Prepared In Cooperation with the Department of Geosciences at the

University of Arkansas 

U.S. Geological Survey Karst Interest Group 

Proceedings, Fayetteville, Arkansas,

April 26-29, 2011

Scientific Investigations Report 2011-5031

U.S. Department of the Interior

U.S. Geological Survey

i

Prepared in Cooperation with the Department of Geosciences at the University of Arkansas

U.S. Geological Survey Karst Interest Group Proceedings, 

Fayetteville, Arkansas, April 26–29, 2011

Edited By Eve L. Kuniansky

Scientific Investigations Report 2011–5031

U.S. Department of the Interior

U.S. Geological Survey
ii

U.S. Department of the Interior

KEN SALAZAR, Secretary

U.S. Geological Survey

Marcia K. McNutt, Director

U.S. Geological Survey, Reston, Virginia 2011

For product and ordering information:

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its natural and living resources, natural hazards, and the environment:

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Suggested citation:

Kuniansky,E.L., 2011, U.S. Geological Survey Karst Interest Group Proceedings, Fayetteville, Arkansas, April 26-29, 

2011, U.S. Geological Survey Scientific Investigations Report 2011-5031, 212p. 

Online copies of the proceedings area available at: 

http://water.usgs.gov/ogw/karst/

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endorsement by the U.S. Government.

Although this report is in the public domain, permission must be secured from the individual 

copyright owners to reproduce any copyrighted material contained within this report.
iii

Contents

Introductions and Acknowledgments vi

Agenda U.S. Geological Survey Karst Interest Group and University of Arkansas Hydro Days Workshops 3

Speloegenesis 7

Paleokarst of the USA: A Brief Review by Arthur N. Palmer and Margaret V. Palmer 7

Karst in coastal Zones 17

Influence of Tides and Salinity on the Discharge at the Coastal Spring Creek Springs Group and the Connection to 

Discharge at the Inland Wakulla Springs, Wakulla County, Florida by J. Hal Davis and Richard J. Verdi 17

Revised Hydrogeologic Framework for the Floridan Aquifer System in the Northern Coastal Areas of Georgia and 

Parts of South Carolina by Harold E. Gill and Lester J. Williams 18

Horizontal Bedding-Plane Conduit Systems in the Floridan Aquifer System and Their Relation to Saltwater Intrusion in 

Northeastern Florida and Southeastern Georgia by Lester J. Williams and Rick M. Spechler 20

Geochemistry and Contaminant Transport in Karst Systems 21

Water-Quality Changes and Dual Response in a Karst Aquifer to the May 2010 Flood in Middle, Tennessee by 

Michael W. Bradley and Thomas D. Byl 21

Interaction between Shallow and Deep Groundwater Components at Fay Spring in the Northern Shenandoah Valley 

Karst by Daniel H. Doctor, Nathan C. Farrar, and Janet S. Herman 25

Aqueous Geochemical Evidence of Volcanogenic Karstification: Sistema Zacatón, Mexico by Marcus O. Gary, Daniel 

H. Doctor, and John M. Sharp, Jr. 35

The Influence of Land Use and Occurrence of Sinkholes on Nitrogen Transport in the Ozark Plateaus in Arkansas and 

Missouri by Timothy M. Kresse, Phillip D. Hays, Mark R. Hudson, James E. Kaufman 36

Seasonal Carbon Dynamics in a Northwestern Arkansas Cave: Linking Climate and Cave Conditions by Erik D. 

Pollock, Katherine J. Knierim, and Phillip D. Hays 37

Characterization of Urban Impact on Water Quality of Karst Springs in Eureka Springs, Arkansas by Renee Vardy 44

The Relation Between Dissolved Oxygen and other Physicochemical Properties in Barton Springs, Central Texas by 

Barbara J. Mahler and Renan Bourgeais 54

Geochemistry, Water Sources, and Pathways in the Zone of Contribution of a Public-Supply Well in San Antonio, 

Texas by Lynne Fahlquist, MaryLynn Musgrove, Gregory P. Stanton, and Natalie A. Houston 55

Temporal Stability of Cave Sediments by Eric W. Peterson, and Kevin Hughes 57
iv

Evaluating the Stormwater Filters at Mammoth Cave, Kentucky by Ashley West, Rickard Toomey, Michael W. Bradley 

and Thomas D. Byl 66

Using Labeled Isotopes to Trace Groundwater Flow Paths in a Northwestern Arkansas Cave by Katherine J. Knierim, 

Erik D. Pollock, and Phillip D. Hays 67

Interpreting a Spring Chemograph to Characterize Groundwater Recharge in an Urban, Karst Terrain by Victor 

Roland, Carlton Cobb, Lonnie Sharpe, Patrice Armstrong, Dafeng Hui, and Thomas D. Byl 74

Investigation of Nitrate Processing in the Interflow Zone of Mantled Karst, Northwestern Arkansas by Jozef Laincz 75

Karst Mapping, Dye Tracing, and Geographic Information Systems 84

Using a Combination of Geographic Information System Techniques and Field Methods to Analyze Karst Terrain in 

Selected Red River Sub-Watersheds, Tennessee and Kentucky by David E. Ladd 84

Parellelism in Karst Development Suggested by Quantitative Dye Tracing Results from Springfield, Missouri by 

Douglas R. Gouzie and Katherine Tomlin 85

Geologic Controls on Karst Landscapes in the Buffalo National River Area of Northern Arkansas: Insights Gained 

from Comparison of Geologic Mapping, Topography, Dye Tracers and Karst Inventories by Mark R. Hudson, Kenzie 

J. Turner, Chuck Bitting, James E. Kaufmann, Timothy M. Kresse and David N. Mott 86

Determining the Relation between Water Loss from Beaver Creek and Kamas Fish Hatchery Springs, Samak, Utah 

by Lawrence E. Spangler 87

Using Geographic Information System to Identify Cave Levels and Discern the Speleogenesis of the Carter Caves 

Karst Area, Kentucky by Eric Peterson, Toby Dogwiler, and Lara Harlan 94

Surface Denudation of the Gypsum Plain, West Texas and Southeastern New Mexico by Melinda G. Shaw, Kevin W. 

Stafford, and B.P. Tate 104

Karst Program Updates 113

National Cave and Karst Research Institute: Growing Capabilities and Federal Partnerships by George Veni 113 

The National Karst Map: an Update on its Progress by David J. Weary, and Daniel H. Doctor 118 

Ozarks Plateau Karst 119

Karst Hydrogeology of the Southern Ozarks by Van Brahana 119

Overview of The Nature Conservancy Ozark Karst Program by Michael E. Slay, Ethan Inlander, and Cory Gallipeau 

 120 

Microbial Effects on Ozark Karst-Water Chemistry—Medicinal Implications by John E. Svendsen 121
v

Karst Aquifer Systems 122

An Integrated Approach to Recharge Area Delineation of Four Caves in Northern Arkansas and Northeastern 

Oklahoma by Jonathan A. Gillip, Rheannon M. Hart, and Joel M. Galloway 122

Groundwater Piracy in Semi-Arid Karst Terrains by Ronald T. Green, F. Paul Bertetti, and Mariano Hernandez 123

Ten Relevant Karst Hydrogeologic Insights Gained from 15 Years of In Situ Field Studies at the Savoy Experimental 

Watershed by Van Brahana 132

Traps Designed to Document the Occurrence of Groundwater Stygofauna at the Savoy Experimental Watershed, 

Washington County, Arkansas by Justin R. Mitchell 142

Assessment of Sinkhole Formation in a Well Field in the Dougherty Plain, near Albany, Georgia by Debbie Warner 

Gordon 149

Detecting Karst Conduits through their Effects on Nearby Monitoring Wells by Fred Paillet, and Terryl Daniels 151

Spring Hydrology of Colorado Bend State Park, Central Texas by Kevin W. Stafford, Melinda G. Shaw, and Jessica L. 

DeLeon 152

Analysis of Long-Term Trends in Flow from a Large Spring Complex in Northern Florida by Jack W. Grubbs 160

Karst Modeling 168

Modifications to the Conduit Flow Process Mode2 for MODFLOW-2005 by Thomas Reimann, Steffen Birk, and 

Christoph Rehrl 168

Comparison of Three Model Approaches for Spring Simulation, Woodville Karst Plain, Florida by Eve L. Kuniansky, 

Josue J. Gallegos, and J. Hal Davis 169

Synthesis of Multiple Scale Modeling in the Faulted and Folded Karst of the Shenandoah Valley, Virginia and West 

Virginia by Kurt J. McCoy, Mark D. Kozar, Richard M. Yager, George E. Harlow, and David L. Nelms 171 

 

A Hydrograph Recession Technique for Karst Springs with Quickflow Components That Do Not Exhibit Simple, ZeroOrder Decay by Darrell W. Pennington and J. Van Brahana 172

Numerical Evaluations of Alternative Spring Discharge Conditions for Barton Springs, Texas, USA by William R. 

Hutchison and Melissa E. Hill 182

Field Trip Guide 191

Geology and Karst Landscapes of the Buffalo National River Area, Northern Arkansas by Mark R. Hudson, Kenzie J. 

Turner, and Chuck Bitting 191
vi

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1

INTRODUCTION AND ACKNOWLEDGMENTS

Karst aquifer systems are present throughout parts of the United States and some of its territories and 

are developed in carbonate rocks (primarily limestone and dolomite) that span the entire geologic time 

frame. The depositional environments, diagenetic processes, and post-depositional tectonic events that 

form carbonate rock aquifers are varied and complex, involving both biological and physical processes 

that can influence the development of permeability. These factors, combined with the diverse climatic 

regimes under which karst development in these rocks has taken place result in the unique dual or triple 

porosity nature of karst aquifers. These complex hydrologic systems often present challenges to scientists 

attempting to study groundwater flow and contaminant transport.

The dissolution of carbonate rocks and the subsequent development of distinct and beautiful 

landscapes, caverns, and springs have resulted in some karst areas of the United States being designated 

as National or State parks and commercial caverns. Karst aquifers and landscapes that form in tropical 

areas, such as the north coast of Puerto Rico, differ greatly from karst areas in more arid climates, such as 

west-central Texas or western South Dakota. Many of these public and private lands also contain unique 

flora and fauna associated with the hydrologic systems in these karst areas. As a result, multiple Federal, 

State, and local agencies have an interest in the study of karst terrains.

Many of the major springs and aquifers in the United States are developed in carbonate rocks and 

karst areas, such as the Floridan aquifer, the Ozark Plateaus in Missouri and Arkansas, and the Edwards 

aquifer in west-central Texas. These aquifers and the springs that discharge from them, serve as major 

water-supply sources and as unique biological habitats. Commonly, there is competition for the water 

resources of karst aquifers, and urban development in karst areas can impact the ecosystem and water 

quality of these aquifers.

The concept for developing a Karst Interest Group evolved from the November 1999 National 

Groundwater Meeting of the U.S. Geological Survey (USGS), Water Resources Division. As a result, the 

Karst Interest Group was formed in 2000. The Karst Interest Group is a loose-knit grass-roots 

organization of USGS employees devoted to fostering better communication among scientists working 

on, or interested in, karst hydrology studies. 

The mission of the Karst Interest Group is to encourage and support interdisciplinary collaboration 

and technology transfer among USGS scientists working in karst areas. Additionally, the Karst Interest 

Group encourages cooperative studies between the different disciplines of the USGS and other Federal 

agencies, and university researchers or research institutes. 

The first Karst Interest Group workshop was held in St. Petersburg, Florida, February 13-16, 2001, in 

the vicinity of the large springs and other karst features of the Floridan aquifer system. The proceedings 

of that first meeting, published in Water-Resources Investigations Report 01-4011, are available online at: 

http://water.usgs.gov/ogw/karst/kig/

The second Karst Interest Group workshop was held August 20-22, 2002, in Shepherdstown, West 

Virginia, in proximity to the carbonate aquifers of the northern Shenandoah Valley. The proceedings of 

the second workshop were published in Water-Resources Investigations Report 02-4174, which is 

available online at the previously mentioned website.

The third workshop of the Karst Interest Group was held September 12-15, 2005, in Rapid City, 

South Dakota, which is in proximity to karst features in the Madison Limestone in the semi-arid Black 

Hills of South Dakota and Wyoming, including Wind Cave National Park and Jewell Cave National 

Monument. The proceedings of the third workshop were published in Scientific Investigations Report 

2005-5160, which is available online at the previously mentioned website. 
2

The fourth workshop was held at the Hoffman Environmental Research Center and Center for Cave 

and Karst Studies at Western Kentucky University in Bowling Green, Kentucky, near Mammoth Cave 

National Park and karst features of the Chester Upland and Pennyroyal Plateau. The proceedings of the 

fourth workshop were published in Scientific Investigations Report 2008-5023, also available online. 

This fifth workshop is a joint workshop of the USGS Karst Interest Group and University of Arkansas 

HydroDays workshop, sponsored by the USGS, the Department of Geosciences at the University of 

Arkansas in Fayetteville. Additional sponsors are: the National Cave and Karst Research Institute, the 

Edwards Aquifer Authority, San Antonio, Texas, and Beaver Water District, northwest Arkansas. The 

majority of funding for the proceedings preparation and workshop was provided by the USGS 

Groundwater Resources Program, National Cooperative Mapping Program, and the Regional Executives 

of the Northeast, Southeast, Midwest, South Central and Rocky Mountain Areas. The University of 

Arkansas provided the rooms and facilities for the technical and poster presentations of the workshop,

vans for the field trips, and sponsored the HydroDays banquet at the Savoy Experimental Watershed on 

Wednesday after the technical sessions.

The session planning committee for this fifth workshop included Van Brahana, University of 

Arkansas, and Tom Byl, Allan K. Clark, Daniel Doctor, Brian Katz, James Kaufmann, Eve Kuniansky, 

Kurt McCoy, Larry Spangler, Chuck Taylor, and Patrick Tucci (retired) of the USGS. The field trip 

committee included Van Brahana, University of Arkansas, Chuck Bitting, National Park Service, and 

Mark R. Hudson, Phil Hays, and Patrick Tucci (retired) of the USGS. We sincerely hope that this 

workshop promotes future collaboration among scientists of varied backgrounds and improves of our 

understanding of karst systems in the United States and its territories. 

The extended abstracts of USGS authors were reviewed and approved for publication by the U.S. 

Geological Survey. Articles submitted by university researchers and other Federal agencies did not go 

through the USGS review process, and therefore, may not adhere to our editorial standards or 

stratigraphic nomenclature. All articles were edited for consistency of appearance in the published 

proceedings. The use of trade names in any article does not constitute endorsement by the U.S. 

Government. The USGS, Office of Ground Water, provides financial support for the Karst Interest Group 

website and public availability of the proceedings from these workshops.

The cover illustration was designed by Ann Tihansky, USGS, St. Petersburg, Florida, for the first 

Karst Interest Group workshop.

Eve L. Kuniansky

Karst Interest Group Coordinator
3

AGENDA U.S. GEOLOGICAL SURVEY KARST INTEREST GROUP AND UNIVERSITY OF 

ARKANSAS HYDRO DAYS WORKSHOP

April 26–29, 2011, Fayetteville, Arkansas, University of Arkansas Campus

TUESDAY, APRIL 26

Registration

Start at 7:45 am-- All day – pick up name tags and proceedings

Welcome and Introductions

8:00 – 8:20 Eve Kuniansky, U.S. Geological Survey, Karst Interest Group Coordinator and Dr. Van 

Brahana, Department of Geosciences, University of Arkansas

Keynote/Speleogenesis (Eve Kuniansky moderator)

8:20 – 9:00 Paleokarst of the USA: A brief overview by Arthur N. Palmer and Margaret V. Palmer

Karst in Coastal Zones (Eve Kuniansky moderator)

9:00 – 9:20 Influence of tides and salinity on the discharge at the coastal Spring Creek Springs group 

and the connection to discharge at the inland by J. Hal Davis and Richard J.Verdi

9:20 – 9:40 Revised hydrogeologic framework for the Floridan Aquifer System in the northern 

coastal areas of Georgia and parts of South Carolina by Harold E. Gill and Lester J. 

Williams

9:40 – 10:00 Horizontal bedding-plane conduit systems in the Floridan Aquifer System and their 

relation to saltwater intrusion in northeastern Florida and southeastern Georgia by Lester 

J. Williams and Rick M. Spechler

10:00 – 10:40 BREAK

Geochemisty and Contaminant Transport of Karst Systems (Zelda Bailey, moderator)

10:40 – 11:00 Water-quality changes and dual response in a karst aquifer to the May 1-2, 2010 flood in 

middle, Tennessee by Michael W. Bradley and Thomas D. Byl

11:00 – 11:20 Interaction between shallow and deep ground water components in the northern 

Shenandoah Valley karst by Daniel H. Doctor, Nathan C. Farrar, and Janet S. Herman 

11:20 – 11:40 Aqueous geochemical evidence of volcanogenic karstification: Sistema Zacatón, Mexico

by Marcus Gary, Daniel H. Doctor, and John M. Sharp, Jr.

11:40 – 1:00 LUNCH ON YOUR OWN
4

Geochemisty and Contaminant Transport of Karst Systems continued (Tom Byl, moderator)

1:00 – 1:20 The influence of land use and occurrence of sinkholes on nitrogen transport in the Ozark 

Plateaus in Arkansas and Missouri by Timothy M. Kresse, Phillip D. Hays, Mark R. 

Hudson, and James E. Kaufman

1:20 – 1:40 Seasonal carbon dynamics in a northwestern Arkansas cave: linking climate, nutrients 

and cave conditions by E. D. Pollock, K.J. Knierim, and P.D. Hays

1:40 – 2:00 Characterization of Urban Impact on Water Quality of Karst Springs in Eureka Springs, 

Arkansas by Renee Vardy

2:00 – 2:40 BREAK

Geochemisty and Contaminant Transport of Karst Systems continued (Brian G. Katz, moderator)

2:40 – 3:00 Factors affecting dissolved oxygen concentrations in Barton Springs, Austin, Texas by 

Barbara Mahler and Renan Bourgeais

3:00 – 3:20 Geochemistry, Water Sources, and Pathways in the Zone of Contribution of a PublicSupply Well in San Antonio, Texas by Lynne Fahlquist, MaryLynn Musgrove, Gregory 

P. Stanton, and Natalie A. Houston

3:20 – 3:40 Temporal stability of cave sediments by Eric W. Peterson and Kevin Hughes

Karst Mapping and Geographic Information Systems (Brian G. Katz, moderator)

3:40 – 4:00 Using a combination of geographic information system techniques and field methods to 

analyze karst terrain in selected Red River sub-watersheds, Tennessee and Kentucky by 

David E. Ladd

4:00 – 6:00 Poster Session at University of Arkansas

WEDNESDAY, APRIL 27

Karst Program Updates (Kurt J. McCoy, moderator)

8:00 – 8:20 National Cave and Karst Research Institute: Growing capabilities and Federal 

partnerships by George Veni

8:20 – 8:40 The National karst map: an update on its progress by David J. Weary and Daniel H. 

Doctor

Ozarks Plateau Karst (Kurt J. McCoy, moderator)

8:40 – 9:00 Karst Hydrogeology of the Ozarks by Van Brahana

9:00 – 9:20 Overview of The Nature Conservancy Ozark Karst Program by Michael E. Slay, Ethan 

Inlander, and Cory Gallipeau

9:20 – 9:40 Microbial Effects on Ozarks Karst Chemistry -- Medicinal Implications by John E. 

Svendsen

9:40 – 10:20 BREAK

Karst Aquifer Systems continued (Allan K. Clark, moderator)

10:20 – 10:40 An integrated approach to recharge area delineation in northern Arkansas and 

northeastern Oklahoma by Jonathan A. Gillip, Rheannon M. Hart, Joel M. Galloway

10:40 – 11:00 Groundwater Piracy in Semi-Arid Karst Terrains by Ronald T. Green, F. Paul Bertetti, 

and Mariano Hernandez
5

11:00 – 11:20 Sequential spring hydrology, karst valleys, and transition conduits draining Kirby 

watershed in south-central Indiana by Garre Connor

Karst Mapping and Geographic Information Systems (Allan K. Clark, moderator)

11:20 – 11:40 Identification and classification of karst using a pseudospectral method by James E. 

Kaufmann

11:40 – 1:00 LUNCH ON YOUR OWN

Karst Modeling (Daniel H. Doctor, moderator)

1:00 – 1:20 Modifications to the Conduit Flow Process Mode2 for MODFLOW-2005 by Thomas 

Reimann, Steffen Birk, Christoph Rehrl, and W. Barclay Shoemaker 

1:20 – 1:40 Comparison of three model approaches for spring simulation, Woodville Karst Plain, 

Florida by Eve L. Kuniansky, Josue J. Gallegos, and J. Hal Davis

1:40 – 2:00 Synthesis of multiple scale modeling in the faulted and folded karst of the Shenandoah 

Valley, Virginia and West Virginia by Kurt J. McCoy, Mark D. Kozar, Richard M. 

Yager, George E. Harlow, and David L. Nelms

2:00 – 2:20 A Hydrograph Recession Technique for Karst Springs with Quickflow Components That 

Do Not Exhibit Simple, Zero-Order Decay by Darrell Pennington and Van Brahana

2:20 – 3:00 BREAK

Geophysical Methods in Karst (Pat Tucci, moderator) 

3:00 – 3:20 New tools for detecting vertical fractures and voids in karst by Jack H. Cole and Phillip 

West

3:20 – 3:40 Mapping epikarst using helicopter electromagnetic methods by David V. Smith, Bruce D. 

Smith, Maryla Deszcz-Pan, and Charles D. Blome

Karst Hydraulics (Pat Tucci, moderator)

3:40 – 4:00 A simple but effective model predicting the effect of a karst conduit on an adjacent 

observation well by Fred Paillet

4:00 – 4:20 KIG Business-planning next meeting Spring 2014

4:20 -6:00 Busses take those that want to attend the HydroDays Banquet at the Savoy 

Experimental Watershed (SEW) 

6:00 - ? HydroDays Banquet of roast razorback and roast goat, in addition to many other tasty 

treats to satiate the full range of palates for which karstophiles are noted. We have 

arranged presentations by recent researchers at SEW.

THURSDAY, APRIL 28

8:00AM – 5:00PM Field Trip Day 1—Geology and Karst Landscapes of the Buffalo National River, 

Northern Arkansas lead by Mark Hudson and Kenzie Turner, U.S. Geological Survey; 

and Chuck Bitting, National Park Service

FRIDAY, APRIL 29

8:00AM –5:00PM Field Trip Day 2—Dominant Hydrogeologic Controls of Karst on the Buffalo National 

River—Float Trip
6

POSTERS

Karst Mapping, Dye Tracing, and Geographic Information Systems

Parallelism in karst development suggested by quantitative dye tracing results from Springfield, Missouri by 

Douglas R. Gouzie and Kati Tomlin

Geologic controls on karst landscapes in the Buffalo National River area of northern Arkansas: Insights gained from 

comparison of geologic mapping, topography, dye tracers and karst inventories by Mark R. Hudson, 

Kenzie J. Turner, Chuck Bitting, James E. Kaufmann, Timothy M. Kresse, and David N. Mott

Sinkholes identified from LiDAR in the Mill Creek area of the Buffalo National River, Arkansas by J. E. Kaufmann 

and A. T. Lingelbach

Investigation of the relation between surface-water losses in the Beaver Creek drainage and springs at the Kamas 

State Fish Hatchery, and implications for transmission of whirling by Larry Spangler

Survey of Springs Issuing from the Trinity Aquifer in the Vicinity of Northern Bexar County, Texas by Allan K. 

Clark, Robert R. Morris and Travis J. Garcia

Using GIS to identify cave levels and discern the speleogenesis of the Carter Caves karst area, Kentucky by Eric W. 

Peterson, Toby Dogwiler, and Lara Harlan

Surface Denudation of the Gypsum Plain, West Texas and Southeastern New Mexico by M. G. Shaw, K. W. 

Stafford, and B.P. Tate

Geochemisty and Contaminant Transport of Karst Systems

Evaluating the Stormwater Filters at Mammoth Cave, Kentucky by Ashley West, Rickard Toomey, Mike Bradley, 

and Tom Byl

Using labeled isotopes to trace groundwater flow paths in a northwestern Arkansas cave by K. J. Knierim, E. D. 

Pollock, and P. D.Hays

Interpreting a spring chemograph to characterize groundwater recharge in an urban karst terrain by Victor Roland, 

Carlton Cobb, Lonnie Sharpe, Patrice Armstrong, Dafeng Hui, and Tom Byl

Investigation of the fate of nitrate in the interflow zone of mantled karst by Jozef Laincz

Alternative approaches to dissolved organic matter characterization in karst aquifers by Terri Brown

Karst Aquifer Systems

Ten Relevant Karst Hydrogeologic Insights Gained from 15 Years of In Situ Field Studies at the Savoy 

Experimental Watershed by Van Brahana

Traps Designed to Document the Occurrence of Groundwater Stygofauna at the Savoy Experimental Watershed, 

Washington County, Arkansas by Justin Mitchell

Assessment of Sinkhole Formation in a Well Field in the Dougherty Plain, near Albany, Georgia by Debbie Warner 

Gordon

Detecting karst conduits through their effects on nearby monitoring wells by Fred Paillet and Terryl Daniels

Spring Hydrology of Colorado Bend State Park, Central Texas by K.W. Stafford, M.G. Shaw, and J.L. DeLeon

Analysis of long-term trends in flow from a large spring complex in northern Florida by J.W. Grubbs

Karst Modeling

Use of MODFLOW-CFP and MT3DMS to simulate karst flow in a laboratory karst analog model and the Woodville 

Karst Plain by Josue Gallegos, Bill X. Hu, Hal Davis, Eve Kuniansky, and Barclay Shoemaker

Numerical Evaluations of Alternative Spring Discharge Conditions for Barton Springs, Texas, USA by W.R. 

Hutchison and M.E. Hill
7

SPELEOGENESIS

Paleokarst of the USA: A Brief Review

By Arthur N. Palmer1 and Margaret V. Palmer2

1

Department of Earth Sciences, State University of New York, Oneonta, NY 13820-4015 

2

619 Winney Hill Road, Oneonta, NY 13820

Abstract

Paleokarst consists of solutional features from a prior geomorphic phase that have been preserved by 

burial or by a substantial change in local environment. The two major paleokarst horizons in North 

America are of Early-Middle Ordovician (post-Sauk) age and Mississippian-Pennsylvanian (postKaskaskia) age. There are also several less extensive paleokarst zones. They all differ in detail but 

typically include remnants of surface karst features, caves, breccias, hypogenic porosity, and related 

mineral suites. In places they provide high-permeability zones significant to water supply, or serve as 

hosts to petroleum, ores, and later karst development. Studies of paleokarst give significant evidence for 

past geologic and hydrologic conditions, both surficial and deep seated. 

INTRODUCTION 

Paleokarst is an important but often 

overlooked aspect of geology (Fig. 1). It is 

typically formed by lengthy exposure of soluble 

rock to karst processes, followed by burial beneath younger rocks or sediment, which preserve 

the karst features more or less intact. Karst features that are left as relics by changes in climate 

or hydrologic regime can also be considered 

paleokarst, although this use of the term is not so 

common. Paleokarst provides clues to former 

geologic and hydrologic conditions, as well as 

insight into global tectonics and changes in 

climate and sea level.

The major American paleokarst horizons are 

briefly described here, with attention to their 

physical characteristics and economic importance. Also described is the impact of paleokarst 

on later caves and karst, as well as the problems 

of recognizing and mapping paleokarst. For a 

broader perspective see books by James and 

Choquette (1988), which mainly concerns the 

USA, and the worldwide perspective by Bozák 

(1989), which includes chapters on the USA (M. 

Palmer and A. Palmer) and Canada (Ford). 

POST-SAUK PALEOKARST 

The Sauk cratonic sequence (latest Precambrian – Early Ordovician) terminates upward 

mainly in carbonate formations that extend over 

most of the USA (Sloss, 1963). These strata 

were exposed to erosion because of relatively 

low sea level for up to tens of millions of years, 

during which an extensive karst surface developed. This surface was buried by Middle Ordovician strata, which also include many carbonates. The resulting paleokarst is extensive but 

poorly exposed at the surface. In the areas of 

greatest exposure, the surface is disturbed either 

by later tectonism (e.g., in the Appalachians) or 

obscured by weathering residuum (e.g., around 

the Ozark Dome). Most information on the unconformity is from mines and drillholes. 

Figure 1. Example of paleokarst: Mississippian 

sinkhole still partly filled with in-place Pennsylvanian sediment, Bighorn Canyon in Wyoming. 

Cliff face is about 20 m high.

In the eastern USA, this stratigraphic interval was greatly deformed by all phases of the 
8

Appalachian Orogeny, so it is often difficult to 

distinguish paleokarst from mechanical disruption of beds. Mussman et al. (1988) describe the 

paleo-karst in the upper Knox–Beekmantown 

Group of the southern Appalachians. They 

report erosional relief of more than 100 m, with 

filled sinkholes and caves extending to more 

than 65 m below the unconformity and intrastratal breccias down to 300 m. Paleosinkholes 

are filled with carbonate breccias and gravels 

with a fine detrital matrix. Caves are filled with 

breccia and laminated dolomite. Stratiform 

breccias containing dolomite clasts are present 

in intervals up to 35 m thick. Paleocaves appear 

to have formed by rapid conduit flow, but 

possibly also by dissolution in mixing zones 

(evidence is sparse). Further brecciation took 

place during deep burial, as did dissolution and 

dolomitization, and sulfide ore emplacement by 

migration of late Paleozoic warm saline basinal 

brines. This paleokarst is also well developed in 

parts of eastern Canada, with a relief up to 10 m 

that includes buried sinkholes and karren, as 

well as mineralized vuggy and cavernous 

porosity (Ford, 1989; Dix et al., 1998). 

The most notable aspect of the post-Sauk

paleokarst is its role in hosting lead-zinc ores 

(e.g., Tennessee, Missouri) and oil and gas fields 

(e.g., west Texas). It also provides high-permeability zones in carbonate aquifers, especially 

along the Cincinnati Arch in Tennessee and 

Kentucky. These are located in the top 60 m of 

the Knox Group, at and beneath the unconformity, along bedding-guided solution zones and 

fractured, vuggy dolomites and breccias. Where 

shallow, these zones provide much groundwater, 

and where deep they have been targets for injection wells. Artesian conditions are present in 

some areas.

The most extensive views of the post-Sauk 

paleokarst are in lead-zinc mines of Tennessee 

and Missouri (see Ohle, 1985; Sangster, 1988). 

But it is curious that a buried karst surface is not 

particularly clear. Most of the mines are located 

well below the unconformity, and their mineralized zones may be related to older paleokarst. 

Some breccias, mineralized or not, extend 

upward to the unconformity, and even above it. 

Are these breccias the indirect result of paleokarst, where subsidence propagated upward 

from solution voids well after the overlying beds 

were deposited? 

Breccias in the Tennessee mining districts 

form tabular bodies in a zig-zag pattern with 

irregular tops and bottoms, and their stratigraphic relationships are obscured by folding 

and faulting. The older breccia occupies the 

upper 150 m of the Knox Group beneath the 

unconformity. The younger breccia is of unknown age and in places extends upward across 

the unconformity. It formed around the earlier 

breccia and was apparently guided by it. The 

ores are present only in the younger breccia (Fig. 

2). These features are not directly related to the 

paleokarst but were probably influenced by 

tectonic and hypogenetic processes. The leadzinc ore fills spaces between angular carbonate 

clasts, and evidence for bedrock replacement is 

minimal, except in high-grade deposits. Regarding the origin of the breccias, Sangster (1988) 

notes their similarity to known paleokarst 

solution collapse breccias and the paucity of 

wall-rock alteration, both of which may relate to 

meteoric karst. (This viewpoint is controversial.)

