River reversals into karst springs: A new mechanism for cave formation in
eogenetic karst aquifers
Jason Gulley, Jonathan B. Martin, Elizabeth J. Screaton, Paul J. Moore* Department of Geological Sciences
P.O. Box 112120 University of Florida
Gainesville, FL 32611-2120
*Now at: Exxon Mobil Exploration Company, 233 Benmar Dr.; Houston, TX 77060 Keywords: Karst, Springs, Dissolution, Conduits, Caves, Estevelle
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Most conceptual models of epigenic conduit development assume that conduits
sourcing karst springs form due to flow of undersaturated water from recharge to
discharge point. This process is not possible in springs sourced from aquifer matrix
porosity and fed by distributed recharge, such as in unconfined eogenetic aquifers.
Diffusely recharged water has a long residence time within the aquifer, and thus would
have equilibrated with the aquifer rocks prior to discharge to the conduits. Many springs
in the high matrix permeability upper Floridan aquifer lack discrete inputs of
undersaturated allogenic water in their recharge areas, thereby necessitating another
explanation for their formation. During flooding of the Suwannee River (north-central
Florida), recharge of highly undersaturated water into upper Floridan aquifer springs is
common and solution scallops oriented upstream on conduit walls suggests most
dissolution occurs during these flow reversals. During a single event at the Peacock
Spring system, flood water was capable of dissolving up to 3.4 mm of the conduit wall
rock reflecting a maximum dissolution rate of 1.6 x 10-4 m/d. Dissolution would occur as
flow reversals exploit and expand pre-existing features such as joints and paleo-water
table caves. Lack of speleothems in upper Floridan aquifer conduits has been used as
evidence that the caves formed in the phreatic zone; however, flooding would dissolve
any speleothems that may have formed during previous subaerial exposure. Flow
reversals reflect dissolution initiating from the downstream end of groundwater flow
systems and should be an important driver of dissolution in any karst aquifers with
outflow to surface water subject to flooding. Flow reversals would introduce dissolved
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organic carbon and oxygen into the groundwater, providing important energy sources for
cave ecosystems and altering redox chemistry of the aquifer water.
Introduction
Conduits form in low matrix-permeability epigenic limestone aquifers (Vacher
and Mylroie, 2002) when undersaturated allogenic runoff flows into discrete recharge
points such as sinkholes or swallets. This recharge dissolves the rock along flowpaths in
joints and bedding planes, thereby expanding these preferential flow paths into conduits
(Palmer, 1991). Fully mature conduits thus often link recharge and discharge points in
these systems.
In contrast, processes forming conduits remain poorly understood in high matrix
porosity limestone (termed eogenetic by Vacher and Mylroie, 2002). These limestones
typically occur in tropical marine settings and have not undergone burial diagenesis that
would occlude the primary depositional porosity and permeability. Because the high
permeability matrix allows rapid infiltration of recharge as diffuse flow through the
surface (e.g., Ritorto et al., 2009), point recharge at sinking streams is less common than
in telogenetic karst aquifers and generally only occurs where streams flow off confining
layers onto the carbonate aquifer (e.g., Screaton et al., 2004). Where sinking streams do
exist, their potential for focused dissolution is greatly diminished because of the large
volumes of water stored in the matrix porosity, which is commonly equilibrated with
carbonate minerals of the aquifer (Moore et al., 2009).
The lack of focused recharge water in eogenetic karst aquifers led to the proposal
that some caves form in phreatic zones through mixing of waters with different chemical
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compositions (Mylroie and Carew, 1990). Such mixing causes undersaturation with
respect to carbonate minerals because of the non-linear relationship of saturation state
with respect to calcite, coupled with a linear change in the composition of the mixed
water, during mixing (Dreybrodt et al., 2009; Mylroie and Carew, 1990; Plummer, 1975).
Undersaturation and dissolution also results from carbonic acid that forms as CO
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introduced to the water from biological processes in the vadose zone. The introduction of
CO2 has recently been suggested is more important to dissolution than mixing of waters
with different compositions (Moore et al., 2009; Whitaker and Smart, 2007). Because the
source of CO2 is from the vadose zone, dissolution tends to occur at the top of the water
table, thereby forming water table caves.
Little allogenic recharge occurs in the eogenetic karst of the Upper Floridan
aquifer, except at the boundary between where it is confined and unconfined.
Nonetheless, many of Florida’s springs discharge from laterally extensive phreatic
conduit systems (Florea and Vacher, 2007; Martin and Gordon, 2000). Because of their
distance from the coast, these conduits could not have formed from mixing of fresh and
sea water as has been proposed for caves in the eogenetic limestone of the Yucatan
(Smart et al., 2006) and the Bahamas (Mylroie and Carew, 1990). The lack of allogenic
recharge limits input of undersaturated water into pre-existing permeable zones at the
upstream end of the conduits. These conduits are assumed to have formed in the phreatic
zone and not been subject to past subaerial exposure because most lack speleothems,
unlike conduits in the Yucatan and the Bahamas. Dissolution has been proposed to occur
in the phreatic zone from “headward sapping” in which high permeability zones act as
low resistance drains and cause flow paths to converge and concentrate dissolution,
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further focusing flow and dissolution (c.f., Rhoades and Sinacori, 1941; White, 2001).
