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River reversals into karst springs: a model for cave enlargement 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
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Abstract
Most conceptual models of epigenic conduit development assume that conduits
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sourcing karst springs form as water that is undersaturated with respect to carbonate
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minerals flows from recharge to discharge points. This process is not possible in springs
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fed by distributed recharge that is transmitted through aquifer matrix porosity, such as
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unconfined aquifers in eogenetic carbonate rocks. Diffusely recharged water has a long
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residence time within the aquifer, and thus would have equilibrated with the aquifer rocks
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prior to discharge to the conduits. The Upper Floridan aquifer (UFA) has high matrix
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permeability (~10-13 m2) and many springs lack discrete inputs of undersaturated
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allogenic water in their recharge areas. Consequently, another explanation for their
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development is necessary. During flooding of the Suwannee River in north-central
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Florida, water highly undersaturated with respect to carbonate minerals commonly
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recharges the UFA through spring vents, and solution scallops oriented away from the
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vents suggests most dissolution along conduit walls occurs during these flow reversals.
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During a single flow reversal at the Peacock Spring cave system, flood water was capable
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of dissolving up to 3.4 mm of the conduit wall rock. Dissolution occurs as flow reversals
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follow pre-existing features that include joints and paleo-water table caves. Lack of
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speleothems in conduits in the UFA has been used as evidence that the caves formed in
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the phreatic zone; however, flooding would dissolve any speleothems that may have
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formed during previous subaerial exposure. Conduit enlargement during flow reversals
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suggests that dissolution can progress in the normal upstream directions and this process
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may be an important driver of dissolution in any karst aquifer with outflows to surface
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water that are subject to flooding. Flow reversals would also introduce dissolved organic
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carbon and oxygen into the groundwater and provide important energy sources for cave
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ecosystems as well as altering redox chemistry of the aquifer water.
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Introduction
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Conduits commonly form in diagenetically mature carbonate aquifers with low
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matrix porosity and permeability (termed telogenetic karst by Vacher and Mylroie, 2002)
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when undersaturated allogenic runoff flows into discrete recharge points such as
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sinkholes or swallets. This recharge dissolves the rock along joints and bedding planes,
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thereby expanding these preferential flow paths into conduits (Palmer, 1991). Fully
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mature conduits thus often link recharge and discharge points in these systems.
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In contrast, processes forming conduits remain poorly understood in aquifers with
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high matrix porosity and permeability (termed eogenetic karst by Vacher and Mylroie,
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2002). These aquifers typically occur in tropical marine settings and have not undergone
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burial diagenesis that would occlude the primary depositional porosity and permeability.
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Because the high permeability matrix allows rapid infiltration of recharge as diffuse flow
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through the surface (e.g., Ritorto et al., 2009), point recharge (i.e. allogenic recharge) at
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sinking streams is less common than in telogenetic karst aquifers and generally only
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occurs where streams flow off confining layers onto the carbonate aquifer (e.g., Screaton
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et al., 2004). Where sinking streams do exist, their potential for focused dissolution is
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greatly diminished because of the large volumes of water stored in the matrix porosity,
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which is commonly equilibrated with carbonate minerals of the aquifer (Moore et al.,
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2009).
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Little allogenic recharge occurs in the eogenetic karst of the Upper Floridan
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aquifer (UFA), except for where streams flow off the edge of the confining layer and into
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the unconfined aquifer. Nonetheless, many of Florida’s springs discharge from laterally
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extensive phreatic conduit systems (Florea and Vacher, 2007; Martin and Gordon, 2000).
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Because of their distance from the coast, these conduits could not have formed from
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mixing of fresh and saline water as has been proposed for caves in the eogenetic
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limestone of the Yucatan (Smart et al., 2006) and the Bahamas (Mylroie and Carew,
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1990). The general lack of allogenic recharge limits input of water undersaturated with
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respect to carbonate minerals into pre-existing high-permeability zones at the upstream
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end of the conduits. These conduits are assumed to have formed in the phreatic zone and
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not been subject to past subaerial exposure because most lack speleothems, unlike
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conduits in the Yucatan and the Bahamas. Conduit dissolution has been proposed to
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occur in phreatic zone from “headward sapping” in which high permeability zones act as
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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).
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However, water from the UFA is generally saturated with respect to calcite (Martin and
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Gordon, 2000; Moore , 2009), and thus headward sapping is unlikely to form the conduits
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found there. Consequently, the origin of submerged eogenetic karst conduits remains
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unresolved.
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In this paper, we use legacy data and new observations from the Suwannee River
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watershed in north-central Florida to suggest that reversals of springs during flood events
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provide a mechanism to form or enlarge conduits. Spring flow reverses when river stage
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increases faster than the hydraulic heads in the aquifer. Although backflooding of air-
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filled caves has been observed in telogenetic karst regions (White and White, 1989) and
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some Florida springs have been reported to reverse (Opsahl et al., 2007), the importance
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of chemical processes such as dissolution during spring reversals has not been evaluated.