Figure 2. Zinc ore (yellow sphalerite, medium 

gray in photo) from a mine in Ordovician carbonate breccia associated with the post-Sauk 

paleokarst, Jefferson City, Tennessee. White = 

saddle dolomite (slightly thermal). Pen for scale.

A great deal of information on the post-Sauk 

paleokarst is available from oil and gas wells. 

The Ellenburger Formation of western Texas is a 

typical example (Kerans, 1993). Most of the reservoir rock is brecciated and dolomitized (Fig. 

3). There is considerable debate about the role of 

paleokarst in providing the reservoir porosity. 

Some attribute it almost entirely to paleokarst 

(e.g., Loucks and Handford, 1992). Others 
9

invoke deep fracture zones related to basementrooted trans-tensional faults (e.g., Smith, 2009). 

According to the latter view, the breccias formed 

in spaces created by the faulting, and dissolution 

was performed mainly by hydrothermal fluids 

rising along faults. It is likely that both views are 

justified to varied degrees, depending on the 

local geomorphic setting and structural history.

Figure 3. Dolomite breccia below the post-Sauk 

unconformity, Ellenburger Formation. The white 

rinds are saddle dolomite (thermal), and the 

matrix is recrystallized internal dolomite sediment. This exposure is in the wall of a quarry in 

central Texas. Pen for scale.

POST- KASKASKIA PALEOKARST 

The post-Kaskaskia paleokarst is much more 

clearly defined than the post-Sauk and merits a 

more detailed description. The Kaskaskia sedimentary sequence extends from Middle Devonian to the end of the Mississippian (Sloss, 1963). 

Carbonates are abundant, especially near the top. 

These strata are host to the most extensive caves 

of the USA, as well as the best-exposed paleokarst (Figs. 1 and 4). Some Cenozoic caves, 

such as Wind and Jewel Caves in South Dakota, 

have inherited much of their pattern from these 

earlier features.

The widespread post-Kaskaskia erosion 

surface was well preserved by deltaic and 

transgressive Pennsylvanian deposits. Because 

of the onlappng pattern of strata, the exact time 

of burial varied from place to place. Valleys and 

karst depressions were filled, as were most of 

the late Mississippian caves. Paleokarst fill is 

most commonly sand and clay colored deep red 

by hematite. 

Karst preservation was more complete than 

in the post-Sauk paleokarst, and exhumed relics 

of Mississippian sinkholes, fissures, and caves 

are well exposed in many areas. Erosional relief 

on the Mississippian strata ranged up to about 

100 m, allowing extensive karst to develop in 

carbonate rocks. The paleokarst is best exposed 

in areas of Laramide uplift and diminishes in 

intensity eastward, where the Mississippian 

carbonates are overlain by thick detrital strata. 

The easternmost extent of well-developed 

post-Kaskaskia paleokarst appears to be in 

northern Illinois, where the erosion surface 

extended down to Ordovician and Silurian carbonates (Plotnick et al., 2009). Clearly developed tubular caves up to 10 m high and 18 m 

wide were filled with clastic sediment, which 

fossil pollen show to be Middle Pennsylvanian. 

Paleokarst is especially well developed in 

the Northern and Middle Rocky Mountains, 

where uplift near the end of the Mississippian 

was accelerated by the Antler Orogeny farther 

west. In this region, local relief on the paleokarst 

is typically about 20–40 m. In places its relief is 

sharp, with deep depressions and underlying 

caves. Elsewhere it is subdued and undulatory. 

In places there is evidence for buried surface 

stream valleys (Sando, 1988).

Figure 4. Post-Kaskaskia paleocaves exposed in 

the Bighorn Canyon, Wyoming. These and related caves are concentrated about 15 m below 

the top of the Madison Limestone (Mississippian) and are filled with poorly lithified Pennsylvanian sediments. Also see Figure 1.

Widespread breccias are also common in the 

early Meramecian carbonates of this area. The 
10

breccias are laterally correlative with evaporite 

deposits in structural basins (e.g., Williston 

Basin) and appear to represent collapse caused 

by intrastratal solution in exposed areas 

(Roberts, 1966; Sando, 1988). They consist 

mainly of angular, unstratified, and unsorted 

carbonate and chert fragments in a red silt-clay 

matrix. Breccias are present only in domal areas, 

so evaporite dissolution must have taken place 

during and/or after uplift. The post-Kaskaskia 

erosion surface truncates most of the breccias, so 

it is likely that these areas were relatively high 

long before Laramide uplift. Farther down the 

flanks of uplifts, strata were exposed to abundant circulation of meteoric groundwater only 

after large-scale Laramide tectonism. 

Lead-zinc-silver mineralization has taken 

place along the post-Kaskaskia paleokarst in 

central Colorado, and paleovalleys and caves 

have also been mapped (Tschauder and Landis, 

1985; DeVoto, 1988), as well as buried tower 

karst (Maslyn, 1976). Correlative breccias, 

breccia pipes, and air-filled pockets lined with 

calcite crystals are abundant in the Mississippian 

limestones of the Grand Canyon, Arizona

(Troutman, 2004). 

OBSERVATIONS IN THE BLACK HILLS 

The various aspects of the post-Kaskaskia 

paleokarst are most clearly displayed in caves of 

the Black Hills of South Dakota. On the basis of 

field mapping and petrographic study of Wind 

Cave and Jewel Cave (Palmer and Palmer, 1995, 

2008), distinct stages of paleokarst and related 

cave development are described here in chronological sequence. Similar paleokarst exposures 

are also common in outcrops and caves in the 

Northern Rockies, although local details differ.

1. Early Diagenetic Karst Processes

The earliest phase of the post-Kaskaskia 

paleokarst in the future Black Hills involved the 

interaction between sulfates and carbonates in 

the Mississippian Madison (locally Pahasapa) 

Formation. Plastic flow, hydration of anhydrite, 

and dissolution of sulfates formed carbonate 

breccias in stratiform bodies and local 

discordant dikes, mainly in the middle and upper 

Madison (Fig. 5). These zones correlate with 

evaporites in the neighboring Williston Basin. 

Many clasts were wedged apart and lifted 

against gravity by crystal growth. In places, 

fracturing was less displacive and produced 

mosaic breccias with calcite veins. 

Meanwhile, sulfate reduction fostered the 

growth of pyrite and associated dolomite 

crystals in the walls of fractures and pores. The 

source of organic compounds as a reducing 

agent is uncertain, although migration from 

nearby petroleum deposits, deep-seated methane, and organics in connate water are all possible. Isotopic evidence has been disrupted by 

later mineral replacement. Saddle dolomite, an 

indicator of thermal conditions, is absent.

Figure 5. Early diagenetic breccia in Madison 

Limestone in Jewel Cave, South Dakota. Bedrock clasts are surrounded by a matrix of 

yellow-brown calcite coated with a thin layer of 

deep-burial scalenohedral calcite. Width of 

photo is approximately 25 cm.

The effect of the Antler Orogeny in the 

Black Hills area was limited to minor uplift and 

block faulting, which probably involved reactivation of Precambrian structures (Sando, 1988). 

During the middle and late Mississippian an 

increasing influx of meteoric water took place as 

the Madison Formation was gradually exposed 

at the surface. 

Oxidation of pyrite (and possibly of hydrogen sulfide from sulfate reduction) produced 

sulfuric acid that attacked the carbonate bedrock 

and breccia clasts, producing porosity and secondary gypsum. Textures similar to those in 

active sulfuric acid caves are still present in 

places, and microscopic remnants of gypsum 

and anhydrite crystals and pseudomorphs are 
11

visible in thin sections. With the increasing 

influx of meteoric water and a rise in water 

table, the gypsum was gradually replaced by 

calcite through the common-ion effect. Iron 

oxides and fossil filaments of iron-oxidizing 

bacteria are abundant in the calcite and impart a 

distinct yellow-brown color to the calcite. Some 

of the oxides are pseudomorphic after pyrite. 

The calcite is rather uniform in texture, with 

poorly defined and irregular growth patterns, 

which is compatible with sulfate replacement, 

rather than deposition on the walls of waterfilled voids. It also contains floating bedrock 

residue and sparse anhydrite inclusions. U/Pb 

dating of this calcite confirms a MississippianPennsylvanian age (Palmer et al., 2009). 

Dissolution of the carbonate bedrock by 

sulfuric acid produced voids up to ~2 m in 

diameter lined by calcite-cemented breccia, and 

with large brown calcite scalenohedra projecting 

into the openings. The openings were probably 

not extensive enough to be considered caves. 

Most have been almost entirely assimilated by 

later dissolution, but they can still be seen more 

or less intact where bedrock collapse has exposed them in the present caves. 

Figure 6. Boxwork in Wind Cave, South Dakota, 

is composed of veins of the yellow-brown calcite identical to the breccia matrix shown in 

Figure 5. The veins shown here project about 20 

cm from the cave wall.

Closely spaced calcite veins developed in 

many of the dolomitic strata in the middle 

section of the Madison and project from the 

present cave walls as boxwork, especially in 

Wind Cave (Fig. 6). These veins show 

considerable scatter in orientation, even in the 

same location. The small scale, high spatial 

density, and diversity of fracture patterns 

required multi-directional stresses that were 

local, rather than regional, and hydration and 

plastic flow of sulfates was the likely cause. At 

first the fractures were filled with gypsum. The 

dolomitic bedrock was later altered by sulfuric 

acid, which had virtually no effect on the 

gypsum veins. Still later, the gypsum veins were 

replaced by calcite as the influx of meteoric 

water increased. The boxwork seen in the caves 

today has been accentuated by later removal of 

the bedrock and reinforcement of the calcite 

veins by accretion of calcite, aragonite, and internal carbonate sediment. This type of boxwork 

is rare elsewhere in the world. It also accompanies certain sulfide ores, which might suggest 

a hydrothermal origin; but the essential factor 

seems to be sulfur redox processes, rather than 

high temperature.

2. Late Mississippian Karst and Caves

Meteoric caves formed in local zones at an 

average of 15–30 m below the highs in the karst 

surface. Most were limited to the upper massive 

limestone of the Madison, above a prominent 

zone of bedded chert. A few solutional fissures 

extended to greater depths, but most were enlarged during the Cenozoic and do not retain 

their original form. Exposures in canyon walls 

and in the present caves show that the Mississippian caves were mainly low arched rooms or 

galleries up to roughly 8 m wide and high (Fig. 

4). They are densest where the Madison contains 

overlying sinkholes and fissures, but a clear relation between the surface and subsurface features is difficult to distinguish in outcrops. The 

caves appear to have survived deep burial and 

Laramide exhumation with little modification. 

Most of the depressions and fissures were 

formed by water infiltrating from small local 

catchments, with some of their growth aided by 

collapse into earlier voids. Sulfate-related 

breccias surround and underlie many of the 

depressions, and it is apparent that either 

dissolution of sulfates contributed to the 

collapse, or the depressions were guided by 

existing breccias. A combination of the two 

scenarios is likely.
12

The paleocaves resemble those that presently form in zones of mixing between fresh water 

and saline water. This process is common in 

coastal caves. Walls of paleokarst depressions 

show no evidence of flowing water (e.g., vadose 

flutes) and no clearly vadose passages, such as 

canyons, extend from them. Depressions and 

caves contain few if any coarse-grained fluvial 

deposits. Groundwater recharge must have been 

constrained to small local catchments. 

3. Burial of the Mississippian Karst

All of the surface depressions and most of 

the caves in the future Black Hills area were 

completely filled with Pennsylvanian sediment 

of the Minnelusa Formation, which consists 

mainly of red sandstones with lesser amounts of 

shale, siltstone, carbonates, and evaporites. The 

Minnelusa directly overlies the carbonate rocks 

of the Madison Formation. Some paleocaves 

have only thin fine-grained deposits of this 

material on their floors, which implies that the 

caves were fairly isolated from the overlying 

depositional surface. Most of those exposed in 

canyon walls are blocked with Pennsylvanian 

sediment a short distance inside.

Figure 7. Poorly indurated Pennsylvanian fill in 

Jewel Cave. A = brecciated bedrock, B = red 

paleofill, C = Oligocene-Miocene calcite crust, 

partly fallen away. The sediment originally filled 

Mississippian paleocaves, and as those openings were enlarged after Laramide uplift, much 

of the sediment subsided into the growing voids. 

Some continues to do so today.

The earliest detrital cave sediments were 

autochthonous carbonate sands derived from 

weathering of overlying beds. These are overlain 

by laminated non-carbonate yellow clay and silt, 

and, in turn by poorly bedded red sand (Fig. 7) 

with clasts of carbonates and chert. In places, 

either or both of the first two sediment types is 

absent. The presence of a clearly defined 

sequence is important in determining whether 

the sediment is in place and not disrupted or 

redistributed by later tectonics or sapping into 

lower voids.

Sediment continued to be deposited from the 

Pennsylvanian through the Cretaceous, burying 

the paleokarst to depths up to 2 km. During deep 

burial, a thin coating of white scalenohedral 

calcite was deposited on the walls of solution 

pockets and caves that were not completely 

filled with sediment. It also lined pockets in the 

paleofill. U/Th dating of this calcite has so far 

been unsuccessful because of low U content.

4. Laramide Uplift and Cave Enlargement

Laramide uplift during the late Cretaceous 

and early Cenozoic allowed the sedimentary 

rocks to be removed from the central Black 

Hills. The truncated edges of these strata are 

exposed around the perimeter of the uplift. Recharge along this boundary fed substantial 

groundwater flow through the limestone, and 

eventually reached the situation seen today, 

where most of the Madison groundwater rises to 

surface springs through the overlying Minnelusa 

Formation along fracture zones, especially along 

anticlinal crests. The paleokarst was partly exhumed, and paleocaves were enlarged to the size 

seen in the present caves (Fig. 8). Not all the 

presently accessible cave passages had earlier 

precursors, but they follow the paleokarst 

breccias and paleocaves rather faithfully. Few 

cave passages are exposed in surface canyons, 

despite the dense array of passages seen in the 

accessible caves. This suggests that the caves are 

far from uniformly distributed and are 

concentrated in paleokarst zones and, one step 

further back, to bodies of Mississippian sulfates. 

Caves show little evidence for rapid 

groundwater flow except for local incursions of 

coarse cobbles and boulders. Most of the cave 

enlargement took place in slow phreatic flow, 

where mixing of several sources could take 

place along the Mississippian solution pockets 
13

and caves, which afforded the greatest permeability. Presently accessible caves are highly 

constrained. They do not extend indefinitely 

down-dip into the heart of the Madison aquifer. 

Only a couple of passages in Wind Cave reach 

the water table. And only a few minor passages 

extend up-dip into areas where the overlying 

Minnelusa has been stripped off the limestone. 

No passages reach the base of the limestone 

(Jewel Cave occupies only the upper half of the 

Madison), and very few small domes and 

chimneys extend up to the base of the Minnelusa. On the other hand, the caves have their 

greatest extent along the strike of the beds, 

maintaining a somewhat narrow vertical range 

(Fig. 8). This pattern suggests that dissolution 

was concentrated in zones of converging 

groundwater – a combination of water sinking 

along the eroded edge of the limestone, infiltrating from above through the Minnelusa, and 

probably also rising from below. Each source 

would have had a different chemical character 

(e.g., CO2 partial pressure), and the mixture 

would be more solutionally aggressive than any 

individual source.

Figure 8. Map of Jewel Cave, South Dakota, as 

of 2003. Its NW-SE extension follows the mean 

strike of the Madison beds. The actual structure 

is more varied than indicated on the map. (Map 

courtesy of National Park Service.)

The early erosion of the Black Hills was 

rapid because of active tectonic uplift and a 

warm, humid climate with much runoff. The 

present caves formed late in the erosional history 

of the area, apparently around the end of the 

Eocene, when the topography had reached a 

stage nearly identical to that of today. Most cave 

walls show evidence that after their main phase 

of enlargement they were exposed to prolonged 

weathering above the water table.

Figure 9. Cross section of Oligocene-Miocene 

calcite wall crust 15 cm thick overlying weathered dolomite in Jewel Cave. The crust was 

naturally broken when part of it fell away from 

the weathered surface.

5. Late Cenozoic Reburial and Exhumation

The weathered bedrock surfaces in the Black 

Hills caves were later coated with subaqueous 

calcite (Fig. 9). In Jewel Cave these crusts 

average about 15 cm thick. U/Pb dates on this 

crust give a range of 26 Ma at the base and 14.7 

Ma at the outer edge. This validates the hypothesis that the crust was deposited during occlusion of the springs and burial of the caves by 

Oligocene-Miocene continental deposits. Since 

the mid-Miocene these deposits have been eroding from the Black Hills and the (presumed) 

original groundwater flow paths have been 

reinstated, while the caves have drained once 

more. Wind Cave is still draining and its lowest 

wall crusts are of Pleistocene age. The midCenozoic topography around the caves is still 

being exhumed and is well preserved in today’s 

semi-arid climate. This late-Cenozoic burial and 

re-exposure represents yet another phase of 

paleokarst development. The cumulative result 

of all these events, from mid-Mississippian 

onward, is a karst system that may be more 

complex in history and details than any other 

documented in the world.
14

OTHER PALEOKARST HORIZONS 

A few other American paleokarst zones 

should be mentioned. Skotnicki and Knauth 

(2007) describe an extensive paleokarst on the 

Mescal Limestone of middle Proterozoic age in 

central Arizona preserved by overlying quartzite 

and predating the intrusion of 1.1 Ga diabase 

sills. Weathering of overlying basalt led to 

silicification of the paleokarst. Dissolution of the 

carbonate led to accumulation of early diagenetic chert while the remaining carbonate underwent nearly complete replacement by secondary 

silica. In paleocaves, silicified collapse breccias, fills of sandstone and thinly laminated 

cave-floor siltstone, and flowstone indicate 

extensive cave development and filling during 

the karst event. These represent some of the 

best-preserved silicified karst and Precambrian 

cave flowstone ever described.

Paleokarst is locally significant in the 

Permian carbonates of New Mexico and west 

Texas. From drill-core and related data, Craig 

(1988) describes solutional voids up to 6 m in 

diameter that reveal distinct levels of paleokarst. 

His interpretation is that they formed during the 

Late Permian by freshwater-seawater mixing 

along the flanks of low-relief carbonate islands. 

In the nearby Guadalupe Mountains, caves 

such as Carlsbad Cavern display early paleokarst features of both Permian and Laramide age 

that partly guided later caves that were formed 

mainly by sulfuric acid (Hill, 1987; Queen, 

2009). These early features include fissures and 

small caves lined by large scalenohedral spar 

(Fig. 10). At the surface, Koša and Hunt (2005) 

have mapped many of the paleokarst fissures 

and breccias and their relation to faulting. 

In Arkansas, deep-seated dissolution related 

to the Oachita Orogeny (late Paleozoic) has 

produced cavities lined by calcite crystals up to 

1.5 m long (Brahana et al., 2009). Some have 

been intersected by presently active caves. The 

best-preserved cavities were apparently encountered by vadose cave streams, which dropped to 

the floors of the fissures without dissolving 

away the calcite crystals, although there has 

been some etching and mud coating by periodic 

flooding. 

Tertiary-Quaternary fluctuations in sea level 

have periodically exposed the carbonates of 

Florida to karst development and burial. Buried 

paleokarst horizons have been mapped as far 

down as 90 m below sea level. In southern Florida a strata-bound horizon of cavernous porosity, 

called the “boulder zone,” lies as much as 900 m 

below sea level. Rather than a relic of low sea 

level, this is probably the result of dissolution of 

sulfates.

The USA contains other paleokarst zones of 

local importance, but they are of limited extent.

Figure 10. Large calcite crystals lining a paleokarst fissure in Carlsbad Cavern, New Mexico. 

The calcite is apparently early Laramide in age 

(ca. 87–98 Ma; Lundberg et al., 2000). The fissure may be as old as Permian. Carbide lamp for 

scale = 12 cm tall.

EFFECT OF PALEOKARST ON LATER 

CAVES

Several examples of how paleokarst can 

guide later cave development have been 

described here. Other less extreme examples are 

cited in this paper. Nearly all of these caves have 

a hypogenic origin – i.e., they represent rising 

water or deep-seated processes in which 

solutional aggressiveness is poorly related to 

meteoric groundwater flow. Ford (1995) gives 

field examples that show that hypogenic caves 

are more likely to be influenced by paleokarst 

than epigenic caves. Why? A paleokarst zone 

should be just as permeable to descending water 

as to rising water. 
15

The tendency for hypogenic caves to follow 

paleokarst is partly due to the fact that both are 

most abundant (or at least most recognizable) in 

the semi-arid West. Although some epigenic 

caves in humid regions follow paleokarst horizons, examples are fewer and less clear. There is 

also an inherent reason why buried karst should 

favor hypogenic cave development. Epigenic 

caves are fed by water that collects at the surface 

and drains underground at specific points. These 

points are governed mainly by topography and 

distribution of rock types. Paleokarst zones have 

no preferred orientation relative to this recharge 

pattern. But rising groundwater concentrates 

along the most efficient routes – i.e., the widest 

available paths – with little regard for their 

location or orientation.

MAPPING PALEOKARST

Note that this paper does not include a map 

showing the distribution of paleokarst. Such 

maps have been attempted in the past (e.g., 

Palmer and Palmer, 1989), but the results are so 

general and boundaries so vague that they can be 

misleading without extensive annotation. 

But mapping the distribution and character 

of paleokarst is worthwhile. Besides its economic value, paleokarst provides a record of 

former geomorphic settings, climates, tectonics, 

and global sea-level stands. The difficulty of 

mapping paleokarst lies in its highly variable 

character and degree of development, and the 

fact that most of it is buried. Paleokarst horizons 

also overlap in many places, so their pattern is

three-dimensional. Distinguishing paleokarst 

features of meteoric origin from those formed by 

deep-seated processes is a further complexity. 

Finally, it is often difficult to determine whether 

paleokarst is even present at all. 

With these handicaps, a karst mapper finds it 

difficult to portray the presence and nature of 

paleokarst and to delineate boundaries. But even 

attempting to do so provides a great amount of 

insight and is a worthwhile goal.

ACKNOWLEDGMENT

Many thanks to Drs. David Weary and 

Daniel Doctor, of the U.S. Geological Survey, 

Reston, Va., for their helpful reviews.

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Palmer, M.V., and Palmer, A.N., 1989, Paleokarst of 

the United States, in Bosák, P., ed., Paleokarst – A 

systematic and regional review: Elsevier and 

Academia, Amsterdam and Prague, p. 337–363.

Plotnick, R., Kenig, F., Scott, A., Glasspool, I., Eble, 

C., and Lang, W., 2009, Pennsylvanian paleokarst 

and cave fills from northern Illinois, USA: Palaios, 

v., 24, p. 627–637.

Queen, J.M., 2009, Geologic setting, structure, tectonic history and paleokarst as factors in speleogenesis in the Guadalupe Mountains, New Mexico 

and Texas, USA: Kerrville, Texas, Proceedings of 

15th International Congress of Speleology, v. 2, p. 

952–957.

Roberts, A.E., 1966, Stratigraphy of the Madison 

Group near Livingston, Montana, and discussion of 

karst and solution-breccia features: U.S. 

Geological Survey Professional Paper 526.B, 23 p.

Sando, W.J., 1988, Madison Limestone (Mississippian) paleokarst: A geologic synthesis, in James, 

N.P., and Chouquette, P.W., Paleokarst: SpringerVerlag, New York, p. 256–277.

Sangster, D.F., 1988, Breccia-hosted lead-zinc 

deposits in carbonate rocks, in James, N.P., and 

Chouquette, P.W., Paleokarst: Springer-Verlag, 

New York, p. 102–116.

Skotnicki, S., and Knauth, P., 2007, The Middle 

Proterozoic Mescal paleokarst, central Arizona, 

USA: Karst development, silicification, and cave 

deposits: Journal of Sedimentary Research, v. 77, 

p. 1046–1062.

Sloss, L.L, 1963, Sequences in the cratonic interior of 

North America: Geological Society of America 

Bulletin, v. 74, p. 93–114.

Smith, L.B., 2009, Upper Ordovician Trenton –

Black River hydrothermal dolomite reservoirs of 

eastern North America, in Stafford, K., Land, L., 

and Veni, G., eds., Advances in hypogene karst 

studies: National Cave and Karst Research Institute, Carlsbad, N.M., Symposium 1, p. 99–110.

Troutman, T., 2004, Reservoir characterization, 

paleogeomorphology, and genesis of the Mississippian Redwall Limestone paleokarst, Hualapai 

Indian Reservation, Grand Canyon area, Arizona: 

MS thesis, University of Texas, Austin, 221 p.

Tschauder, R., and Landis, G., 1985, Late Paleozoic 

karst development and mineralization in central 

Colorado, in DeVoto, R., ed., Sedimentology, 

dolomitization, karstification, and mineralization of 

the Leadville Limestone (Mississippian), central 

Colorado: Field trip guidebook, Society for 

Economic Paleontologists and Mineralogists, 

Section 6, p. 89-91.
17

KARST IN COASTAL ZONES

Influence of Tides and Salinity on the Discharge at the Coastal Spring 

Creek Springs Group and the Connection to Discharge at the Inland 

Wakulla Springs, Wakulla County, Florida

By J. Hal Davis1 and Richard J. Verdi1

1

U.S. Geological Survey, 2639 North Monroe Street, Suite A-200, Tallahassee, FL 32303

Abstract

The Floridan aquifer water budget in north-central Florida has been poorly described in the past because 

Spring Creek Springs Group, one of the major discharge points, is in a tidal estuary and discharge has been 

measured only a few times. To improve the understanding of the water budget, the U.S. Geological Survey and 

the Florida Department of Environmental Protection conducted a cooperative study to measure discharge in 

Spring Creek Springs Group. 

The Spring Creek Springs Group is composed of 13 springs located in a tidal estuary in north-central Florida. 

In June 2007, the U.S. Geological Survey installed a gaging station to monitor stage, water velocity, precipitation, 

and conductance in Spring Creek, which connects these springs to the Gulf of Mexico. Salinity was calculated 

from conductance and used to determine the freshwater and saltwater components of flow. Spring flow varied 

from 2007 to 2009, but could be grouped into periods of similar characteristics, each lasting several months. 

There was no net outflow of freshwater from Spring Creek Springs Group for the three periods of June 2007 to 

March 2008, June to September 2008, and July to September 2009. In contrast, there was a net outflow of 

freshwater from the aquifer of about 800 ft3

/s for the periods from March to June 2008 and September 2008 to 

July 2009. 

Salinity in Spring Creek Springs is strongly influenced by inflows into nearby sinkholes, especially Lost 

Creek Sink, which contribute flow to the springs. Flow into Lost Creek Sink is generally low, less than 50 cubic 

feet per second (ft3

/s), except during periods of heavy rainfall when inflows can briefly peak at 2,000 to 3,000 

ft3

/s. This influx of freshwater causes the salinity in Spring Creek Springs to drop from near seawater values of 

35,000 parts per million (ppm) to values as low as 5,000 ppm. Extended periods of low rainfall result in low 

flows into Lost Creek Sink and the increase of salinity in Spring Creek Springs to near seawater values. 

When salinity increases in Spring Creek Springs, the increase in water density results in a higher equivalent 

freshwater head. This higher head exceeds the head in the more inland Wakulla Springs, causing increased flow 

at Wakulla Springs and decreased freshwater flow at Spring Creek Springs. When salinity is low, the equivalent 

freshwater head at Spring Creek Springs is lower than at Wakulla Springs, and relatively higher freshwater 

discharge shifts back to Spring Creek Springs. An estimated 9-inch rise in sea level since the 1930’s may be 

causing long-term salinity increases in Spring Creek Springs and the cause of an observed increase in flow at 

Wakulla Springs over this same period. 
18

Revised Hydrogeologic Framework for the Floridan Aquifer System in

the Northern Coastal Areas of Georgia and Parts of South Carolina

By Harold E. Gill (retired) and Lester J. Williams

U.S. Geological Survey, 3039 Amwiler Rd, Suite 130, Atlanta, GA 30360

Abstract

The hydrogeologic framework for the Floridan aquifer system was revised for eight northern coastal 

counties in Georgia and five coastal counties in South Carolina (fig. 1) as part of a regional assessment of 

water resources by the U.S. Geological Survey (USGS) Groundwater Resources Program. In this study, 

selected well logs were compiled and analyzed to determine the vertical and horizontal continuity of 

permeable zones that make up the aquifer system, and define more precisely the thickness of confining 

beds that separate individual aquifer zones. The results of the analysis indicate that permeable zones in 

the Floridan aquifer system can be divided into (1) an upper group of extremely transmissive zones that 

correlate to the Ocala Limestone in Georgia and the Parkers Ferry Formation in South Carolina, and (2) a 

lower group of zones of relatively lower transmissivity that correlates to the middle part of the Avon Park 

formation in Georgia and updip clastic equivalent units of South Carolina (fig. 2). This new subdivision 

simplifies the hydrogeologic framework originally developed by the USGS in the 1980s and helps to 

improve the understanding of the physical geometry of the system for future modeling efforts. Revisions 

to the framework in the Savannah–Hilton Head area are particularly important where permeable beds

control the movement of saltwater contamination. The revised framework will enable water-resource 

managers in Georgia and South Carolina to assess groundwater resources in a more uniform manner and 

help with the implementation of sound decisions when managing water resources in the aquifer system.

REFERENCE

Williams, L.J., and Gill, H.E., 2010, Revised 

hydrogeologic framework of the Floridan 

aquifer system in the northern coastal area of 

Georgia and adjacent parts of South Carolina: 

U.S. Geological Survey Scientific Investigations

Report 2010–5158, 103 p., 3 plates.

Figure 1. Study area in eight northern coastal 

counties of Georgia and five southern coastal 

counties of South Carolina. (Structural 

features: Gulf Trough from Applied Coastal 

Research Laboratory, 2002; Beaufort Arch 

from Clarke and others, 1990; modified from 

Williams and Gill, 2010).
19

Figure 2. Hydrogeologic units and confining beds of the Floridan aquifer system showing representative log response 

and location of permeable zones and mapping horizons (modified from Williams and Gill, 2010; Fm., formation) 
20

Horizontal Bedding-Plane Conduit Systems in the Floridan Aquifer 

System and Their Relation to Saltwater Intrusion in Northeastern 

Florida and Southeastern Georgia

By Lester J. Williams1 and Rick M. Spechler2

1

U.S. Geological Survey, 3039 Amwiler Rd, Suite 130, Atlanta, GA 30360

2

U.S. Geological Survey, 12703 Research Parkway, Orlando, FL 32826

Abstract

Acoustic televiewer (ATV) images, flowmeter, and borehole geophysical logs obtained from the open 

intervals of deep test wells were used to develop a revised conceptual model of groundwater flow for the 

Floridan aquifer system in northeastern Florida and southeastern Georgia. Borehole information was used 

to identify and map the types and distribution of highly-transmissive production zones in the Floridan 

aquifer system. The ATV images and flowmeter traverses indicate that water produced from most wells 

is largely derived from a system of highly-transmissive solution zones formed along bedding planes and 

major formational contacts. These “horizontal bedding-plane conduit systems” may locally influence the 

movement of brackish and saline water in the Floridan aquifer system. 

A modified conceptual model of regional flow in the Floridan aquifer system is proposed that incorporates locally interconnected horizontal conduit systems within the largely porous matrix rock (fig. 1). 

Each of the conduit systems represents a highly-transmissive zone along which water can move 

preferentially through the aquifer system. These may or may not be laterally continuous across the area.