Water from the Upper Floridan aquifer is generally saturated with respect to calcite
(Martin and Gordon, 2000; Moore et al., 2009), and thus headward sapping is unlikely to
form the conduits found there.
In this paper, we use legacy data and new observations from the Suwannee River
watershed in north-central Florida to suggest that reversals of springs during flood events
could provide a mechanism to form or enlarge conduits. Spring flow reverses when river
stage increases faster than the hydraulic heads in the aquifer. Although backflooding of
air-filled caves has been observed in telogenetic karst regions (White and White, 1989),
and some Florida springs have been reported to reverse (Opsahl et al., 2007), the
importance of chemical processes such as dissolution during spring reversals has not been
evaluated. We establish that surface water is undersaturated with respect to calcite during
high discharge events and use observations of solution scallop direction in two water-
filled conduit systems to support that dissolution occurs during flow reversals. We collect
high-resolution specific conductivity records within two conduit systems and
geochemical data during reversal of one spring to document influx of highly
undersaturated water during flooding. These data allow an assessment of the magnitude
of dissolution by flood waters.
Study Locations
The Suwannee River watershed in north-central Florida is entirely underlain by
the Floridan Aquifer System (FAS), a thick sequence of limestone and dolomite that is
subdivided into the Upper Floridan aquifer (UFA) and, where they exist, a middle
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confining unit and the Lower Floridan aquifer (Miller, 1986). The Cody Scarp generally
marks the boundary between the confined and unconfined regions of the UFA and
separates the Northern Highlands and Gulf Coastal Lowlands physiographic areas (Fig.
1). In the Northern Highlands, the UFA is overlain by the siliciclastic Hawthorn Group,
which acts as a confining unit, and the Surficial Aquifer System. Water sources to the
Suwannee River include the Surficial Aquifer System and runoff, which provide tannic-
rich water due to organic matter contributions from wetlands. Downstream of the Cody
Escarpment, the Suwannee River watershed transitions to being sourced by the UFA,
including discharge from more than 100 springs (Rosenau et al., 1977; Scott et al., 2004).
These springs include 9 of Florida’s 27 first magnitude springs, which are defined as
having a discharge of > 2.8 m
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3/sec, (i.e., > 100 cfs; Meinzer, 1927). At baseflow, these
springs discharge water that is saturated with respect to calcite, reflecting equilibration
with the aquifer rocks.
Groundwater of the UFA has higher specific conductivity than surface water
because of the dissolution of carbonate minerals, but it has lower dissolved organic
carbon concentrations, and thus does not have the characteristic tannic stain of water
from the surficial aquifer system or surface water draining the Northern Highlands.
Differences in specific conductivity and staining between groundwater and surface water
are particularly strong during floods (e.g., Moore et al., 2009) providing natural tracers
that allow separation of flood water flowing off of the Northern Highlands from
groundwater of the UFA. Flooding is common in winter and spring from rainfall
associated with cold fronts and in late summer and fall from tropical storms (Grubbs and
Crandall, 2007). These floods frequently elevate river water levels that have their
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headwaters in the Northern Highlands above the hydraulic head of the unconfined UFA
(Martin and Dean, 2001; Martin et al., 2006; Ritorto et al., 2009). The changing hydraulic
gradients cause springs to reverse, as shown by whirlpools at spring vents and changes in
water levels and tannins in wells up to 4.8 km from the river (Crandall et al., 1999).
In this paper, we investigate flow and chemical composition of water at two
springs (Madison Blue and Peacock) that discharge within the Suwannee River watershed
in north-central Florida (Fig. 1). Madison Blue Spring is classified as a first magnitude
spring that contributes baseflow discharge of 2.0 to 3.9 m3/sec to the Withlacoochee
River via a short spring run (Rosenau et al., 1977; Scott et al., 2004). The Withlacoochee
River is a major tributary to the Suwannee River, and Madison Blue Spring is located
about 12 km upstream of the convergence of the two rivers (Fig. 1). Madison Blue
Spring reverses flow during floods up to several times per year (e.g., Fig. 2). More than 8
km of passages have been mapped in Madison Blue Spring (Fig. 2A). Additional
passages have been observed or explored but have not yet been surveyed.
Peacock Spring is located ~67 km downstream of Madison Blue Spring (Fig. 1)
and 2.3 km north of the Suwannee River. It lacks a conduit connection or a spring run to
the Suwannee River and is technically a group of water-filled sinkholes (karst windows)
that lead to 7.5 km of mapped conduits (Fig 2B). The Suwannee River periodically
floods, inundating the spring along a normally dry channel connecting the river and the
entrance to the conduit system. The channel contains a sill that restricts direct infiltration
of river water into conduits to times when river water elevation exceeds ~8 masl (Rick
Owen, Florida Department of Environmental Protection, Personal Communication).