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We establish that surface water is undersaturated with respect to carbonate minerals
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during high discharge events and use observations of solution scallop direction in two
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water-filled conduit systems to support the concept that dissolution occurs during spring
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reversals. We collected high-resolution specific conductivity records within two conduit
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systems and geochemical data during reversal of one spring to document the influx of
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highly undersaturated water during flooding. These data allow an assessment of the
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magnitude of dissolution by flood waters.
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Study Locations
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The Suwannee River watershed in north-central Florida is entirely underlain by
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the Floridan Aquifer System (FAS), a thick sequence of limestone and dolomite that is
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subdivided into the UFA, a middle confining unit (where it exists), and the Lower
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Floridan aquifer (Miller, 1986). The Cody Scarp generally marks the boundary between
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the confined and unconfined regions of the UFA and separates the Northern Highlands
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and Gulf Coastal Lowlands physiographic areas (Fig. 1). In the Northern Highlands, the
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UFA is overlain by the confining siliciclastic Hawthorn Group and the Surficial Aquifer
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System. Water sources to the Suwannee River include the Surficial Aquifer System and
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runoff, which provide tannic-rich water due to organic matter contributions from
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wetlands. Downstream of the Cody Scarp, the Suwannee River watershed transitions to
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being sourced by the UFA, including discharge from more than 100 springs (Rosenau et
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al., 1977; Scott et al., 2004). These springs include 9 of Florida’s 27 first magnitude
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springs, which are defined as having a discharge of > 2.8 m3/sec, (i.e., > 100 cfs:
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Meinzer, 1927). At baseflow, these springs discharge water that is saturated with respect
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to carbonate minerals, reflecting equilibration with the aquifer rocks.
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Groundwater of the UFA has higher specific conductivity than surface water
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because of the high dissolved load of carbonate minerals but it has lower dissolved
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organic carbon concentrations, and thus does not have the characteristic tannic stain of
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water from the surficial aquifer system or surface water draining the Northern Highlands.
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Differences in specific conductivity and staining between groundwater and surface water
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are particularly strong during floods (e.g., Moore et al., 2009), providing natural tracers
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that allow separation of flood water flowing off of the Northern Highlands from
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groundwater of the UFA.
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Flooding is common in winter and spring from rainfall associated with cold fronts
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and in late summer and fall from tropical storms (Grubbs and Crandall, 2007). These
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floods frequently elevate river water levels that have their headwaters in the Northern
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Highlands above the hydraulic head of the unconfined UFA (Martin and Dean, 2001;
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Martin et al., 2006; Ritorto et al., 2009). The changing hydraulic gradients cause springs
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to reverse, as shown by whirlpools at spring vents and changes in water levels and
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tannins in wells up to 4.8 km from the river (Crandall et al., 1999).
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We investigate flow and chemical composition of water at two springs (Madison
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Blue and Peacock) that discharge within the Suwannee River watershed in north-central
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Florida (Fig. 1). Madison Blue Spring is classified as a first magnitude spring that
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contributes baseflow discharge of 2.0 to 3.9 m3/sec to the Withlacoochee River via a
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short stream connected to the spring vent known as a spring run (Rosenau et al., 1977;
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Scott et al., 2004). The Withlacoochee River is a major tributary to the Suwannee River,
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and Madison Blue Spring is located about 12 km upstream of the convergence of the two
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rivers (Fig. 1). More than 8 km of passages have been mapped in Madison Blue Spring
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(Fig. 2A). Additional passages are known but have not yet been surveyed.
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Peacock Spring, located about 67 km downstream of Madison Blue Spring and
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2.3 km north of the Suwannee River (Fig. 1), lacks a conduit connection or a spring run
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to the Suwannee River. The ‘spring’ is technically a group of water-filled sinkholes (karst
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windows) that lead to 7.5 km of mapped conduits (Fig 2B). The Suwannee River
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periodically floods and inundates the spring along a normally dry channel that connects
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the river to the entrance of the conduit system. The channel contains a sill that restricts
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direct infiltration of river water into conduits to times when river water elevation exceeds
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~8 masl (Rick Owen, Florida Department of Environmental Protection, Personal
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Communication).
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We have also made cave-diving observations of conduit wall morphology at
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Madison Blue and Peacock Springs, as well as at two other locations, Little River and
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Cow Springs, that were not sampled for this study. Little River Spring is located about
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18 km downstream from Peacock Spring and Cow Spring is located a few hundred
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meters north of the Suwannee River about 3 km southeast of Peacock (Fig. 1).
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Methods
Legacy flow data were collected by the Suwannee River Water Management
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District (SRWMD) and the United States Geological Survey (USGS) for the
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Withlacoochee River near Madison Blue Spring at Lee, Florida (USGS station 02319394)
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and the Suwannee River near Peacock Spring near Luraville, Florida (USGS station
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02320000). The Lee station is approximately 10 km downstream from Madison Blue
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Spring and the Luraville station is approximately 3 km upstream from Peacock Springs
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(Fig. 1). Discharge data from Madison Blue Spring were provided by the USGS from
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continuous velocity measurements from a current meter (USGS station 02319302).
