from the
theoretical slope by 6%and 8 8 , which is within a reasonable range of emr considering
the measurement ermr introduced in obtaining field data.
Figure AS. Refraccometer measurement [ppt] vs conductivity probe measurement
[mSlcm] for a t h d c a l quation and water samples on 7/21/03 and 8/14/03.
168
Appendix B
FEFLOW Model Descriptions
B.1 Model Attributes and Parameters
The FEFLOW shell allows the user to select various options to control the simulation.
This includes the overall class of problem, time steps, iteration methods, free-surface
constraints, approximations, and upwinding options in addition to the initial and
boundary conditions and parameter values that define the problem. The problem
attributes selected in the six theoretical models (Section 5.2) and the Waquoit Bay model
(Section 6.4) are listed below under the menu that contains them for comparison to other
models. A more detailed description of each option can be found in the FEFLOW
documentation (Diersch 1998).
Problem Class:
-
Unsaturated or variably saturated media
Transient flow, transient mass transport
Vertical problem projection
Temporal and Control Data:
-
-
Automatic time step control via predictor-correcter schemes
Forward Euler/backward Euler time integration scheme
Error tolerance: lx10-3 applied to Euclidian L2 integral (RMS) norm
Maximum interations per time step = 12
Adaptive mesh error = lx10-2
A posteriori error estimator, Onate-Bugeda
No upwinding
Specific option Settings:
169
·
·
·
·
·
·
·
·
·
·
·
Neglect fluid viscosity dependencies
Default Boussinesq approximation applied to density coupling
Constant thermal expansion of fluid density
Lumped mass (rather than consistent mass)
Standard Gauss Quadrature
All free surfaces unconstrained
Richards flow equation used in head-based (standard) form
Evaluation of capacity term with chord slope scheme
Influence coefficient method with central weighting
Picard iteration scheme
Does not check capillary head and saturation errors
Flow Data:
Boundary conditions are discussed in Section 5.2.3 and 6.4.1.
Flow materials (see Table 5.1 for values of K, D, and thickness of Chapter Five
models, see Table 6.3 for Chapter Six model geometry and K values):
· Chapter Five models: anisotropy factor = 1; Chapter Six model: anisotropy
factor = 1 - 0.05 (Table 6.3)
·
·
Angle = 0
Density ratio = 0.020
·
Compressibility = lx10
4
m- 1
Transport Data:
Boundary conditions are discussed in Section 5.2.3 and 6.4.1.
Transport materials:
·
·
·
Aquifer thickness (width) = 1 m
Porosity = 0.32
Diffusion coefficient = lx10 - 9 m2 /s
Table B.1. Theoretical model mesh size and maximum time step (Chapter Five).
Model
# Nodes
1
299,992
597,638
0.5
thick, med K, low D
2
177,974
352,641
0.5
thin, med K, low D
3
47,767
93,532
1
thin, med K, high D
4
148,685
295,534
1
thick, med K, high D
5
100,088
198,818
1
thick, high K, high D
6
94,462
187,579
1.5
thick, low K, high D
W.B.
28,358
55,988
0.75
# Elements Maximum At [d]
170
Description
variable
B.2. Analysis of Model Output
Once pseudo-steady-state, or dynamic equilibrium, was reached during each simulation,
time steps were saved every 10 d over 365 d, with a 5 d step at the end. The FEFLOW
Budget Analyzer was used to confirm the total fluid mass balance by calculating the total
inward and outward flux through each boundary type (head or flux) over the year along
the entire border. The error in the total fluid balance for the Chapter Five models ranged
from 0.11% (model 3) to 0.70% (model 6). The mass balance in the Waquoit Bay model
was higher, 5%, possibly due to density instabilities. The flux across the constant-head
boundary (representing the sea floor) also gives the total saline circulation over the year,
since the only inflow across the boundary is saline, and the total flux out is equal to that
amount of saline circulation plus the freshwater discharge, which equals the inflow across
the flux boundary (and the landward head boundary in the Waquoit Bay model).