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We have also made cave-diving observations of conduit wall morphology at
Madison Blue and Peacock Springs, as well as at two other locations, Little River and
Cow Springs, that were not sampled for this study. Little River Spring is located about
18 km downstream from Peacock Spring and Cow Spring is located about 3 km southeast
of Peacock and a few hundred meters north of the Suwannee River (Fig. 1).
Methods
Legacy flow data were collected by the Suwannee River Water Management
District (SRWMD) and the United States Geological Survey (USGS) for the
Withlacoochee River near Madison Blue Spring at Lee, Florida (USGS station 02319394)
and the Suwannee River near Peacock Spring near Luraville, Florida (USGS station
02320000). The Lee station is approximately 10 km downstream from Madison Blue
Spring and the Luraville station is approximately 3 km upstream from Peacock springs
(Fig. 1). Discharge data from Madison Blue Spring were provided by the USGS using
continuous velocity measurements from a current meter (USGS station 02319302).
Chemical composition data was also collected by the SRWMD for water discharging
from Madison Blue Spring and from the rivers at the Lee and Luraville stations. Similar
flow and chemistry data are unavailable for Peacock Spring. We used the chemistry data
to calculate calcite saturation indices (SIcal) using PHREEQC with the LLNL database
(Parkhurst and Appelo, 1999). We define SI values here as the log of the ion activity
product divided by the equilibrium constant for calcite dissolution reaction.
Two observation periods of spring reversals are reported here: one in fall 2008
and the other in spring 2009. During fall 2008, Tropical Storm Fay passed through the
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area causing a minor flood. This event was recorded by a Schlumberger Conductivity-
Temperature-Depth (CTD)-diver that was installed at the entrance to Peacock Spring to
make time-series measurements of specific conductivity (SpC) and temperature (T).
Specific conductivity and T were also monitored at 20-minute intervals during and
following major flooding in April 2009 with CTD-divers installed at the entrances to
Peacock and Madison Blue springs and at six locations within conduits sourcing these
springs (Fig. 2). CTD-divers were installed within the conduits at penetrations of 152 m
(Martz Sink), 610 m (Courtyard) and 1097 m (Back Section) in the Madison Blue Spring
conduits, and at penetrations of 214 m (between Pothole and Olsen sinks), 884 m
(Challenge Sink), and 1067 m (Distance Tunnel) in the Peacock Spring conduits. CTD-
divers have accuracies for T of
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+ 0.1º C and SpC of +1% of the measured value. The
CTD-divers record pressure to a maximum depth of 10 m of water, which was exceeded
during most of the flood and thus we have no data for water depths. CTD-divers were
also installed at the Lee and Luraville stations during spring 2009, but the flood covered
both sensors with sediment, preventing data collection.
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To complement the SpC and T data, water was collected six times (16 and 24
April, 1, 8, and 15 May, and 14 July 2009) during the April 2009 flood and its recession
at the Luraville station and from two sinkholes, Challenge and Orange Grove sinks, that
intersect Peacock Spring conduits (Fig. 2B, Table 1). Samples could not be collected at
Madison Blue Spring during the April 2009 flood because roads to the spring were
submerged, making the spring inaccessible. Water was collected by extending a PVC
tube from the banks to directly above the center of the sinkholes and to about 5 m from
the banks of the river. The tubing was connected to a peristaltic pump, which drew water
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into an overflow cup. The water was monitored for its SpC, T, dissolved oxygen (DO)
concentration, and pH using a calibrated YSI model 566 multi-parameter field meter, and
pumping continued until all values stabilized. Following stabilization, samples were
collected in PVC bottles for analyses of alkalinity and major element concentrations and
kept chilled until measurement. Samples for measurements of cation concentrations were
preserved with concentrated nitric acid. Alkalinity was titrated within a day of collecting
the samples using the Gran method (e.g., Drever, 1997) and the major element
concentrations were measured using a Dionex Model 500DX ion chromatograph (IC) in
the Department of Geological Science, University of Florida.
Most of the samples collected during the first three weeks of the flood had very
low solute concentrations and were at or near the detection limit of the IC. Consequently,
their charge balance errors are large, averaging around 17%. Charge balance errors
(CBEs) are less for samples collected during the flood recession, averaging around 4%.
We use PHREEQC with the LLNL database (Parkhurst and Appelo, 1999) to calculate
calcite saturation indices (SIcal) based on these data. Charge balance was alternately
forced on Ca and alkalinity concentrations to assess the impact of CBEs on calculated
calcite saturation index. Forcing charge balances changes the SI less than 1 SI unit for
most of the dilute samples with high CBEs and for most samples changes the SI less than
0.l SI unit.
The April 2009 flood destroyed the station at Madison Blue Spring, and Peacock
Spring is not gaged because it lacks a spring run. We thus estimate the rate and volume
of river water intruding into the springs by dividing the distance between CTD divers by
the time it took for flood water to pass the CTD-divers, as estimated from changes in SpC
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of the water. We assume river water flowed into the conduits during the time of
decreasing and sustained low conductivity and that springs began to discharge when the
SpC rose following the maximum flood elevation. To estimate the total amount of
recharge during a reversal, we assumed flow was maximum at the start of the reversal
and decreased linearly until the reversal stopped. We convert the flow rate to a volume
of water based on an estimated average conduit diameter of 3 meters, which is consistent
with conduit surveys (Michael Poucher, personal communication; Fig 2) as well as our
observations of the water-filled conduits. More accurate assessment of influx volumes
would require either continuous velocity measurements or head data from the conduits
and surface water, which were not available.