Flow paths within the system are restricted vertically by local or regional confining units except where these 

are breached by collapse features or vertical fractures. Near major pumping centers, water probably moves 

preferentially along the horizontal conduits to reach the discharging well. The source of water moving into 

the transmissive open conduits is either derived from upward migration along vertical discontinuities in the 

rock or from diffuse leakage from adjacent porous rock units. Some trapped relict water in adjacent lowerpermeability units may locally contribute to the higher chloride concentrations observed in some wells.

Figure 1. Conceptual model of for the Floridan aquifer system from the outcrop area to the offshore area

(modified from Krause and Randolph, 1989; Spechler, 1994).

REFERENCES

Krause, R.E., and Randolph, R.B., 1989, Hydrology of the Floridan aquifer system in southeast Georgia and 

adjacent parts of Florida and South Carolina: U.S. Geological Survey Professional Paper 1403-D, 65 p., 18 pl.

Spechler, R.M., 1994, Saltwater intrusion and the quality of water in the Floridan aquifer system, northeastern 

Florida: U.S. Geological Survey Water-Resources Investigations Report 92–4174, 76 p.
21

GEOCHEMISTRY AND CONTAMINANT TRANSPORT OF KARST 

SYSTEMS

Water-Quality Changes and Dual Response in a Karst Aquifer to the 

May 2010 Flood in Middle, Tennessee

By Michael W. Bradley and Thomas D. Byl

U.S. Geological Survey, Suite 100, 640 Grassmere, Nashville, Tennessee 37211

Abstract

During May 1 and 2, 2010, an area stretching from Memphis through Nashville received about 14 to 

19 inches of rain. Extensive flooding occurred and the Cumberland River in Nashville was above flood 

stage from May 1 through May 7, 2010. Water-quality data were collected at a spring and an observation 

well at Tennessee State University in Nashville, Tennessee (figure 1) during the flood. Changes in 

groundwater quality were observed, including a dual response to the heavy rainfall at the spring and the 

well. The dual response observed in the water-quality data consists of more rapid changes in water 

quality at the spring and minimal water-quality changes in the well. The spring issues from the 

Ordovician Bigby-Cannon Limestone. The observation well is 202 feet deep with water-bearing zones in 

the lower Bigby-Cannon Limestone.

Water level elevations in the well and in the spring rose in response to the heavy rainfall (figure 2). 

The spring area was flooded by water from the Cumberland River and the water-level elevations of the 

spring match the elevation of the Cumberland River during the flood (figure 2). Groundwater elevations 

in the observation well are typically higher than the Cumberland River. However, during the flood, the 

gradient reversed and the level of the Cumberland River was higher than the groundwater elevations 

(figure 2).

Specific conductance records show the contrasting responses to rainfall and flooding at the spring and 

the well (figure 3). The specific conductance of water in the observation well shows little response to the 

flood event with values fluctuating from about 700 to 705 microSiemens per centimeter (figure 3). The 

specific conductance of water from the spring fluctuates rapidly over a range of about 300 to 500 

microSiemens per centimeter during and after the flood (figure 3). The water quality in the spring 

responded quickly to the large amounts of precipitation. The water-quality parameters in the observation 

well, however, showed very little response to the flood event, even though water levels in the well rose 

about 23 feet (figures 2 and 3).
22
23
24
25

Interaction between Shallow and Deep Groundwater Components at 

Fay Spring in the Northern Shenandoah Valley Karst

By Daniel H. Doctor1, Nathan C. Farrar2, and Janet S. Herman2

1

U.S. Geological Survey, 12201 Sunrise Valley Drive, MS 926A, Reston, VA, 20191 

2

University of Virginia, Dept. of Environmental Sciences, P.O. Box 400123, Charlottesville, VA 22904

Abstract

The fractured carbonate rock aquifer of the Shenandoah Valley has been karstified to depths 

exceeding 300 meters (m). Springs in the region integrate flow through both shallow and deep flowpaths 

within solutional conduits which tend to be concentrated in areas of bedrock structural weakness, such as 

along lineaments and fault zones, and through fractured bedrock. In order to investigate the local 

dynamics of the aquifer, continuous monitoring of discharge, conductivity and temperature was combined 

with frequent geochemical sampling and dye-tracing at Fay Spring, located in Winchester, Va. This 

perennial spring shows rapid response in discharge following large rain events with a concomitant 

decrease in conductivity; however, snowmelt leads to an increase in conductivity, indicating recharge by 

surface water carrying seasonally applied road salt. Quantification of local, surface-water input to the 

spring was sought through a dye trace in July 2009. Rhodamine WT injected at the terminal sink point of 

an ephemeral stream approximately 1 kilometer west of Fay Spring provided evidence of initial rapid dye 

transit (greater than 500 m/day), yet overall low total mass recovery (12 percent after 5 weeks). Water 

samples were collected by an automatic sampler every six hours to daily during the trace period, and also 

from January-April, 2010 during snowmelt of the largest snowfall on record for the region. Chemographs 

revealed dilution in NO3, Cl, SO4, and Mg with most large rain events; however, Cl increased during peak 

snowmelt with little change in other ion concentrations. Principal components analysis (PCA) was used 

on major ions (Mg, SO4, NO3, Cl, Sr, and Si) in an effort to distinguish components of deeper, regional 

flow and shallow, local flow within the spring discharge. The PCA results show that two components 

cumulatively explain 72 percent of the variance of the data, permitting a mixing model comprised of three 

end-members. The end-members were representative of 1) a Ca-Mg-HCO3 type water of relatively low 

NO3, SO4 and Cl concentrations, 2) an “agriculturally influenced” type water of high NO3 with moderate 

SO4 and Cl concentrations, and 3) an “urban influenced” type water with high Cl and moderate SO4 and 

NO3 concentrations. The results of the end-member mixing model indicate that Fay Spring is dominated 

by a regional flow component with elevated background levels of NO3 (10-12 mg/L as NO3), SO4 (45-50 

mg/L) and Cl (35-40 mg/L). The background concentrations of SO4 and NO3 are similar between summer 

and winter; however, Cl concentration is higher in the winter (up to 80 mg/L). The spring chemistry is 

noticeably diluted only after large rain events by water interpreted to be derived from shallow diffuse 

groundwater storage; direct sinking surface runoff is a lesser component. Overall, the impacts of rapid 

surface runoff are muted at Fay Spring compared with springs in other Appalachian karst regions. A 

combination of approaches including dye tracing and interpretation of natural geochemistry is helpful for 

elucidating the nature of discharge from karst springs in fractured rock carbonate aquifers.

INTRODUCTION 

The carbonate aquifer system of the northern 

Shenandoah Valley in Virginia and West 

Virginia is a fractured rock aquifer that has 

undergone karstification at depth far below the 

water table. Maximum depth of groundwater 

circulation is unknown; however, circulation 

through conduits exceeds depths of 300 meters

(m) as evidenced by a small number of highyield deep wells. For example, Cady (1936) 

described a well in the City of Winchester (well 

no.W 161, drilled in 1931) that is 436 m deep 

and produced 560 liters/minute (L/min), or 148 

gallons/minute (gal/min). Water bearing cavities 

intersected by this well are reported at depths of 

30, 91, 213, and 335-365 m below the surface. 

Cady (1936) reported three other high-yield 

wells in the Shenandoah Valley at depths greater 

than 300 m below the surface that have 

intersected cavities, all drilled in Cambrian and 

Ordovician carbonate rocks. Recent drillers’ 
26

accounts corroborate the older information. A 

well constructed approximately 12 km north of 

Winchester near Green Spring was mostly dry 

until a depth of 410 m, at which point the well 

hit a cavernous zone and produced 750 L/min 

(198 gal/min); a second well in the vicinity hit a 

cavernous zone at a depth of approximately 245 

m, and emanated sulfurous water and black finegrained sediment that caused the well to be 

abandoned (G. Payne, Payne Well Drilling, pers. 

comm., 2008).

McCoy and Kozar (2007) correlated highyield wells and springs with cross-strike 

fractures and faults where upward-convergent 

groundwater flow may occur. Kozar and others 

(2007a,b) and Jones (1991, 1997) interpreted the 

groundwater system as being a solutionally 

modified fractured-rock carbonate aquifer, with 

interspersed solution conduits serving as primary 

drains for the karst groundwater system. Kozar 

and others (2007 a,b) point out that such 

conduits typically coincide with strike-parallel 

thrust faults and cross-strike extensional or 

strike-slip faults. Conduit flow is evident from 

springs often found along high angle normal or 

reverse faults that cut across the dominant strike 

of bedding, or along thrust faults where 

carbonates are brought into contact with 

overlying shales. Springs are common where 

cross-strike faults intersect thrusts within 

carbonate units (Perry and others, 1979; 

Orndorff and Harlow, 2002). 

The karst aquifer discharges to numerous 

artesian springs that range in flow rate over three

orders of magnitude. Cady (1936) reported 95 

springs with discharge ranging from 27 to 

29,000 m3

/day (0.01 to 11.9 cfs), and an average 

flow of 363 m3

/day (0.15 cfs). Some larger 

springs exhibit a delayed response to rainfall 

events, with long hydrograph recession periods 

and muted chemical variability (Vesper and 

others, 2008). Spring discharge accounts for 60

percent to 97 percent of stream flow (Harlow 

and others, 2005; Nelms and Moberg, 2010). 

Figure 1. Map of study area near Winchester, Virginia showing sampling sites.
27

With the majority of the surface streamflow 

derived from karst aquifer discharge, questions 

remain as to the mechanisms and proportions of 

recent aquifer recharge expressed in the 

discharge at springs. Here, we investigated the 

interaction between shallow, recently recharged 

water at a stream sink and water derived from 

deeper, phreatic flow at one artesian spring site 

in the Shenandoah Valley.

Study site: Fay Spring 

Fay Spring is located within the City of 

Winchester, Virginia (fig. 1). The spring is 

situated on a northwest-trending strike-slip fault 

cutting Lower Ordovician dolomite and 

limestone (Orndorff and others, 2004). The 

spring is owned by the City of Winchester, and 

was once used as a part of the municipal water 

supply (Harlow and others, 2005). 

The spring was instrumented by the Virginia 

Water Science Center of the U.S. Geological 

Survey for continuous monitoring of discharge, 

water temperature, and specific conductance 

from late April 2007 to October 2010. For the 

period of record, the mean discharge was 2,592 

m3

/day (1.09 ft3/s) and the mean temperature 

was 13.1 oC. The continuous record for this 

perennial spring shows a broad seasonal rise and 

fall in water temperature and discharge, 

punctuated by rapid increases in flow with 

concomitant decreases in conductivity (fig. 2). 

An increase in spring discharge corresponding to 

a decrease in conductivity generally indicates 

the rapid addition of less-mineralized water from 

the surface to the spring during recharge events, 

and is commonly observed in karst settings 

(White, 1988). At Fay Spring, the water 

temperature does not respond to discharge in as 

flashy a manner as conductivity, and relatively 

little temperature change is observed even 

during periods of large changes in flow (fig. 2). 

Conversely, conductivity does not show a 

pronounced seasonal trend like water 

temperature, and tends not to vary greatly when 

above the long-term mean value of 711 μS/cm 

(fig. 2). However, a winter precipitation event in 

January 2008 caused an unusual dramatic 

increase in conductivity, with no corresponding 

change in discharge or water temperature at the 

spring (fig. 2). This increase in conductivity was 

hypothesized to result from sinking surface 

water carrying road salt to the spring, and may 

have been a delayed response to events ranging 

from several days up to two weeks earlier. In 

order to test this hypothesis, quantification of the 

influence of local, surface-water input to the 

spring was sought through a dye trace. 

Figure 2. Continuous monitoring at Fay Spring, 

from 2007 to 2010. An unusual conductivity 

increase in January 2008 is highlighted in the 

green box and expanded in the right-hand 

panels. A similar, but more gradual conductivity 

excursion was observed during a major period of 

recharge as a result of above normal snowfall in 

2010. Precipitation data obtained from National

Weather Service Coop Station 449181 at 

Winchester, Virginia.

Dye trace to Fay Spring

On June 30, 2009, approximately 2 kg of 20

percent Rhodamine WT dye solution was 

injected at the terminal sink point of Sunnyside 

Run, an ephemeral stream ~1 km west of Fay 

Spring (fig. 3). Water samples for dye and major 

ion analysis were collected by an automatic 

sampler with a frequency of every six hours to 

once per day during the trace period. Dye was 

recovered at two springs (Fay Spring and 

Sempeles Spring, fig. 1) located 300 m apart 

along a single fault, and within the channel of 

Redbud Run immediately upstream of its 

confluence with the spring run of Sempeles 

Spring; dye was not recovered at any of the 

other sampling sites. 
28

Figure 3. Geologic map of study area showing water table contours and traced connections. 

The dye breakthrough curves at Fay Spring and 

Sempeles Spring are shown in figure 4. The 

results of this tracing test provided evidence of 

initial rapid dye transit (>500 m/day) and low 

overall mass recovery (~12 percent) that 

includes sampling at both springs over the 

course of five weeks. In spite of the low tracer 

recovery in the water samples, monitoring by 

passive activated charcoal samplers verified a 

positive recovery of the dye at each of the three 

sites. The amount of dye recovered at Sempeles 

Spring was approximately ten times greater than 

that at Fay Spring; however, Sempeles Spring 

was not instrumented for continuous monitoring, 

so our discussion focuses on Fay Spring. The 

amount of dye recovered at Redbud Run was 

less than at either spring based on the passive 

charcoal sampler results.

Snowmelt monitoring at Fay Spring

Chemical sampling continued at Fay Spring 

throughout the winter and spring of 2009-2010. 

The results of the continuous monitoring during 

this period revealed a gradual rise in 

conductivity following the recession of a rainfall 

event in late January. Then in early February 

2010, the largest snowfall on record for the 

region occurred. Melting of this snow began on 

February 7, and during the melt period the rise 

in discharge was closely correlated to a rise in 

conductivity (fig. 5). Rapid increases in 

discharge occurred with rainfall both prior to 

and following the snowmelt period, and resulted 

in concomitant decreases in conductivity. 

However, during the large snowmelt event, only 

a gradual, damped discharge increase was 

observed that corresponded with a similar 

damped increase in conductivity.

Water samples were collected by an 

automatic sampler every six hours to daily from 

January-April 2010, and a complete record of 

daily samples was obtained during the snowmelt 

period. Chemographs revealed dilution in NO3, 
29

Cl, SO4, and Mg with most large rain events; 

however, Cl increased during peak snowmelt 

with little change in other ion concentrations 

(fig. 6). In comparison to the results of chemical 

sampling during the summer, the background 

concentrations of SO4 and NO3 are similar in the 

winter; however, Cl concentration is higher in 

the winter (up to 80 mg/L). This added chloride 

component may be due to the impact of road salt 

in the winter months. 

Principal components analysis and estimation 

of end-members 

We used principal components analysis 

(PCA) on major ions (Mg, SO4, NO3, Cl, Sr, and 

Si) in an effort to distinguish possible endmember components of deeper, regional flow 

and shallow, local flow that are mixed within the 

spring discharge. PCA has been shown to be an 

effective means for screening hydrochemical 

data in order to identify possible end-member 

compositions contributing to mixed samples 

(Christopherson and Hooper, 1992), and has 

previously been successfully applied in a karst 

aquifer setting (Doctor and others, 2006).

PCA reduces the entire dataset into a 

smaller set of factors that account for the 

greatest variance in the data. In this case, the 

PCA results from the winter samples show that 

two components cumulatively explain 72

percent of the variance of the data. The first 

component (54 percent) is most positively 

weighted on Mg and NO3; the second 

component (18 percent) is most positively 

weighted on SO4, and negatively weighted on Sr 

and Cl. 

An end-member mixing analysis (EMMA) 

was performed using the results of the PCA. 

Since two principal components account for the 

majority of the variance of the data, only three 

end-members are necessary to account for the 

compositions of the mixtures, subject to the 

constraint that the proportions of all endmembers in each observed mixture sum to 1 

(Christopherson and Hooper, 1992). For the 

EMMA, it is assumed that end-members mix 

linearly and that the geochemical constituents 

behave conservatively. For this reason, the most 

non-conservative chemical parameters (Ca and 

HCO3) were excluded.

Figure 4. Dye recovery at Sempeles Spring and 

Fay Spring. Note pulsed recovery of dye 

released from shallow groundwater storage 

during later storm events.

Figure 5. Record at Fay Spring before and after 

snowmelt in winter and spring 2010. Note the 

muted, gradual discharge increase in response to 

snowmelt versus the greater, more rapid 

discharge increase in response to rainfall events. 

Temperature decrease in response to rainfall is 

evident, though damped, and almost nonexistent 

in response to snowmelt.
30

Figure 6. Major ion chemistry at Fay Spring during summer and winter events. Left panel shows 

chemographs during the dye-trace monitoring in July 2009; right panel shows chemographs during the 

snowmelt period of 2010. Note that the ranges of each chemical constituent are shown on equivalent 

vertical axes for both sampling periods. Except for chloride and silica, the chemistry varies around 

average background values that are remarkably consistent between the two seasons, despite nearly the full 

range of observed discharge at Fay Spring being represented.

Reducing the data matrix to a set of two 

factors allows projection of the transformed data 

into a 2-dimensional mixing space (fig. 7). This 

data projection is convenient for exploring the 

data cloud in relation to possible end-member 

compositions that are likewise projected into the 

principal components mixing space. The endmember compositions might be estimated by the 

extremes of the data cloud if candidate samples 

are not available, but this can result in 

ambiguous and non-unique solutions that may 

be unrealistic (Christopherson and Hooper, 

1992). Therefore, it is best to obtain samples of 

candidate end-member sources according to a 

conceptual model of the system in order to test 

the mixing model. In this study, samples of 

surrounding springs and wells were obtained 

periodically during the course of the monitoring 

at Fay Spring. These samples were projected 

into the same principal components space as the 

Fay Spring samples, and of those samples the 

three that fell at the outer extremes and enclosed 

the data cloud of the Fay Spring data were 

chosen as possible end-members (fig. 7). 

The three estimated end-members were: 1) the 

composition of the well water at the Woodbine 

Assembly Church, 2) the composition of 

Washington Spring and 3) the ephemeral flow in 

Sunnyside Run (see fig. 1 for localities).
31

Table 1. Hypothesized geochemical end-member compositions used in the mixing model for Fay Spring.

Hypothesized End-Members1 Cl NO3 SO4 Mg Si Sr

EM1: Woodbine Well (WW)

low NO3, low SO4 , low Cl 4.0 7.0 18.0 15.1 8.7 0.2

EM2: Washington Spring (WS)

high NO3, mod. SO4 , mod. Cl 18.0 21.1 22.4 20.5 8.0 0.4

EM3: Sunnyside Run (SR)

mod. NO3, high SO4 , high Cl 47.4 14.0 48.0 18.5 7.0 0.2

1

Concentrations are in units of mg/L, or ppm; nitrate is as NO3.

Table 1 shows the chemical composition of 

these hypothesized end-members. Each of these 

waters is a Ca-Mg-HCO3 type composition, with 

an additional component of Cl, NO3 and SO4. 

The Woodbine Well end-member has the most 

dilute chemical composition in relation to the 

other end-members, with the lowest NO3, SO4

and Cl values among them. This well is located 

10 km outside of the Winchester city limits in a 

pastoral setting of low development and 

moderate agricultural use. Washington Spring is 

also located outside of the city limits; however, 

it is surrounded by land used for cattle pasture 

and a dairy operation. Due to restricted access, 

the spring run was sampled approximately 500 

m downstream of the rise pool; no tributary 

flows enter the spring run upstream of the 

sampling point. The composition of Washington 

Spring shows the highest measured NO3, with 

moderately high SO4 and Cl values. Sunnyside 

Run has its source within the city limits, and 

flows through an industrial park; it therefore 

represents an urban influence with the highest Cl 

and SO4 of the end-members. It is important to 

emphasize that listing of these hypothetical endmembers does not necessarily mean that the 

actual water from these locales is being 

expressed at Fay Spring; rather, water samples 

from these locales exhibit a chemistry of a type 

that satisfies the end-member mixing model, and 

are feasible within the conceptual model of the 

aquifer system.

Assuming simple linear mixing among the 

end-members and conservative behavior among 

the chemical constituents, the proportions of the 

end-members in each of the mixed samples were

subsequently calculated. The results of the 

mixing calculation are shown in time series in 

figure 8. The shifting end-member (EM) 

proportions are evident in the different sizes of 

the bars that comprise each sample collected at 

Fay Spring. One can observe large increases in 

the proportion of EM1 during rainfall events, 

with a limited increase during snowmelt. This 

shift is reflective of more dilute water affecting 

Fay Spring during these events. It is noteworthy 

that the overall trend is a gradual decrease in the 

proportion of this dilute end-member as 

baseflow discharge gradually increases in

between discrete events. Also, decreases in 

water temperature accompany increases in the 

proportion of EM1 during rainfall events (fig. 5). 

Figure 7. Principal components mixingspace diagram indicating Fay Spring samples 

fully enclosed by the three hypothesized endmembers. 

During the snowmelt period, the proportion 

of the high-Cl end-member (EM3) is observed to 

rise and fall, primarily at the expense of the 

higher NO3 end-member (EM2). If Cl is indeed 

a tracer of recent recharge from the surface 

during snowmelt, then this behavior may be 

reflective of the input of recently recharged 

snowmelt water. The high NO3 component 

represented by EM2 appears to provide the 
32

background chemical composition within the 

spring discharge, with shifting proportions 

between the low-Cl, low-NO3 composition of 

EM1 and high-Cl, high-SO4 composition of 

EM3 reflective of the dynamic changes in 

aquifer components expressed at the spring. 

Figure 8. Estimated end-member proportions in the Fay Spring samples collected during the winter and 

spring of 2010.

DISCUSSION

In many karst areas, focused recharge from 

sinking surface runoff into the karst aquifer 

comprises a large proportion of the discharge at 

springs. In the Shenandoah Valley, however, the 

karst aquifer seems to have a large component of 

deeper, phreatic flow in its discharge. In spite of 

the evidence of sinking surface runoff rapidly 

impacting Fay Spring through conduitdominated flowpaths documented by the dye 

trace results, the amount of dye recovered was 

low, and repeated pulses of dye recovery during 

subsequent storm events indicate significant dye 

retention within the shallow portions of the 

aquifer.

Chemical variability at Fay Spring

Frequent sampling of ambient geochemistry 

combined with continuous monitoring of 

discharge, temperature and conductivity also 

reveals functioning of the aquifer. First, only 

rainfall events above a threshold magnitude 

cause a discharge response at Fay Spring. This 

threshold is not well defined, and is dependent 

upon antecedent conditions, but appears to be in 

the range of a 0.4 in (10 mm) rainfall event. 

When a rainfall event of sufficient magnitude 

does cause a response at Fay Spring, the pattern 

is generally one of conductivity decrease that is 

a nearly direct reflection of the discharge 

hydrograph, and a muted or negligible change in 

water temperature. Second, only during brief 

periods in winter does the conductivity rise with 

an increase in discharge, yet during these 

periods the water temperature has also shown 

little change.

Based upon geochemical sampling in both 

summer and winter seasons, Fay Spring seems 

to be dominated by a diffuse flow component 

with elevated background levels of NO3
33

(10-12 mg/L as NO3), SO4 (45-50 mg/L) and Cl 

(35-40 mg/L) that is noticeably diluted only after 

large rain events. The majority of the 

background at Fay Spring is a similar water type 

as that sampled at Washington Spring (EM2). 

Greater chemical variability was observed 

during larger discharge events in the winter and 

spring than in the summer, but chemical 

variability was relatively muted overall. For 

example, conduct-ivity values at Fay Spring do 

not drop below 600 μS/cm, even during the 

largest events. Sinking surface runoff might be a 

reasonable source of the water that causes

dilution of the spring chemistry during peak 

flows, but this component does not dominate the 

spring discharge. The dye trace resulted in a 

demonstrable conduit connec-tion between the 

sinking surface stream of Sunnyside Run and the 

spring, yet the overall mass recovery was low, 

and a large proportion of that recovery occurred 

with subsequent storm events long after the 

initial dye breakthrough. 

During the snowmelt period of 2010, Fay 

Spring showed a higher conductance and 

temperature than the water sinking at Sunnyside 

Run throughout the melt period. It is possible 

that another, as yet unidentified, source of 

sinking surface runoff affects Fay Spring. 

Mixing of end-member waters at Fay Spring

An end-member mixing analysis (EMMA) 

using samples collected from nearby sources 

was performed. The chemical composition of 

Fay Spring was modeled as a mixture among 

three end-members using the compositions 

shown in table 1, and all the Fay Spring samples 

fell inside the range of compositions generated 

by mixing among these three end-member types. 

EM1, represented by the Woodbine Assembly 

well, reflects karst groundwater that has been 

impacted least by NO3, Cl, or SO4. Our 

preliminary interpretation is that this water type 

represents a component within relatively shallow 

fractures and solutional voids. During larger rain 

events, this water is mobilized into conduits as 

connections are made among normally isolated 

solutional pockets. We observe temperature 

decreases at Fay Spring after the snowmelt in 

2010, but these temperature responses are quite 

damped (at most 0.6 oC), indicating the cooler 

component is probably shallow groundwater and 

not sinking surface water. EM2, represented by 

Washington Spring, reflects karst groundwater 

that has likely been impacted by agricultural 

activities with the highest NO3 concentrations of 

the end-members and moderate Cl and SO4

concentra-tions. Being a rather large spring with 

flow on the order of that at Fay Spring, this 

spring can be interpreted as part of the broader 

regional conduit flow system, comprised itself of 

mixed sources. EM3 is the most impacted by the 

urban environment, and is represented by the 

ephemeral sinking surface runoff of Sunnyside 

Run. When sampled during the snowmelt 

period, it had high Cl and SO4 concentrations 

with moderately high NO3 concentration; 

however, it is expected that the composition of 

this end-member would not be constant through 

time. 

The changing proportions of the 

hypothesized geochemical end-members shown 

in time series illustrate dilution of the spring 

water across hydrographs in response to large 

rainfall events; however, during winter the 

spring water is impacted by a high Cl 

component instead. It is difficult to define the 

actual sources of water responsible for these 

contrasting effects, because the composition of 

any single end-member is unlikely to be constant 

in time. For example, although Sunnyside Run 

(EM3) showed high Cl values during the 

snowmelt period in which samples were 

collected, this ephemeral flow may show a 

different composition in another season. Further 

sampling is needed to account for this seasonal 

variability. 

In conclusion, a combination of approaches 

including dye tracing and interpretation of 

natural geochemistry is helpful for elucidating 

the nature of discharge for karst springs in 

fractured carbonate regions. In this case study, 

the chemistry of Fay Spring appears to show

dominance of a deeper aquifer component at 

base flow, and dominance of shallow groundwater as opposed to sinking surface water runoff

during events. The geochemistry of the deeper 

component indicates impacts by both 

agricultural and urban activities, resulting in 

elevated background levels of NO3 and Cl in the 

spring discharge greater than that observed in 

groundwater from a well farther from the urban 

center. 
34

ACKNOWLEDGMENTS

We thank George Harlow and Roger 

Moberg of the USGS Virginia Water Science 

Center (VWSC) for installing and maintaining 

the continuous monitoring station at Fay Spring. 

Jim Lawrence of Opequon Watershed, Inc. 

provided valuable field assistance and use of 

equipment. Chemical analyses were performed 

by Mike Doughten at the USGS National 

Research Program (NRP) lab in Reston, Va. 

Helpful insight and discussions were provided 

by Niel Plummer (USGS, NRP) and David 

Nelms (USGS, VWSC). Finally, we thank 

Clearbrook Fire & Rescue and Whit Wagner for 

assisting with the dye trace. Partial funding for 

this project was provided to N.C. Farrar through 

the Harrison Undergraduate Research Award 

from the University of Virginia.

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Cady, R.C., 1936, Ground-water resources of the 

Shenandoah Valley, Virginia. Virginia 

Geological Survey Bulletin 45. 137 pp., 2 plates.

Christopherson, N., and Hooper, R.P., 1992, 

Multivariate analysis of stream water chemical 

data: the use of principal components analysis 

for the end-member mixing problem. Water 

Resources Research, vol. 28, no. 1, p. 99-107

Doctor, D.H., Alexander, E.C., Jr., Petric, M., 

Kogovsek, J., Urbanc, J., Lojen, S., and Stichler, 

W., 2006, Quantification of karst aquifer 

discharge components during storm events 

through end-member mixing analysis using 

natural chemistry and stable isotopes as tracers, 

Hydrogeology Journal, vol. 14, p. 1171-1191

Harlow, G.E. Jr., Orndorff, R.C., Nelms, D.L., 

Weary, D.J., and Moberg, R.M., 2005, 

Hydrogeology and Ground-Water Availability in 

the Carbonate Aquifer System of Frederick 

County, Virginia, U.S. Geological Survey 

Scientific Investigations Report 2005-5161, 30 p.

Jones, W.K., 1991, The carbonate aquifer of the 

Northern Shenandoah Valley of Virginia and 

West Virginia, in Kastning, E.H., and Kastning, 

K. M., editors, Proceedings of the Appalachian 

Karst Symposium, Radford Virginia, March 23-

26, 1991, p. 217-222.

Jones, W.K., 1997, Karst hydrology atlas of West 

Virginia, Special Publication no. 4, Karst Waters 

Institute, Charles Town, West Virginia, 111 pp.

Kozar, M.D., McCoy, K.J., Weary, D.J., Field, M.S., 

Pierce, H.A., Schill, W.B., and Young, J.A., 

2007a, Hydrogeology and Water Quality of the 

Leetown Area, West Virginia. U.S. Geological 

Survey Open File Report 2007-1358, 99 pp., 6 

appendices. 

Kozar, M.D., Weary, D.J., Paybins, K.S. and Pierce, 

H.A., 2007b, Hydrogeologic setting and groundwater flow in the Leetown area, West Virginia, 

U.S. Geological Survey Scientific Investigations 

Report 2007-5066, 80 pp.

McCoy, K.J., and Kozar, M.D., 2007, Use of 

sinkholes and specific capacity distributions to 

assess vertical gradients in a karst aquifer, 

Environmental Geology, DOI: 10.1007/s00254-

007-0889-1.

Nelms, D.L. and Moberg, R.M., 2010, Hydrogeology 

and Groundwater Availability in Clarke County, 

Virginia, U.S. Geological Survey Scientific 

Investigations Report 2010-5112, 119 p.

Orndorff, R.C. and Harlow, Jr., G.E., 2002, 

Hydrogeologic Framework of the Northern 

Shenandoah Valley Carbonate Aquifer System. 

Proceedings of the U.S. Geological Survey Karst 

Interest Group, Shepherdstown, West Virginia, 

August 20-22, 2002 (E. Kuniansky, ed.), Field 

Trip Guide, USGS Water Resource 

Investigations Report WRI-02-4174, p 81-89.

Orndorff, R.C, Weary, D.J. and Parker, R.A., 2004, 

Geologic Map of the Winchester Quadrangle, 

Frederick County, Virginia. U.S. Geological 

Survey Open-File Report 03-461, 1 plate.

Perry, L.D., Costain, J.K., and Geiser, P.A., 1979, 

Heat flow in western Virginia and a model for 

the origin of thermal springs in the folded 

Appalachians. Journal of Geophysical Research, 

v. 84 no. B12, p. 6875-6883.

Vesper, D.J., Grand , R.V., Ward, K. and Donovan, 

J.J., 2009, Geochemistry of a spring-dense karst 

watershed located in a complex structural 

setting, Appalachian Great Valley, West 

Virginia, USA. Environmental Geology, vol. 58, 

no. 3, p. 667-678.