The temporal variation in discharge across the constant head (sea floor) boundary was
determined in the idealized models by writing velocity and salt concentrations to a file
every 30 d over the year at 200 points along two border segments: from the coastline to 8
m offshore, and from 8 m to 200 m offshore. The two lengths were chosen to provide
greater resolution where velocity and salinity change the most. In the Waquoit Bay
model, the values at each node were written to a file. This was possible because there
were fewer nodes in this model than in the Chapter Five models. The velocity at two or
three nodes (depending on the velocity and location of the third) to the left of the
shoreline was also determined at each 30 d time step using the fluid flux analyzer in order
to capture fresh outflow through the seepage face. The direction of flow at each point was
then determined, a positive value assigned for outflow, and negative for inflow. The
relative amount of fresh and saline fluid flux was calculated by multiplying the velocity
at each point by its percent salinity and the length (or area) between points. Time series
of head values along a line segment at one point in time were obtained in a similar
manner.
171
172
Appendix C
Well Logs
The Zinch diameter wells were installed by GZA Drilling, Inc. from an offshore drilling
vessel (Figure C,1) by driving a larger steel casing into the ground and flushing out the
cuttings approximately every 3 meters. The flushed material was collected in plastic bags
and the descriptions recorded. The depths are approximate due to compaction, mixing
within the casing, and estimation of the depth of the flushing instrument. The sediment
descriptions and depths of wells 1,4,5, and 8 rve listed here.
Mgure C.1, Offshore well drilling system. Outer metal casing is driven into the sediment
by r e p t d y dropping it using the rope, puiley, and winch. Aquifer material was flushed
out of the casing and collezted for the well logs.
Well 1
Location: 41°34.813 ' N, 70°31.474 ' W, at the head of Waquoit Bay
Depth [m]
surface
4
Description
pebbles, medium sand, fine material
finer sand, some pebbles, fine material
4.5
finer sand
6.4
fine to medium brown sand
8
fine to medium brown sand
9
fine to medium brown sand
9.1
coarser sand and pebbles
9.8
reddish brown fine sand and silt
10.3
reddish brown fine sand and silt
11
reddish brown fine sand and silt
11.5
gray silt, some very fine sand
12.2
gray silt, some very fine sand
12.5
gray silt, some very fine sand
12.8
gray silt, some very fine sand
13
gray silt, some very fine sand
13.4
gray silt, some very fine sand
13.7
gray fine sand and silt
Well 5
Location: 18 m south of Well 1, offshore
Depth [m]
surface- 1
1-1.5
Description
blackish gray fine sand
coarser gray sand, less black
1.5-2.5
medium-coarse
2.5-3.5
medium-coarse gray sand
3.5-4.3
medium-coarse gray sand
4.3-4.6
medium-coarse gray sand
4.6-5.2
gray fine sand with organics, mica
5.2-5.8
fine-coarse sand, brownish, fines
5.8-6.4
fine-coarse sand, brownish, fines
gray sand
174
6.4-7
medium-coarse brownish sand, some fines
7-7.5
medium-coarse brownish sand, some fines
8
medium-coarse brownish sand, some fines
8.2
fine mixed with coarse brown sand
8.5
mostly fine brown sand
9.1
fine brown sand
Well 8
Location: 36 m south of Well 1, offshore
Depth [m]
surface-0.6