Results
SICal response to elevated river discharge. Legacy data demonstrate that the
Withlacoochee and Suwannee rivers have an inverse exponential relationship between
discharge and SIcal (Fig. 3). High river discharge approaches 500 m3/sec in the
Withlacoochee River at the Lee station and 1000 m3/sec in the Suwannee River at the
Luraville station. Water during these high flow events reach SIcal values of < -4 at both
stations.
Discharge data from Madison Blue Spring reflect frequent reversals (Fig. 4).
Over the period of record, the volume of backflow is around 7% of discharge from
Madison Blue Spring. Most water sampled from the spring falls within about 0.2 SI units
of saturation with respect to calcite, with the lowest values of SIcal only -0.6, much closer
to saturation than the < -4 values found for the river during floods (Fig. 3A). Spring
samples are not generally collected during peak flood times because of limited access to
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the spring, so that sampling times for the spring water shown in figure 4 do not
correspond to sampling times for the river shown in figure 3A.
Observational evidence for conduit enlargement during reversals. During cave
dives at Little River and Cow springs we observed well-developed scallops on conduit
walls where there are constrictions in the conduits (Fig. 5). Scallops were not observed
during cave dives at Peacock and Madison Blue Springs but no systematic search has
been made for them. Additionally, passages we have explored in these springs generally
lack significant constrictions and conduits tend to be larger with slower flow velocities
than at Little River or Cow springs.
Springs’ Response to Floods – SpCond and discharge. Tropical Storm Fay
raised the Suwannee River level from ~5.4 m to ~7.7 meters above sea level (masl) at the
Luraville station, below the ~8 masl threshold required to flood overland into the Peacock
Spring karst windows. Nonetheless, tannic water was observed in the conduit system and
the CTD-diver installed in the spring vent recorded a decrease of SpC of around 60
µS/cm during the flood (Fig. 6). This drop in specific conductivity occurred at the same
time as an increase in temperature of around 0.3º C. A nearly instantaneous drop in SpC
from ~400 µS/cm to 357 µS/cm occurred at the time of the maximum river stage. This
minimum SpC value occurred approximately 6 days after the maximum flood elevation
in the river.
During the April 2009 flood, the stage of the Withlacoochee River at the Lee
station rose from about 9.3 to 19.5 masl in 13 days and the Suwannee River at the
Luraville station rose from 5.71 to 14.24 m in 15 days, causing overland flow to Peacock
Spring (Fig. 7A). Recession from the flood peak required more than 6 weeks at both
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gauges. At the entrance to Peacock Spring, SpC initially increased from 400 to 414
μS/cm between 12:51 h and 15:51 h on 4 April as the flood reached the spring before
dropping to minimum values of 24 μS/cm at 23:51 h on 6 April at the entrance (Fig 7B).
Minimum SpC values varied among the CTD-divers installed in the conduits. These
values were not as low as at the entrance, and occurred several days after the minimum at
the entrance. Minimum values were 52 μS/cm at 18:01 h on 6 April at Distance Tunnel,
50 µS/cm at 16:20 on 9 April at Challenge Sink, and 46 µS/cm at 07:11 on 9 April
between Pothole and Olsen sinks. At Peacock Spring, the rate of the propagation of the
low conductivity water from the entrance CTD-diver to the CTD-diver farthest from the
entrance (Challenge Sink) indicates a flow rate of approximately 0.02 m/s. Specific
conductivity first increased gradually after the minimum was reached at all CTD-divers
until around 27 April
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when the rate of increase became greater. At Distance Tunnel, the
rapid increase in SpC continued to a maximum of about 350 µS/cm at 07:31 h on 3 May
when conductivity decreased again to the end of the record. The other CTD-divers do not
record a similar maximum, but do show an inflection as the increase in SpC slows.
There is considerably more structure to the SpC records at Madison Blue Spring
than at Peacock Spring (Fig. 7A). The CTD-diver at the entrance to Madison Spring
dropped from a background value of around 286 μS/cm to a flood value of around 50
μS/cm in 45 hours, between 00:03 h on 1 April 2009 and 21:18 h on 2 April 2009 (Fig.
7A). This drop in SpC occurred as the river stage increased from around 9.5 to 10.9
masl, elevations much below the flood peak at 19.5 masl. All of the other CTD-divers
record a similar rapid drop, although they lag the CTD-diver at the entrance. After the
minimum value was reached at the entrance, SpC continued to decrease at each CTD-
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diver, but at different rates. Specific conductivity at Back Section reached a minimum of
47 µS/cm on 8 April, and values at Courtyard fluctuated between around 45 and 60
µS/cm until 03:24 h on 13 April. Between this time and 03:00 h on 15 April, SpC at both
Courtyard and Back Section rapidly increased to 170 µS/cm over a period of
approximately 48 hours. At Martz Sink, SpC slowly decreased to a minimum value of
slightly less than 40 µS/cm on 6 April and then gradually increased to a value of about 85
µS/cm at 01:03 h on 17 April and approximately six hours later at 07:18 h on 17 April at
the entrance. After these sites reached a value of around 85 µS/cm, the conductivity SpC
increased rapidly to around 250 µS/cm at 15:50 on 17 April at Martz Sink and about six
hours later at 21:18 on 17 April at the Entrance. The increases in SpC were not
monotonic. A second, smaller reversal began 18 April at the entrance and Martz Sink.