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Chemical composition data was also collected by the SRWMD for water discharging
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from Madison Blue Spring and from the rivers at the Lee and Luraville stations. Similar
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flow and chemistry data are unavailable for Peacock Spring. We used the chemistry data
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to calculate calcite saturation indices (SIcal) using PHREEQC with the LLNL database
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(Parkhurst and Appelo, 1999). We define SI values here as the log of the ion activity
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product divided by the equilibrium constant for calcite dissolution reaction.
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Two observation periods of spring reversals are reported here: one in fall 2008
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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-
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Temperature-Depth (CTD)-diver that was installed at the entrance to Peacock Spring to
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make time-series measurements of specific conductivity (SpC) and temperature (T).
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In spring 2009, specific conductivity and T were also monitored at 20-minute
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intervals during and following major flooding in April 2009 with CTD-divers installed at
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the entrances to Peacock and Madison Blue springs and at six locations within conduits
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sourcing these springs (Fig. 2). CTD-divers were installed within the conduits at
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distances from the main spring vent of 152 m (Martz Sink), 610 m (Courtyard) and 1097
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m (Back Section) in the Madison Blue Spring conduits, and at distances of 214 m
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(between Pothole and Olsen sinks), 884 m (Challenge Sink), and 1067 m (Distance
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Tunnel) in the Peacock Spring conduits. CTD-divers have accuracies for T of + 0.1º C
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and SpC of +1%. The CTD-divers record pressure to a maximum depth of 10 m of
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water, which was exceeded during most of the flood and thus we have no data for water
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depths. CTD-divers were also installed at the Lee and Luraville stations during spring
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2009 but the flood covered both sensors with sediment and prevented any data collection.
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To complement the SpC and T data, water was collected six times (16 and 24
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April, 1, 8, and 15 May, and 14 July 2009) during the April 2009 flood and its recession
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at the Luraville station and from two sinkholes, Challenge and Orange Grove sinks, that
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intersect Peacock Spring conduits (Fig. 2B, Table 1). Samples could not be collected at
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Madison Blue Spring during the April 2009 flood because roads to the spring were
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submerged, making the spring inaccessible. Water was collected by extending a PVC
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tube from the banks to directly above the center of the sinkholes and to about 5 m from
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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)
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concentration, and pH using a calibrated YSI model 566 multi-parameter field meter, and
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pumping continued until all values stabilized. Following stabilization, samples were
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collected in PVC bottles for analyses of alkalinity and major element concentrations and
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kept chilled until measurement. Samples for measurements of cation concentrations were
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preserved with nitric acid. Alkalinity was titrated within one day of collecting the
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samples using the Gran method (e.g., Drever, 1997) and the major element concentrations
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were measured using a Dionex Model 500DX ion chromatograph (IC) in the Department
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of Geological Science, University of Florida.
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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,
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their charge balance errors are large, averaging around 17%. Charge balance errors
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(CBEs) are less for samples collected during the flood recession, averaging around 4%.
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We used PHREEQC (Parkhurst and Appelo, 1999) to calculate SIcal based on these data.
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Charge balance was alternately forced on Ca2+ and alkalinity concentrations by
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increasing or decreasing the concentration of Ca2+ and alkalinity within the calculations
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until charge balance was achieved to assess the impact of CBEs on calculated calcite
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saturation index. Forcing charge balances changes the SI less than 1 SI unit for samples
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with high CBEs and for most samples changes the SI less than 0.l SI unit. While an error
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margin of 1 SI is large, the sample with largest CBE was still had an upper calcite SI of
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nearly -4 and reflects that the water is capable of dissolving considerable amounts of
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calcite.
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While some data was lost due to equipment being damaged by flooding or
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logistical constraints, data that were collected clearly demonstrate large volumes of
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undersaturated river water flow into springs during springs reversals. Missing data
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include discharge at Madison Blue spring for the April 2009 flood, which destroyed the
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USGS discharge gauging station at Madison Blue Spring. Peacock Spring is not gauged
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because it lacks a spring run. We thus estimated the rate and volume of river water
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intruding into the springs by dividing the distance between CTD divers by the time it
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took for flood water to pass them, as estimated from changes in SpC of the water. We
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assume river water flowed into the conduits during the time of decreasing and sustained
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low conductivity and that springs began to discharge when the SpC rose following the
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maximum flood elevation. To estimate the total amount of recharge during a reversal,
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our calculations assume flow was maximum at the start of the reversal and decreased
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linearly until the reversal stopped. We convert the flow rate to a volume of water based
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on an estimated average conduit diameter of 3 m, which is consistent with our
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observations of the water-filled conduits. There are large variations in conduit diameter
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and cross-sectional morphology in both systems, 3 m is considered to be a rough average.
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More accurate assessment of influx volumes would require detailed conduit cross section
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measurements and either continuous velocity measurements or head data from the
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conduits and surface water, which were not available.
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Results
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SICal response to elevated river discharge. Legacy data demonstrate that the
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Withlacoochee and Suwannee rivers have an inverse exponential relationship between
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discharge and SIcal (Fig. 3). Highest river discharges approach 500 m3/sec in the
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Withlacoochee River at the Lee station and 1000 m3/sec in the Suwannee River at the
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Luraville station. Water during high flow events can reach SIcal values of < -4 at both
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stations across a wide range of discharges and most likely reflects the importance of
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antecedent aquifer heads as a control on river water chemistry.