Description
medium grain sand and fine organics
1.2
less black, more sandy: coarser medium grains
2.5
light gray sand, mostly medium grained but some coarse and fine
3.5
light gray sand, mostly medium grained but some coarse and fine
4.6
light gray sand with some bigger grains and gravel, quartz chips
5.5
medium-coarse sand with fines: brown and gray
6.1
medium-coarse sand with fines: brown and gray
6.7
medium-coarse (coarser than above) gray sand
7
7.3
8
medium-coarse brown sand with fines (big change in color)
very fine brown sand with some black, brown fines
gray and black fine sand, gray fines
8.5
gray and black fine sand (slightly finer than 8m depth), gray fines
9.1
gray and black fine sand (slightly finer than 8m depth), gray fines
Well 4
Location: 8.8 m north of Well 1, onshore
Depth [m]
0-13.0
Description
fine-medium brown sand
13.0-13.7
coarser band of sand
13.7-18.3
gray silt with very fine sand
18.3
reddish brown fine sand
175
176
Appendix D
Seepage Meter Data
D.1 Seepage Meter Flux [m/d]: Head of the Bay Experiments: August
1999 and July 2000
August, 1999
TIME:
Meter
1
2
3
4
5
6
7
8
9
10
11
12
13
8:51
10:51
12:40
14:31
16:25
18:27
Sample
0.000
0.077
-0.013
0.268
0.127
0.005
0.102
0.264
0.044
0.000
0.158
0.060
-0.006
Sample2
0.149
0.069
0.000
0.236
0.074
0.190
0.428
0.100
0.048
0.005
0.124
0.079
0.007
0.005
0.507
0.229
0.108
0.067
0.012
0.012
0.037
0.050
0.038
0.057
0.184
0.574
Sample3
0.154
Sample4
0.187
0.060
0.025
0.010
0.054
0.073
0.617
Sample5
0.180
0.094
0.027
0.000
0.034
0.049
-0.031
-0.033
0.013
0.000
0.117
0.036
0.032
0.023
0.534
0.362
0.054
0.063
0.004
0.000
0.058
0.076
0.049
0.000
0.013
0.438
Sample6
0.089
0.043
0.019
0.254
0.020
0.030
0.468
0.000
0.043
0.005
0.115
0.028
0.000
0.005
0.904
0.266
0.080
0.029
0.015
0.015
0.010
0.034
0.037
0.000
0.058
0.551
14
15
0.481
16
21
0.208
0.052
0.026
0.020
0.007
0.052
22
0.061
23
0.035
24
0.021
25
0.036
0.423
17
18
19
20
26
0.041
0.009
0.009
0.038
0.038
0.556
0.126
0.005
0.000
0.121
0.046
0.006
0.023
0.290
0.210
0.060
0.072
0.052
0.005
0.036
0.036
0.022
0.022
0.205
0.220
0.027
0.000
0.143
0.053
0.000
0.022
0.538
0.388
0.085
0.030
0.036
0.005
0.060
0.077
0.043
0.022
0.032
0.620
177
0.505
Average
0.127
0.064
0.011
0.130
0.058
0.064
0.446
0.071
0.030
0.002
0.130
0.050
0.006
0.016
0.542
0.277
0.073
0.048
0.023
0.007
0.042
0.056
0.037
0.020
0.088
0.471
0.163
27
28
29
30
-0.011
0.029
0.000
0.000
31
0.061
32
33
34
0.037
0.068
35
0.281
36
0.040
0.080
0.000
0.013
0.000
37
38
39
40
I
0.459
0.000
0.024
-0.061
0.013
0.046
0.046
0.069
0.065
0.148
0.118
0.000
0.007
0.007
0.250
0.005
0.029
-0.030
0.048
0.031
0.037
0.055
0.217
-0.005
0.614
0.176
0.000
0.058
0.000
0.000
0.025
-0.049
0.073
0.038
0.071
0.010
0.041
-0.005
-0.005
-0.005
0.005
0.491
0.020
0.036
0.235
0.038
0.399
0.005
0.051
0.005
0.010
0.006
0.006
0.005
0.330
0.040
0.498
0.005
0.000
0.005
0.005
0.045
0.501
0.226
0.001
0.020
-0.032
0.048
-0.023
0.084
0.060
0.010
0.024
0.023
0.252
0.009
0.014
-0.028
-0.015
0.020
0.369
0.004
0.003
0.007
July, 2000
TIME:
7:51
10:02
12:03
13:55
15:57
18:10
Meter
Sample 1
Sample 2
Sample 3
Sample 4
Sample 5
Sample 6
Average
1
0.293
0.570
0.067
0.456
0.137
0.116
0.860
0.274
0.015
0.030
0.338
0.227
0.159
0.123
0.199
0.049
0.569
0.062
-0.004
0.020
0.116
0.139
0.137
0.107
0.027
0.479
0.689
0.096
0.020
0.012
0.488
0.207
0.223
0.093
0.127
0.916
0.333
0.013