These small reversals do not occur farther back in the conduit system at the Courtyard or
Back Section CTD-divers.
Peacock Springs - SICal response during reversal. Chemical compositions
measured on samples collected from Peacock Spring represent the first systematic
sampling of a spring through a flood reversal (Fig. 7B). At the peak of the April flood on
16 April, approximately 9 days after low conductivity water entered the conduit system at
Peacock Spring, the SIcal at Orange Grove and Challenge sinks were found to be around -
5.0, slightly lower than the value of -4.5 found for the Suwannee River (Fig. 6B; Table
1). Seven days later on 24 April the SIcal had increased to –3.3 at Challenge Sink, but
remained low at –4.6 at Orange Grove, while this value had increased to -2.7 in the river
water. SIcal values at both locations increased to around -1 on 1 May simultaneously with
the rapid increase in SpC at Orange Grove and Challenge sinks, while SIcal estimated for
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the river water was slightly lower at -1.5. After 1 May, the SIcal values slowly increased
to around –0.5 for both Challenge and Orange Grove sinks on 14 July, and the Suwannee
river value approached equilibrium with calcite with a value of SI
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cal of -0.2 (Table 1).
Discussion
The differences in chemical compositions of water sources to the Suwannee River
in the Northern Highlands and Gulf Coastal Lowlands affect the correlation between
discharge and SIcal (Fig. 3). The best correlation at our two sites between discharge and
SIcal (r2 = 0.86) exists for the Withlacoochee River near Madison Blue Spring (Fig 3A).
This site is located near the Cody Scarp (Fig. 1), and thus river water has had little
opportunity to react with carbonate minerals or mix with groundwater that has
equilibrated with carbonate minerals. In contrast, there is more scatter in the correlation
downstream at the Luraville station (r2 = 0.67) (Fig 3B). This station is located nearly 70
km downstream of the Cody Scarp and thus changes in saturation state are influenced by
variable contributions from the UFA. The scatter of SIcal with discharge at the Luraville
station reflects the control of groundwater over the SIcal of the river water. The fractional
volumes of these two sources can vary depending on runoff and the magnitude and
direction of the head gradient between the aquifer and the river.
The low SIcal values calculated from legacy data, along with low SIcal values
estimated from samples collected during flood events in karst windows far from the river
(Fig. 7B), indicate that dissolution will occur during flow reversals. Flow direction
during times of dissolution can be estimated based on the orientation of scallops on
conduit walls, which form by dissolution in back-eddy currents in turbulent flow and are
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strongly asymmetrical in the direction of flow (Bretz, 1942; Curl, 1974). The orientation
of the scallops at Little River (Fig. 5) and Cow springs indicate flow was into the cave
during times of dissolution as would be expected from the low SI
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cal values of river water
during flood times (Fig. 3). We have observed whirlpools forming over the conduit
entrance of Little River Spring during floods, reflecting rapid flow into the conduit. The
only known allogenic recharge to Cow Spring is during flooding and thus dissolution
would only occur then because matrix water in the Floridan Aquifer is at equilibrium
with calcite.
Mixing of flood water and pre-flood groundwater. Observations from the karst
window at Peacock Spring during flooding in August and September, 2008 caused by
Tropical Storm Fay indicate river flooding can affect regions of the aquifer that do not
have direct conduit or overland flow connections to the river (Fig. 6). Although the river
level did not exceed the sill elevation at Peacock Spring, the decrease in conductivity at
the conduit entrance approximately 6 days after the flood peak indicates dilute flood
water flowed into the cave. The 6 day lag for the decrease in SpC is about 3 times longer
than the time for the decrease in conductivity observed for the April 2009 flood in which
the sill depth was exceeded (Fig. 7B). The greater lag time when the sill was not flooded
reflects slower influx of flood water, probably through secondary permeability features
such as joints, rather than rapid influx of surface water. Although we have no samples
from flooding during Tropical Storm Fay, specific conductivity was measured to be 358
µS/cm. Assuming that specific conductivity can be used as a proxy for SIcal, we can
assess dissolution potential from a sample collected on 14 July 2009 at Challenge Sink
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which had an identical SpC and a SIcal of -0.58. Dissolution was likely if water during
Tropical Storm Fay had a similar undersaturation with respect to calcite.
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The sharp drop in SpC clearly indicates that flood waters entered conduits on 1
April at Madison Blue Spring and on 5 April at Peacock spring. The return of pre-flood
water to the conduits (resumption of groundwater outflow) is more difficult to estimate.