White, W.B, 1988, Geomorphology and Hydrology 

of Karst Terrains, New York: Oxford University 

Press, 464 p.
35

Aqueous Geochemical Evidence of Volcanogenic Karstification: 

Sistema Zacatón, Mexico

By Marcus O. Gary1, Daniel H. Doctor2, and John M. Sharp, Jr.3

1 Zara Environmental LLC, 1707 West FM 1626, Manchaca, TX 78652 

2 U.S. Geological Survey,12201 Sunrise Valley Drive, MS 926A, Reston, VA 20192-0002

3 Department of Geological Sciences, Jackson School of Geosciences The University of Texas, Austin, 

TX 78712-0254

Abstract

The Sistema Zacatón karst area in northeastern Mexico (Tamaulipas state) is limited to a relatively 

focused area (20 km2

) in a carbonate setting not prone to extensive karstification. The unique features

found here are characteristic of hydrothermal karstification processes, represent some of the largest 

phreatic voids in the world, and are hypothesized to have formed from interaction of a local Pleistocene 

magmatic event with the regional groundwater system. Aqueous geochemical data collected from five 

cenotes of Sistema Zacatón between 2000 and 2009 include temperature (spatial, temporal, and depth 

profiles), geochemical depth profiles, major and trace ion geochemistry, stable and radiogenic isotopes, 

and dissolved gases. Interpretation of these data indicates four major discoveries: 1) rock-water 

interaction occurs between groundwater, the limestone matrix, and local volcanic rocks; 2) varying 

degrees of hydrogeological connection exist among cenotes in the system as observed from geochemical 

signatures; 3) microbially-mediated geochemical reactions control sulfur and carbon cycling and 

influence redox geochemistry; and 4) dissolved gases are indicative of a deep volcanic source. Dissolved 

87Sr/86Sr isotope ratios (mean 0.70719) are lower than those of the surrounding Cretaceous limestone 

(0.70730-0.70745), providing evidence of groundwater interaction with volcanic rock, which has a 

87Sr/86Sr isotope ratio of 0.7050. Discrete hydraulic barriers between cenotes formed in response to 

sinkhole formation, hydrothermal travertine precipitation, and shifts in the local water table, creating 

relatively isolated water bodies. The isolation of the cenotes is reflected in distinct water chemistries 

among them. This is observed most clearly in the cenote Verde where a water level 4-5 meters lower than 

the adjacent cenotes is maintained, seasonal water temperature variations occur, thermoclines and 

chemoclines exist, and the water is oxic at all depths. The surrounding cenotes of El Zacatón, Caracol, 

and La Pilita show constant water temperatures both in depth profile and in time, have similar water 

levels, and are almost entirely anoxic. A sulfur (H2S) isotope value of δ34S = -1.8 ‰ (CDT) in deep water 

of cenote Caracol, contrasted with two lower sulfur isotopic values of sulfide in the water near the surface 

of the cenote (δ34S = -7 ‰ and -8 ‰ CDT). These δ34S values are characteristic of complex biological 

sulfur cycling where sulfur oxidation in the photic zone results in oxidation of H2S to colloidal sulfur near 

the surface in diurnal cycles. This is hypothesized to result from changes in microbial community 

structure with depth as phototropic, sulfur-oxidizing bacteria become less abundant below 20 m. Unique 

microbial communities exist in the anoxic, hydrothermal cenotes that strongly mediate sulfur cycling and 

likely influence mineralization along the walls of these cenotes. Dissolved CO2 gas concentrations ranged

from 61-173 mg/L and total dissolved inorganic carbon (DIC) δ13C values measured at cenote surfaces 

ranged from -10.9 ‰ to -11.8 ‰ (PDB), reflecting mixed sources of carbon from carbonate rock 

dissolution, biogenic CO2 and possibly dissolved CO2 from volcanic sources. Surface measurements of 

dissolved helium gas concentrations range from 50 nmol/kg to 213 nmol/kg. These elevated helium 

concentrations likely indicate existence of a subsurface volcanic source; however, helium isotope data are 

needed to test this hypothesis. The results of these data reflect a speleogenetic history that is inherently 

linked to volcanic activity, and support the hypothesis that the extreme karst development of Sistema 

Zacatón would likely not have progressed without groundwater interaction with the local igneous rocks.
36

The Influence of Land Use and Occurrence of Sinkholes on Nitrogen 

Transport in the Ozark Plateaus in Arkansas and Missouri

By Timothy M. Kresse1, Phillip D. Hays1, Mark R. Hudson2, James E. Kaufman3

1

U.S. Geological Survey, 401 Hardin Road, Little Rock, AR, 72211 2

U.S. Geological Survey, Box 25046, Denver Federal Center, Denver, CO, 80225-0046 

3 U.S. Geological Survey, 1400 Independence Road, Rolla, MO 65401-2602

Abstract

The presentation will focus on results from a 4-year study that occurred in 3 phases. Phase 1 

included a data mining exercise that resulted in over 850 sites (wells and springs) with nitrate-N 

concentrations for 8 counties in northwest Arkansas. Statistical analysis showed a strong correlation to 

increasing nitrate-N and increasing cleared (agricultural) land use. Using only spring data, rather than 

combined springs and wells, resulted in better correlation and higher r2 values. The Ozarks are divided 

into three hydrologic units: the Western Interior Plains confining system of the Boston Mountains, the 

Springfield Plateau aquifer, and the Ozark aquifer (Salem Plateau). The Springfield Plateau aquifer was 

observed to have statistically significant greater mean NO3-N concentrations than the other two units. 

Using this information, only springs were selected for sampling in the Buffalo River watershed in areas 

with outcropping Boone Formation (Springfield Plateau aquifer) and in areas with dominantly agricultural

land use for the phase 2 study. By controlling the influence of these variables, shown to be dominant 

variables affecting nitrate-N concentrations in the phase 1 study, approximately half of the fifty-six 

springs sampled for the study were in areas with a high density of sinkholes and the other half in areas 

devoid of mapped sinkholes to test the influence of karst features on nutrient transport. Nitrate-N 

concentrations were significantly higher in areas with high sinkhole density using non-parametric 

statistics. Because an inspection of the data revealed a strong correlation between nitrate-N concentration 

and agricultural land use using regression analysis, an analysis of covariance was used to test the level of 

importance of sinkhole density as a secondary variable. Results demonstrated that at a 10 percent 

significance level, nitrate-N concentrations were higher for increasing agricultural land use for springs in 

areas of high sinkhole density. Results from the first two phases were compared to available data from the 

Ozark Scenic Riverways in southeastern Missouri (phase 3), where data in regard to mapped sinkholes 

and defined spring recharge areas are much more detailed than that from Arkansas. Using water quality 

available for 9 springs in this area showed a strong relation (r2 = 0.86) between nitrate-N concentrations 

and agricultural land use, with a lower, but positive relation between nitrate-N concentrations and 

sinkhole density (r2 = 0.19). The data set was too small for use of multivariate analysis, but indicates that 

the occurrence of sinkholes influences nitrate transport, although possibly to a lesser degree than 

percentage of agricultural land use. Results from this study provide the Buffalo River Park Service and 

other land use mangers with a potentially strong management tool for identifying areas within a given 

watershed that are more vulnerable to nutrient sources based on the occurrence of sinkholes. An 

interesting finding from this study is that nitrate-N concentrations exceed regional mean concentrations 

where agricultural land use exceeds approximately 30-40 percent of total land use. This suggests that 

shifting of land use from dominantly forested to agricultural land use exceeding 30-40 percent may result 

in elevated NO3-N concentrations and consequently may increase the potential for degradation of 

groundwater resources. With further corroboration, this finding additionally could be an important 

criterion for guiding land-management planning.
37

Seasonal Carbon Dynamics in a Northwestern Arkansas Cave: 

Linking Climate and Cave Conditions

By Erik D. Pollock1, Katherine J. Knierim2, and Phillip D. Hays3

1

University of Arkansas Stable Isotope Lab, Department of Biological Sciences, University of Arkansas, 

Fayetteville Arkansas 72701

2

Environmental Dynamics Program, University of Arkansas, Fayetteville, Arkansas 72701 3 U.S. Geological Survey, Arkansas Water Science Center, Fayetteville, Arkansas 72701

Abstract

The transport of carbon (inorganic and organic and solid, aqueous, and gaseous phases) was 

examined in a simple, well-characterized cave system to understand the seasonal variations and 

meteorological effects on carbon cycling. Soil- and cave-gas and water samples were collected monthly 

from August 2008 until August 2009 and cave atmosphere carbon dioxide was continuously monitored

from April 2009 until September 2009. Water samples were analyzed for the concentration and isotopic 

composition (δ13C) of dissolved inorganic and organic carbon. Gas samples were analyzed for carbon 

dioxide concentration and isotopic composition (δ13C). Seasonal variations were apparent in cave carbon 

dioxide concentration, ranging from near atmospheric concentrations in the winter (580 parts per million) 

to over 4,200 parts per million in the summer. Carbon dioxide δ13C ranged from -10.79 per mil (‰) in 

the winter to -23.18‰ in the summer. Water sample dissolved inorganic carbon concentrations ranged 

from 2.9 to 6.9 milligrams per liter with δ13C values from -24.19 to -9.47‰. Dissolved organic carbon 

ranged from 1.1 to 4.4 milligrams per liter with 13δ C values from -29.78 to -24.71‰. Cave-atmosphere 

carbon dioxide concentration followed a seasonal pattern with high concentrations in the summer, 

indicating soil gas (~10,000 parts per million) infiltration into the cave air, and low concentrations in the 

winter as atmospheric carbon dioxide (~380 parts per million) ventilated the cave. Surface air 

temperature was positively correlated with carbon dioxide concentrations in the cave. The carbon 

isotopic composition of both the dissolved carbon species and carbon dioxide did not follow a simple 

seasonal pattern. The system varied from more negative carbon isotope values in the summer, 

representative of soil gas input, to more positive values during the spring and fall. However, during the 

winter a more negative excursion occurred in dissolved inorganic carbon and carbon dioxide isotopic 

compositions. This excursion is interpreted as the dynamic mixing of soil air and atmospheric air.

INTRODUCTION

The transport of carbon (inorganic and 

organic and solid, aqueous, and gaseous phases)

was examined in a simple, well-characterized 

cave system to understand the seasonal 

variations and meteorological effects on carbon

(C) cycling. In karst systems, the welldeveloped interaction between surface and 

groundwater and soil, cave, and atmosphere gas

masses and the interplay of biotic and abiotic 

processes provide multiple, bidirectional 

pathways for C movement between organic and 

inorganic pools resulting in a system of 

considerable complexity (Fairchild and others, 

2006). Thorough understanding of C processing 

in these systems is critical to the development of 

effective land-use and agricultural management 

practices, as well as protection of sensitive karst 

ecosystems (Davis and others, 2000). Important 

inorganic processes (such as bedrock dissolution 

and mineral precipitation) and biogeochemical 

processes (such as nutrient attenuation) are 

dependent on the quantity and quality of C in 

these distinct pools (Spötl and others, 2005; 

Winston, 2006). For example, calcite 

precipitation in caves is controlled by the 

concentration gradient between dissolved 

inorganic carbon (DIC) in infiltrating water and 

carbon dioxide ([CO2

]) in the cave air (Baldini 

and others, 2006). The [CO2

] in the cave 

atmosphere is a function of the relative 

contribution of soil gas and atmospheric CO2

(Spötl and others, 2005). In one study on the 

interflow zone (the soil zone immediately above 

the epikarst surface), dilution decreased solute 

concentration during storm events, but 

biological processing also contributed to lower 
38

nitrate (NO3

-

) and dissolved organic carbon 

(DOC) concentrations (Winston, 2006). 

Although water residence times were shorter 

during storm events, an influx of labile organic 

matter to the epikarst zone may have increased 

rates of denitrification and decomposition of 

organic matter (Winston, 2006).

For this study, air and water samples from a 

small, well-characterized cave and the overlying 

soil in northwestern Arkansas were analyzed for 

aqueous C species and gaseous CO2 to better 

understand the movement of C on a seasonal 

time scale. 

STUDY SITE

The study site is located in the southern part 

of the Ozark Plateaus Province, which receives 

an average of 112 centimeters (cm) of 

precipitation annually (Adamski and others, 

1995). According to the National Oceanic and 

Atmospheric Administration’s (NOAA) 

National Climatic Data Center for the 

northwestern Arkansas division for the years 

1895-2009 precipitation peaks in May at an 

average of 13.8 centimeters per month and a 

second, lower maximum typically occurs in 

October with 10.5 centimeters per month. 

Annual average air temperature is 14.1°C 

(NOAA, 2009). 

Jack’s Cave (36°17’16”N, 93°40’10”W) is 

representative of the physical and chemical 

hydrogeology of caves and karst in northwestern 

Arkansas. The morphology of most caves in the 

Ozarks comprises single passages controlled by 

joints that terminate in narrowing sedimentfilled rooms (Taylor and others, 2009). The area 

is capped by Mississippian limestone of the St. 

Joe Limestone Member of the Boone Formation 

and other members of the Boone Formation. 

The entrance to the cave is a fracture in the 

Middle Ordovician Kings River Sandstone 

Member of the Everton Formation, and Jack’s 

Cave is developed in an underlying dolomite 

unit of the Everton Formation. The soil is a very 

gravelly silt loam and thickness varies 

depending on the slope (National Cooperative 

Soil Survey, 2008).

Meteoric water recharges the shallow 

groundwater system at the site. In Jack’s cave, 

infiltrating water collects and flows along the 

cave passage towards the southwest. Direct 

hydraulic connections were observed at the 

study site between the soil and cave, but two dye 

traces failed to identify the discharge point(s) of 

the cave (Knierim, 2009).

Figure 1. The location and cave map of Jack’s

Cave showing the cave pool sampling site. Soil 

lysimeters were installed at the surface 

approximately 75 m southwest of the cave 

entrance.

METHODS

Two soil lysimeters were placed upslope 

from the cave at depths of 80 and 90 cm. A 

continuous CO2 monitor was placed in the cave 

approximately 20 meters (m) from the entrance. 

Water was sampled from lysimeters and a small 

cave pool located near the continuous monitor 

(fig. 1). Water and gas samples were collected 

approximately monthly from August 2008 

through August 2009. 

Water samples were collected in 40-

milliliter (mL) precombusted total organic 

carbon (TOC) vials and filtered through 0.45-

micrometer (μm) filters. DOC samples were 
39

treated with 40 microliters of 3.6 M sodium 

azide, and all samples were stored at 4°C until 

analysis. Gas samples were collected in 

precombusted 100-mL serum vials purged with 

helium. Water samples were analyzed for the 

concentration and isotopic composition (δ13C) of 

DIC and DOC at the Colorado Plateau Stable 

Isotope Laboratory using continuous flow 

isotope ratio mass spectrometry (IRMS), TOCIRMS (±0.75‰; St-Jean, 2003). [DIC] and 

[DOC] were analyzed using an Aurora IO TOC 

analyzer (accuracy ±0.1 milligram per liter

(mg/L); St-Jean, 2003). Discrete samples of 

[CO2] were measured monthly using a Vaisala 

CARBOCAP ® handheld CO2 meter (accuracy 

±1.5% of the range + 2% of the reading; Vaisala, 

2008), and a LI-COR 840 CO2/H2O gas analyzer 

(accuracy ± 1.5% of the reading; LI-COR, Inc., 

2003) was installed in the cave for continuous 

measurements. Gas samples were analyzed for 

[CO2] and δ13C-CO2 at the University of 

Arkansas Stable Isotope Laboratory using gas 

chromatography-combustion IRMS (±0.56‰). 

Additionally, samples of TOC in soil (n=12) and

carbonate in bedrock (n=5) were analyzed for 

carbon isotope ratios. Soil samples were 

collected at approximately 15-cm intervals from 

the holes in which the lysimeters were installed. 

Bedrock samples were collected from the St Joe 

Limestone Member of the Boone Formation 

above the cave and from dolomite of the 

Ordovician Everton Formation within the cave.

All isotope ratios are reported using delta 

notation (δ) in per mil (‰) relative to VPDB 

(Vienna Peedee Belemnite). The δ notation 

indicates the ratio of the heavier to the lighter 

isotope (13C/12C) relative to VPDB (Clark and 

Fritz, 1997). 

Daily temperature and precipitation data 

from NOAA’s weather station at Huntsville, AR 

(located approximately 23 km southwest of the 

study location; NOAA, 2010) were averaged for 

the 14-day period prior to and including the date 

of sample collection. The 14-day averages 

eliminated short-term fluctuations in 

temperature and precipitation.

RESULTS

Seasonal changes occurred in C species in 

soil and cave gas and water C isotope ratios and 

concentrations. Soil gas [CO2] ranged from 

1,400 parts per million (ppm) (April 2009) to 

19,450 ppm (July 2009) (data not shown). The 

[CO2] in the atmosphere of Jack’s Cave ranged 

from 4,200 ppm in the summer to near outside 

atmospheric concentrations (580 ppm) in the 

winter (fig. 2). Continuous data were averaged 

to provide daily readings, and the concentrations 

recorded by the continuous monitor were in 

good agreement with discrete monthly 

concentrations. Continuous data only were 

recorded for a part of the study period (April 

2009 to September 2009) because of equipment 

problems. The daily averages recorded highfrequency variation of [CO2] (fig. 3); daily 

changes of several hundred ppm in [CO2] were 

common (fig. 3). The cave air temperature was 

nearly constant over the study period (fig. 2), 

and high frequency changes in CO2

concentration were not related to precipitation 

events (fig. 3).

Figure 2. Cave air carbon dioxide concentration 

([CO2]) and cave air temperature over the course 

of the study compared to surface air temperature

from August 2008 to August 2009. 

The aqueous carbon concentrations in Jack’s 

Cave did vary though not in a simple seasonal 

pattern. DOC concentration ranged from 1.1 to 

4.4 milligrams per liter (mg/L) and DIC 

concentration ranged from 2.9 to 6.9 mg/L (data 

not shown). The highest DOC concentration 

was recorded in January 2009 and the lowest 

concentration was in December 2008 (fig. 4). 
40

Figure 3. Daily average cave air carbon dioxide 

concentration ([CO2]) recorded by the 

continuous monitor and precipitation from April, 

2009 to September, 2009. 

Figure 4. Dissolved organic carbon (DOC) 

concentration and isotopic composition of DOC 

(δ13C-DOC) from September 2008 to August 

2009 in Jack’s Cave.

The isotopic ratios of both the cave air CO2

and the DIC showed similar but not identical 

patterns. The cave-air isotopic composition of 

CO2 (δ13C-CO2) ranged from -23.18‰ in the 

summer to -10.79‰ in the winter (fig. 5). The 

data exhibit bimodal negative periods. During 

the warm season, the CO2 isotopic compositions 

show a long period of more negative isotope 

ratios with values all less than -17‰. During 

the cool season, a second period of more 

negative ratios occurred, from January through 

March 2009, with consistent δ13C values of 

approximately-15.6 ±0.3‰ VPDB. The C 

isotope ratios of DIC in Jack’s Cave varied from 

-24.19 to -9.47‰ VPDB (fig. 5). More negative 

isotopic ratios were observed during the warm 

season. A period of negative isotope ratios also 

occurred during the winter, separated by a 

preceding and a following more positive periods 

(fig. 5). The negative excursion is similar to that 

observed in the δ13C-CO2 of the cave-air. The 

difference between the δ13C-CO2 and the δ13CDIC was variable (-5‰ to 12‰) (fig. 5).

Figure 5. The isotopic composition (δ13C) of 

dissolved inorganic carbon (DIC) and carbon 

dioxide (CO2) in Jack’s Cave compared to [CO2] 

from September, 2008 to August 2009. The 

solid lines represent equilibrium enrichment 

between DIC and CO2 at 15°C.

The isotopic composition of DOC (δ13CDOC) ranged from -29.78 to -24.71‰ (fig. 4) 

and showed a similar but inverse pattern to that 

of the cave CO2. During the cool season, a 

period of more positive isotope ratios was 

observed from February to April. However, the 

timing of the cool-season light ratio period for 

δ13C DOC response was delayed relative to the 

δ13C-CO2 response (February for DOC 

compared to January for CO2).

Bedrock samples from the overlying St. Joe 

Limestone Member of the Boone Formation had 

a C isotope ratio of -1.36‰ ±0.01 (n=2). 

Dolomite samples from the Everton Formation

collected within the cave had C isotope ratio of 

-3.7‰ ±0.43 (n=2). Soil samples collected 
41

from the lysimeter holes had an average δ13C of 

-25.15‰ ±0.44 (n= 9). 

DISCUSSION

The [CO2] in the cave displayed a seasonal 

signature as recorded by monthly readings. A 

highly variable [CO2] was measured by the 

continuous monitor. The [CO2] was lowest 

during the colder months and was correlated to 

surface air temperature. As surface air 

temperature increased, [CO2] increased 

exponentially (r2 = 0.82, data not shown) for the 

monthly data. The decrease in [CO2] during 

periods of cooler surface temperatures has been 

observed in other cave systems (Baldini and 

others, 2008; Spötl and others, 2005). Seasonal 

temperature variations change the density 

contrast between the relatively static cave air 

and the outside air, which causes ventilation 

patterns to alternate seasonally. During the 

cooler months, dense, isotopically heavy, low 

[CO2] surface air is pulled into the cave. As the 

surface temperature warms, this ventilation shuts 

down. Without this ventilation the cave air is 

dominated by CO2-rich, isotopically more 

negative soil gas as a result of direct gas 

diffusion. The relation of the daily reading for 

[CO2] to surface air temperature was less clear 

(r2

= 0.45, data not shown). The changes in daily 

[CO2] also were not explained by changes in 

precipitation (r2

= 0.04). For some settings, the

rewetting of soils has been shown to cause the 

release of a pulse of CO2 (Xiang and others, 

2008). Other research suggests that high or low 

soil water content can retard the production of 

CO2 (Davidson and others, 1998). As such, the 

high frequency variation in the daily readings is 

likely a result of changes in the soil temperature 

and moisture, which are not currently resolved 

by our data. 

The concentrations of DOC (fig. 4) and DIC 

did vary though not in clear seasonal pattern. 

The isotopic composition of DIC is a mixture of 

organically derived soil-gas CO2 and bedrock 

carbonate, and varies seasonally. High soil gas 

contributions of relatively negative 

(approximately -22‰) CO2 during the warm 

season decreased δ13C-DIC values (fig. 5). Cool 

season δ13C-DIC ratios were higher, showing an 

increase in the relative bedrock (-1.36‰ to 

-3.7‰) contribution to the DIC pool.

The difference between the δ13C-DIC in the 

cave pool and δ13C-CO2 in the cave air showed 

that the system is in isotopic disequilibrium; the 

difference in the two isotopic compositions is 

variable (fig. 5) and generally less than the 

-7.77‰ that would be predicted by the relevant 

fractionation factor (Mickler and others, 2004). 

This indicates that the cave-air CO2 is largely 

derived from soil gas during the warm season. 

The time required for isotopic equilibration for 

the DIC-CO2 system is on the scale of hours 

(Zhang and others, 1995), sufficiently rapid for 

equilibrium to be established in the relatively 

low-flow conditions found in Jack’s Cave. The 

lack of equilibrium suggests that the cave air is 

not from evolved DIC, but is from direct input 

of soil CO2. 

Winter [CO2] were lower and δ13C-CO2

values were heavier in Jack’s Cave than during 

warmer periods (fig. 2); this is indicative of a 

change in the cave air ventilation as relative 

surface-cave air density values changed with 

cooling temperature (Spötl and others, 2005). 

The reversal allowed for surface air, relatively 

CO2-poor (approximately 380 ppm) and 

isotopically heavy (approximately -7.7‰) to 

enter the cave. However, the heaviest δ13C-DIC 

values were not consistently observed during the 

period with the lowest [CO2] (figs. 2 and 5). 

Both the δ13C-CO2 and δ13C-DIC values in the 

cave were slightly lighter between January and 

February (fig. 5). The respiration of an organic 

C pool would provide a source for the lighter 

isotopic ratios observed in the cave. The 

addition of soil gas to the cave air would explain 

the higher values; however no increase in the 

[CO2] was observed (soil [CO2] averaged 6,880 

ppm over the period). The concentration and 

isotopic ratios of the DOC support the idea of 

organic respiration. Biological processing of 

DOC would preferentially use the labile, 

isotopically higher C (Santrucková and others, 

2000). This would have two effects on the DOC 

pool: (1) the pool itself would shrink and (2) the 

remaining pool would become progressively 

enriched in the heavier isotope (13C) (Garten and 

others, 2000). Over the same time period that 

the cave-air δ13C CO2 exhibits a negative 

departure, DOC concentration is decreasing and 

the isotopic composition is becoming higher. 

The change in DOC occurs during the cooler 

and drier portion of the year when the water 
42

movement between the surface and the cave is 

lower and less able to transfer mass. The soil 

system is also less able to provide those 

materials in dry, cold, low-productivity 

conditions, which decrease surface biological 

processing and subsequent respiration. 

However, the reduction in the DOC content is 

not sufficient to account for the more negative 

isotope ratios alone; insufficient C exists in the 

organic C pool. The δ13C-CO2 in the winter is a 

mixture of atmospheric, soil gas and respired 

DOC, and the respired DOC accounts for the 

smallest contribution to cave CO2. 

CONCLUSION

The [CO2] in the cave displayed not only a 

seasonal signature, but also a highly variable 

daily signal. The [CO2] in the atmosphere of 

Jack’s Cave ranged from 4,200 ppm in the 

summer to near atmospheric levels (580 ppm) in 

the winter. During the cooler months, dense, 

isotopically heavy, low-CO2 concentration 

surface air was pulled into the cave. The high 

frequency variation in the daily readings is likely 

a result of changes in the temperature and soil 

moisture, which are not currently resolved by 

our data. The difference between the δ13C-DIC 

in the cave pool and δ13C-CO2 in the cave air 

showed that the system is in isotopic 

disequilibrium. The lack of equilibrium 

suggests that the cave air is not from evolved

DIC, but is from direct input of soil CO2. The 

δ13C-CO2 and δ13C-DIC values in the cave were 

slightly lighter between January and February. 

The addition of soil gas to the cave air would 

explain the higher values; however no increase 

in [CO2] was observed. The respiration of an 

organic C pool would provide a source for the 

lighter isotopic ratios observed in the cave. 

However, the reduction in the DOC content is 

not sufficient to account for the more negative 

isotope ratios alone; insufficient C exists in the 

organic C pool. The δ13C-CO2 in the winter is a 

mixture of atmospheric, soil gas, and respired 

DOC, and the respired DOC accounts for the 

smallest contribution to cave CO2.

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44

Characterization of Urban Impact on Water Quality of Karst Springs in 

Eureka Springs, Arkansas

By Renee Vardy

Environmental Dynamics, 113 Ozark Hall, University of Arkansas, Fayetteville, AR 72701

Abstract

Shallow karst aquifers in northwest Arkansas are highly vulnerable to urban land use; many 

springs have shown contamination from dissolved constituents and by indicator bacteria. Eureka Springs, 

Arkansas is located on the karst terrain of the Ozark Plateau and more than 100 springs have been 

identified within the city limits. Through collaborative efforts to characterize spring water quality and 

potentially identify contaminant sources, the U. S. Geological Survey, University of Arkansas, City of 

Eureka Springs, and Arkansas Department of Environmental Quality have partnered in three recent 

studies. For these studies, spring water samples were collected and analyses were conducted for Organic 

Wastewater Constituents (OWC’s), indicator bacteria (E. coli and fecal coliform), selected geochemical 

water quality parameters, and flow. OWC’s have been detected in both stream sites and one of seven 

springs. Indicator-bacteria colony counts in the springs ranged from 2 (estimated(e)) to 500(e) cfu 

(colony forming units)/100mL with an average of 115 cfu/100mL for E. coli, and from 8(e) to 130 

cfu/100mL with an average of 59 cfu/100mL for fecal coliform. Dissolved nitrate, and chloride 

concentrations in the springs are greater than natural background levels, with averages of 2.75 mg/L and 

11.42 mg/L respectively. These data suggest the influence of urban contaminants in the karst aquifer, 

highlighting the vulnerable character of the karst hydrologic system. 

INTRODUCTION

Accounting for approximately 25% of 

worldwide water needs, karst aquifers are a 

vital water source for many areas of the 

world (Ford and Williams, 2007). When

urbanized areas encroach on karst recharge 

areas, these aquifers are extremely 

vulnerable to degradation (Jiang et al.,

2009). As urban populations continue to 

grow, karst areas will be subjected to 

greater development pressures, and

potentially urban contaminants. Protection 

of karst aquifers from impacts of 

urbanization is imperative if the water 

needs of millions of people are to be met, 

which will be increasingly difficult as 

urbanization expands (Showers et al., 

2008). 

Karst aquifers are composed of a 

framework of interconnected conduits,

fractures, voids and pores formed by 

dissolution of the host rock resulting in a 

complex flow system with intimate 

connection between surface water and 

groundwater. Urban contamination is a 

common source of degradation of many 

karst aquifers, not only in the southern 

Ozarks (Adamski, 1997), but also 

throughout the world (Jiang et al., 2009; 

Murray et al., 2007; Zhao et al., 2010). 

These aquifers are especially vulnerable to 

contamination for several reasons. For one, 

rather than being composed of a finegrained medium capable of effective 

filtration, as is the case for granular-media 

aquifers, the open conduits in karst aquifers 

typically allow unimpeded, unfiltered 

channel-type flow. In addition, karst areas 

generally have relatively thin soils,

providing minimal filtration between 

surface runoff and the aquifer (Ford and 

Williams, 2007). Sinkholes and other 

solutional features also provide direct 

connections between surface runoff and 

aquifer. Underground, leaky storage tanks 

or damaged piping can lead to direct 

contamination of the aquifer. These are 

just a few of the many contamination 

pathways and sources; others include 

surface runoff, septic-system effluent, road 
45

salt, automotive fluids, household products 

and industrial chemicals. (Davis et al., 

2000; Zhao et al., 2010; Liu et al., 2006; 

Panno et al., 2001). 

As the world population nears 7 billion, 

the finite water resources available will be 

stressed to accommodate this growing 

population. By 2025, close to five billion 

people are expected to live in urban areas, 

nearly double the 1995 statistic (United 

Nations, 1997). The amount of water 

required for these urban populations would 

significantly increase, escalating in 

conjunction with projected agricultural 

demands. This leads to a dilemma, to serve 

the basic domestic water needs, or to serve 

the agriculture food source requirements

(Postel, 2000). Areas of the world, stressed 

by water usage today, will only experience 

increasing turmoil as the struggle between 

food production, urban water needs, and 

monetary gain of water rights sales 

continues (Postel, 2000). Owing to the 

overwhelming trend of increasing 

population (United Nations, 2009) with 

access to finite resources, even the small, 

currently unusable, water sources cannot be 

ignored. The restoration and development 

of these abused and unused water sources 

will be increasingly important as the 

present finite water resources are allocated. 

The project described in the following 

pages has been designed to identify sources 

of urban contamination, suggest options to 

alleviate the problems, and to restore the 

shallow karst aquifer to a usable condition 

in hopes of reestablishing a viable water 

resource for the city of Eureka Springs. 

Additionally, it is hoped this case study 

will provide valuable information on the 

structure and vulnerability of karst aquifers 

in similar settings elsewhere. 

Study Area

Eureka Springs is a city of ~2,500 

residents, in northwestern Arkansas (Figure 

1). Of the known springs, none have 

acceptable water quality to serve as a 

public drinking water supply. Recent 

studies have shown that many of the 

springs continue to be contaminated by 

nitrate, organic wastewater constituents 

(OWC’s), and indicator bacteria (Hays, 

personal communication; Eureka Springs 

Springs Committee, personal 

communication). 