0.025
0.380
0.187
0.260
0.036
0.327
-0.025
0.307
-0.005
0.009
0.334
0.079
0.373
0.138
0.103
0.262
0.200
0.078
0.430
0.371
0.073
0.087
0.061
1.932
0.331
0.084
1.656
0.083
1.197
0.014
0.399
0.014
0.317
0.014
2
3
4
5
6
7
8
9
10
11
12
13
14
15
16
17
18
19
20
21
22
23
24
25
26
27
28
29
0.131
0.226
0.158
0.098
0.487
0.102
-0.012
0.009
0.015
0.066
0.138
0.074
0.026
0.661
0.980
0.095
0.067
0.069
0.030
0.412
0.058
0.042
1.813
0.003
0.274
0.016
0.289
0.012
0.021
0.033
0.021
0.021
0.194
0.052
0.067
0.233
0.296
0.095
0.605
0.060
0.156
0.019
0.060
0.214
0.013
0.075
0.502
0.055
-0.028
0.023
0.074
0.082
0.112
0.064
0.159
0.194
0.025
0.237
0.003
0.045
0.163
0.103
0.098
0.092
0.480
0.018
0.093
0.061
0.073
0.452
0.021
0.015
0.523
0.413
0.287
0.566
1.113
0.301
0.034
0.023
0.031
0.212
0.049
-0.002
0.019
0.086
0.069
0.325
0.082
0.039
0.548
0.694
0.244
0.016
0.051
0.373
0.081
-0.001
0.028
0.083
0.075
0.121
0.071
0.029
0.350
0.017
0.062
0.339
0.052
0.000
0.021
0.070
0.069
0.096
0.058
0.485
0.270
0.240
0.117
0.063
-0.008
0.020
0.074
0.083
0.155
0.076
0.026
0.521
0.729
0.216
0.032
178
--
--
30
31
32
33
34
35
36
37
38
39
40
0.020
0.312
0.135
0.,009
0.020
0.214
0.056
0.062
0.188
0.096
0.270
0.733
0.243
0.012
0.,018
0.011
0, 168
0.,001
0,121
0,307
0 ..565
0.455
0.015
0.020
0.126
0.009
0.029
0.018
0.130
0.019
0.002
0.098
0.058
0.612
0.785
0.757
0.013
0.015
0.005
0.047
0.045
0.105
0.065
0.207
broken
0.692
0.007
0.015
-
0.011
0.214
0.088
0.067
0.133
0.060
0.505
0.653
0.246
0.024
0.018
0.105
0.050
0.656
0.555
0.104
0.018
0.025
0.017
0.167
0.059
0.062
0.105
0.075
0.426
0.658
0.416
0.014
0.017
August 1999 and July 2000
shoreline
Qj
V.z
a
Q
Q
(a
D
0a
(ad
(a
(a
G
0a
Q
(a
a
(9
0
Figure D.1. Seepage meter numbering map for 1999 and 2000 head of the bay
experiments. Not to scale.
179
D.2 Seepage Meter Flux [m/d]: Single Transect Experiments: 2002 and
2003
Intertidal seepage meters, 2002
Distance South
of Well 1 [m]
Time *
7:12
8:44
11:42
13:50
-1.47
-0.9
-0.9
-0.25
0.24
0.22
0.06
-0.25
0.06
15:43
17:37
Average
0.07
-0.46
0.12
0.42
0.42
0.30
0.33
0.24
0.07
0.21
0.19
0.11
0.29
0.15
0.29
0.34
0.30
0.29
0.40
0.75
0.66
0.34
0.07
0.75
0.41
0.44
0.30
0.04
2
0.42
0.49
0.33
-0.08
2
0.58
0.63
0.27
0.14
data is only displayed here if inside and outside water levels matched
Submerged seepage meters, 2002
Meter
7:35
Sample 1
9:47
Sample 2
11:46
Sample 3
13:42
Sample 4
15:43
Sample 5
17:38
Sample 6
21
22
23
24
25
26
27
28
29
0.379
0.160
0.156
0.155
0.407
0.148
0.048
0.019
0.744
0.349
0.324
0.166
0.125
0.550
0.165
0.046
0.026
0.609
0.026
0.013
0.058
0.038
0.908
0.209
0.024
0.014
0.532
0.010
-0.015
-0.009
0.027
0.543
0.181
0.024
0.009
0.433
0.000
-0.015
-0.015
-0.005
0.194
0.172
0.009
0.005
0.506
-0.004
-0.012
0.000
-0.004
1.178
0.135
0.012
0.012
30
0.021
0.015
0.033
0.000
0.011
0.008
31
32
33
34
35
36
37
0.553
0.238
0.244
0.242
0.017
0.142
0.463
0.696
0.434
0.011
0.196
0.159
0.478
0.439
0.180
0.088
0.322
0.134
0.397
0.000
0.456
0.051
-0.005
0.258
0.119
0.429
0.309
0.424
0.000
-0.015
0.084
0.256
0.046
0.479
38
39
-0.005
0.023
0.005
0.011
-0.004
-0.009
-0.011
-0.005
40
0.005
0.016
0.000
0.000
180
0.534
0.090
0.075
0.059
0.036
0.630
0.168
0.027
0.014
0.015
0.221
0.216
0.000
0.240
0.513
0.150
0.078
0.190
0.157
0.195
0.328
-0.019
-0.015
-0.016
0.000
-0.008
0.001
0.000
0.008
0.005
0.000
Intertidal seepage meters, 2003
Time *
io ^
00
-1O
o
0
.