We interpret the rapid increase in conductivity that occurred on 13 April at Back Section
and Courtyard, 17 April at Martz Sink and Entrance of Madison Blue Spring, and on 27
April at Peacock (Fig. 7) to represent the time when mostly pre-flood ground water
returned to the conduits. Even after the pre-flood water enters the conduits, none of these
sites return to the background SpC values following the rapid rise, indicating the conduits
retain a fraction of flood water for at least six weeks after the flood peak (Fig 7). The
gradual increase in conductivity and the increase in the saturation state of the water prior
to this time could represent reactions of flood water with the aquifer rocks, mixing with
pre-flood ground water that had previously equilibrated with the aquifer rocks, or both.
Floodwater that had infiltrated the matrix porosity thus appears to discharge slowly to the
river allowing long periods of time to react with the aquifer material.
At Madison Blue Spring, changes in SpC following the flood event are similar at
Martz and the entrance, but the response at these two locations differs from the responses
at Back Section and Courtyard. These responses at Madison Blue Spring also differ from
responses at Peacock Spring where changes in conductivity are similar at all locations
throughout the flood, with only slight variations at Distance Tunnel during the recession.
The differences between Madison Blue Spring and Peacock Spring suggest that the
numerous entrances into the Peacock Spring conduit system (Fig. 2B) allow rapid and
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complete mixing of the conduit and flood waters once overland flow is established to the
river. In contrast, river water infiltrates and discharges primarily through the central
spring vent at Madison Blue Spring, and thus discharge of the flood water is controlled
by branching and constriction in the conduit as well as permeability variations in the
matrix rocks that receive flood waters. The more gradual return to background SpC
values at Courtyard and Back Section than at Entrance or Martz Sink suggests these
locations have less groundwater exchange and mixing with the matrix than at Entrance or
Martz Sink.
Estimated volumes of recharged flood waters. In the Madison Blue Spring
conduit system, the passage of the low conductivity water could be tracked at each CTD-
diver location. The drop in conductivity associated with the influx of flood water
occurred progressively later at CTD-divers located farther from the entrance. The time
lag of the drop in SpC with distance into the conduits reflects flow into the conduits at a
rate of around 0.03 m/s. Using an estimated average cross sectional area of the conduits
(roughly estimated from highly variable conduit cross-sections during cave diving
observations) of 7 m2, the initial intrusion of water into the conduits is estimated to be
around 0.18 m3/s. We assume the minimum value in SpC that occurred around April 8
indicates the final influx of flood water, and thus the spring was reversed for 7.5 days.
We use this length of time for reversals to estimate that about 5.8 x 104 m3 of water
flowed past the most distal CTD diver. Although we have no data for hydraulic head of
the groundwater during the flood, the river stage continues to increase through 10 April,
about 9 days after the initial drop in conductivity and thus the hydraulic gradient between
the river and groundwater may have continued to be reversed during that time.
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Alternatively, the gradient may flatten or reverse in the farthest part of the conduit around
8 April, when the SpC at the Back Section CTD-diver begins to increase (Fig. 7A). We
chose to use the SpC minimum on 8 April as representing the end of the reversal rather
than the river stage maximum on 10 April because it is the more conservative value.
In the Peacock Springs system we estimate river water was flowing into the
conduit for about 20.5 days based on the duration of low conductivity water in the system
and when SpC began its sharp increase on 21 April. The rate of the propagation of the
low conductivity water from the Entrance CTD-diver to the CTD-diver furthest from the
entrance (Challenge Sink) indicates an initial flow rate of 0.14 m3 s-1. With this flow
rate, the total volume of water to infiltrate past the deepest CTD was also around 1.2 x
105 m3. These estimates are likely to be minimum values because they include only
water that flowed through a single entrance to the point where the CTD-diver is farthest
from the entrance. The estimate thus neglects water that may have flowed into other
conduit branches or river water that may have intruded into the conduits at other
entrances. Furthermore, as shown by the response of SpC at Peacock Spring following
Hurricane Fay (Fig. 6), flow may also enter the conduits through matrix porosity and
fractures.
Estimates of dissolution rates. Samples collected during the April 2009 flood at
Peacock Spring allow order-of-magnitude estimates of the amount of dissolution that
might occur during spring reversals. Chemical compositions are similar for samples from
all locations at the peak of the flood on 16 April (Table 1). We thus use the SIcal value
for the sample from Challenge Sink, which is central to the Peacock conduit system, to
calculate dissolution, on the assumption it best represents water in the conduits. If all of
19
the undersaturated flood water reaches equilibrium with calcite of the UFA, the water
would have dissolved 4.52 mmol/L of calcite (4.44 mmol/L and 7.03 mmol/L of calcite if
charge balance is forced on calcium and alkalinity respectively). This estimate is a
maximum because it is unlikely that all the water equilibrated with calcite of the aquifer.