Eureka Springs was founded in 1879 as 

a resort town known for the healing powers 

of its springs. Although earliest water 

quality analyses show little uniqueness to 

the water, tourists were drawn to the city 

and its economy flourished as bathhouses, 

hotels, and other attractions opened to 

accommodate the guests. By late 1879, the 

population was over 10,000 and it was the 

fourth largest city in Arkansas. Today the 

population has dropped and tourism is the 

towns largest industry (Eureka Springs 

Historical Society, 2010). Although there 

are other attractions within the city that 

draw tourists, the springs are an integral 

part of the allure. At their current levels of 

contamination, signs warn of the danger of 

drinking the water. Several of the springs 

are even considered public hazards by the 

Arkansas Department of Environmental 

Quality owing to their high indicator 

bacteria counts. Iron gates now restrict 

access to many springs and none of them 

are suitable for potable consumption 

without treatment.

Figure 1: Distribution of sinkholes and aquifer 

outcrops in northwestern Arkansas (Renken, 1998)
46

For this area, available background 

values show springs in the study area are 

impacted by an external contaminant 

source(s). Nearby studies of pristine 

springs have shown nitrate concentrations 

are < 1 mg/L, typically about 0.3 mg/L 

(Adamski, 1997). In the study site, nearly 

all nitrate samples analyzed are above 1 

mg/L (Table 1). Also, OWC’s have been 

detected in one of seven springs and both 

stream sites at low flow. Acceptable levels 

for many OWC’s have not been 

established. Detections of OWC’s in 

northwest Arkansas are primarily 

associated with wastewater treatment plant 

(WWTP) effluent (Galloway et al., 2004). 

OWC’s have been detected in areas of low 

population density in northwest Arkansas, 

but they are less prevalent than in sites 

receiving WWTP effluent. In Eureka 

Springs, OWC’s are detected both above 

and below the WWTP, as well as in one 

spring at low flow. Bacteria also are a 

common contaminant in karst systems in 

northwest Arkansas (Davis et al., 2000). 

Their presence can also be highly variable 

in streams and springs in the area. Low 

impact areas, even within the study area, 

can display values as low as 2 cfu/100 mL 

E. coli, and 8 cfu/100 mL fecal coliform 

(Table 1). Other sites within the study 

area show counts of as many as 1200 

cfu/100 mL E. coli, and 38,360 cfu/100 mL 

total coliform (Table 1). These data 

suggest that the springs and karst aquifer in 

Eureka Springs is influenced by urban 

contamination. 

Hydrologic Framework

Eureka Springs is located in an area of 

thinly mantled karst terrain on the Ozark 

Plateau, a 48,000 mi2 area covering parts of 

Arkansas, Missouri, Kansas and Oklahoma 

(Imes and Emmett, 1994). The city lies on 

a steep escarpment between the Springfield 

and the Salem Plateau. The geologic 

formations are the Mississippian aged St. 

Joe Limestone member of the Boone 

Formation and older, Ordovician dolomites 

(Haley et al., 1993; Renken, 1998). These 

two units host the two regional aquifers in 

the area. Springs in this area have 

discharge variable with season and location 

generally ranging from <1 to 5 cfs. 

The St. Joe member of the Boone 

Formation is included in the Springfield 

Plateau aquifer. The St. Joe is a relatively 

pure phase limestone, with little chert. It is 

Table 1. Water Quality Attributes Indicating Anthropogenic Impacts on Springs in Eureka Springs, Arkansas

Pathogens and Microbes: Range Average

E. coli <1 to 1200 cfu/100 mL 115 cfu/100 mL

Fecal Coliform 8 to 130 cfu/100 mL 59 cfu/100 mL

Total Coliform 6 to 38,360 cfu/ 100 mL 2,782 cfu/100 mL

Dissolved Constituents: Range Average

Nitrate 0.63 to 4.63 mg/L 2.75 mg/L

Chloride 2.34 to 34.60 mg/L 11.42 mg/L

Organic Wastewater Constituents: # of Detections

Pesticides 1

Synthetic Musks 2

UV Light Protectant 1

Fragrance/Pharmaceutical Additive 1

Food Additive 1

(Hays, personal 

communication; Eureka Springs 

Springs Committee, personal 

communication)

Flame Retardant 4

Dye/Resins/Asphalt 1

Topical Insect Repellent 2
47

known throughout northwest Arkansas as a 

prominent cave former. Caves and karst in 

the St. Joe are largely aligned with 

solutionally enlarged fractures, which are 

arranged orthogonally and potentially were 

caused by regional uplift. The resulting 

karst terrain allows surface runoff to 

quickly mix with groundwater with little 

filtration (Ford and Williams, 2007, Winter 

et al., 1998; Davis et al., 2000). 

The Chattanooga Shale comprises the 

Ozark confining unit, which separates the 

Springfield Plateau aquifer from the deeper 

Ozark aquifer. The Ordovician aged Cotter 

and Jefferson City Formations (Ozark 

aquifer) lie beneath the Chattanooga. 

Many prominent springs in the area flow 

from the Ozark aquifer. Locally, water 

from the Springfield Plateau aquifer passes 

though the Ozark confining unit by way of 

fractures. These fractures, along with 

direct infiltration into exposed Cotter and 

Jefferson City Formations, cause the Ozark 

aquifer to also be susceptible to 

contamination. 

Historical Studies

Several studies provide an excellent 

water quality database and hydrogeologic 

framework for Eureka Springs. The most 

significant of these are an extensive 

geologic mapping project by Purdue and 

Miser (1916), and an exfiltration study by 

Husain and Aley (1981). 

Although several other geologic maps 

of the area are available, none are as 

detailed as the original map produced by 

Purdue and Miser. In this map localized 

faults and fractures are delineated along 

with precise locations of outcrops. This 

aids in analyzing potential flow within the 

aquifer, and suspected mixing between 

aquifers along fracture zones. 

 The exfiltration study was designed to 

identify leakage in aging sewer lines 

throughout the city. It was funded by a 

group of concerned citizens, who would 

later become the Eureka Springs Springs 

Committee, through a grant from the EPA. 

In this study performed by the 

Comprehensive Planning Institute, optical 

brighteners, chemicals found in most 

laundry detergents, were utilized to identify 

sewage effluent and classify spring 

contamination levels. Dye tracing, as well 

as knowledge of the geology of the aquifercontaining units and the structures present, 

were used to delineate the recharge areas 

for each of the springs (Husain and Aley, 

1981). These data have proven to be a 

great resource for comparisons over the last 

30 years. 

Recent Studies

Since the exfiltration study in 1981, 

several other studies have been conducted 

on the springs of Eureka Springs. These 

include bacterial investigations, student dye 

tracing, hydrologic investigations, general 

geochemistry, and organic wastewater 

constituent sampling (Eureka Springs 

Springs Committee, personal 

communication; Husain and Aley, 1981; 

Brahana et al., 1993; Hays, personal 

communication).

Indicator Bacteria

Many different types of bacteria are 

found in spring water systems, several of 

which are not dangerous for human 

consumption. Those associated with 

wastewater and more specifically, fecal 

matter, can be harmful to human health. 

Several types of fecal bacteria are present 

in the digestive tracts of mammals, and 

passed to the outside environment through 

excrement. Indicator bacteria, easily 

identifiable bacterial species that are 

associated fecal matter, are often used to 

signal fecal input (Davis et al., 2005; 

Carson et al., 2001). At the study site, 

indicator bacteria analysis is used to signify 

the presence of fecal matter and thus 

potentially more dangerous bacterial 

species. 

In Eureka Springs, indicator bacteria 

have been recorded for various springs 

since 2008 (Figure 2). The USGS in 

conjunction with the Springs Committee 

initiated sampling for the common 

indicator bacteria E. coli, fecal coliform, 
48

and total coliform. Counts for these 

bacteria range from <1 to 1,200 cfu/100 

mL (n=38, average=100 cfu/100 mL) for E. 

coli, 8 to 130 cfu/100 mL (n=9, average=59 

cfu/100mL) for fecal coliform and from 6 

to 38,360 cfu/100 mL (n=31, 

average=2,782 cfu/100 mL) for total 

coliform (Table 1, Figure 2; Springs 

Committee, personal communication; 

Hays, personal communication).

Bacteria counts in these springs vary 

significantly with precipitation patterns, 

spring location, and spring aquifer source. 

Precipitation increases runoff, which in 

turn causes a flush of sediment and bacteria 

from the surface into the aquifer or 

potentially from storage within in the 

aquifer (Davis et al., 2005). In a study 

done on E. coli growth and survival 

bacteria were observed to survive in a 

dormant form in the aquifer for up to four 

months. Therefore, if the precipitation 

events flushing fresh bacteria into the karst 

system occur more often than every four 

months, a viable bacterial community 

likely is present in the aquifer at all times, 

capable of being flushed out into the spring 

during rain events (Davis et al., 2005). 

Many of the spring orifices’ are altered 

by decorative additions meant to showcase 

the pride of Eureka Springs, their springs. 

The construction of houses, businesses, 

roadways, parking lots, and drainage 

structures in springs recharge areas, has 

caused a general decrease in flow (Eureka 

Springs Springs Committee, personal 

communication) as well as increased the 

potential for contamination from urban 

sources, such as wastewater. As a 

prominent source of bacteria, wastewater is 

a likely culprit of the high indicator 

bacteria counts in several of the springs. 

Wastewater in the city has two major 

potential sources: leaking sewer lines, or 

improperly functioning septic systems. 

Much of the city’s sewer infrastructure was 

designed and installed prior to 1900. These 

original structures were constructed from 

vitrified clay piping. Many of these pipes 

have been replaced, but several original 

sections within the city limits are still in 

use. These aged pipes are suspected of 

leakage in some areas of the city and could 

be contributing wastewater input to the 

springs. On one occasion, preliminary dye 

tracing discovered connection between a 

toilet and spring within minutes of dye 

introduction (Hays, personal 

communication). 

The other major source of potential 

wastewater contamination in the springs are 

improperly functioning septic systems. 

Properly designed and functioning septic 

systems require the presence of soils to 

disperse the nutrient and bacteria-rich 

wastewater evenly to aid in degradation 

(McQuillan, 2004). In the karst terrain at 

Eureka Springs, soils are generally thin 

leading to little filtration of the wastewater 

between the septic leach lines and the open 

conduits of the karst aquifer. 

Distinct chemical differences are 

observable between the Springfield Plateau 

and Ozark aquifers and the springs that 

recharge from each (Figure 4). Many small 

springs discharge from the Springfield 

Plateau aquifer, which on average show 

higher bacteria counts than their Ozark 

aquifer counterparts. This is likely due to 

the large outcrop of the Springfield Plateau 

aquifer unit in the central part of town 

allowing increased exposure to potential 

contamination sources. The Ozark aquifer 

also shows impacts of bacterial 

contamination in many springs, but the 

Figure 2: Indicator bacteria concentrations for 

selected springs, 2008-2010 (Eureka Springs 

Springs Committee, personal communication)
49

degree of contamination is not as great as 

the Springfield Plateau aquifer, due to the 

presence of the Ozark confining unit that 

separates the two aquifers and the relatively 

small outcrop of the Ozark aquifer in the 

center of the city (Purdue and Miser, 1916). 

Organic Wastewater Constituents

Organic Wastewater Constituents 

(OWC’s) are a relatively new set of 

chemical parameters to be examined in 

water systems because the ability to detect 

them at the concentrations in which they 

typically are found in the environment has 

only recently been developed. This group 

of chemicals includes any substance that is 

passed into wastewater after its intended 

use including antibiotics, pharmaceuticals, 

hormones, personal care products, caffeine, 

over the counter medications, and many 

others. Although the long-term effect of 

these products beyond their intended use 

requires further research, the potential 

effects of exposure to OWC’s include

reproductive impairment, carcinogenic 

effects, and the eventual development of 

antibiotic resistant bacteria (Galloway et 

al., 2004; Kolpin et al., 2002). 

In northwest Arkansas, OWC’s have 

been detected in many streams (Galloway 

et al., 2004). Common OWC’s found in 

the area are caffeine, flame-retardants, 

byproducts of pharmaceutical degradation, 

personal care products, pesticides, and 

insect repellents. The presence of these 

constituents in northwest Arkansas leads to 

concern of their presence at Eureka 

Springs. In September 2009, the USGS, 

Eureka Springs Springs Committee, and the 

Arkansas Department of Environmental 

Quality analyzed water samples for OWC’s 

in Eureka Springs. Of the 62 OWC’s

analyzed, 12 were found in at least one of 

the sampling sites (Hays, personal 

communication). Although the potential 

effects of exposure to OWC’s are 

uncertain, in Eureka Springs, their presence 

could impact a much larger area than 

Eureka Springs (Galloway et al., 2004). 

Leatherwood Creek, the main stream 

draining Eureka Springs, is a tributary to 

the White River. The White River, a 

nationally protected water body and 

municipal water supply to thousands, is 

located downstream of Eureka Springs. 

Not only could contamination of these 

springs lead to a decrease in the livelihood 

of Eureka Springs by discouraging tourism, 

but it could also affect a major regional

water supply. 

General Geochemistry

This study utilizes three separate sets of 

available geochemical data for selected 

springs in the study area. The first data set 

includes total nitrogen, ammonia, nitrate, 

total phosphorous, and reactive 

phosphorous for selected springs from 2008 

to present (Figure 3). 

In Eureka Springs a majority of 

nitrogen detections in groundwater are in 

the form of dissolved nitrate. Since 2008, 

the values found have all been under the 

Environmental Protection Agency 

maximum contaminant level for drinking 

water of 10 mg/L (EPA, 2009), but they 

have also almost all been above the pristine 

levels of >1 mg/L (Adamski, 1997). Basin, 

Harding, and Sweet springs show the 

highest levels of nitrate of any sampled in 

this dataset. These springs are all located, 

and recharged in the central part of town, 

potentially affected by wastewater effluent. 

They are also all sourced from the 

Springfield Plateau aquifer. Magnetic 

Spring is consistently one of the more 

pristine springs. It is sourced from the 

Ozark aquifer with a largely undeveloped 

recharge area on the outskirts of town. 

Leatherwood Creek is the main water body 

draining the central part of town and has 

also historically shown relatively low 

nitrate values (Eureka Springs Springs 

Committee, personal communication). 
50

Figure 3A: Average nitrate concentration for selected 

springs and Leatherwood Creek 2008-2010

Figure 3B: Average ammonia and reactive 

phosphorous concentrations for selected springs and 

Leatherwood Creek 2008-2010

Other contaminants of 

concern, ammonia and 

reactive phosphorous, are 

also detected in area 

springs and streams. 

Basin spring shows the 

highest ammonia 

concentration at 0.05 

mg/L. As with nitrate, a 

potential source of 

ammonia is wastewater 

effluent leaking from 

damaged underground 

piping. Phosphorous, 

also potentially from 

wastewater effluent, is a 

contaminant of major 

concern in northwest 

Arkansas (Soerens et al., 

2003; Benson et al., 

2008).

 A second geochemical data set was 

collected simultaneously with the OWC 

samples. This includes data from 7 springs 

and 2 stream sites within the study area. 

Parameters included in this analysis 

included major cations and anions, metals 

and nutrients (Figure 4).

Through analysis of these geochemical 

data, much important information is gained. 

The characteristic Ozark aquifer spring, 

Magnetic, displays a relatively high 

concentration of Mg2+, as a reflection of its 

dolomitic host rock. The Springfield 

Plateau aquifer springs, with their 

limestone host rock, show high Ca2+

concentration as in the Ozark aquifer, but 

lack Mg2+. The Ozark aquifer also has 

generally lower Cland nitrate 

concentrations than the Springfield Plateau 

aquifer, potentially indicating less affect 

from urban sources. This reflects the 

greater isolation from the anthropogenic 

activities on the surface.

Figure 4: Geochemistry of selected spring and stream sites (Hays, 

personal communication; Purdue and Miser, 1916)
51

Ozark aquifer springs vary in 

chemistry, depending on location. 

Sycamore Spring is an Ozark aquifer 

spring, but it appears to be influenced by 

Springfield Plateau aquifer as indicated by 

the concentrations of Na+ and SO4

2-

. The 

major dissolved constituents look similar to 

that of the Springfield Plateau aquifer 

springs. This apparent mixing suggests a 

fault or other open connection across the 

Chattanooga Shale near this location. 

An integrated mixing of surface runoff 

and the springs can be observed in 

Leatherwood Creek prior to the addition of 

WWTP effluent. Below the WWTP outlet 

there is a distinct change in the chemistry 

noted by elevated levels of Cl-

, Na+, and 

SO4

2-

. These are all urban contaminants 

common in wastewater effluent (Figure 4; 

McQuillan, 2004). 

The final set of geochemical data was 

gathered Fall 2010. This project was 

designed to analyze nitrate, chloride, 

conductivity, and temperature though a 

rainfall event in order to characterize the 

“flush” of urban contaminants in the karst 

aquifer system. As suspected, nitrate 

varied with water level (Figure 5B). With 

the flush of water from the storm pulse, 

nitrate levels were elevated. Conductivity 

also showed an influence by the water level 

pulse, but in a different way. Prior to the 

nitrate surge, a conductivity pulse occurred 

as the dissolved solid laden aquifer water 

was forced out by the incoming runoff. 

Chloride showed no recognizable response 

to the storm event and the concentrations 

did not suggest anthropogenic chloride 

input. 

Flow Data

Numerous spring discharge 

measurements have been recorded in 

Eureka Springs, and most reflect low-flow 

discharges of several liters per minute. 

Several are known to go dry during the 

summer months, whereas others have 

perennial flow. After intense precipitation, 

many of these springs show obvious 

increases in discharge. From known 

records, no long term continuous 

monitoring of the springs in Eureka Springs 

has been undertaken. 

During the fall 2010, the water level of 

Magnetic Spring was monitored throughout 

a storm event. In this analysis two 

transducers were set up within the 

Magnetic Spring pool along with a control 

transducer, which was used to correct the 

spring stage data for changes in barometric 

pressure. Several precipitation events were 

recorded over the 10-day span the 

transducers were in place. The average 

spring response lag time was approximately 

Figure 5: 

A. Spring stage variation in response to precipitation 10/23/10.

B. Nitrate variation in response to spring stage 10/23/10.

C. Spring stage variation in response to precipitation for the period 

10/20/10-10/30/10

D. Spring stage variation and water temperature for the period 

10/20/10-10/30/10

A. 

B. 

C. 

D. 
52

1 hour (Figures 5Aand 5C). A diurnal 

fluctuation in temperature and flow was 

recognized as well (Figures 5C and 5D). 

The diurnal fluctuation in flow showed an 

increase in flow at night relative to 

daylight. The recharge area of Magnetic 

Spring is heavily forested and thus, evapotranspiration is suspected as the cause for 

this periodic signal. 

Future Work

Within the study area there are many 

pieces of the puzzle that are missing. 

Community groups are providing an 

excellent base of information, but gaps are 

present in the data. A continuous suite of 

flow, indicator bacteria, and water 

chemistry data from representative springs 

would lend immense amounts of data on 

the variations of the springs throughout the 

year, relative to contaminant sources. 

Isotopes from water and nitrate can also 

be helpful in determining sources. The 

two elements in water, hydrogen and 

oxygen, have varying isotopic ratios that 

can signify the evaporation exposure, and 

season of precipitation of the recharge 

water. Also, as nitrogen and oxygen are 

incorporated into nitrate, the original 

isotopic signatures of these elements are 

retained. These signatures distinguish 

whether the nitrate was originated from 

sewage, soil organic matter, agriculture, or 

the atmosphere (Clark and Fritz, 1997; 

Cascotti et al., 2002, Steffy and Kilham, 

2004). Thus, invaluable information on 

contaminant sources is available through 

isotopic studies of Eureka Springs. 

Conclusion

Although there have been several 

excellent studies centered on Eureka 

Springs, there are still many basic questions 

left unanswered. Answering these 

questions is imperative to the preservation 

and restoration of this valuable 

groundwater resource. When the city was 

founded, the springs were the crowning 

jewel, and now their impurity is an 

embarrassment. The citizens of Eureka 

Springs are dedicated to the preservation 

and restoration of the springs. With the 

appropriate resources, these springs can 

once again be the pride of Eureka Springs. 

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water of the Ozark Plateaus in Arkansas, 

Kansas, Missouri, and Oklahoma, U. S. 

Department of the Interior, U. S. Geological 

Survey, Water-Resources Investigations Report 

96-4313.

Benson, V. W., Baffaut, C., Robinson, D., 

Intarapapong, W., Farrand, T., Rogers, W., & 

Nam, L., 2008, Upper White River Watershed 

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A., 1993, Hydrogeologic Data for Carroll 

County, Arkansas, U. S. Geological Survey,

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2001, Identification of Fecal Escherichia coli 

from Humans and animals by ribotyping, 

Applied and Environmental Microbiology, Apr. 

2001, p. 1503–1507.

Casciotti, K., Sigman, D., Galanter Hastings, M., 

Bohlke, J., Hilkert, A., 2002, Measurement of 

the Oxygen Isotopic Composition of Nitrate in 

Seawater and Freshwater Using the Denitrifier 

Method, Analytical Chemistry, 74, 4905-4912. 

Clark, I., and Fritz, P., 1997, Environmental Isotopes 

in Hydrogeology, CRC Press, New York.

Davis, R., and Brahana, J., Johnston, J., 2000, 

Groundwater in northwest Arkansas: 

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Resources Center, Publication No. MSC-288.

Davis, R. K., Hamilton, S., and Brahana, J. V., 2005, 

Escherichia Coli survival in mantled karst 

springs and streams, northwest Arkansas 

Ozarks, USA, Journal of American Water 

Resources Association. vol. 41:6

Environmental Protection Agency, 2009, National 

Primary Drinking Water Standards, EPA 816-F09-004

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10/27/2010, Historical, 

http://www.eurekasprings.com/history/

Ford, D., and P. Williams, 2007, Karst Hydrogeology 

and Geomorphology, Wiley. 

Galloway, J. M., Haggard, B. E., Meyers, M. T., & 

Green, W. R., 2004, Occurrence of 

Pharmaceuticals and Other Organic Wastewater 

Constituents in Selected Streams in Northern 
53

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Scientific Investigations Report 2005-5140.

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Woodward, M., and Zachry, D., 1993, Geologic 

Map of Arkansas, Arkansas Geological 

Commission and U. S. Geological Survey, scale 

1:500,000

Husain, K., and Aley, T., 1981, An Exfiltration Study 

for Eureka Springs, AR, Comprehensive 

Planning Institute.

Imes, J.L., and Emmett, L.F., 1994, Geohydrology of 

the Ozark Plateaus Aquifer System in parts of 

Missouri, Arkansas, Oklahoma, and Kansas: 

U.S. Geological Survey Professional Paper 

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Jiang, Y., Wu, Y., and Yuan, D., 2009, Human 

Impacts on Karst Groundwater Contamination 

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System in Yunan, China, Environ. Sci. Technol. 

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Thurman, E. M., Saugg, S. D., Barber, L. B., 

and Buxton, H. T., 2002, Pharmaceuticals, 

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from on-site septic systems, Proceedings, 

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from septic systems as tracers, Environmental 

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using isotopic and chemical indicators, 

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groundwater on an urbanized dairy farm, 

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42, no. 13. 

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Quality in the Illinois River: Conflict and 

Cooperation between Oklahoma and Arkansas, 

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15N in 

stream biota in areas with septic tank systems in 

and urban watershed, Ecological Applications, 

14(3), pp. 637-641.

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SW China, Journal of Hydrology 388 (2010) 

121–130
54

The Relation between Dissolved Oxygen and other Physicochemical 

Properties in Barton Springs, Central Texas

By Barbara J. Mahler1 and Renan Bourgeais2

1

U.S. Geological Survey, 1505 Ferguson Lane, Austin, TX 78754 

2 Ecole Nationale du Génie de l’Eau et de l’Environnement de Strasbourg, 1 Quai Koch, B.P. 61039, 

67070 Strasbourg, France

Abstract

Dissolved oxygen (DO) is a critical component of karst water quality and is necessary for the health 

and reproduction of aquatic biota at karst springs. To better understand the relation between DO and 

other physicochemical properties, 6 years of daily DO data collected during January 2004–April 2010 

from Barton Springs in Austin, Tex., were analyzed. Barton Springs is the only known habitat of the 

endangered Barton Springs salamander, Eurycea sosorum, whose probability of survival decreases by

50% at a DO concentration of less than 3.4 mg/L. For the 6 years examined, DO ranged from 4.0 to 8.5 

mg/L. A two-segment multiple linear regression model was developed using DO, spring discharge (Q), 

and temperature (T), with a breakpoint of Q=70 ft3

/s separating the segments of the model. Q explained 

most of the variability in DO when Q was less than 70 ft3

/s (positive correlation with Q) (r2=0.88), and T 

explained most of the variability in DO when Q was greater than 70 ft3

/s (inverse correlation) (r2=0.67). 

When only the data collected after a change to an optical DO sensor (November 2006) were considered, 

an improved relation between DO and a combination of Q and T was obtained when Q was more than 70 

ft3

/s (r2=0.86). The relations between DO, Q, and T did not hold, however, in the hours to days following 

storm recharge. For example, DO was observed to increase by as much as 55 percent in 18 hours 

following one storm, and was closely correlated with turbidity and specific conductance, indicating that 

simple mixing with recharging surface water likely controls the DO concentration. These results indicate 

that species that depend on sufficient amounts of DO in Barton Springs for survival, such as the Barton 

Springs salamander, might be under increased stress if discharge decreases, temperature increases, or 

other processes that reduce DO concentrations intensify.
55

Geochemistry, Water Sources, and Pathways in the Zone of 

Contribution of a Public-Supply Well in San Antonio, Texas

By Lynne Fahlquist, MaryLynn Musgrove, Gregory P. Stanton, and Natalie A. Houston

U.S. Geological Survey, 1505 Ferguson Lane, Austin, TX 78754-4733

Abstract

In 2001, the National Water-Quality Assessment (NAWQA) Program of the U.S. Geological Survey 

initiated a series of studies on the transport of anthropogenic and natural contaminants (TANC) to publicsupply wells in representative aquifers of the United States with the goal of better understanding source, 

transport, and receptor factors (such as well field management practices) that affect contaminant 

movement to public-supply wells (PSWs). One of the selected TANC regional study areas is within the

San Antonio segment of the Edwards aquifer (hereinafter the regional aquifer), an important water 

resource in a rapidly urbanizing region in south-central Texas. Like many karst aquifers, the Edwards 

aquifer responds rapidly to changes in hydrologic conditions and is highly susceptible to contamination. 

A local-scale study area was delineated on the basis of a flow-modeled zone of contribution to a selected

PSW and its associated well field in the larger study area. The local-scale study area is representative of 

the regional aquifer, including unconfined (recharge) and confined zones of the aquifer. Results for a 

variety of geochemical constituents and tracers were compared for groundwater samples from regional 

wells, the well field (wellhead samples), nearby monitoring well clusters, an overburden well, and 

selected depths within the selected PSW collected under varying pumping conditions. Temporal samples 

that were collected from the selected PSW and selected monitoring wells in response to a recharge event 

also were compared.

Although hydrogeologic zones of preferential flow were determined for the selected PSW, 

groundwater samples from different hydrogeologic zones were not geochemically distinct. The 

geochemistry of groundwater from the selected PSW was similar to that of confined groundwater samples 

collected from other PSWs throughout the regional aquifer. These results indicate that water from the 

selected PSW and its associated well field is generally representative of the regional confined aquifer. All 

samples collected from the well field and specifically those collected from different depths in the selected 

PSW under a variety of pumping conditions were relatively homogeneous and well-mixed for numerous 

geochemical constituents, with the notable exception of tritium/helium-3 age tracers.

Geochemical and isotopic data are useful tracers of recharge, groundwater flow, fluid mixing, and 

water-rock interaction processes affecting water quality at the well field. Results for stable isotopes of 

hydrogen and oxygen for groundwater samples indicated that groundwater is meteoric in origin. Sources 

of dissolved constituents to Edwards aquifer groundwater and the selected PSW include dissolution of 

and geochemical interaction with overlying soils and calcite and dolomite minerals that compose the 

aquifer. Molar ratios of magnesium to calcium and strontium to calcium in carbonate groundwater 

typically increase along flow paths; and strontium isotope values in Edwards aquifer groundwater 

typically decrease along flow paths, approaching values similar to those of the Cretaceous limestone 

aquifer rocks. Molar ratios of magnesium to calcium and strontium to calcium, along with strontium 

isotope values, were used to assess relative extents of water-rock interaction associated with shorter and 

longer flow paths.

One of the other well field PSWs and four selected monitoring wells were sampled several times in 

response to a storm event that recharged the regional aquifer. The measured temporal variations in water 

chemistry were consistent with the observation that response to individual recharge events in the confined 

aquifer is likely attenuated by mixing processes along regional flow paths, unless flow paths intersect 

conduits.
56

Anthropogenic sources that might influence groundwater quality at the selected PSW include septic 

systems, leakage from municipal water and wastewater systems, and general industrial, commercial, or 

residential use of pesticides and volatile organic compounds. Constituents of concern for the long-term 

sustainability of the groundwater resource include anthropogenic organic compounds and the nutrient 

nitrate. While the urban San Antonio environment is a likely source of anthropogenic chemicals of 

concern to the selected PSW, the data collected indicate a few organic chemicals of concern are widely 

distributed in the regional Edwards aquifer at low concentrations, which is indicative of widely 

distributed sources. These are the pesticide atrazine, its degradate deethylatrazine, the drinking water 

disinfection byproduct trichloromethane (chloroform), and the solvent tetrachloroethene (PCE). Detected 

concentrations of all organic constituents regionally and at the well field of the selected PSW were low 

(<1 μg/L), consistent with previous NAWQA findings. The frequent detection of these anthropogenic 

chemicals of concern indicates that the entire aquifer is susceptible to contamination, and water quality is 

likely affected by anthropogenic activities.

Nitrate concentrations for samples collected during this study had a similar median value (1.95 mg/L) 

compared to regional confined groundwater samples collected in previous studies (1.67 mg/L) and were 

below drinking water standards. Nitrogen isotope results from PSWs in the well field and monitoring 

wells indicate that soil organic nitrogen is the dominant source of nitrate in the regional aquifer. A 

comparison with historic nitrogen isotope values, however, suggests that inputs of nitrate from biogenic 

sources might have increased during the past 30 years (1980–2010).

Apparent groundwater ages (determined using a piston flow model) for samples from the well field

and monitoring wells ranged from 0.8 to 41.3 years, with a median of 16.6 years. Although a piston-flow 

model might not be valid for karst aquifers, model ages provide a reference point for comparison. Results 

for depth-dependent samples from the selected PSW in San Antonio yielded a range of apparent ages (0.8 

to 21.7 years) but with no consistency with respect to depth or stratigraphic unit contributing to the PSW.

Selected geochemical constituents that might provide independent information about groundwater 

residence time or the extent of effect from anthropogenic activities were compared with piston flow 

model ages. Results of these comparisons were not consistent with the range of apparent groundwater 

ages. Mixing processes in this aquifer are complex; whereas age tracers might provide insight into the 

generally young nature of the groundwater, they might not readily allow for distinguishing the mixing 

history of water at a PSW or be well-suited to distinguish relatively small differences in groundwater age.

Nonetheless, groundwater apparent age interpretations indicating that water supplied to the selected PSW 

in San Antonio is young, and these interpretations were consistent with particle tracking results from a 

companion hydrogeologic modeling study. 
57

Temporal Stability of Cave Sediments

By Eric W. Peterson, and Kevin Hughes

Illinois State University, Department of Geography-Geology, Campus Box 4400, Normal, IL 61790

Abstract

Sediments within cave systems have been examined concerning source, mineralogy, and transport 

potential. While sediments in the thalweg are very mobile, the entrainment potential of the sediment piles 

has not been examined. Over the course of nine months, sediment piles in a Missouri cave were sampled 

to determine the stability of the piles and of the sediment properties. Sediment cores were analyzed for 

dry bulk density (ρd), porosity (n), volumetric wetness (θ), organic content (O.C.), hydraulic conductivity 

(K), and sediment particle size distribution. Observational evidence, deposited sediment and deformation 

of previous sample holes, suggests that despite elevated flows, the sediment piles were not mobilized over 

the course of the nine months. Physical properties remained constant at each location, but varied among 

the various locations. Dry bulk density values ranged between 1.2 g/cm3 to 1.5 g/cm3

. Porosity values 

were 0.42 to 0.57. Volumetric wetness showed similar variation ranging from 0.38 to 0.53. Organic 

content had the highest variation among the parameters ranging from 1.24 percent at one site to 4.84 

percent at another. Hydraulic conductivity ranged from 2.13×10-7 m/s to 3.10×10-7 m/s.