-1.34
-0.5
0.43
-0.4
0.41
-0.26
0.6
0.79
0.94
0.8
0.15
0.66
0.39
0.30
0.32
0.72
0.33
1.1
2
0.64
0.18
0.29
0.18
2.15
0.75
3.02
0.65
8 °7t
00
O
In
0.18
0.23
0.20
0.55
0.40
0.46
0.56
0.49
0.52
0.44
0.94
0.31
0.19
0.15
0.95
1.15
3.16
0.99
3.2
0.60
0.32
0.31
2.19
0.94
0.47
0.34
0.66
0.52
0.53
0.65
0.99
0.34
0.47
3.45
0.32
4.24
0.75
1.19
Submerged seepage meters, 2003
7:35
9:47
11:46
13:42
15:43
9
0.000
0.510
0.467
0.326
10
0.315
0.386
0.367
0.356
0.492
0.366
0.363
11
12
13
1.070)
0.420
0.316
0.286
0.179
0.068
0.020
14
0.107
0.065
0.010
0.038
15
0.065
0.382
-0.015
0.000
16
0.831
0.534
0.518
17
18
19
20
21
22
23
24
25
0.112
0.000
0.000
0.009
1.115
0.424
1.138
0.234
0.871
0.067
0.290
0.355
0.089
0.379
0.037
-0.005
-0.015
-0.010
0.552
0.019
0.220
0.019
0.199
-0.005
0.005
0.32
0.97
* intertidal seepage meters were moved throughout the experiment and sampled at different times
data is only displayed here if inside and outside water levels matched
Meter
U
17:38
19:04
Average
0.562
0.758
0.602
0.353
0.127
0.035
0.080
0.190
0.016
0.041
0.061
0.048
-0.005
0.030
0.035
0.070
0.573
0.566
0.647
0.041
-0.004
0.000
0.000
0.518
0.019
0.552
0.013
0.299
0.063
0.005
0.000
-0.005
0.557
-0.004
0.510
0.042
0.172
0.083
0.000
0.049
0.005
0.679
0.127
0.509
0.062
0.206
0.153
0.008
0.063
0.009
0.598
0.272
0.514
0.121
0.320
0.080
0.001
0.013
0.002
0.670
0.164
0.543
0.083
0.349
0.612
26
0.315
0.131
0.168
0.235
0.247
0.344
0.429
0.267
27
28
0.175
0.056
0.054
0.247
0.039
0.336
0.048
0.082
0.023
0.045
0.057
0.190
0.057
0.179
0.048
29
30
0.071
-0.005
0.059
0.005
0.015
-0.018
0.036
-0.005
0.038
-0.005
0.000
-0.005
0.000
0.009
0.031
-0.003
181
D.3 Seepage Meter Flux [m/d]: Washburn Island, 2000
Meter
9:46
Sample 2
11:45
Sample 3
13:46
Sample 4
11
0.112
0.072
0.092
12
13
0.145
0.079
0.111
0.193
0.110
0.155
0.111
0.124
0.114
0.167
0.115
0.164
0.110
0.140
0.123
0.079
0.188
0.185
0.174
0.286
0.178
0.100
0.227
0.218
0.167
0.180
0.266
0.188
0.127
0.124
0.223
0.216
0.190
0.226
0.340
0.209
0.178
0.156
0.197
0.257
0.196
0.136
0.165
0.169
0.147
0.265
0.070
0.140
0.185
0.154
0.201
0.170
0.199
0.191
0.084
0.126
0.234
0.223
0.193
0.212
0.177
0.211
0.114
0.236
0.238
0.216
0.274
0.468
0.198
0.222
0.152
0.280
14
15
16
17
18
19
20
21
7:43
Sample 1
22
23
24
25
26
27
28
29
30
182
15:49
Sample 5
17:42
Sample 6
0.122
0.051
0.173
0.181
0.177
0.159
0.238
0.159
0.122
0.099
0.123
0.142
0.138
0.195
0.156
0.162
0.153
0.182
0.163
0.173
0.239
0.172
0.135
0.140
0.156
0.141
0.188
0.208