This estimated value of dissolved calcite was multiplied by the volume of water
calculated to flow into the conduit (1.2 x 10
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5 m3) to estimate the maximum volume of
limestone dissolved during the flood. Assuming a density and a molar volume of calcite
of 2,710 kg m3 and 36.934 cm3 mol-1, respectively (Robie et al., 1984) and a porosity of
30 percent for the Ocala Limestone, the amount of calcite dissolved was about 28.6 m3
(28.1 m3 or 53.4 m3 of calcite if charge balance was forced on calcium and alkalinity,
respectively) This volume of calcite was divided by conduit area (8.33 x 103 m2) using
the surveyed distance from Peacock Entrance to Challenge Sink of 884 m and an
estimated average conduit diameter of 3 m to convert dissolution to wall retreat. The
volume of calcite that could have been dissolved during the reversal equates to a wall
retreat rate of 3.4 mm (or an average value of 1.6 x 10-4 m/d). If charge balance is forced
on Ca, wall retreat rates are not significantly reduced. If charge balance is forced on
alkalinity, wall retreat is calculated to be 5.34 mm (2.6 x 10-4 m/d). This value is nearly
two orders of magnitude higher than wall retreat rates reported from the nearby sink-rise
system at Oleno State Park, Florida (Moore, 2009) and estimates of maximum wall
retreat in telogenetic conduits (3 x 10-6 to 3 x 10-7 m/d; Palmer, 1991). Our estimated rate
of retreat is likely to be a maximum since a fraction of the water would have reacted with
the surfaces surrounding porosity in the matrix rocks rather than the conduit walls and if
the flood water does not completely equilibrate with calcite. Alternatively, dissolution
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within the matrix is likely to weaken the matrix material thus allowing physical erosion
of the walls during rapid flow in floods, thus increasing wall retreat. Despite the degree
of uncertainty associated with these calculations, there is clearly the potential for
significant amounts of dissolution during spring reversals.
We suggest that spring reversals in eogenetic rock would dissolve more rock than
would occur during flood conditions at sink-rise systems. In sink-rise systems the amount
of dissolution that occurs during flood pulses is limited by relatively short conduit
residence times and because most water discharges from the spring before reacting to
equilibrium (e.g., Martin and Dean, 1999; Martin and Dean, 2001). In contrast, where
flood pulses cause spring reversals, water must flow to matrix porosity because there is
no spring outlet for the floodwaters. Once in the matrix, long residence times and large
surface areas of the matrix porosity would enhance the extent of chemical reactions,
increasing the amount of calcite that would dissolve. The growth of conduits from
reversals of springs should thus provide a powerful enlargement mechanism, and is also
probably more effective than the headwater-sapping hypothesis of Rhoades and Sinacori
(1941).
Origin of Florida Springs. Scallop direction and spring chemistry data support
the hypothesis that the springs in our study have been significantly enlarged by the input
of undersaturated water during reversals. During spring reversals, water likely exploited
previously-existing high permeability features, including vertical joints or caves formed
at older, lower water tables. Water-filled conduits have been mapped at distinct levels
across the Florida carbonate platform, and by their correspondence with marine terraces,
have been proposed to reflect the elevation of the water table at the time the caves formed
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(Florea et al., 2007). Their formation is hypothesized to be caused by the diffuse input of
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2 from the vadose zone during times of low sea level when the water table was also
low (c.f., Moore, 2009). As water tables increased with the Holocene rise in sealevel, the
paleo water-table caves would have been flooded, and isolated water-table caves could
have become connected to surface water as rivers incised into joints. Once this
connection was established, the laterally extensive paleo-water table caves would have
captured undersaturated water during periods of flooding, with subsequent enlargement
of high-permeability zones into conduits feeding the springs. This model would also
explain the lack of speleothems in most water-filled caves in Florida. Undersaturated
flood water would dissolve any speleothems that may have been formed if the caves were
exposed during times of lower sea level.
The enlargement of pre-existing joints or caves that formed at paleo-water tables
by reversing river water allows phreatic caves to be created at the spring entrance without
discrete allogenic inputs in upstream regions. High surface area of the porous rock
making up the Floridan aquifer would allow extensive water-rock interactions as water is
discharged from the conduit into the matrix (Moore, 2009) resulting in “spongework”
cave morphology described for porous karst aquifers (Palmer, 1991). Dissolution during
spring reversals may also occur in telogenetic aquifers with low matrix porosity. Low
matrix porosity of the telogenetic aquifers would limit exchange of water between the
conduit and the matrix, and thus flood waters may extend farther into the conduits along
pre-existing dissolutional voids. Because telogenetic aquifers are also characterized by a
large amount of allogenic recharge of undersaturated water, characteristics related to
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dissolution during reversals may be overprinted by standard mechanisms of dissolution
resulting from capture of undersaturated water in a downstream direction.