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Discharge data from Madison Blue Spring reflect frequent reversals (Fig. 4).
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Over the period of record, the volume of backflow is around 7% of discharge from
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Madison Blue Spring. Most water sampled from the spring’s head pool is within about
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0.2 SI units of saturation with respect to calcite. During the summer of 2007, Madison
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Blue did not experience any reversals (Fig. 4) and SIcal was, with one exception, either
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saturated or slightly supersaturated. The lowest values of SIcal are reported from sampling
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trips near the end of a reversal and are only -0.6, much closer to saturation than the < -4
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values found for the river during floods (Fig. 3A). Samples during these times probably
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reflect a combination of river water and groundwater. Spring samples were not generally
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collected during peak flood times because of limited access to the spring, so that
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sampling times for the spring water shown in figure 4 do not correspond to sampling
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times for the river shown in figure 3A.
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Observational evidence for conduit enlargement during reversals. During cave
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dives at Little River and Cow springs we observed well-developed scallops on conduit
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walls where there are constrictions in the conduits (Fig. 5). Scallops were not observed
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during cave dives at Peacock and Madison Blue Springs but no systematic search has
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been made for them. In addition, the portions of passages we have observed in these
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springs generally lack significant constrictions and conduits tend to be larger with slower
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flow velocities than at Little River or Cow springs.
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Response of Springs to Floods - SpCond and discharge. The passing of Tropical
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Storm Fay resulted in the Suwannee River rising from ~5.4 m to ~7.7 meters above sea
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level (masl) at the Luraville station, which is below the ~8 masl threshold required to
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flood overland into the Peacock Spring karst windows. Nonetheless, we observed tannic
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water in the conduit system and the CTD-diver installed in the spring vent recorded a
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decrease of SpC of around 60 µS/cm during the flood (Fig. 6). This drop in SpC
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occurred at the same time as an increase in temperature of around 0.3º C. A nearly
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instantaneous drop in SpC from ~400 µS/cm to 357 µS/cm occurred at the time of the
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maximum river stage. This minimum SpC value occurred approximately 6 days after the
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maximum flood elevation in the river.
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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
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Luraville station rose from about 5.7 to 14.2 m in 15 days, resulting in overland flow to
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Peacock Spring (Fig. 7A). Recession from the flood peak required more than 6 weeks at
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both gauges.
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At the entrance to Peacock Spring, SpC initially increased from 400 to 414 S/cm
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between 12:51 h and 15:51 h on 4 April as the flood reached the spring before dropping
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to a minimum value of 24 S/cm at 23:51 h on 6 April at the entrance (Fig 7B). SpC
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decreased rapidly to minimum values over a few hours at each CTD location with the
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onset of the flood. Each minimum displays a small time lag related to increasing distance
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from the cave entrance. Minimum values of SpC were lowest at the entrance (24 S/cm)
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but were slightly higher with increasing distance from the entrance. SpC reached a
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minimum of 46 µS/cm at 07:11 on 9 April between Pothole and Olsen sinks (214 m from
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the entrance), 50 µS/cm at 16:20 on 9 April at Challenge Sink (884 m from the entrance)
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and 52 S/cm at 18:01 h on 6 April at Distance Tunnel (1067 m from the entrance). The
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rate of propagation of the low conductivity water from the entrance CTD-diver to the
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CTD-diver farthest from the entrance (Distance Tunnel) indicates a flow velocity of
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approximately 0.02 m/s.
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SpC records during flood recession at Peacock suggest complex interactions
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between pre and post flood waters in the conduit and matrix. There does not appear to be
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any consistent relationship between the rate of SpC rebound and distance from the
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entrance. SpC in the Distance Tunnel (1067 m from the entrance) rebounded most
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quickly, reaching a maximum of 352 S/cm at 07:31 h on 3 May before beginning a
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gradual decline. None of the other CTD-divers record similar maximum but they do show
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an inflection as the rate of increase in SpC slows.
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There is considerably more structure to the SpC records at Madison Blue Spring
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than at Peacock Spring (Fig. 7A). The CTD-diver at the entrance to Madison Blue
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Spring dropped from a background value of around 286 S/cm to a flood value of around
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50 S/cm in 45 hours, between 00:03 h on 1 April 2009 and 21:18 h on 2 April 2009
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(Fig. 7A). This drop in SpC occurred as the river stage increased from around 9.5 to 10.9
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masl, elevations much below the flood peak at 19.5 masl. All of the other CTD-divers
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record a similar rapid drop, although they lag the CTD-diver at the entrance.