0.219
0.215
0.238
0.155
0.188
0.086
0.171
0.197
0.067
0.208
0.193
0.137
0.174
0.111
0.187
0.210
0.169
0.209
0.242
0.169
0.209
0.103
0.190
D.4 Seepage Meter Flux [m/d]: Multiple Tidal Cycle Experiment, 2001
7/17/2001
#
11
12
13
14
15
16
17
18
19
20
21
22
23
24
25
26
27
28
Time
1 7:07
9:01
10:53
12:55
14:58
16:55
18:49
20:56
23:00
0.274
0.048
0.190
0.076
0.070
0.092
0.059
0.012
0.004
0.011
0.175
0.087
0.051
0.129
0.035
0.118
0.030
0.241
0.147
0.059
0.034
0.000
0.022
0.042
0.078
0.030
0.055
0.048
0.015
0.015
0.108
0.103
0.043
0.060
0.043
0.148
0.074
0.048
0.013
0.000
0.045
0.047
0.128
0.039
0.087
0.046
0.022
0.009
0.061
0.132
0.076
0.072
0.095
0.082
0.143
0.156
0.004
-0.004
0.015
0.086
0.136
0.058
0.113
0.056
0.031
0.066
0.190
0.146
0.091
0.081
0.078
0.091
0.156
0.183
0.004
0.000
0.000
0.127
0.188
0.078
0.137
0.125
0.068
0.102
0.190
0.112
0.061
0.068
0.077
0.105
0.185
0.188
0.005
0.000
0.005
0.086
0.176
0.045
0.109
0.175
0.053
0.134
0.196
0.066
0.039
0.043
0.023
0.083
0.041
0.085
0.176
0.179
0.085
0.169
0.194
0.132
0.196
0.216
0.044
0.125
0.026
0.106
0.115
0.046
0.126
0.235
0.023
0.035
0.016
0.080
0.156
0.039
0.110
0.202
0.083
0.114
0.047
0.143
0.184
0.050
0.186
0.253
0.060
0.017
0.024
0.075
7/18/2001
Time
7/17-
#
0:59
2:59
5:23
7:57
9:53
11:53
13:53
15:41
16:57
18:22
11
12
13
14
15
16
17
18
19
20
21
22
23
24
25
26
27
28
0.129
0.095
0.084
0.151
0.098
0.136
0.157
0.123
0.141
0.228
0.245
0.064
0.237
0.150
0.181
0.135
0.204
0.248
0.165
0.088
0.159
0.077
0.129
0.141
0.218
0.004
0.079
0.247
0.095
0.167
0.180
0.027
0.113
0.230
0.190
0.237
0.151
0.212
0.138
0.079
0.169
0.265
0.068
0.053
0.084
0.062
0.083
0.122
0.123
0.000
0.000
0.005
0.074
0.139
0.095
0.112
0.108
0.054
0.166
0.175
0.078
0.039
0.062
0.061
0.074
0.120
0.130
-0.004
0.000
0.021
0.105
0.158
0.126
0.108
0.124
0.051
0.124
0.159
0.098
0.092
0.076
0.077
0.083
0.117
0.142
0.000
0.006
-0.011
0.140
0.188
0.132
0.147
0.097
0.055
0.123
0.179
0.088
0.085
0.067
0.077
0.067
0.092
0.126
0.017
0.000
0.008
0.140
0.150
0.085
0.122
0.128
0.048
0.112
0.157
0.038
0.047
0.043
0.052
0.114
0.095
0.099
0.005
0.000
0.021
0.066
0.085
0.043
0.076
0.083
0.035
0.091
0.133
0.141
0.181
0.082
0.181
0.173
0.063
0.179
0.274
0.250
0.309
0.128
0.206
0.192
0.116
0.235
0.111
0.201
0.186
0.173
0.092
0.079
0.196
0.273
183
7/18
A1
0.110
0.072
0.090
0.067
0.107
0.150
0.155
0.008
0.001
0.013
0.106
0.156
0.080
0.130
0.120
0.046
0.114
0.189
Seepage Meter Distance South of Low Tide Mark, Diagram (not to scale)
.