Summary and Implications
Most conceptual models of epigenic conduit development assume that conduits
sourcing karst springs form due to flow of undersaturated water from recharge to
discharge point, a process which is not possible in springs supplied primarily by
distributed recharge from aquifer matrix porosity such as in unconfined eogenetic
aquifers. Diffusely recharged water has a long residence time within the aquifer, and thus
would have equilibrated with the aquifer rocks prior to discharge to the conduits. Where
springs are subject to flow reversals during river floods, undersaturated flood water can
dissolve conduits from the spring entrance. Dissolution during reversals likely enlarges
pre-existing void spaces such as joints and horizontal caves that were likely formed at
paleo-water tables and subsequently connected to the river along joints. Previously, lack
of speleothems was used as evidence that the underwater caves in Florida formed below
the water table. The model proposed here indicates that the laterally-extensive horizontal
conduits of many of the underwater caves in Florida may have initially formed at the
water table similar to present-day water-table caves (Florea et al., 2007). These conduit
systems would have begun to function as springs when channel incision exposed a joint
that intersected the conduit or conduit roof collapse created a connection. The conduits
were later modified and enlarged by dissolution during spring reversals. We suggest that
spring reversals would also lead to dissolution of speleothems and thus obscure evidence
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of water-table formation and past subaerial exposure. Laterally extensive, horizontal
galleries are common in water-filled caves and may reflect an origin as water-table caves.
Consequently, they could be used to reconstruct changes in past water-table positions in
response to glacioeustasy and climate change (e.g., Florea et al., 2007).
Spring reversals also have implications for aquifer contamination and
geochemistry. Contaminant flow into karst systems is generally considered as originating
as flow into sinkholes and swallets. Spring reversals would provide another mechanism,
besides diffuse or swallet recharge, for the injection of water with distinct chemical
compositions. Floodwater chemistry typically has organic carbon and oxygen
concentrations elevated over those of groundwater. These differences in compositions of
floodwater and groundwater should lead to shifts in redox conditions within the aquifer
from initially oxic to anoxic conditions as the organic carbon is remineralized following
reversals. These changes in oxidation –reduction potential would influence diagenetic
reactions as well as calcite dissolution. For example, injection of organic carbon into the
aquifer may be an important energy source for ecosystems in the typically oligotrophic
environment of the aquifer. Reversals of springs would allow more time for microbes to
oxidize the organic carbon than would be typical in flow-through systems from sinks to
springs (e.g., Martin and Dean, 1999; Martin and Dean, 2001). Microbes use various
terminal electron acceptors such as oxygen, nitrate, and metal oxides in the oxidation of
organic carbon and thus these reactions should also influence nitrogen and metal
concentrations of the flood water. Understanding these processes will require detailed
time-series analyses of chemical composition of water as it flows into and from reversing
springs.
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Acknowledgements
We acknowledge staff at Suwannee River Water Management District for
supplying much of the geochemical data used in this paper. We thank cave divers Kelly
Jessop, Rick Crawford, Agnes Milowka, James Toland, Bill Huth for assistance with the
installation of CTDs and Jim Wyatt, Jeff Hancock and Wayne Kinard for additional
support. We thank Ken Clizbe and Bonnie Stelzenmuller for the photograph of Little
River Spring. Two Sisters BBQ in Mayo, Florida, is acknowledged for providing key
motivational support for fieldwork. We acknowledge support from the National Science
Foundation in the form of research grants EAR-0510054 and EAR-0910794, and for a
Graduate Research Fellowship to JG.
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Figure 1: The Suwannee River Basin, showing locations of springs (filled squares) and
gauging stations (filled circles) used in this study. The heavy black line represents the Cody Scarp, and marks the boundary between the confined (northeast) and unconfined (southwest) Floridan aquifer system. Inset map shows the location of the Suwannee River Basin relative to the outline of the state of Florida.
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Figure 2: (A) Conduit system at Madison Blue Spring. Breaks shown in the conduit walls to the west of Courtyard represent unexplored and unsurveyed conduits opening to the surveyed portion of the system. (B) Conduits system at Peacock springs. Locations of CTD-divers are represented by the stars. Depths at the floor and ceilings of the conduits are also shown. Peacock spring has several connections to the surface through sinkholes, but Madison Blue Spring is only connected at its entrance and a single sinkhole (Martz Sink).
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Figure 3: Calcite SI is plotted against discharge on the Withlacoochee River near (A) Lee, Florida (USGS 02319394) and (B) near Luraville, Florida (USGS 02320000) for SRWMD legacy data. Samples were collected approximately monthly from 2003 to 2008.
Figure 4: Discharge (solid line) and calcite saturation index (circles) at Madison Blue Spring. Positive discharge indicates outflow from the spring and negative discharge indicates river water flowing into the spring.
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Figure 5 – A cave diver swims upstream in the cave attached to Little River Spring. Solution scallops on the walls indicate formation during spring reversals, not during normal spring discharge. Photo courtesy of Ken Clizbe.
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Figure 6: Temperature (solid black line) and specific conductivity (solid gray line) within
the conduit at Peacock Spring versus discharge (dashed line) on the Suwannee River at Luraville during the flooding caused by Tropical Storm Fay.
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Figure 7: Specific conductance time series from CTD loggers installed in A. conduits at
Madison Blue Spring along with river stage on the Withlacoochee River at Lee (USGS 02319394), and B. conduits at Peacock Springs along with river stage on the Suwannee River at Luraville (USGS 02320000). Open square and filled
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circles represent the SIcal values measured from Challenge and Orange Grove sinks, respectively.
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1 2
Table 1. Major element concentrations from April 2009 flood and recession.
Date Sample pH Temp SpC Na K Mg Ca Cl SO4 Alkalinity Charge SIcal