INTRODUCTION

Cave sediment work has focused on many 

different themes. White (1977) focused on 

finding the source of sediment. Herman and 

others (2007) examined the sediment mineralogy 

of the suspended sediment. Murray and others

(1993) determined the sediment age and rate of 

accumulation. Granger and others (2001) and 

Anthony and Granger (2004) employed the 

dating of cave sediments to develop histories of 

cave systems. Krekeler and others (1997)

examined the role of sediment in landscape 

evolution and what the sediments can reveal 

about surface-weathering conditions. Engel and 

others (1997) analyzed cave sediments to gain 

information about reversals of the geomagnetic 

field. Springer (2002) used cave sediments for 

paleoflood reconstruction. 

Examining the mobility of sediment within

the thalweg facies, Dogwiler and Wicks (2004)

reported that fluviokarst systems can transport 

up to 85 percent of the substrate during bankfull 

conditions. In the systems they investigated, 

flows capable of transporting the d50 and d85

sized particles occurred at intervals of 2.4 and 

11.7 months, respectively. Whereas Dogwiler 

and Wicks (2004) examined the sediment within 

the channel, they did not examine the sediment 

piles adjacent to the cave streams. 

Sediment piles serve as zones of low 

discharge and high storage, for water, solutes, 

and bacteria. Peterson and Wicks (2003)

analyzed the physical and hydraulic properties 

concluding that the sediment possessed 

hydraulic properties similar to the matrix and 

that the cave sediments were an extension of the 

bedrock for modeling purposes. Lines of 

evidence exist suggesting that the sediment piles 

are periodically entrained. White (1988) states 

that if cave streams were incapable of flushing 

sediment from the conduits, then the conduits 

would quickly clog. The large number of caves 

with traversable passages stands as a testament 

to the fact that flushing occurs. Herman and 

others (2008) reported that flow thresholds had 

to be exceeded to mobilize the sediment in the 

sediment piles. Thus, the sediment is mobile and 

can be reworked. 

Hence, the question of frequency of 

entrainment for sediment piles needs to be 

examined. This work examines the stability, 

both temporally and spatially, of cave sediments. 

While, the work done by Peterson and Wicks 

(2003) examined the sediment piles, their work 

was a one-time sampling and did not examine 

the potential temporal changes. This work 

examines the role of high-flow events and the 

possibility that the hydraulic properties could 

change after the sediments are disturbed.

FIELD AREA

The investigation centered on Berome 

Moore Cave of the Moore Cave System of the 
58

Perryville Karst Area located in Perry County, 

Missouri (fig. 1) The strata of the Perryville 

Karst Area are comprised of three stratigraphic 

units of Middle Ordovician age. The basal unit is 

the St. Peter Sandstone, a well-sorted quartz 

arenite, marking the lower limit of karstification

(Panno and others, 1999). Unconformably 

overlying the St. Peter Sandstone is the Joachim 

Dolomite; a yellow-brown, silty dolomite 

containing limestone interbeds and minor shale 

(Panno and others, 1999) with intense cavern 

development. The Joachim Dolomite is about 76 

meters thick in the study area (Panno and others, 

1999). Overlying the Joachim Dolomite, the 

Plattin Formation is even bedded, fine-grained to 

sublithographic limestone that has intervals of 

chert nodules and thin beds of shale (Dean, 

1977; Martin and Wells, 1966). This formation 

is about 106 meters thick and forms most of the 

bedrock of the eastern portion of the Perryville 

Karst Area (Panno and others, 1999). The Plattin 

Formation forms a geomorphic surface where 

there is intense sinkhole formation. Overlying 

the Plattin Formation, are thick loess deposits

(Panno and others, 1999). Geochemical analysis 

of the cave sediment reveals that the sediment 

composition is primarily silicates and are loess 

derived (unpublished data); entering the conduit 

system through the many sinkholes.

Figure 1: Location of the Moore Cave System, Perry 

County Missouri

Sinkholes of the Plattin Formation serve as 

discrete groundwater recharge points, allowing 

surface runoff and transported sediments to enter 

the cave system. After penetrating the Plattin 

Formation, the water and sediments enter the 

karst aquifer. Flow through the Moore Cave 

system is dominated by conduit flow that has 

been steady as indicated by small scallops along 

the main cave stream. 

METHODS AND PROCEDURES

Sediment collection

Sediment piles from four sites were chosen

within Berome Moore Cave for sampling. The 

Waterfall Passage site was located at the edge of 

a shallow plunge pool below a perennial 

waterfall. After entering the plunge pool, the 

water flows out to two small streams that 

eventually merge into a single stream. Collection 

was from a sediment pile between the two 

streams emerging from the plunge pool. 

Hydraulically, the site represented a waterfall. 

The Middle Main Stream Passage occurred 

within a conduit with a constant flow of water. 

In many places, the channel facies was almost 

non-existent; existing only as a laminated clay 

layer on top of bedrock. Collection occurred 

from a point bar along the main channel. 

Hydraulically, the site represented steady flows 

of moderate to high discharge. The Drum 

Passage has water pools and a small stream that 

was only centimeters across and millimeters 

deep. During high flow events, the main cave 

stream has the ability to overflow its banks and 

backflood the Drum Passage area. Sediment was 

collected from a point bar. Hydraulically, the 

site represented a point of stagnation with 

occasional high flows and backflooding. The No 

Name Passage was located in a small side 

conduit with a diameter of less than two meters. 

The roof of the conduit had fine sediment and 

plant debris deposited on the edges and along 

the apexes of dome structures. The deposits 

indicated that the conduit experienced pipe-full 

conditions and that the water source carried 

surface materials into the system. The collection 

site was a sediment pile on the edge of the 

conduit. Hydraulically, the site represented pipefull flow conditions.

During each sampling trip, three cores were 

collected from each location to determine dry 

bulk density (ρd), porosity (n), volumetric 

wetness (θ), and organic content (O.C.). A bulk 

density sampler was used to collect individual 

sediment cores of known volume. Each core was 

wrapped in aluminum foil, placed in a collection 
59

bag, and stored in a cooler with ice to preserve 

water content. Three additional sediment cores 

were collected in 1.4 cm internal diameter 

collection tubes to determine hydraulic 

conductivity (K). Finally, sediment for grainsize analyses was collected with a hand-trowel 

and placed into sample bags. The sample size 

was small, six cores per site, to accommodate 

the small area and immobility of the sediment 

piles. If more samples were collected, then the 

integrity of sediment piles would have been 

destroyed and the temporal objectives of this 

study would have been lost. 

Sediment Properties

Dry bulk density values were determined 

using the procedure presented by Blake (1986). 

Porosity (n) was determined from the formula:

p

d n

ρ

ρ = 1− (1)

where ρd is the dry bulk density and ρp is the 

sediment particle density, which for this work a 

value of 2.65 g/cm3 was used. A sediment 

particle density of 2.65 g/cm3 represents clays 

and quartz (Blake, 1986), which is the dominant 

mineralogical composition of the sediment as 

determined by SEM analysis. The volumetric 

wetness (θ) was determined for each core using 

the method presented by Gardner (1986). The 

organic content (O.C.) of the sediment was 

determined following the method presented by 

Schulte and Hopkins (1996).

Determination of Grain size distribution

For each sample, the grain-size distribution 

was determined using a combination of sieve 

and hydrometer analysis (ASTM D 422). The 

sieve and hydrometer data were used to generate 

a cumulative frequency curve. 

Determination of Hydraulic Conductivity

Hydraulic conductivity (K) was determined 

by conducting a falling head test on the intact 

sediment cores that were collected using a 

plastic tube with a known cross-sectional area. 

Hydraulic conductivity was found using the 

method presented by Lee and Fetter (1994).

RESULTS

Six sampling events occurred over a nine 

month period from February to October 2006 

(fig. 2). With the exception of the first sampling 

trip, all cores were successfully collected and 

analyzed. During the first sampling trip, all of 

the cores collected for the sediment property 

analyses from the Drum Passage were destroyed 

during transport out of the cave.

Figure 2: Distribution of daily rainfall (black bars) in Perry 

County (National Climatic Data Center, 2011). Gray 

dashed lines indicate sample collection dates.

Over the course of this study, the sediment 

piles appeared to be permanent fixtures of the 

cave. During each consecutive sampling trip, a 

thin layer of sediment, primarily clay, had been 

deposited in past sampling holes, indicating that 

deposition was active at each of these sites. High 

flow events occurred as a result of rain events 

(fig. 2). The presence of turbid pools of water 

located on shelves between three to five meters 

above the cave stream base level attests to the 

elevated flows. Even with the higher flows, the 

only noticeable alteration to the sediment piles 

was plastic deformation, as indicated by the 

elongation of past sampling points along the axis 

parallel to stream flow.

Each sediment pile was underlain by gravel 

sized particles. Overlying the gravel were layers 

of varying lithologies, but primarily fine 

grained, silt and clay sized, particles. One layer 

of note was an organic rich layer composed 

largely of forest litter from a conifer forest, 

predominant land-use for the area is agriculture. 

This layer was another indicator that the 

sediment piles had been immobile for quite 
60

some time. The top layer of all piles was a clay 

layer. This top clay layer would, with its strong 

electrostatic attraction, armor the sediment piles 

and make them hard to move. Sampling was 

conducted on the layers overlying the gravel. 

Although multiple layers were present, the 

collection methods for analysis of the physical 

properties and hydraulic conductivity did not 

allow for individual analysis of the layers. Thus, 

the results are representative of the material 

overlying the gravel as whole. 

Sediment Properties

Bulk Density Porosity, Volumetric Water Content, 

and Organic Carbon Content

Waterfall Passage (fig. 3a) had a ρd that 

ranged from 1.16 g/cm3 to 1.25 g/cm3 over the 

course of this study. Dry bulk density standard 

deviations ranged from 0.01 g/cm3 to 0.08 

g/cm3

. The Middle Main Stream Passage (fig. 

3a) has ρd values that ranged from 1.46 g/cm3 to 

1.53 g/cm3 with a standard deviation that ranged 

from 0.01 g/cm3 to 0.06 g/cm3

. The ρd for 

sediment at the Drum Passage ranged from 1.32 

g/cm3 to 1.38 g/cm3

, with a standard deviation 

that ranged from 0.02 g/cm3 to 0.22 g/cm3 (fig. 

3a). Dry bulk density for the No Name Passage 

ranged from 1.15 g/cm3 to 1.29 g/cm3 with a 

standard deviation ranging from 0.01 g/cm3 to 

0.23 g/cm3 (fig. 3a).

Porosity of the sediment at the Waterfall 

Passage (fig. 3b) ranged from 0.53 to 0.56 with a 

standard deviation range from 0.00 to 0.02. 

Middle Main Stream Passage (fig. 3b) had 

porosity ranging from 0.42 to 0.45 with a 

standard deviation ranging from 0.00 to 0.02. 

Drum Passage sediments had a porosity range of 

0.48 to 0.49 with a standard deviation ranging 

from 0.00 to 0.08 (fig. 3b). At No Name 

Passage, porosity ranged from 0.51 to 0.57 with 

a standard deviation ranging from 0.01 to 0.09 

(fig. 3b).

Volumetric wetness at the Waterfall Passage 

(fig. 3c) ranged from 0.50 to 0.53 with the 

standard deviations ranging from 0.01 to 0.03. 

Middle Main Stream Passage had a θ, ranging 

from 0.38 to 0.43, with standard deviations 

ranging from 0.00 to 0.02 (fig. 3c). For Drum 

Passage, θ ranged from 0.43 to 0.47 with a 

standard deviation ranging from 0.01 to 0.06 

(fig. 3c). The θ at No Name Passage ranged 

from 0.46 to 0.49 with a standard deviation 

ranging from 0.00 to 0.03 (fig. 3c). 

Figure 3: Sediment properties at the four sampling 

locations. a) Dry Bulk Density b) Porosity c) Volumetric 

wetness d) Organic Carbon Content

Compared to the other properties, the 

organic carbon content exhibited a greater range 

of values at each site (fig. 3d). At the Waterfall 

Passage, O.C. ranged from 3.12 to 3.53 percent 

with a standard deviation range of 0.07 to 0.52 

percent. O.C. for the Middle Main Passage 

ranged from 1.24 to 1.61 percent with a standard 

deviation range of 0.07 to 1.11percent. The 

Drum Passage experienced the greatest range in 

O.C. with values from 2.45 to 4.84 percent and a 

standard deviation range of 0.11 to 0.62 percent. 

No Name Passage exhibited some of the highest 

O.C. values ranging from 3.28 to 4.84 percent 

with standard deviation values ranging from 

0.05 to 0.73 percent. 

For each site and each sediment property an 

ANOVA was performed to test whether the 

values of the sediment properties changed from 

sampling event to sampling event (Table 1). A 

quick examination of figure 3 would suggest that 

the individual sediment property values are 

rather stable at each location. This is confirmed 

by the ANOVA analyses (Table 1). With the 

exception of the θ at the Water Fall Passage and 

Middle Main Passage, and the O.C. at the Drum 

Passage and No-Name Passage, the values were 
61

statistically similar for all sampling events. 

Note that the ρd and n values did not statistically 

vary over the duration of the study for each 

location. 

The values for the individual sediment 

properties vary among the sample locations 

(figs. 3 and 4). Waterfall Passage had the lowest 

ρd value with the Main Middle Passage having 

the highest ρd value. Given the relationship 

between ρd and porosity, Waterfall Passage had 

the highest porosity value with the Main Middle 

Passage having the lowest porosity value. The θ

values had a similar relationship as porosity. 

The O.C. values exhibited some of the larger 

differences among the sites. Drum Passage and 

No Name Passage had higher O.C. than 

Waterfall Passage and Middle Main Passage. 

Middle Main Passage had considerably less O.C.

than the other three locations. To test whether 

the variation in the individual properties were 

statistically different among the locations, an 

ANVOA was performed for each property. 

Among the locations, there is significant 

variation for each of the sediment properties 

(Table 2).

Grain Analysis

Sediment particle size distributions reveal 

similar distributions among the four sites and

show relatively stable particle size distributions 

at the sites over time (fig. 5). A noticeable 

exception is that the Drum passage has a smaller 

percentage of sand size particles, 12 percent as 

compared to the other sites that have between 17 

percent at Middle Main to 21percent at No 

Name. At all locations clay sized particles 

comprised 2.5 percent or less of the overall 

material. No Name had the highest amount at 

2.5 percent, followed by Middle Main (2 

percent), Waterfall (1.8 percent), and Drum (1.6 

percent). 

Sieving of sediment samples revealed that 

more than geosediments composed the sediment 

piles. Plant debris, pine needles, and the shells of 

small mollusks were observed with the 

sediment. These materials account for the higher 

O.C. values measured for the sediments. 

Figure 4: Comparison of the sediment properties among 

the four passages. The ends of the boxes represent the 25th 

and 75th percentiles with the solid line at the median; the 

error bars depict the 10th and 90th percentiles and the points 

represent the 5th and 95th percentiles.

Figure 5: Grain size distribution curves for the sediment 

particles at the four sampling locations. The points 

represent the mean value over the course of the sampling 

and the error bars represent the standard deviation over the 

course of the sampling.

Hydraulic Conductivity (K)

The cores collected at each location during 

each sampling event provided statistically 

similar K values for each location (Table 1 and 

fig. 6). The Waterfall Passage had the highest K 

values, which ranged from 3.04×10-7 m/s to 3.28 

×10-7 m/s, with a standard deviation range of 

1.65×10-9 m/s to 1.63 ×10-8 m/s. Middle Main 

Passage had the lowest K values that ranged 

from 2.01×10-7 m/s to 2.17 ×10-7 m/s, with a 

standard deviation range of 2.45×10-10 m/s to 

1.02 ×10-8 m/s. The Drum Passage had K values 

that ranged from 2.31×10-7 m/s to 2.45 ×10-7
62

m/s, with a standard deviation range of 4.12×10-

10 m/s to 1.49 ×10-8 m/s. K values at No Name 

Passage ranged from 3.03×10-7 m/s to 3.16 ×10-7

m/s, with a standard deviation range of 9.12×10-

10 m/s to 2.10 ×10-8 m/s.

Figure 6: Hydraulic conductivity values for the sediments. 

a) comparison of the individual sampling events at the four 

sampling locations. b) Comparison of the collective K 

values among the four passages. The ends of the boxes 

represent the 25th and 75th percentiles with the solid line at 

the median; the error bars depict the 10th and 90th percentiles 

and the points represent the 5th and 95th percentiles.

The K values from the different sampling 

events were pooled to provide an overall 

comparison of the K values among the sites (fig. 

6b). There exists a statistically significant 

variation among the K values among the sites 

(Table 2). Overall, Waterfall Passage and No 

Name Passage had the highest K values at 

3.13×10-7 ±1.26×10-8 m/s and 3.10×10-7

±1.27×10-8 m/s. Middle Main had the lowest K 

value at 2.13×10-7 ±7.61×10-8 m/s and Drum 

Passage had a slightly higher K value at 

2.38×10-7

±9.28×10-9 m/s.

DISCUSSION

During the nine-month period of sampling, the 

sediment properties remained stable, resulting in 

consistent values among the various sediment 

properties (Table 1 and fig. 3). Exceptions to 

the stable properties include θ for the sediments 

at Waterfall Passage and Middle Main Passage 

and O.C. at Drum Passage and No Name 

Passage. A noticeable difference in the θ occurs 

in the May sampling event, which followed a 

precipitation event. Both Waterfall and Middle 

Main sediments had higher θ values, which may 

be a direct result of the wetter conditions. For 

the variations in O.C. values two possible 

reasons exist. First, there was a large vertical 

heterogeneity as dictated by the different beds 

that make up the sediment profile. If each 

sample did not capture the exact same beds, then 

a difference would result even if the individual 

beds were homogenous. Second, there was a 

large horizontal heterogeneity, meaning that as 

the sample points moved along the sediment 

pile, the heterogeneity caused deviation in the 

data points and could be interpreted as a 

temporal shift. Small variation in the ρd, n, 

particle size distributions, and K values existed, 

but since the sediment piles were immobile, 

there is reduced likelihood that these properties 

were changing. 

A previous study reports that organic carbon 

content of karst sediments are less than 

1.5percent (Peterson and Wicks, 2003). The only 

site that was in-line with the reported data is the 

Middle Main Stream Passage. The common trait 

that the Middle Main Stream Passage has 

compared to the previous study is that there is a 

large volume of water constantly moving by the 

sediment pile. The less than 1.5 percent organic 

carbon content may be due to the washing effect 

of this moving water. Organic debris caught in 

this swift water may not have the opportunity to 

deposit with the sediment. This is further 

supported by the lack of the forest-litter rich 

layer that was found at the other sites. 

Compared to the Middle Main Stream Passage 

and previous studies mentioned, the Waterfall 
63

Passage, Drum Passage, and No Name Passage 

have more than twice the organic content (fig. 

3). The Waterfall Passage had a very strong 

presence of the forest litter rich layer. Some core 

samples had a simple black layer while other 

cores were rich in recognizable pine needle 

remains. It is likely that the waterfall feeds in 

from a sinkhole. Forest litter at the bottom of the 

sinkhole was washed in and deposited around 

the plunge pool. The Drum Passage showed 

greater variation in organic carbon content, 

which may be related to the back flooding and 

presence of water pools. Near the pools, more 

organic carbon has been deposited. Away from 

the pools, less organic material is deposited. The 

most revealing part of the organic carbon 

content was revealed in the No Name Passage. 

Here, a lot of insect activity was observed in the 

form of beetles and crickets. Earth worms crawl 

through the sediment piles and fresh plant 

debris, including tree seeds that attempt to 

germinate in the dark cave environment, is 

carried in. The rich organic nature of the 

sediment, almost 5percent, is likely due to the 

fresh inputs.

The clay armor seems to make the sediment 

piles a somewhat permanent feature of the cave 

streams and only very high flow events are 

capable of altering them. The permanence of the 

sediment piles was demonstrated in August 

when high flow conditions only caused plastic 

deformation of previous sampling points. 

Although the sediment piles are composed of 

grains that should become mobile according to 

the study done by Dogwiler and Wicks (2004), 

the finer sized particles make the sediment piles 

behave like larger particles.

The controlling factor on the hydraulic 

conductivity of the sediment piles is likely the 

hydraulic conductivity of the silt. Silt tends to 

have a hydraulic conductivity that ranges from 

1x10-9 m/s to 1x10-5 m/s. The K values 

measured for the sediments, on the order of 10-7

m/s, at each location were within this range and 

were similar to those measured by Peterson and 

Wicks (2003). The data presented here further 

support the claim by Peterson and Wicks (2003) 

that the cave sediments can be modeled as an 

extension of the matrix rock because the 

hydraulic conductivity of the sediment piles are 

similar to the hydraulic conductivity of 

limestone. A limitation to this claim would be if 

the sediment piles are highly mobile. 

Compared to other published work, the 

sediment piles appear to be less mobile than the 

sediment within the karst streams. Dogwiler and 

Wicks (2004) observed movement of the d50

particle on a 2.4 month interval; over the nine 

months, there was no movement of the sediment 

pile in Berome Moore Cave. While comparison 

is made between the studies, caution must be 

exercised. Dogwiler and Wicks (2004) were 

examining larger bedload particles, the d50

particle was 70 mm, located in the thalweg of 

the stream. The sediment piles are finer grained, 

d50 particle ~0.04 mm, but require similar stream 

velocities to become entrained as the larger 

particles. Additionally, the sediment piles are 

located along the banks of the stream where 

velocities are lower than in the thalweg. Thus, 

erosion is limited as a result of the need for 

higher velocities to entrain the sediment and the 

lower velocities that occur along the edges of the 

stream.

Finer sized particles do move through 

systems. Herman and others (2008) indentified 

flow thresholds that needed to be exceeded 

before deposited sediment in the conduits would 

be entrained and moved through the system. For 

the systems examined by Herman and others 

(2008), the entrainment thresholds were reached 

as a consequence of extremely high flows, 

greater than two orders of magnitude above 

baseflow, associated with precipitation from 

remnant hurricanes. During the nine months, the 

Berome Moore system did not experience 

extreme flows remotely close to the levels 

observed by Herman and others (2008). This 

work was not able to address the point when the 

flow threshold for entrainment was exceeded.

CONCLUSIONS

Over the course of the nine-month sampling 

period, the sediment piles within Berome Moore 

Cave did not move. The hydrologic properties, 

n and K, did not change over the study, 

confirming the stability of the piles. While the 

sediment was not entrained and the properties 

remained stable at the individual sampling 

locations, the locations behaved differently. 

Particle size distributions were similar among 
64

the sites, but ρd, n, θ, O.C., K., statistically 

varied among the sites. This variability appears 

to be a result of the difference in flow regime 

among the sampling locations.

Table 1. ANOVA results examining temporal variation of the sediment properties at the four locations.

Parameter Waterfall Passage Main Middle 

Passage

Drum Passage No-Name Passage

Dry Bulk Density F(5, 12) = 2.12, p = .13 F(5, 12) = 2.03, p = .15 F(4, 10) = 0.14, p = .96 F(5, 12) = 0.10, p = .99

Porosity F(5, 12) = 2.44, p = .09 F(5, 12) = 2.02, p = .15 F(4, 10) = 0.15, p = .97 F(5, 12) =0.11, p = .99

Volumetric wetness F(5, 12) = 3.61, p = .04 F(5, 12) = 5.20, p = .01 F(4, 10) = 1.31, p = .33 F(5, 12) = 0.71, p = .99

Organic Carbon 

Content F(5, 12) = 0.55, p = .73 F(5, 12) = 0.25, p = .93 F(5, 12) = 22.32, p< .01 F(5, 12) = 6.74, p < .01

Hydraulic

Conductivity1 F(5, 12) = 3.09, p = .06 F(5, 12) = 2.03, p = .17 F(5, 12) = 0.95, p = .49 F(5, 12) = 0.45, p = .80

1

Data were subjected to a logarithmic transformation

Table 2. ANOVA results examining the spatial variation of the sediment properties among the four locations.

Parameter

Dry Bulk Density F(3, 65) = 53.04, p <0.01

Porosity F(3, 65) = 52.16, p <0.01

Volumetric wetness F(3, 65) = 58.61, p <0.01

Organic Carbon Content F(3, 68) = 68.66, p <0.01

Hydraulic Conductivity1 F(3, 68) = 449.69, p <0.01

1

Data were subjected to a logarithmic transformation

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66

Evaluating the Stormwater Filters at Mammoth Cave, Kentucky

By Ashley West1, Rickard Toomey2, Michael W. Bradley3 and Thomas D. Byl1,3

1

College of Engineering, Tennessee State University, 3500 John A Merritt Blvd, Nashville, TN 37209

2 Geography/Geology Dept., Western Kentucky University, Bowling Green, KY 42101-1066

3 U.S. Geological Survey, 640 Grassmere Park, Suite 100, Nashville, TN 37211

Abstract

Mammoth Cave is home to many unique animals that could be harmed by contaminants carried into 

the cave system during storm events. This project was conducted to determine if leaf-pack filter-systems 

attenuated storm runoff at seven parking lots in Mammoth Cave National Park. Grab samples were 

collected at the inlet and outlet of the filter systems, and analyzed for oil and grease, sediments, turbidity, 

gasoline compounds, nitrate, ammonia, fecal bacteria, dissolved iron, and chemical oxygen demand 

(COD). For the first sampling round, the filters had not been serviced for 8 years and did very little to 

remove contaminants. The contaminant concentrations at the outlet were similar to those at the inlet, with 

the exception of removing 20 to 70 percent of the oil and grease. After replacing leaf packs and cleaning 

out debris, the re-conditioned filters did not remove oils and greases and did little to remove copper and 

ammonia from runoff waters. However, the re-conditioned filters removed up-to 99 percent of the 

benzene, toluene, ethyl-benzene and xylene, and, up to 90 percent of the turbidity, E. coli, COD, and iron 

from the storm runoff. These results indicate that well-maintained filtration systems are more effective 

than clogged filters at removing many, but not all, contaminants from parking lot runoff.
67

Using Labeled Isotopes to Trace Groundwater Flow Paths in a 

Northwestern Arkansas Cave

By Katherine J. Knierim1, Erik D. Pollock2, and Phillip D. Hays3

1

Environmental Dynamics Program, 113 Ozark Hall, University of Arkansas, Fayetteville, AR 72701 2

University of Arkansas Stable Isotope Laboratory, 116 Ferritor Hall, University of Arkansas, Fayetteville, 

AR 72701

3 U.S. Geological Survey, Arkansas Water Science Center, 700 W. Research Center Blvd MS36,

Fayetteville, AR 72701

Abstract

Tagged water was used to investigate the conditions of carbon and nitrogen cycling over small spatial 

and temporal scales while mimicking natural conditions (that is, small concentrations and slow addition 

of tracers) in a small (< 200 meter) and shallow (< 9 meter) cave system in northwestern Arkansas. 

Approximately 950 liters of isotopically tagged water with labeled glucose (13C6H12O6) and potassium 

nitrate (K15NO3) was released to a trench above the cave at a rate of 2 liters per minute to simulate 

infiltration by precipitation. The signature of the labeled glucose was observed in the carbon dioxide and 

dissolved inorganic carbon isotopic compositions of soil water samples closest to the trench. The labeled 

tracers were not detected in the cave stream after 2 months, as evidenced by the isotopic compositions of

carbon dioxide, dissolved inorganic carbon, dissolved organic carbon, and nitrate in the cave stream. 

Although the labeled tracers were not detected in the cave stream, stored soil water did infiltrate into the 

cave, as evidenced by changes in the pH and conductance of the cave stream and the concentration of 

inorganic and organic carbon in the cave stream. Most of the heavy isotopic signature was likely respired 

as carbon dioxide and passed out of the soil zone. The nitrate was also processed in the soil zone, but 

gaseous nitrogen was not monitored during the tracing experiment. 

INTRODUCTION 

Northwestern Arkansas is characterized by 

karst, leading to water-quality degradation in 

cave and spring systems because of the large 

concentration of agricultural operations and 

population growth in the region (Davis and 

others, 2000; Gillip, 2007; Peterson and others, 

2002). Nationally, Arkansas is ranked second in 

poultry production and the three top counties for 

agricultural sales in Arkansas are located in the 

northwestern region of the state (U.S. 

Department of Agriculture, 2010). Nitrate 

(NO3

-

) is often the dominant and most mobile 

form of nitrogen (N) in agricultural settings 

(Böhlke, 2002) and can contribute to 

eutrophication of surface water and degradation 

of karst ecosystems. Concentrations and stable 

isotopic compositions of NO3

- (δ15N and δ18O) 

and dissolved inorganic/organic carbon (δ13C) 

have been used to characterize denitrification –

an important process for removal of NO3

-

– in 

natural systems because of the interaction 

between NO3

- and carbon (C) species (Aravena 

and Robertson, 1998; Einsiedl and others, 2005; 

Nascimento and others, 1997). Stable isotopes 

provide an economical and powerful tool for 

tracing biogeochemical cycling of nutrients in 

groundwater systems because of the observable 

and predictable changes in isotopic compositions 

as compounds are cycled between reservoirs 

(Kendall, 1998).

This study investigated C and N cycling for 

a short temporal scale experiment (2 months) 

while mimicking natural conditions (that is, 

small concentrations and slow addition of 

tracers) in a small (< 200 meter (m) long) and 

shallow (< 9 m deep) cave system in 

northwestern Arkansas. We used isotopically 

tagged water with labeled glucose (13C6H12O6) 

and potassium nitrate (K15NO3) as tracers to 

monitor C and N cycling from the soil zone into 

the cave to better understand how nutrients and

organic substrate may be utilized in shallow 

karst settings.
68

STUDY SITE

Jack’s Cave is located in Madison County, 

Arkansas in the Ozark Plateaus (Ozarks). The 

cave is in the Sneeds Dolomite Member of the 

Middle Ordovician Everton Formation and is 

capped by the Kings River Sandstone Member 

of the Everton Formation. Meteoric water 

infiltrates into a very gravelly silt loam (National 

Cooperative Soil Survey, 2008), which is 

developed from colluvium deposited on the 

Kings River Sandstone Member. Gravel is 

predominantly remnant chert of the Early 

Mississippian Boone Formation. After 

infiltrating through the silt loam, water enters 

the cave through fractures in the sandstone cap 

rock, collects in a main room of Jack’s Cave 

(location C1, Fig. 1), and flows southwest along 

the cave passage towards C2 and C3. The cave 

ends in a sediment-filled constriction; the exit 

point(s) for the cave stream were not identified. 

The cave extends an additional 55 m beyond C3.

Figure 1. Site map showing the locations of soil sampling sites on the surface and cave 

sampling sites along the trace of Jack’s Cave. Water and gas samples were collected 

from the soil sampling sites (L4 through L8). Water and gas samples were collected 

from the cave at C1, and only water samples were collected at C2 and C3. Topographic 

contours are elevation in feet above mean sea level. 
69

METHODS 

The position of the main room in Jack’s 

Cave and a fracture trace trending along the cave 

were located on the surface by surveying from 

the cave map to define the boundaries for the 

experiment. A trench (4.3 m long, 0.3 m wide, 

0.2 m deep) was dug upslope from the fracture 

trace and five suction lysimeters were installed 

for collecting soil-water samples downslope

from the trench at varying depths over a 6-m by 

9-m area (Fig. 1). The shallow trench was 

designed to temporarily pond the tagged water 

and ensure infiltration into the soil. Water 

introduced to the trench mimicked groundwater 

flow paths where meteoric water infiltrates into 

the soil and travels along fractures into the cave. 