Meter
Distance [m]
11
9.1
12
15.2
13
21.3
14
27.4
15
33.5
16
39.6
45.7
51.8
57.9
64.0
7.6
17
18
19
20
21
22
9.1
23
10.7
24
9.1
25
42.7
44.2
44.2
44.2
26
27
28
Transect E
shoreline
0
0
aO
.
184
I).5 Seepage Meter Flux [m/d]: Cluster Experiments: 1999
-
-
Seepage Meter 29
Seepage Meter 11
low tide
0.063
0.081
2h after low
2h before high
high
2hafterhigh
2h before low
Average
0.058
0.104
0.052
0.069
low tide
0.081
0.075
0.052
0.058
0.066
0.069
0.098
0.115
0.052
0.084
0.063
0.098
0.104
0.088
0.075
low tide
0.052
2h after low
0.086 2h before high
0.092
high
2hafterhigh
2h before low
Average
0.076
Seepage Meter 25
0.075
2h after low
0.081
2h before high
0.069
0.127
0.109 0.058
0.058
0.063 0.058
0.121 0.127
0.127 0.121
0.138
0.161
0.184
0.115
0.138
0.173
0.196
0.144
0.173
0.105
0.150
0.169
0.125
0.091
Seepage Meter 30
0.086 0.058
0.127
0.063
low tide
0.058
2h after low
0.098
2h before high
0.075
0.069
0.075
high
0.075
0.063
0.058
0.173
0.086
0.046 0.046
0.069 0.046 0.167
0.161
2 h after high 0.081
0.052
0.196
0.230
2 h after high 0.086
0.081
0.150
0.161
2h before low 0.092
0.075
0.098
0.062
0.135
0.154
2h before low 0.098 0.063
Average
0.079 0.062
0.075
0.094
0.098
high
Average
0.081
0.098
Seepage Meter 26
low tide
0(.075 0.052
2h after low
0.071
0.065
2h before low 0.092
0.058
0.069
0.081
0.072
0.078 0.059
0.111
0.058
0.092
0.092
Average
0.109 0.075
2h after low
2h before high
high
2hafterhigh
0.121
0.121
0.069 0.069 0.127
0.115 0.127 0.144
0.132
0.132
2h before low ().109 0.121
Average
0.092
low tide
0.086
2h after low
0.129
0.190
0.132
0.104
0.104
0.173
0.098
0.113
0.094
0.167
0.138
0.063
2h before high
0.121
0.132 0.138
high
0.190
0.253
0.184
0.276
2 h after high 0.282
0.259
0.242
0.265
2h before low 0.207
0.305
0.374
0.217
0.265
0.156
0.144
0.101 0.098 0.134
0.063
Seepage Meter 32
Seepage Meter 27
low tide
0.086
0.040
0.069
0.127
0.052 0.058 0.081
high
2 h after high (0.075 0.069 0.138
Average
2h after low
0.144
0.092
0.138
0.046
2h before low ().081 0.081
low tide
0.058
2h before high
0.069
0.052 0.023
high
2hafterhigh 0.086 0.104
0.058
2h before high
0.104
Seepage Meter 31
0.081 0.052
0.114
Average
0.203
0.242
0.202
0.305
0.177
Seepage Meter 28
low tide
0.075 0.086
TN
0.075
0.081
0.081
0.046
0.115
0.092
2 h after high 0.098 0.075
2h before low 0.092 0.132
0.109
0.132
0.092
0.104
0(.079 0.085
0.108
0.092
0.101
2h before high
high
Average
3
0.156
2h after low
0.052
* this data was obtained by Jonathan Lubetsky for his 1999 summer Undergraduate Research Opportunity
(UROP) project
185
References
Diersch, H. J. G. (1998) FEFLOW finite element subsurface flow and transport
simulation system - user's manual/reference manual/white papers. Release 4.9.
WASY Ltd, Berlin.
Holzbecher, E. (1998) Modeling Density-Driven Flow in Porous Media: Principles,
Numerics, Software. Berlin, Springer-Verlag.
186
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