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Specific conductivity at Back Section (1097 m from the entrance) reached a
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minimum of 47 µS/cm on 8 April, and values at Courtyard (610 m from the entrance)
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fluctuated between around 45 and 60 µS/cm until 03:24 h on 13 April. Between this time
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and 03:00 h on 15 April, SpC at both Courtyard and Back Section rapidly increased to
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170 µS/cm over a period of approximately 48 hours. At Martz Sink (152 m from the
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entrance), SpC slowly decreased to a minimum value of slightly less than 40 µS/cm on 6
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April and then gradually increased to a value of about 85 µS/cm at 01:03 h on 17 April
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and approximately six hours later at 07:18 h on 17 April at the entrance. After these sites
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reached a value of around 85 µS/cm, the SpC increased rapidly to around 250 µS/cm at
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15:50 on 17 April at Martz Sink and about six hours later at 21:18 on 17 April at the
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Entrance. A second, smaller reversal began 18 April at the entrance and Martz Sink but
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these small reversals do not occur farther back in the conduit system at the Courtyard or
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Back Section CTD-divers.
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Peacock Springs - SICal response during reversal. Chemical compositions
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measured on samples collected from Peacock Spring represent the first systematic
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sampling of a spring through a flood reversal (Fig. 7B). At the peak of the April flood on
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16 April, approximately 9 days after low conductivity water entered the conduit system at
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Peacock Spring, the SIcal at Orange Grove and Challenge sinks were found to be around -
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5.0, slightly lower than the value of -4.5 found for the Suwannee River (Fig. 6B; Table
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1). Seven days later on 24 April the SIcal had increased to –3.3 at Challenge Sink, but
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remained low at –4.6 at Orange Grove, while this value had increased to -2.7 in the river
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water. SIcal values at both locations increased to around -1 on 1 May, simultaneously
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with the rapid increase in SpC at Orange Grove and Challenge sinks, while SIcal
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estimated for the river water was slightly lower at -1.5. After 1 May, the SIcal values
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slowly increased to around –0.5 for both Challenge and Orange Grove sinks on 14 July,
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and the Suwannee river value approached equilibrium with calcite with a value of SIcal of
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-0.2 (Table 1).
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Discussion
The differences in chemical compositions of water sources to the Suwannee River
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in the Northern Highlands and Gulf Coastal Lowlands affect the correlation between
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discharge and SIcal (Fig. 3). The best correlation at our two sites between discharge and
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SIcal (r2 = 0.86) exists for the Withlacoochee River near Madison Blue Spring (Fig 3A).
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This site is located near the Cody Scarp (Fig. 1), and river water has had little opportunity
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to react with carbonate minerals or mix with groundwater that has equilibrated with
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carbonate minerals (e.g., Moore, 2009). In contrast, there is more scatter in the
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correlation downstream at the Luraville station (r2 = 0.67) (Fig 3B). This station is
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located nearly 70 km downstream of the Cody Scarp and thus changes in SIcal are
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influenced by variable contributions from the UFA. The scatter of SIcal with discharge at
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the Luraville station reflects the control of groundwater over the SIcal of the river water.
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The fractional volumes of these two sources can vary depending on runoff and the
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magnitude and direction of the head gradient between the aquifer and the river (e.g.,
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Moore et al., 2009).
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The low SIcal values calculated from legacy data, along with low SIcal values
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estimated from samples collected during flood events in karst windows far from the river
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(Fig. 7B), indicate that dissolution occurs as a result of the recharge of undersaturated
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river water during spring reversals. Flow direction during times of dissolution can be
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estimated based on the orientation of solution features such as scallops on conduit walls,
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which form by dissolution in back-eddy currents in turbulent flow and are strongly
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asymmetrical in the direction of flow (Bretz, 1942; Curl, 1974). The orientation of the
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scallops at Little River (Fig. 5) and Cow springs indicate that flow was into the cave
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during times of dissolution, as would be expected from the low SIcal values of river water
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during flood times (Fig. 3). We have observed rapid flow into conduits as whirlpools
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formed over the spring vent of Little River Spring during floods. The only known
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allogenic recharge to Cow Spring is from backflooding during high river stages.
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Dissolution would only occur during these backflooding events because matrix water in
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the UFA is at equilibrium with calcite.
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Mixing of flood water and pre-flood groundwater. Observations from the karst
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window at Peacock Spring during flooding caused by Tropical Storm Fay in August and
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September 2008 indicate river flooding can affect regions of the aquifer that do not have
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direct conduit or overland flow connections to the river (Fig. 6). Although the river level
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did not exceed the sill elevation at Peacock Spring, the decrease in conductivity at the
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conduit entrance approximately 6 days after the flood peak indicates dilute flood water
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flowed into the cave. The 6 day lag for the decrease in SpC is about 3 times longer than
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the time for the decrease in conductivity observed for the April 2009 flood in which the
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sill depth was exceeded (Fig. 7B). The greater lag time when the sill was not flooded
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reflects slower influx of flood water, probably through secondary permeability features
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such as joints, rather than rapid influx of surface water. Although we have no samples
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from flooding during Tropical Storm Fay, SpC was measured to be 358 µS/cm. Because
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groundwater in this part of the UFA is of the Ca-HCO3 type, SpC can be used as a
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qualitative proxy for SIcal (Krawczyk and Ford 2006). We can therefore estimate calcite
365
saturation from a sample collected on 14 July 2009 at Challenge Sink which had an
366
identical SpC of 358 µS/cm and a SIcal of -0.58. Dissolution was likely if water during
367
Tropical Storm Fay had a similar undersaturation with respect to calcite.