Groundwater flow paths between the soil surface 

and the cave were tested by releasing 

approximately 900 liters (L) of untagged, 

surface water into the trench. Prior to release of 

water at the surface, no water was infiltrating 

along fractures into Jack’s Cave because of dry 

conditions at the site. The water arrived in the 

main room of the cave approximately 23 

minutes following release at the surface, as 

evidenced by water flowing through fractures in 

the sandstone cap rock. Water flowing along 

fractures from the land surface is considered the 

source of water flowing in the cave stream; 

therefore, cave water samples were collected 

from the cave stream where water dripping from 

fractures immediately pooled in the cave (C1).

Water and gas samples for C and N isotopic 

analyses were collected between September and 

November 2009. Background sampling at all 

sampling sites was conducted prior to the release 

of labeled water. Water samples from the cave 

stream (C1, C2, and C3) and soil lysimeters (L4 

through L8) were collected in 40-milliliter (mL) 

precombusted vials and filtered through 0.45-

micrometer (μm) filters. The concentration and 

isotopic composition (δ13C) of dissolved 

inorganic carbon (DIC) and dissolved organic 

carbon (DOC) were analyzed at the Colorado 

Plateau Stable Isotope Laboratory using 

continuous flow isotope ratio mass spectrometry 

(IRMS) (St-Jean, 2003). Nitrate concentration 

([NO3

-

]) and isotopic composition (δ18O and 

δ15N) were analyzed at the University of 

Arkansas Stable Isotope Laboratory using the

bacterial denitrifier method for analysis on a Gas 

Bench II interfaced to an IRMS (Casciotti and 

others, 2002; Sigman and others, 2001). Gas 

samples from the main room of Jack’s Cave 

(C1) and soil atmosphere pumped from the soil 

lysimeters (L4 through L8) were collected in 

precombusted 100-mL serum vials purged with 

helium. Gas samples were analyzed for [CO2] 

and isotopic composition of CO2 (δ13C-CO2) at 

the University of Arkansas Stable Isotope 

Laboratory using gas chromatographycombustion IRMS. All isotope ratios are 

reported using delta notation (δ) in parts per 

thousand, or per mil (‰). The δ notation 

represents the ratio of the heavier to lighter 

stable isotope (13C/12C, 15N/14N, or 18O/16O) 

relative to a standard (Vienna Peedee Belemnite, 

air, or Standard Mean Ocean Water, 

respectively) (Clark and Fritz, 1997). 

The labeled tracers used in this experiment 

were applied at small concentrations to avoid 

stimulating the biologic system. Approximately 

950 L of water tagged with 0.2713 grams (g) of

13C6H12O6 (99 atom %) and 1.134 milligrams 

(mg) of K15NO3 (99 atom %) was released to the 

trench from a tank beginning on 9/12/09 at 

15:00 at a rate of approximately 2 L/min to 

simulate infiltration by precipitation

(approximately 2 mm/hr). Water temporarily 

ponded in the trench and on the soil surface, and 

then infiltrated into the soil. Periodic sampling

was conducted over the next 2 days at all sites

following the release. Approximately 2 months 

following the addition of the tracer, a final 

sampling event was conducted at all sites. 

Water samples were kept at 4°C until analysis; 

water temperature, conductivity, and pH were 

measured on site.

RESULTS AND DISCUSSION 

Background samples could not be collected 

at every sampling site because of dry conditions 

at the study location prior to addition of the 

tagged water. In the soil, there was insufficient 

water for DIC analysis. The isotopic 

composition of DOC (δ13C-DOC) averaged 

-25.0‰ and the isotopic composition of NO3

-

(δ15N-NO3) averaged 6.3‰ at L5 and L6. Soil 
70

δ13C-CO2 values averaged -19.2‰ for all sites. 

Background values in the cave are only provided 

for site C1; δ13C-DIC was -14.2‰, δ13C-DOC

was -23.3‰, and δ15N-NO3 was 5.4‰. During 

the summer and fall, δ13C-CO2 values in the 

atmosphere of Jack’s Cave averaged -20.4‰ 

(Knierim, 2009). 

Water samples collected from the trench 

immediately following release of the tagged

water from the tank were enriched in 15N of 

NO3

- and 13C of DIC/DOC because of addition 

of the labeled glucose and potassium nitrate. 

Trench water had a δ15N-NO3 value of 135.8‰ 

and δ13C values of 208.7‰ for DIC and 

1,007.8‰ for DOC. 

In the soil, δ13C-DIC values ranged between 

-0.4‰ and -24.3‰ (n = 10) and varied by 

lysimeter; the heaviest values occurred closest to 

the trench. At sites L4 and L5 (Fig. 1), the 

heaviest δ13C-DIC values occurred on 9/13/09 at

13:30 (-3.9‰) and on 9/15/09 at 17:00 

(-0.44‰), respectively (Fig. 2). The δ13C-DOC 

values showed a smaller range between 

lysimeters (-23.7‰ to -27.4‰, n = 16), but 

isotopic composition varied by location. In the 

cave, δ13C-DIC values ranged between -11.1‰ 

and -15.0‰ (n = 7) and δ13C-DOC between -

21.9‰ and -26.1‰ (n=7). Relative minimum 

and maximum concentrations of DIC and DOC, 

respectively, were observed at C1on 9/13/09 at

12:30 (Fig. 2, DIC data only).

Nitrate isotopic composition showed a 

greater range in δ15N and δ18O values in the soil 

(n=18) than in the cave (n=18). The δ15N-NO3

values were between -7.7‰ and 12.8‰ in the 

soil and 0.7‰ and 6.4‰ in the cave (Fig. 3). 

Nitrate concentrations as N ([NO3

-

]) ranged from

0.80 to 2.36 milligrams per liter (mg/L) in the 

soil and 0.35 to 3.15 mg/L in the cave. Similar 

to DIC concentration, a relative minimum in 

[NO3

-

] (0.83 mg/L) was observed on 9/13/09 at

12:30 at C1.

Figure 2. Concentration and isotopic 

composition of dissolved inorganic carbon 

(DIC) in the soil (L4 and L5) and in the cave 

stream in the main room of Jack’s Cave (C1)

over time. Note how the δ13C-DIC values in the 

soil are heavier than in the cave due to the heavy 

signature from the labeled glucose.
71

Figure 3. Isotopic composition of nitrate (NO3

-

)

in the soil (L4 through L7) and the cave (C1 

through C3). 

Concentration of CO2 ([CO2]) in L4, L5, and 

L6 was greater than 20,000 parts per million 

(ppm) (instrument range). In the cave, [CO2] 

averaged 2,950 ppm in September (n=7) and 

decreased to 750 ppm in November. During the 

warmer months of the summer and fall, [CO2] 

was found to increase in Jack’s Cave because 

respired CO2 from the soil zone moved into the 

cave through gaseous and aqueous phases 

(Knierim, 2009). The δ13C-CO2 values ranged 

between 1.7‰ and -22.4‰ in the soil (n=13) 

and -20.0‰ and -22.7‰ in the cave (n=5). In 

L4 and L5, δ13C-CO2 values varied over time 

and in L4 the heaviest value (1.8‰) occurred on 

9/13/09.

The isotopic signature of the labeled tracer 

was observed in the CO2 and DIC isotopic 

compositions in L4 and L5 because the signature 

of the labeled glucose was converted to 

inorganic C species through respiration. L4 was 

closest to the trench (approximately 1 m) and the 

heaviest δ13C-CO2 values were observed 18.5 

hours after releasing the tagged water into the 

trench. The isotopic compositions of CO2 and 

DIC began to approach background values 

during the experiment, but had not returned to 

the typical range of values 5 days after adding 

the tracer. From a previous study conducted at 

Jack’s Cave, soil δ13C-CO2 averaged -21.3‰ 

over a one-year time period (Knierim, 2009). 

Most of the heavy isotopic signature was likely 

respired as CO2 and fluxed out of the soil zone

because heavy δ13C-DOC values were not 

observed during the study.

The tracer was not detected in the two 

lysimeters further from the trench (L6 and L7). 

During the experiment the volumes of water 

collected from L6 and L7 were much smaller 

than L4 and L5, likely because the tagged water 

did not reach the more distant lysimeters located 

between 3 to 4 m from the trench, and any water 

collected from L6 and L7 had a meteoric source. 

The lysimeter installed farthest from the trench 

(L8) did not yield water during the experiment; 

L8 was purposefully installed on the downslope 

side of the infiltration fracture. Any infiltrating 

water (either meteoric or tagged) could have 

flowed into the fracture before reaching L8, or 

the lysimeter installed at L8 may not have 

developed a proper seal with the surrounding 

soil.

The isotopic compositions of NO3

- were not 

outside the range of typical compositions for 

NO3

- in soil or precipitation (Clark and Fritz, 

1997); therefore, the 15N signal from the labeled 

water could not be confirmed. Although the 

heavy N signature was not observed during the 

experiment, processing of NO3

- in the soil zone 

did occur, as evidenced by the variation in the 

concentration and isotopic composition of NO3

-

in L4. Mixing between two NO3

- sources (small

concentration and heavier composition with 

large concentration and lighter composition) also 

accounted for some of the variation, but the 

effects of mixing compared to processing could 

not be separated. 

No labeled tracers were detected in the cave 

stream or atmosphere, as evidenced by the 

isotopic compositions of CO2, DIC, DOC, and 

NO3

- in the cave, which were all within the 

range of values observed previously in Jack’s 

Cave (Knierim, 2009) or the literature for C and 

N species (Kendall,1998; Spötl and others, 

2005). The isotopic compositions of CO2 were 

typical of values for Jack’s Cave during the 

warmer periods of summer and fall (Knierim, 

2009). Although the δ13C-DIC values were 

heavier than δ13C-CO2 values, the composition 

of DIC was likely controlled by interaction with 

carbonate bedrock which would provide a 

heavier, inorganic source of C. Additionally, the 
72

heaviest δ13C-DIC value that occurred on 

11/8/09 (-11.1‰) coincided with the occurrence 

of the smallest CO2 concentration (750 ppm); 

the heavier composition is because of surface 

atmosphere being drawn into the cave and 

interaction between gaseous and aqueous C 

species. Atmospheric [CO2] is approximately 

380 ppm and has a δ13C-CO2 value of -7.7‰ 

(Faure and Mensing, 2005), which would 

provide a source of isotopically heavier and 

smaller [CO2]. 

Although the tagged water was not detected 

in the cave environment, stored soil water did 

infiltrate into the cave. On 9/13/09 at 12:30, the 

concentration of NO3

- and DIC in the cave (C1) 

reached relative minima, DOC concentration 

peaked (2.8 mg/L), and the heaviest δ13C-DOC 

value (-21.9‰) was observed. The changes in 

the concentration and isotopic compositions of C 

and N species in the cave stream were not 

outside of the range of background values and, 

therefore, not caused by tagged water entering 

the cave. Following release of the tagged water, 

precipitation fell on the morning of 9/13/09, 

although the rainfall was not recorded at nearby 

weather stations. The decrease in pH and 

specific conductance observed in the cave on

9/13/09 at 12:30 provided further evidence that 

stored soil water was moved into the cave 

through fractures in the sandstone cap rock 

following infiltration of labeled water and 

precipitation. The stored soil water was 

relatively more acidic, lower in NO3

- and DIC, 

and higher in recalcitrant DOC than the cave 

water. Although the labeled tracers were not 

detected in the cave, the pulse of stored soil 

water provided evidence that allochthonous C 

infiltrates into the cave environment during 

precipitation events. Many caves are 

oligotrophic, and the external source of C could 

provide an important organic substrate for 

organisms in the cave environment. The 

decrease in NO3

- concentration because of the 

pulse of stored soil water was delayed at the 

downstream sites (C2 and C3) compared to the 

upstream site (C1). The small water flow in the 

cave and abundance of clastic sediments may 

provide micro-environments suitable for 

denitrifcation, although NO3

- processing cannot 

be confirmed in the cave because the labeled 

tracers did not reach the cave stream. However, 

DOC did enter the cave from the soil zone and 

could serve as an electron donor during 

denitrification.

SUMMARY 

Glucose (13C6H12O6) was likely utilized in 

the soil and the isotopic signature was rapidly 

converted to inorganic C species through 

respiration. Most of the heavy isotopic signature 

was likely respired as CO2 and fluxed out of the 

soil zone. The K15NO3 was also utilized in the 

soil zone, but N2 was not monitored during the 

tracing experiment and the heavy isotopic 

signature was not observed in soil or cave NO3

-

. 

Monitoring soil respiration gases (N2, CO2, and 

likely N2O) is important for understanding 

biological processing during isotopic tracing 

experiments. 

Although the labeled tracer was not detected 

in the cave environment, stored soil water did 

infiltrate into the cave through fractures in the 

sandstone caprock. The stored soil water was 

relatively more acidic, lower in NO3

- and DIC, 

and higher in recalcitrant DOC than the cave 

water. The pulse of stored soil water provided 

evidence that allochthonous C infiltrates into the 

cave environment during precipitation events. 

ACKNOWLEDGMENTS 

Support for this project was provided by the 

Cave Research Foundation.

REFERENCES

Aravena, R., and Robertson, W.D., 1998, Use of 

multiple isotope tracers to evaluate 

denitrification in ground water: Study of nitrate 

from a large-flux septic system plume: Ground

Water, v. 36, p. 975-982.

Böhlke, J.K., 2002, Groundwater recharge and 

agricultural contamination: Hydrogeology 

Journal, v. 10, p. 153-179.

Casciotti, K. L., Sigman, D. M., Hastings, M. G., 

Bohlke, J. K., Hilker, A., 2002, Measurement of 

the oxygen isotopic composition of nitrate in 

seawater and freshwater using the denitrifier 

method: Analytical Chemistry, v. 74, p. 4905-

4912.

Clark, I., and Fritz, P., 1997, Environmental Isotopes 

in Hydrogeology: Boca Raton, N.Y., Lewis 

Publishers.
73

Davis, R. K., Brahana, J. V., and Johnston, J. S., 

2000, Ground water in northwest Arkansas: 

Minimizing nutrient contamination from nonpoint sources in karst terrane: Arkansas Water 

Resources Center, Publication Number MSC288, http://www.uark.edu/depts/awrc/

pdf_files/MSC/MSC-288.pdf.

Einsiedl, F., Maloszewski, P., and Stichler, W., 2005, 

Estimation of denitrification potential in a karst 

aquifer using the 15N and 18O isotopes of NO3

-

: 

Biogeochemistry, v. 72, p. 67-86.

Faure, G., and Mensing, T.M., 2005, Carbon, in 

Isotopes: Principles and Applications: Hoboken, 

N.J., John Wiley and Sons, Inc., p. 753-802. 

Gillip, J. A., 2007, The effects of land-use change on 

water quality and speleogenesis in Ozark cave 

systems – a paired cave study of Civil War and 

Copperhead Caves, northwestern Arkansas [M.S. 

thesis]: Fayetteville, University of Arkansas, 77 

p. 

Kendall, C., 1998, Tracing nitrogen sources and 

cycling in catchments, in (Eds.) C. Kendall and 

J. J. McDonnell, Isotope Tracers in Catchment 

Hydrology: Amsterdam, Elsevier, p. 519-576.

Knierim, K. J., 2009, Seasonal variation of carbon 

and nutrient transfer in a northwestern Arkansas 

cave [M.S. thesis]: Fayetteville, University of 

Arkansas, 141 p.

Peterson, E.W., Davis, R.K., Brahana, J.V., and 

Orndorff, H.A., 2002, Movement of nitrate 

through regolith covered karst terrane, northwest 

Arkansas: Journal of Hydrology, v. 256, p. 35-

47. 

Nascimento, C., Atekwana, E.A., and Krishnamurthy, 

R.V., 1997, Concentrations and isotope ratios of 

dissolved inorganic carbon in denitrifying 

environments: Geophysical Research Letters, v. 

24, p. 1511-1514.

National Cooperative Soil Survey, 2008, United 

States Department of Agriculture, Custom Soil 

Resource Report for Madison County, Arkansas, 

17 p.

Sigman, D. M., Casciotti, K. L., Andreani, M., 

Barford, C., Galanter, M., Bohlke, J. K., 2001, A 

bacterial method for the nitrogen isotopic 

analysis of nitrate in seawater and freshwater: 

Analytical Chemistry, v. 73, p. 4145-4153.

Spötl, C., Fairchild, I.J., and Tooth, A.F., 2005, Cave 

air control on dripwater geochemistry, Obir 

Caves (Austria): Implications for speleothem 

deposition in dynamically ventilated caves: 

Geochimica et Cosmochimica Acta, v. 69, 

p. 2451-2468. 

St-Jean, G., 2003, Automated quantitative and 

isotopic (13C) analysis of dissolved inorganic 

carbon and dissolved organic carbon in 

continuous-flow using a total organic carbon 

analyser: Rapid Communications in Mass 

Spectrometry, v. 17, p. 419-428.

U.S. Department of Agriculture, 2010, State Facts 

Sheets: Arkansas, http://www.ers.usda.gov/

StateFacts/AR.htm, [accessed February 26, 

2010].
74

Interpreting a Spring Chemograph to Characterize Groundwater 

Recharge in an Urban, Karst Terrain 

By Victor Roland1, Carlton Cobb1, Lonnie Sharpe1, Patrice Armstrong2, Dafeng Hui2, 

and Thomas D. Byl3

1

College of Engineering, Tennessee State University, 3500 John A Merritt Blvd, Nashville, TN 37209

2

Biology Dept., Tennessee State University, 3500 John A Merritt Blvd, Nashville, TN 37209 3 U.S. Geological Survey, 640 Grassmere Park, Suite 100, Nashville, TN 37211

Abstract

Karst aquifers in urban locations are considered particularly vulnerable to contamination for three 

reasons. First, karst features and hydraulic processes tend to promote rapid movement of surface waters 

into the sub-surface with little or no filtration. Second, urban settings tend to have a higher density of 

contaminant sources. Third, urban settings have a greater amount of impervious surfaces that were 

frequently designed to direct surface runoff to sinkholes or losing streams with no filtration. The 

objective of this project was to better understand the vulnerability of Nashville’s shallow aquifer to 

contamination by using tracers to estimate residence time in the shallow karst aquifer. Two years of near 

continuous temperature and conductivity monitoring was done at Tumbling Rock Spring; located on 

Tennessee State University’s campus. Additional synoptic sampling was done to augment the electronic 

monitoring. The data include dO2, pH, NO3, SO4, iron, discharge, and turbidity. Attention was also paid 

to the weather to ascertain patterns in the spring chemographs associated with rain events. The data show

that groundwater discharging from the spring maintained a temperature of 17.5o centigrade (C) plus or 

minus 1 degree C year round. Specific conductance generally dropped during the drier summer months 

and then rose during the wet winter season. This pattern was punctuated by sharp peaks and valleys 

associated with rain events. Based on tracer studies using sodium chloride, up to 10 percent of the 

spring’s flow has an estimated aquifer residence time of less than 1 month.
75

Investigation of Nitrate Processing in the Interflow Zone of Mantled 

Karst, Northwestern Arkansas

By Jozef Laincz

Environmental Dynamics, University of Arkansas, 113 Ozark Hall, Fayetteville, AR 72701

Abstract

Anthropogenic nitrate contamination of groundwater is a common problem in vulnerable terrains 

dominated by karst topography. Elucidation of in-situ nitrate dynamics is important to the design of 

sustainable land-management practices in these terrains. A field-scale study was conducted at a manureamended, mantled karst site in the Ozark Highlands to characterize multiple potential sources and 

processes affecting nitrate in the interflow zone. Increased water residence time and water-matrix contact 

may favor nitrate attenuation processes, such as denitrification, in the interflow zone, which is situated 

between the soil and focused-flow (bedrock) zones. Groundwater samples were collected along the 

hydrologic gradient in and below the study plot and analyzed for reactive species concentration (nitrate), 

conservative species concentration (chloride), nitrate isotopic compositions (δ15N and δ18O), and 

dissolved organic carbon concentration and bioavailability. δ15N and δ18O indicated a mixed soil organic 

matter/manure origin for nitrate. Mass-balance calculations indicated that although mixing was the 

primary process decreasing nitrate concentration along flowpaths through the interflow zone, up to 33 

percent of nitrate moving through the interflow zone may have been removed through microbial 

processing. The magnitude of this processing varied spatially and temporally. Dissolved organic carbon 

bioavailability was elevated and its concentration decreased in the interflow zone relative to the focusedflow zone, suggesting the availability and utilization of the substrate required for nitrate processing in this 

zone. Overall, the observations suggest that the interflow zone may be important for nitrate attenuation in 

karst systems.

INTRODUCTION

High nitrate (NO3

-

) concentrations in water 

are detrimental to man and the environment. 

High intake of NO3

- may result in formation of 

potentially carcinogenic compounds in the 

human gastric system (Tenovuo, 1986) as well 

as low oxygen levels in blood of infants, a 

potentially fatal condition and the reason for the 

U.S. Environmental Protection Agency to 

regulate NO3

- in drinking water by establishing 

maximum contaminant level of 10 mg/L NO3

- as 

nitrogen (N) (Fan and Steinberg, 1995). Large 

amounts of NO3

- discharging from agricultural 

watersheds have been implicated in the 

development of hypoxic zones around the world 

including in the Gulf of Mexico (Rabalais and 

others, 1996; Goolsby and Battaglin, 2001). 

Intensive animal production in Northwest 

Arkansas generates large volumes of manure 

rich in NO3

- precursors, organic N compounds. 

Counties in Northwest Arkansas rank among the 

top 2 percent of counties in the nation which 

together generate 12 percent of the nation’s total 

manure nitrogen. With the mass of manure 

generated, these counties do not have sufficient 

crop and pasture land to assimilate the N 

generated (Gollehon and others, 2001). 

Moreover, these large volumes of manures are 

typically applied to pastures with soils which 

often are, as is the case in Northwest Arkansas, 

extremely thin with limited nutrient 

bioremediation potential. To accentuate 

problems further, Northwest Arkansas is 

underlain by karst; a carbonate rock terrain made 

up of interconnected and hierarchically enlarged 

fractures, conduits and large voids which can 

readily transport manure-derived contaminants 

into deeper subsurface with little chance for 

attenuation. Unsurprisingly, numerous studies 

reported NO3

- contamination in springs and 

wells throughout Northwest Arkansas (Adamski, 

1997; Steele and McCalister, 1990; Davis and 

others, 2000; Laubhan, 2006) where demand for 

water is rapidly increasing, with the region being 

the sixth most dynamically growing 
76

Metropolitan Statistical Area in the country 

(U.S. Census Bureau, 2001).

To protect groundwater in this vulnerable 

landscape, effective manure management 

practices need to be designed taking into account 

the capacity of the local mantled karst 

environment to attenuate NO3

-

. That, in turn, 

requires elucidation of in-situ NO3

- attenuation 

processes among which primarily are vegetative 

uptake, dilution with waters low in NO3

-

, and 

denitrification, i.e. microbial reduction of NO3

-

to N gases (NO, N2O, N2). Denitrification is the 

most important NO3

- attenuation process in as 

much as it removes NO3

- from the watershed in 

the form of N gases (Martin and others, 1999). 

These gases represent a longer term sink for N 

relative to N immobilization processes that lead 

to generation of organic N which may later be 

subjected to remobilization. Denitrification has 

been successfully measured using a range of 

methods in a wide range of terrestrial, marine, 

and freshwater environments (see review by 

Groffman and others, 2006). The occurrence of 

denitrification generally requires a set of 

conditions including NO3

-

/NO2

-

availability, 

presence of denitrifiers, low O2 concentrations, 

sufficiently long residence time, and a supply of 

bioavailable organic matter (Seitzinger and 

others, 2006).

Karst environments are considered 

unfavorable to denitrification because of high 

flow velocities and low nutrient supply. 

However, many karst terrains, including 

Northwest Arkansas, are mantled by regolith 

containing impermeable layers such as relict 

insoluble chert ledges and fragipan. These along 

with the surface of the underlying limestone 

bedrock divert infiltrating groundwater laterally 

(Al-Rashidy, 1999; Little, 2001), creating a 

hydrologic compartment known as the interflow 

zone. The interflow zone delays the movement 

of water into the deeper, focused-flow 

compartment of the system and facilitates 

contact between water and soil or rock where 

NO3

-

utilizing microbes reside. In addition, the 

interflow zone may contain surface-derived 

bioavailable dissolved organic matter as an 

energy source for denitrifiers. Thus the interflow 

zone may be an important zone of NO3

-

attenuation. This study aimed to investigate the 

potential of this zone for NO3

- attenuation 

including denitrification and other 

biogeochemical as well as physical processes 

affecting the fate of NO3

-

, using a mass balance 

approach involving reactive (NO3

-

) and 

conservative (Cl-

) species, NO3

- δ15N and δ18O 

isotopic composition, and dissolved organic 

carbon concentration (DOC) and bioavailability. 

Previous research of these topics has been 

sparse.

METHODS

Site

The study was conducted at the Savoy 

Experimental Watershed (SEW), near the town 

of Savoy, Arkansas. The area is typical of the 

mantled karst setting of the Springfield Plateau 

of the Ozarks where regolith covers the 

underlying chert-rich limestone. Topography is 

ridge and valley with elevation in the watershed 

ranging from 317 to 376 m. Land cover consists 

of hardwood forest (57%) and pasture (43%). 

The study focused on and around an 

instrumented plot (Fig. 1) in Basin 1, which is 

drained by an ephemeral stream that flows 

towards the southwest and discharges into the 

Illinois River. Average annual rainfall for the 

area is 1,119 mm with mean January and July air 

temperatures of 1.1 and 25.9°C, respectively 

(Owenby and Ezell, 1992).

Sampling

The fate of NO3

- was studied across a 

hydraulic gradient; from an infiltration area 

within and in the vicinity of an instrumented plot 

located on a ridge-top pasture (15% slope) in 

Basin 1, through the subsurface interflow zone, 

to the main drainage points of the basin – Langle 

and Copperhead Spring; special attention was 

given to the interflow zone. Hydrogeology of the 

system was established by previous research 

focusing on various aspects such as runoff and 

infiltration mechanisms (Sauer and others, 2000; 

Sauer and Logsdon, 2002, Leh, 2006), regolith 

and bedrock geophysics (Ernenwein and

Kvamme, 2004), water and solute movement in 

the regolith mantle (Al-Qinna, 2003; Brahana 

and others, 2005; Brahana and others, 2006; 

Laincz, 2007) as well as saturated-flow 

characteristics and controls in the underlying 

bedrock (Brahana, 1995; Brahana, 1997; 
77

Brahana, 1999; Whitsett, 2002; Ting, 2005). 

Based on this research, runoff events are rare 

and rainfall tends to quickly infiltrate into the 

regolith mantle where water moves via diffuse 

and macropore flow in both vertical and lateral 

directions; lateral movement – interflow –

occurs owing to permeability contrasts presented 

by chert ledges and the bedrock surface. 

Macropores including root channels, worm 

holes, and pores created by the loose contact 

between chert fragments and soil matrix are 

abundant and greatly accelerate water and solute 

flow through the mantle. At the study site, 

interflow paths begin in the upland area 

underneath a thin (30-75 cm) soil horizon 

relatively rich in organics. These emerge on the 

side slope down-gradient as springs. During wet 

antecedent conditions the time of travel of a 

solute from the upland plot area to these springs 

(through the interflow zone) can range from 10 

to 40 hours. Down-gradient of the emergence 

points, the flow forms a losing stream, with the 

water entering focused-flow dominated bedrock 

to emerge after approximately 500 meters at 

Langle and Copperhead Springs. The time of 

travel for this focused flow for a conservative 

tracer can be approximately 11.5 hours with a 

peak concentration arriving 15-19 hours after 

tracer injection (Whitsett, 2002; Ting, 2002). 

Three distinct hydrologic components or 

zones of the system were characterized (Fig. 1): 

diffuse flow, interflow, and focused flow. 

Diffuse flow (soil water) was sampled by 8 

suction-cup lysimeters installed 75 cm deep in 

the ground inside the plot; these samples 

required compositing owing to low volume yield 

of individual lysimeters. The interflow zone was 

sampled at an upper and lower point on the 

flowpath in order to evaluate NO3

-

evolution 

within this zone. The upper point (referred to as 

‘interflow in’) was an interceptor trench 

(Smettem and others 1991) and constructed at 

the down-slope boundary of the plot. The second 

point (interflow out) were three seeps (J2, J2b, 

J3) on the slope side, downgradient from the plot 

and the trench, representing the terminus of the 

interflow zone flowpaths. Finally, focused-flow 

sampling points included two main springs 

draining Basin 1, Copperhead and Langle, and 

one smaller spring adjacent to the plot area, J1. 

Dye tracing (Laincz, 2007) confirmed that J1 

has no hydrologic connection to the plot area, 

and distinct physical and chemical parameters of 

its water suggest that it is dominated by focused 

flow. In addition, surface runoff was collected 

for three events. Discharge and field parameters 

including pH, conductivity, temperature, and 

dissolved oxygen were measured. Sampling was 

conducted for seven events between July 2004 

and June 2005, three of which were conducted 

under storm flow and four under base flow 

conditions. The plot area was amended with 

chicken litter in June and October 2004 at a rate 

of 5 tons per acre to simulate common 

agricultural practices in terms of NO3

- loading 

and to increase NO3

- signal.

Analysis

Samples were analyzed for major anions and 

cations using ion chromatography (EPA method 

300.0) and ICP (EPA method 200.7), 

respectively. Stable isotopes of NO3

- (δ15N and 

δ18O) were using the Denitrifier Method 

(Sigman and others, 2001; Casciotti and others, 

2002). The method precision and accuracy for 

δ15N were 0.35‰ and 0.43‰, respectively, and 

for δ18O they were 0.66‰ and 0.78‰, 

respectively. DOC concentration was 

determined using high-temperature combustion 

with a Shimadzu TOC5050 (Benner and Strom, 

1993). Bioavailability of DOC was measured as 

groundwater community respiration normalized 

to DOC concentration and bacterial abundance, 

Figure 1. Study plot with hydrogeologic components 

and sampling instruments. Langle and Copperhead 

springs are located about 500 yards from the site.
78

which was determined using epifluorescent 

microscopy (Porter and Feig, 1980). Details 

pertaining to bioavailability measurement as 

well as to all of the above mentioned sampling 

and analytical procedures can be found in 

Winston (2006).

RESULTS AND DISCUSSION

DOC concentration and DOM bioavailability

DOC comprised 80% of total organic carbon 

in SEW water samples, providing the necessary 

substrate for microbial processes. DOC 

concentration ranged from 0.14 to 22.44 mg/L 

with overall median of 1.43 mg/L. The median 

for high-flow events was slightly higher (3.03 

mg/L) compared to low-flow events (1.10 

mg/L), and DOC concentration of high-flow 

events also exhibited greater variability. Mean 

DOC concentrations for the July and December 

2004 events were elevated relative to the other 

events sampled. DOC concentrations tended to 

decrease downgradient (Fig. 2) suggesting 

possible utilization for microbial metabolic 

processes. 

DOM bioavailability was determined for only 12 

samples all of which were collected from the 

higher discharge sites – J1, J2, Copperhead 

Spring and Langle Spring – due to difficulty 

obtaining sufficient representative sample 

volumes from the rest of the sampling sites. In 

addition, no one sampling event had all of the 

four sites sampled, and the July 2004 event was 

not sampled for bioavailability at all. 