368
The sharp drop in SpC clearly indicates that 2009 flood waters entered conduits
369
on 1 April at Madison Blue Spring and on 5 April at Peacock spring (Fig. 7); however,
370
the return of pre-flood water to the conduits (return of groundwater outflow) is difficult to
371
estimate. We interpret the rapid increase in SpC that occurred on 13 April at Back
372
Section and Courtyard, 17 April at Martz Sink and Entrance of Madison Blue Spring, and
373
on 27 April at Peacock (Fig. 7) to represent the time when mostly pre-flood ground water
374
returned to the conduits. Even after the pre-flood water enters the conduits, none of these
375
sites return to the background SpC values following the rapid rise, indicating the conduits
376
retain a fraction of flood water for at least six weeks after the flood peak (Fig 7). The
377
gradual increase in SpC and the increase in SIcal of the water prior to this time could
378
represent reactions of flood water with the aquifer rocks, mixing with pre-flood ground
17
379
water that had previously equilibrated with the aquifer rocks, or both. Floodwater that
380
had infiltrated the matrix porosity thus appears to discharge slowly to the river, resulting
381
in long periods of time to react with the aquifer material.
382
At Madison Blue Spring, changes in SpC following the 2009 flood event are
383
similar at Martz and the entrance, but the response at these two locations differs from the
384
responses at Back Section and Courtyard. These responses at Madison Blue Spring also
385
differ from responses at Peacock Spring, where changes in conductivity are similar at all
386
locations throughout the flood with only slight variations at Distance Tunnel during the
387
recession. The differences between Madison Blue Spring and Peacock Spring suggest
388
that the numerous entrances into the Peacock Spring conduit system (Fig. 2B) allow rapid
389
and complete mixing of the conduit and flood waters once overland flow is established to
390
the river. In contrast, river water infiltrates and discharges primarily through the central
391
spring vent at Madison Blue Spring, and thus discharge of the flood water is controlled
392
by branching and constriction in the conduit, as well as permeability variations in the
393
matrix rocks that receive flood waters from the conduits. The more gradual return to
394
background SpC values at Courtyard and Back Section than at Entrance or Martz Sink
395
suggests that at Courtyard and Back Section the flood water has mixed more extensively
396
with the pre-flood water in the matrix and has returned to the conduits more slowly than
397
flood water in the front section of the cave. Extensive mixing and slow return to pre-
398
flood conditions could reflect decreasing flow velocities during the flood with distance
399
from the entrance.
400
401
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-
18
402
diver location. The drop in SpC associated with the influx of flood water occurred
403
progressively later at CTD-divers located farther from the entrance. The time lag of the
404
drop in SpC with distance into the conduits reflects flow into the conduits at a rate of
405
around 0.03 m/s, similar to the 0.02 m/s rate estimated for Peacock Spring. Using an
406
estimated average cross-sectional area of the conduits (roughly estimated from highly
407
variable conduit cross-sections observed during cave-diving) of 7 m2, the initial intrusion
408
of water into the conduits is estimated to be around 0.18 m3/s. We assume the minimum
409
value in SpC that occurred around April 8 indicates the final influx of flood water, and
410
thus the spring was reversed for 7.5 days. We use this length of time for reversals to
411
estimate that about 5.8 x 104 m3 of water flowed past the most distal CTD-diver.
412
Although we have no data for hydraulic head of the groundwater during the flood, the
413
river stage continued to increase through 10 April, about 9 days after the initial drop in
414
SpC and thus the hydraulic gradient between the river and groundwater may have
415
continued to be reversed during that time. Alternatively, the gradient may flatten or
416
reverse in the farthest part of the conduit around 8 April, when the SpC at the Back
417
Section CTD-diver begins to increase (Fig. 7A). We chose to use the SpC minimum on 8
418
April as representing the end of the reversal rather than the river stage maximum on 10
419
April.
420
In the Peacock Springs system we estimate river water was flowing into the
421
conduit for about 20.5 days during the 2009 flood based on the duration of low
422
conductivity water in the system and when SpC began its sharp increase on 21 April. The
423
rate of the propagation of the low conductivity water from the Entrance CTD-diver to the
424
CTD-diver farthest from the entrance (Challenge Sink) indicates an initial flow rate of
19
425
0.14 m3 s-1. With this flow rate, the total volume of water to infiltrate past the deepest
426
CTD was also around 1.2 x 105 m3. These estimates are likely to be minimum values
427
because they include only water that flowed through a single entrance to the point where
428
the CTD-diver is farthest from the entrance. The estimate thus neglects water that may
429
have flowed into other conduit branches or river water that may have intruded into the
430
conduits at other entrances. Furthermore, as shown by the response of SpC at Peacock
431
Spring following Hurricane Fay (Fig. 6), flow may also enter the conduits through matrix
432
porosity and fractures.
433
Estimates of dissolution rates. Samples collected during the April 2009 flood at
434
Peacock Spring allow order-of-magnitude estimates of the amount of dissolution that
435
might occur as a result of spring reversals. Chemical compositions are similar for
436
samples from all locations at the peak of the flood on 16 April (Table 1). We thus use the
437
SIcal value for the sample from Challenge Sink, which is central to the Peacock conduit
438
system, to calculate dissolution on the assumption it best represents water in the conduits.
439
If all of the undersaturated flood water reaches equilibrium with calcite, the water would
440
have dissolved 4.52 mmol/L of calcite (4.44 mmol/L and 7.03 mmol/L of calcite if
441
charge balance is forced on Ca2+ and alkalinity, respectively). This estimate is a
442
maximum because it is unlikely that all the water equilibrated with calcite of the aquifer,
443
but it does provide an estimate of the amount of dissolution that could occur from
444
flooding events. This estimated value of dissolved calcite was multiplied by the volume
445
of water calculated to flow into the conduit (1.2 x 105 m3) to estimate the maximum
446
volume of calcite dissolved during the flood. Assuming a density and a molar volume of
447
calcite of 2,710 kg m3 and 36.934 cm3 mol-1, respectively (Robie et al., 1984) and a
20
448
porosity of 30% for the Ocala Limestone (Budd and Vacher, 2004), the amount of
449
bedrock dissolved was about 28.6 m3 (28.1 m3 or 53.4 m3 if charge balance was forced on
450
Ca2+ and alkalinity, respectively).
451
The amount of wall retreat can be estimated from these estimated amounts of
452
dissolution. The conduit area is estimated to be 8.33 x 103 m2 using the surveyed distance
453
from Peacock Entrance to Challenge Sink of 884 m and an estimated average conduit
454
diameter of 3 m. Given these values, the volume of bedrock that could have been
455
dissolved during the reversal equates to a maximum wall retreat of 3.4 mm (or an average
456
value of 1.6 x 10-4 m per day of reversal). Because we lack sufficient constraints on the
457
residence time of river water in the conduits and matrix to use a kinetic model to refine
458
our wall retreat estimates, we adopt the term ‘meters per day of reversal’ to distinguish
459
our estimates from direct measurements of wall retreat rates (cf. Palmer, 1991). While the
460
amount of wall retreat will be overestimated using these boundary conditions, these
461
calculations show that the amount of retreat can be substantial. If charge balance is forced
462
on Ca2+, wall retreat calculations are not significantly reduced. If charge balance is
463
forced on alkalinity, wall retreat is calculated to be 5.34 mm (2.6 x 10-4 m per day of
464
reversal). This value is nearly two orders of magnitude higher than wall retreat rates that
465
were estimated from the nearby sink-rise system at Oleno State Park, Florida (Moore,
466
2009) and estimates of maximum wall retreat in conduits in telogenetic aquifers (3 x 10-6
467
to 3 x 10-7 m/d: Palmer, 1991). Our estimated rate of retreat is likely to be a maximum
468
since a fraction of the water would have reacted with the surfaces surrounding porosity in
469
the matrix rocks rather than the conduit walls and if the flood water does not equilibrate
470
with calcite. Alternatively, dissolution within the matrix is likely to weaken the matrix
21
471
material thus allowing physical erosion of the walls during rapid flow in floods to
472
increase wall retreat (Moore, 2009). Despite the degree of uncertainty associated with
473
these calculations, there is clearly the potential for significant amounts of dissolution
474
during spring reversals.
475
We suggest that spring reversals in eogenetic rock would dissolve more rock than
476
would occur during flood conditions at sink-rise systems. In sink-rise systems the amount
477
of dissolution that occurs during flood pulses is limited by relatively short conduit
478
residence times and because most water discharges from the spring before reacting to
479
equilibrium (e.g., Martin and Dean, 1999; Martin and Dean, 2001). In contrast, where
480
flood pulses cause spring reversals, water must flow to matrix porosity because there is
481
no spring outlet for the floodwaters. Once in the matrix, long residence times and large
482
surface areas of the matrix porosity would enhance the extent of chemical reactions,
483
increasing the amount of calcite that would dissolve (e.g., Moore, 2009). The growth of
484
conduits from reversals of springs should thus provide a powerful enlargement
485
mechanism, and is also probably more effective than the headwater-sapping hypothesis of
486
Rhoades and Sinacori (1941).
487
Origin of Florida Springs. Scallop direction and spring chemistry data support
488
the hypothesis that the springs in our study have been significantly enlarged by the input
489
of undersaturated water during spring reversals. During spring reversals, water likely
490
exploited previously-existing high permeability features, including vertical joints or caves
491
formed at older, lower water tables. Water-filled conduits have been mapped at distinct
492
levels across the Florida carbonate platform, and by their correspondence with marine
493
terraces, have been proposed to reflect the elevation of the water table at the time the
22
494
caves formed (Florea et al., 2007). We hypothesize these caves are formed from the
495
input of soil CO2 that diffuses through high matrix permeability and joints in the vadose
496
zones during times of low sea level when the water table was also low (cf. Moore, 2009).
497
As water tables increased due to late Pleistocene sea level rise, the paleo water-table
498
caves would have been flooded, and isolated water-table caves could have become
499
connected to surface water as rivers incised into joints. Once this connection was
500
established, the laterally extensive paleo-water table caves would have captured
501
undersaturated water during periods of flooding, with subsequent enlargement of high-
502
permeability zones into conduits feeding the springs. This model would also explain the
503
lack of speleothems in most water-filled caves in Florida, since undersaturated flood
504
water would dissolve any speleothems that may have been formed if the caves were
505
subaerially exposed during times of lower sea level.
506
The enlargement of pre-existing joints or caves that formed at paleo-water tables
507
by reversing river water allows phreatic caves to be created at the spring entrance without
508
discrete allogenic inputs in upstream regions. High surface area of the porous rock
509
making up the Floridan aquifer would allow extensive water-rock interactions as water is
510
discharged from the conduit into the matrix (Moore, 2009) resulting in “spongework”
511
cave morphology described for porous karst aquifers (Palmer, 1991). Dissolution during
512
spring reversals may also occur in telogenetic aquifers with low matrix porosity. Low
513
matrix porosity of the telogenetic aquifers would limit exchange of water between the
514
conduit and the matrix and flood waters may extend farther into the conduits along pre-
515
existing dissolutional voids.
516
Summary and Implications
23
517
Most conceptual models of epigenic conduit development assume that conduits
518
sourcing karst springs form by the flow of undersaturated water from recharge to
519
discharge points, a process which is not possible in springs supplied primarily by
520
distributed recharge from aquifer matrix porosity such as in unconfined eogenetic
521
aquifers. Diffusely recharged water has a long residence time within the aquifer, and thus
522
would have equilibrated with the aquifer rocks prior to discharge to the conduits. Where
523
springs are subject to flow reversals during river floods, undersaturated flood water can
524
dissolve conduits from the spring entrance. Dissolution during spring reversals enlarges
525
pre-existing void spaces such as joints and horizontal caves that were likely formed at
526
paleo-water tables and subsequently connected to the river along joints. Previously, lack
527
of speleothems was used as evidence that the underwater caves in Florida formed below
528
the water table. The model proposed here indicates that the laterally-extensive horizontal
529
conduits of many of the underwater caves in Florida may have initially formed at the
530
water table similar to present day water-table caves (Florea et al., 2007). These conduit
531
systems would have begun to function as springs when channel incision exposed a joint
532
that intersected the conduit or conduit roof collapse created a connection. The conduits
533
were later modified and enlarged by dissolution during spring reversals. We suggest that
534
spring reversals would also lead to dissolution of speleothems and thus obscure evidence
535
of water-table formation and past subaerial exposure. Laterally extensive, horizontal
536
galleries are common in water-filled caves and may reflect an origin as water-table caves.
537
Consequently, they could be used to reconstruct changes in past water-table positions in
538
response to glacioeustasy and climate change (e.g., Florea et al., 2007).
24
539
Spring reversals also have implications for aquifer contamination and
540
geochemistry. Contaminant flow into karst systems is generally considered as originating
541
as flow into sinkholes and swallets. Spring reversals would provide another mechanism,
542
besides diffuse or swallet recharge, for the injection of water with distinct chemical
543
compositions. Floodwater chemistry typically has organic carbon and oxygen
544
concentrations elevated over those of groundwater. These differences in compositions of
545
floodwater and groundwater should lead to shifts in redox conditions within the aquifer
546
from initially oxic to anoxic conditions as the organic carbon is remineralized following
547
spring reversals. These changes in redox conditions would influence diagenetic reactions
548
as well as calcite dissolution. For example, injection of organic carbon into the aquifer
549
may be an important energy source for ecosystems in the typically oligotrophic
550
environment of the aquifer. Reversals of springs would allow more time for microbes to
551
oxidize the organic carbon than would be typical in flow-through systems from sinks to
552
springs (e.g., Martin and Dean, 1999; Martin and Dean, 2001). Microbes use various
553
terminal electron acceptors such as oxygen, nitrate, and metal oxides in the oxidation of
554
organic carbon and thus these reactions should also influence nitrogen and metal
555
concentrations of the flood water. Understanding these processes will require detailed
556
time-series analyses of chemical composition of water as it flows into and from reversing
557
springs.
558
559
560
Acknowledgements
25
561
We acknowledge staff at Suwannee River Water Management District for
562
supplying much of the geochemical data used in this paper. We thank cave divers Kelly
563
Jessop, Rick Crawford, Agnes Milowka, James Toland, Bill Huth and Tom Hundley for
564
assistance with the installation of CTDs and Jim Wyatt, Jeff Hancock and Wayne Kinard
565
for additional support. We thank Ken Clizbe and Bonnie Stelzenmuller for the
566
photograph of Little River Spring. We thank Art Palmer and Dave Budd for their helpful
567
reviews of an earlier draft of the paper. Two Sisters BBQ in Mayo, Florida, is
568
acknowledged for providing key motivational support for fieldwork. We acknowledge
569
support from the National Science Foundation in the form of research grants EAR-
570
0510054 and EAR-0910794, and for a Graduate Research Fellowship to JG.
571
572
573
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