Nevertheless, the limited data indicate seasonal 

and spatial variation of DOM quality. DOM 

bioavailability in the interflow and focused-flow 

zones was lower during fall events relative to 

spring events (Fig. 3), suggesting a relatively 

more refractory DOC pool during fall and influx 

of labile organic carbon from growing plants and 

microbial exudates during spring. Spatially, 

comparing the individual sampling points, 

bioavailability was elevated in J2 and Langle 

Spring relative to J1 and Copperhead Spring. 

The sites representing the interflow zone 

exhibited about 2.6 times greater bioavailability 

than the focused-flow zone sites (Fig. 4), 

providing the former a needed substrate for NO3

-

 

microbial processing in this zone.

Figure 2. Mean (± 1 Std. Error) DOC 

concentration across the hydrologic gradient.

Figure 3. Mean (± 1 Std. Error) relative 

bioavailability of DOC by season.

Figure 4. Mean (± 1 Std. Error) relative 

bioavailability of DOC by flow type. Focused 

flow is a mean of average bioavailabilities of 

seep J1, Langle and Copperhead Springs. 

Interflow (out) represents seep J2.
79

Nitrate concentration and isotopic composition

NO3

- concentration ranged from 0.64 to 48 

mg/L throughout the study with a median of 3.3 

mg/L. Average NO3

- concentrations were not 

significantly different among events. Along the 

gradient, NO3

- concentration averages for 

lysimeters, trench, and Copperhead spring were 

10.2, 36.6, and 5.3 mg/L, respectively, which 

was higher relative to the interflow zone seeps 

and Langle spring where concentrations were 

within a relatively narrow range of 2.8 to 3.4 

mg/L.

δ15N-NO3

- ranged from 0.35 to 9.66‰ and 

had a mean of 4.85‰. Seasonally, δ15N values 

appeared to slightly increase during spring and 

fall; averages for spring and fall events were 

elevated relative to the overall mean by about 

1.19 and 0.42‰, respectively. Across the flow 

path, the highest average value was in the 

diffuse-flow zone (8.08‰) and the lowest in the 

runoff samples (3.02‰). δ18O-NO3

- values 

ranged from -0.89 to 19.67‰ with a mean of 

5.13‰. Similar to δ15N, δ18O values also tended 

to increase during spring and fall. Average 

values for spring and fall events were elevated 

relative to the overall mean by about 3.24 and 

0.19‰. Across the flow path, runoff had the 

highest average value (8.01‰) while the 

‘interflow in’ (trench) samples had the lowest 

(3.76‰). The measured δ15N and δ18O values 

are similar to those obtained for ground water 

draining karst (e.g. Panno and others, 2001; 

Einsiedl and others, 2005). Ninety percent of 

both δ 15N and δ 18O values fell between 

approximately 2 and 8‰, a range indicative of a 

mixture of nitrate derived from soil organic 

matter and animal manure (Clark and Fritz, 

1997).

Denitrification is known to cause the δ15N 

and δ18O values of the residual NO3

- to increase 

as a result of kinetic isotopic fractionation 

(Böttcher and others, 1990). In addition to this 

isotopic enrichment of the remaining NO3

- pool, 

the ratio of the enrichment of oxygen to N tends 

to be close to 1:2 (Kendall and McDonnell, 

1998) imparting an additional signal for 

recognition of denitrification. Nitrate δ15N and 

δ18O of the samples from the interflow zonedraining seeps and the adjacent small spring J1 

positively correlate indicating simultaneous 

enrichment in both isotopes likely due to 

denitrification. While the slope deviates from 

0.5 and the correlation is only moderately strong 

(r = 0.4), it may not be reasonable to expect a 

much stronger fit with the model considering the 

geochemical and physical heterogeneity of the 

system, 

Nitrate processing assessment through mass 

balance 

Three post-storm sampling events (July 3, 

December 1 and 7 2004) were conducted under 

hydrologic conditions wet enough to induce 

flow in the trench, and these events thus enabled 

assessing NO3

- evolution between the trench as

an upper interflow zone point and interflow zone 

exit points, the seeps. Concentrations of solutes 

in the seeps were lower as compared to the 

trench. For example, average Cl- concentration 

at the seeps was 75% lower than the average 

concentration at the trench. In addition, for a 

given trench-seep flowpath, the magnitude of the 

decrease was unequal between conservative and 

reactive species; NO3

- decrease was on average 

16% greater than Cldecrease. While the 

decrease in Cl- indicated dilution or mixing 

along the interflow zone flowpaths, the greater 

decrease in NO3

- concentration indicated the 

existence of an additional mechanism of NO3

-

removal – microbial processing. 

Figure 5. Co-linearity between δ15N and δ18O 

of seep and J1 samples.
80

The amount of microbial processing was 

evaluated for the three sampling events by 

modeling NO3

- concentration, NO3- δ15N and 

δ18O, and DOC concentration expected at the 

seeps as a result of mixing (dilution) and 

comparing these model values to measured 

values, which represent the combined effects of 

mixing and microbial processes, such as 

microbial assimilation and denitrification. 

Microbial processing would have decreased 

NO3

- concentration beyond the decrease 

resulting from mixing and would also have 

isotopically enriched the remaining NO3

- pool 

(Kendall and Mcdonell, 1998). 

The evidence of dilution along the trenchseep flowpaths led us to adopt a flow model that 

assumed two types of water mixing along these 

flow paths, producing a mixture discharging out 

of the seeps; one type infiltrating in and carrying 

the chemical and isotopic signature of the upgradient plot, where manure amendment had 

provided a strong solute source, and another 

type infiltrating in and carrying the chemical and 

isotopic signature of the surrounding unamended 

off-plot area. Based on dye-tracing evidence, the 

interceptor trench immediately down-gradient 

from the plot was considered representative of 

the plot water, and spring J1 with lack of 

hydrologic connection to the plot was regarded 

as representative of off-plot water type.

Nitrate and DOC concentrations expected in 

the mixtures of plot and off-plot waters were 

calculated using a common two-component 

mass-balance equation (Faure and Mensing, 

2005; Fry, 2006):

where XM, XP, and XOP are the concentrations of 

DOC or NO3

- in a seep mixture, plot and offplot water components, respectively, and f

represents the mass fractions or mixing 

proportions of these components. The mass 

fractions were calculated by solving the equation 

above for using conservative species (Cl-

) as 

concentrations. The off-plot mass fraction 

was then obtained by subtracting from 1.

Table 1 lists the mass fractions of plot and 

off-plot components calculated for each seep 

mixture, NO3

- concentrations expected in the 

seep mixtures from mixing, corresponding 

measured concentrations, and the difference 

between calculated and measured. Average 

calculated NO3

-

concentration in the seep 

mixture was 6.6 mg/L. Average measured 

concentration, however, was about 3.1 mg/L, 

and so the average 3.47 mg of “missing” NO3

- in 

a liter of the seep mixture was possibly 

denitrified and/or microbially immobilized. 

Since plant roots at the site do not reach into the 

interflow zone, assuming that microbial 

processing is indeed the process responsible for 

the missing NO3

- is reasonable. Performing 

similar calculations with DOC, 0.23 mg/L on 

average of DOC loss in the interflow zone 

mixtures could not be accounted for by plot/offplot mixing and was likely consumed by some 

DOC-utilizing microbial process. Although the 

calculated amount of missing DOC appear 

modest, it may be underestimated. Other 

potential inputs of DOC percolating into the 

interflow zone flow-paths and not accounted for 

in the mixing equation may have increased the 

final measured DOC concentration and thus 

reduced the calculated difference ascribable to 

processes other than mixing. 

CONCLUSIONS

Mass-balance calculations based on twocomponent mixing indicated that most of the 

NO3

- moving through the interflow zone was 

removed by dilution. Between 1 and 33%, 

however, depending on flow-path, was removed 

by additional processes which likely included 

microbial processing since uptake by vegetation 

does not occur in the interflow zone. The δ15N 

and δ18O values of this remnant NO3

- present in 

the seep samples and the adjacent site J1 

exhibited moderately strong positive correlation 

(r = 0.4), indicating denitrification. Nitrate δ15N 

and δ18O also proved useful in determining 

correctly the source of nitrate – a mixture of 

organic matter and animal manure. DOC 

concentration generally decreased along the 

flowpath and mass balance computation showed 

that for the majority of the seep samples dilution 

could not account for this decrease entirely and, 

as a result, a different process such as utilization 

by microbes might have also occurred. Both 

DOC concentration and DOM bioavailability 

were elevated in the interflow zone relative to 
81

focused-flow paths, providing the former with 

increased potential for NO3

- attenuation within 

this karst system. Temporally, DOC as well as 

NO3

- increased during high-flow conditions, 

creating favorable conditions for microbial 

processing of NO3

-

. In the karst watershed as a 

whole, DOM bioavailability appeared to be 

influenced by seasonality, with spring having 

greater DOM bioavailability than fall, 

suggesting an influx of labile organic matter 

during spring. Whereas all measured parameters

varied considerably in both time and space, the 

overall biogeochemical evidence from this study

implies that the interflow zone is a zone of 

significant NO3

-

attenuation by dilution and 

denitrification. It seems that more research 

would be warranted to establish more definitive 

patterns of how these processes change with 

environmental factors, flow path, and land use.

Table 1. NO3

- concentration measured and predicted by the binary mixing 

model, and the difference between the two ascribable to microbial 

processing.

Sampling 

event Seep Endmember fractions Nitrate, mg/L

Nitrate 

Processing, 

% Plot 

(Trench)

Off-plot 

(J1)

Modeled Measured Mod.-

Meas.

3-Jul-04 J2 0.11 0.89 3.84 2 1.84 8

3-Jul-04 J2b 0.39 0.61 9.75 2.42 7.33 33

3-Jul-04 J3 0.09 0.91 3.50 1.4 2.10 9

1-Dec-04 J2 0.01 0.99 3.70 3.26 0.44 1

1-Dec-04 J2b 0.03 0.97 4.30 3.42 0.88 2

1-Dec-04 J3 0.02 0.98 4.00 2.63 1.37 3

7-Dec-04 J2 0.15 0.85 7.96 3.42 4.54 12

7-Dec-04 J2b 0.34 0.66 15.06 7.96 7.10 18

7-Dec-04 J3 0.12 0.88 7.01 1.4 5.61 14

ACKNOWLEDGMENTS

I am grateful to Dr. Van Brahana and Katherine 

Knierim, Univ. of Arkansas, for reviewing this 

manuscript. I also thank my advisor, Dr. Phillip 

Hays, for his guidance throughout this project. 

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(2006). Denitrification across landscapes and 

waterscapes: A synthesis. Ecological 

Applications, 16(6), 2064-2090. 

Sigman, D. M., Casciotti, K. L., Andreani, M., 

Barford, C., Galanter, M., and Böhlke, J. K. 

(2001). A bacterial method for the nitrogen 

isotopic analysis of nitrate in seawater and 

freshwater. Analytical Chemistry, 73, 4145-

4153. 

Smettem, K. R. J., Chittleborough, D. J., Richards, B. 

G., and Leaney, F. W. (1991). The influence of 

macropores on runoff generation from a hillslope 

soil with a contrasting textural class. Journal of 

Hydrology, 122(1-4), 235-252. 

Tenovuo, J. (1986). The biochemistry of nitrates, 

nitrites, nitrosamines and other potential 

carcinogens in human saliva. Journal of Oral 

Pathology, 15(6), 303-307. 

Ting, T. (2005). Assessing bacterial transport, storage 

and viability in mantled karst of Northwest 

Arkansas using clay and Escherichia coli labeled 

with lanthanide series metals. Unpublished Ph.D. 

Dissertation, University of Arkansas, 

Fayetteville. 

U.S. Census Bureau. (2001). Metropolitan areas 

ranked by percent population change: 1990 to 

2000. Retrieved January 1, 2008, from 

http://www.census.gov/population/cen2000/phct3/tab05.pdf 

Whitsett, K. S. (2002). Sediment and bacterial tracing 

in mantled karst at Savoy Experimental 

Watershed, Northwest Arkansas. Unpublished 

M.S. Thesis, University of Arkansas, 

Fayetteville. 

Winston, B. A. (2006). The biogeochemical cycling 

of nitrogen in a mantled karst watershed. 

Unpublished M.S. Thesis, University of 

Arkansas, Fayetteville.
84

KARST MAPPING, DYE TRACING, AND 

GEOGRAPHIC INFORMATION SYSTEMS

Using a Combination of Geographic Information System Techniques 

and Field Methods to Analyze Karst Terrain in Selected Red River 

Sub-Watersheds, Tennessee and Kentucky

By David E. Ladd

U.S. Geological Survey, 640 Grassmere Park, Suite 100, Nashville, TN, 37211 

Abstract

Karst features such as closed depressions and their catchments present challenges to natural-resources 

management and topographic analysis. Diversion and collection of surface-water runoff by these features 

can complicate analyses of stream flow, groundwater recharge and flow, and contaminant transport. In 

karst areas, some component of surface-water runoff is diverted to closed depressions that recharge 

groundwater at various rates, thus affecting the amount of direct runoff reaching streams and providing 

potential pathways for contaminant entry to a groundwater system. 

The U.S. Geological Survey (USGS) Tennessee Water Science Center (WSC), in cooperation with 

the Tennessee Department of Environment and Conservation (TDEC), is applying Geographic 

Information System (GIS) terrain analysis along with field data-collection methods to identify karst 

features, characterize their surface-water/groundwater connectivity, and determine the fate of 

contaminants that enter these features in support of water-resources protection in selected sub-watersheds 

of the Red River in Tennessee and Kentucky.

The analysis focuses on quantifying the number, locations, catchment areas, storage capacities, and 

drainage rates of closed depressions in this watershed. The results are incorporated into analytical, 

statistical, and conceptual models to explore the role of closed depressions and their internal drainage 

characteristics as controls on basin-scale hydrologic response and contaminant delivery to streams and 

springs.

Calculations of depression water storage volume and catchment area provide estimates of the rainfall 

required to fill a closed depression. If the amount of rainfall runoff diverted into a depression during a 

storm event exceeds the depression drainage rate, flooding may occur. A flooded depression that can no 

longer accept water may reject storm flow and re-divert it to runoff. Depressions with low storage 

capacity will flood and overflow quickly if poor connection to a groundwater system exists. Depressions 

with these characteristics are less likely to transmit significant amounts of contaminated runoff to 

groundwater during storm events. In contrast, depressions with high storage capacity are less likely to 

overflow regardless of their ability to transmit runoff as recharge to a groundwater system. These 

depressions have the potential to collect more runoff than low-storage capacity depressions, and they may 

recharge a groundwater system relatively rapidly if they are hydraulically well connected to the 

subsurface. If depression drainage rates can be determined, techniques to define which depressions are 

most likely to quickly transmit the highest quantities of runoff to a groundwater system could be 

developed. Such techniques could be used to categorize depressions based on their potential to contribute 

contaminants to groundwater and surface water, and to focus more-detailed efforts such as dye-tracer tests 

on depressions that have the highest contaminant-transport potential. These methods then could be 

applied in karst areas at a regional scale to help define surface-water/groundwater interaction at lower 

costs than typically labor intensive field-only studies.
85

Parellelism in Karst Development Suggested by Quantitative Dye 

Tracing Results from Springfield, Missouri

By Douglas R. Gouzie and Katherine Tomlin

Department of Geography, Geology, and Planning, Missouri State University, 901 S. National Ave, 

Springfield, MO 65897 

Abstract

Recent and continuing urban development on the southern edge of the city of Springfield, Missouri 

has drawn public attention to typical karst processes such as sinkhole formation and surface stream 

capture by underlying conduit piracy. Four dye traces were performed in the Ward Branch watershed to 

better understand the subsurface karst drainage system in this rapidly urbanizing area – including a trace 

from a sinkhole that opened in 2009 in the stormwater detention basin of a McDonald’s restaurant. 

Quantitative dye tracing involved using automatic samplers in springs to collect discharging groundwater 

and direct-water fluorometry of these samples to analyze for the presence and concentration of dye at 

each of these springs. This extensive data collection using hourly sample data allowed construction of 

well-defined dye-trace break-through curves. The results of these dye traces indicated that dye injected 

into each of four sinkholes travelled to springs in two distinct discharge areas located approximately onequarter mile apart on opposite sides of US 160. The average flow velocities calculated during this study 

ranged from 0.11 to 0.27 feet per second for the springs in the western area, and from 0.22 to 0.35 feet per 

second for the springs in the eastern area. Follow-up work, which is ongoing in this study area, includes 

use of a dye-injection pump to improve discharge determinations along the flowpaths during a dye trace 

and surface geophysics to better delineate likely conduits controlling groundwater flow. Preliminary 

results from the dye-injection system suggest the presence of a conduit connecting the groundwater 

system between the eastern and western springs (within, or essentially beneath, the rapidly developing US 

160 corridor) and that this conduit may be transporting as much as five cubic feet per second (ft3

/s) of 

flow. This volume is similar to the typical 4 to 5 ft3

/s in the surface portions of Ward Branch flowing in 

the study area. 
86

Geologic Controls on Karst Landscapes in the Buffalo National 

River Area of Northern Arkansas: Insights Gained from 

Comparison of Geologic Mapping, Topography, Dye Tracers and 

Karst Inventories

By Mark R. Hudson1

, Kenzie J. Turner1, Chuck Bitting2, James E. Kaufmann3, Timothy M. 

Kresse4 and David N. Mott5

1

U.S. Geological Survey, Box 25046, Mail Stop 980, Denver, CO 80225

2

National Park Service, 402 N. Walnut, Suite 136, Harrison, AR 72601

3

U.S. Geological Survey, 1400 Independence Road, Mail Stop 543, Rolla, MO 65401

4

U.S. Geological Survey, 401 Hardin Road, Little Rock, AR 72211

5

U.S. Geological Survey, 2617 E. Lincolnway, Suite B, Cheyenne, WY 82001

Abstract

The Ozark Plateaus host extensive karst landscapes developed on soluble bedrock. Geologic 

mapping by the USGS at 1:24,000 scale in the western Buffalo River region of northern Arkansas 

highlights the geologic controls on karst features and provides a scientific basis for resource 

management at Buffalo National River (BNR), a 214 kilometer (km) long, river-corridor park.

The western Buffalo River flows eastward through plateau surfaces of the Boston Mountains and 

Springfield Plateau where it has eroded a valley 130 to 400 meter (m) deep. Stratigraphically, the 

watershed exposes a 500 m aggregate thickness of carbonate, sandstone, and shale formations of 

Pennsylvanian, Mississippian, and Ordovician age. As potential karst hosts, limestone and 

dolostone intervals are significant in five of the eight major map formations. Noteworthy are (1) 

the 120 m thick cherty limestone of the Mississippian Boone Formation that has widespread 

surface exposures, forming the Springfield Plateau aquifer, and (2) dolostone and lesser limestone 

intermixed with sandstone in the underlying Ordovician Everton Formation, forming the upper 

part of Ozark aquifer. Structurally, most rocks in the Buffalo River region dip gently (less than 

5°) but are broken by a series of faults and folds that formed during late Paleozoic development 

of the southern Ozark Dome. These structures produce vertical relief of rock units of as much as 

300 m across the region. Multiple joint sets are pervasive in bedrock; most-common strike 

directions are north to south, northeast to southwest, and west-northwest to east-southeast. 

In BNR, comparison of a cave inventory to the geology demonstrates that most caves either 

lie within limestone of the Boone Formation (55 percent), or limestone or dolostone intervals 

within the Everton Formation (38 percent). Large sinkholes are preferentially concentrated near 

the contact of Boone Formation with overlying Batesville Sandstone. Most springs within the 

watershed discharge near the base of the Boone Formation, particularly in its basal Saint Joe 

Limestone Member. The largest of these springs in the Springfield Plateau aquifer are localized in 

structural lows formed by faults and folds; dye-tracer studies demonstrate that springs in three 

structural lows gather interbasin recharge from outside the Buffalo River watershed. A secondary 

concentration of springs, including the largest in BNR, discharge from the lower part of the 

Ordovician Everton Formation. Development of this lower karst aquifer in the Everton Formation 

was facilitated by a thickening of carbonate-rich facies compared to more sand-rich facies of the 

formation farther west. In contrast to relations for the Boone Formation, structural highs localize 

recharge and discharge of this lower Everton karst aquifer. In summary, both stratigraphic and

structural characteristic have had important influence on karst development in the Buffalo River 

watershed.
87

Determining the Relation between Water Loss from Beaver Creek and 

Kamas Fish Hatchery Springs, Samak, Utah

By Lawrence E. Spangler

U.S. Geological Survey, 2329 Orton Circle, Salt Lake City, Utah 84119 

Abstract

On October 13, 2010, fluorescein dye was injected into a sinkhole that developed alongside Beaver 

Creek about 4,000 feet upstream of the Kamas State Fish Hatchery in Samak, Utah. Activated charcoal 

packets used for adsorption of the dye were placed in the north and south hatchery springs as well as in 

several other springs in the vicinity of the hatchery and at selected locations along Beaver Creek. An 

automatic water sampler also was installed in the south spring to collect water samples on a 6-hour 

interval for determining groundwater time of travel from the sinkhole to the spring. The injected dye was 

first detected in a water sample collected at 2200 on October 16, or approximately 3 days and 8 hours 

after injection. The concentration of dye in the spring water peaked in the sample collected at 1000 on 

October 17. The dye also was detected on activated charcoal samples that were placed in both hatchery 

springs as well as those placed in other groundwater discharge points on the hatchery property. In 

addition, dye was detected on charcoal samples placed in the outflow from a private fish hatchery on the 

south side of Beaver Creek, in the outflow from springs discharging in the lower part of the Left-hand 

Canyon drainage, and in the outflow from a spring located about 4,600 feet west of the hatchery. A

hydrologic connection between the sinkhole and the hatchery springs as verified by dye tracing also was 

indicated by measured changes in specific conductance and metered discharge of water from the south 

spring during the dye test, as well as by visual changes in turbidity. Results of the dye tracing indicate that 

only part of the water diverted into the sinkhole discharges from the hatchery springs and that divergent 

flow paths exist to other springs. Groundwater flow from the sinkhole to the hatchery springs may be in 

part, along preferential flow paths within the unconsolidated valley-fill deposits. 

INTRODUCTION

In early September 2010, residents in the 

community of Samak, Utah, about 50 miles east 

of Salt Lake City, noticed that flow in Beaver 

Creek had diminished substantially in the 

vicinity of the Kamas Fish Hatchery, one of 10 

State-operated spring-fed hatcheries. Several 

days later, it was discovered that a sinkhole had 

developed along the south side of the creek 

approximately 4,000 feet upstream of the 

hatchery and was taking most of the flow of the 

creek, reported to have been as much as 1,000 

gallons per minute (gpm). Personnel at the 

Kamas hatchery also noticed that flow of the 

two principal springs that supply the hatchery 

had increased by about 60 to 80 gpm and had 

become increasingly turbid. These relations 

indicated that at least some of the water losing 

into the sinkhole was likely discharging from the 

hatchery springs. More importantly, because fish 

in Beaver Creek upstream of the hatchery tested 

positive for the spore that causes whirling 

disease, a potential existed that water carrying 

these spores could move along fractures or other 

high permeability flow paths from the Beaver 

Creek sinkhole to the hatchery springs. As a 

result, the U.S. Geological Survey (USGS) in 

cooperation with the Utah Division of Wildlife 

Resources conducted a dye-tracing study to 

determine if a hydraulic connection existed 

between the sinkhole and one or both of the 

hatchery springs.

During 2004 and 2006, the USGS also 

conducted investigations at the Kamas Fish 

Hatchery to identify sources of water to the 

north and south hatchery springs (Spangler, 

unpublished data) and to assess the potential for 

spores that can transmit whirling disease to 

move from contaminated surface-water sources 

along underground flow paths to the hatchery 

springs (Spangler and others, 2005). During this 

investigation, the area of focus was on drainages 

along the flank of the adjacent mountains, where 

streamflow losses were known to occur. Results 

of this investigation indicated that streamflow 

losses in Left-hand Canyon and a tributary 

drainage discharged at Left Fork Spring, at the

confluence of these drainages (fig. 1), but did 
88

not discharge from either of the two springs at 

the hatchery. Increases in turbidity in the 

hatchery springs during the annual snowmelt 

runoff period, however, indicated that 

unidentified, discrete (focused) sources of 

recharge to the springs likely existed. 

METHODS

Dye-tracing techniques were used to 

determine if a hydrologic connection existed 

between the sinkhole along Beaver Creek and 

the springs at the hatchery. For this study, 

fluorescein (uranine) dye was selected because 

of its relatively conservative nature in the 

environment, detectability at low concentrations 

and over long distances, ease of analytical 

detection, and very low toxicity. Activated 

charcoal contained in nylon screen packets was 

used for adsorption of the dye and represent an 

integration of dye concentration during the 

period they are deployed. In addition, an 

automatic water sampler at the hatchery’s south 

spring collected water samples on a 6-hour 

interval to refine groundwater time of travel 

between the sinkhole and the spring. Extraction 

of dye from the activated charcoal was done in 

the Utah Water Science Center laboratory by 

using a 5-percent solution of potassium 

hydroxide and isopropyl alcohol, a common 

eluent for extraction of fluorescein dye. Water 

samples also were analyzed in the Center 

laboratory using a Turner Designs Model 10 

filter fluorometer, and selected samples were 

sent to the Edwards Aquifer Authority 

laboratory in San Antonio, Texas, for 

confirmation on a Perkin-Elmer scanning 

spectrofluorometer. Prior to injection of the dye, 

activated charcoal packets were placed in the 

north and south hatchery springs, other selected 

springs in the vicinity of the hatchery, and at 

selected locations along Beaver Creek. A total of 

18 locations were monitored during the test 

(table 1). The automatic sampler began 

collecting water samples one day prior to 

injection of the dye to establish a baseline 

(background) of natural fluorescence in the 

water. The activated charcoal packets were 

collected and replaced initially after 5 days, then 

every week thereafter over the next month. The 

bottles in the automatic sampler were collected 

and replaced every 6 days (4 samples per day on 

a 6-hour interval). In addition, water samples 

were collected from two domestic-use wells 

located near the sinkhole. Landowners of these 

wells collected the samples on a daily basis.

Discharge of the hatchery springs (combined 

outflow) was measured at a concrete weir on the 

south side of the hatchery building that encloses 

the fish runs. Discharge of the south spring was 

obtained from a continuous flow meter located 

inside a degassing unit (to reduce nitrogen 

concentrations) that also allows access to the 

spring vault. Discharge recorded by this meter 

was observed to vary over about 20 gpm. The 

meter at the north spring was not functioning 

during the tracer test but discharge was obtained 

by subtracting the metered flow of the south 

spring from the measured combined flow of both 

springs. Discharge of inflow to the sinkhole 

from Beaver Creek was measured at the time of

the dye injection. Discharge also was measured 

at other selected locations in the vicinity of the 

hatchery (table 1). These measurements were 

used for evaluating the relation between the 

amount of flow going into the sinkhole 

compared with discharge from the hatchery and 

other springs in the area. Flow was measured 

using a standard pygmy current meter and 

wading rod. 

Specific-conductance and temperature 

measurements were taken intermittently at the 

hatchery springs, in Beaver Creek, and at other 

selected springs and outflows in the vicinity of 

the hatchery (table 1). These data were used to 

assess similarities between the water in Beaver 

Creek (and going into the sinkhole) and that 

discharging from the hatchery and other springs 

in the area, and to observe changes in these 

parameters during the dye test. 

In addition, samples of water from Beaver 

Creek at the sinkhole and the south hatchery 

spring were collected and sent to the Utah State 

Health Laboratory for analysis of total and fecal 

coliform bacteria. This was done to determine 

the concentration of bacteria in the creek water 

in comparison to that in the spring water to 

evaluate the potential for a particle-sized 

organism such as the parasitic spore that causes 

whirling disease to be transported underground 

from the sinkhole to the hatchery springs. 
89

Table 1. Field measurements and results of tracer test for monitored sites in the Kamas Fish Hatchery study area, Samak, Utah.

[Map ID, refer to figure 1; gpm, gallons per minute; microS/cm, microSiemens per centimeter @ 25 degrees Celsius (C); NMT, no measurement taken; 

NM, not monitored; NMP, no measurement possible; E, estimate based on visual observation or rough calculation; --, not applicable; <, less than]

Specific Water

Map Discharge Conductance Temperature Detector In Detector Out

Site name ID Date (gpm) (microS/cm) (degrees C) (Date-Time) (Date-Time) 1

Result

Beaver Creek sinkhole inflow 10/13/2010 386 134 8.8 -- -- --

10/18/2010 NMT 138 11 -- -- --

10/24/2010 NMT 135 6.8 -- -- --

10/30/2010 E 300 NMT NMT -- -- --

11/13/2010 269 NMT NMT -- -- --

Beaver Creek at sinkhole 10/13/2010 1,480 NMT NMT -- -- --

10/24/2010 525 NMT NMT -- -- --

Stevens well 1 NMT NMT NMT NMT 2

10/12/10 1900 3

10/30/10 1150 Negative

Worthen well 2 NMT NMT NMT NMT 2

10/13/10 1130 3

10/30/10 1215 Negative

Private hatchery spring source 3 NMT NMT NMT NMT NM NM NM

North hatchery spring (source vault) 4 10/18/2010 NMP 432 11.9 10/12/10 1635 10/18/10 1515 Very positive

10/24/2010 NMP 429 11.8 -- -- --

11/13/2010 NMP 440 12.1 -- -- --

North hatchery spring (degasser vault) 5 NMT NMP NMT NMT 10/12/10 1625 10/18/10 1435 Very positive

South hatchery spring (source vault) 6 10/18/2010 NMP 443 11.6 10/12/10 1645 10/18/10 1525 Very positive

10/24/2010 NMP 438 11.5 -- -- --

11/13/2010 NMP 454 11.6 -- -- --

South hatchery spring (degasser vault) 7 10/12/2010 745 NMT NMT 10/12/10 1610 10/18/10 1405 Very positive

10/18/2010 770 NMT NMT -- -- --

10/24/2010 795 NMT NMT -- -- --

10/30/2010 780 NMT NMT -- -- --

11/13/2010 745 NMT NMT -- -- --

Groundwater inflow (vault) 8 10/12/2010 E 100 NMT NMT 10/12/10 1700 10/18/10 1455 Very positive

Combined hatchery springs outflow 9 10/13/2010 1,500 476 12.2 10/12/10 1655 10/18/10 1450 Very positive

11/19/2010 1,400 NMT NMT -- -- --

Beaver Creek above confluence 10 NMT NMT NMT NMT 10/12/10 1720 10/18/10 1650 Negative

Private hatchery spring outflow 11 10/13/2010 NMT 452 13.6 10/12/10 1555 10/18/10 1720 Positive

10/24/2010 255 424 10.6 -- -- --

11/13/2010 NMT 431 8.3 -- -- --

Left Fork spring 12 10/20/2010 E <100 NMT NMT 10/20/10 2010 11/13/10 1750 Negative

Thompson springs outflow 13 11/19/2010 191 NMT NMT 10/13/10 1240 10/18/10 1735 Positive

Left-Hand Canyon outflow 14 11/19/2010 E 180 NMT NMT 10/18/10 1840 10/30/10 1055 Positive

Spring near Beaver Creek Inn 15 10/24/2010 E 100 394 9.7 10/13/10 1255 10/24/10 1645 Negative

Beaver Creek at Country Lane 16 NMT NMT NMT NMT 10/13/10 1810 10/18/10 1745 Very positive

Gazebo spring 17 10/20/2010 E <100 NMT NMT 10/20/10 1925 10/30/10 1045 Weakly positive

Beaver Creek at Kamas 18 NMT NMT NMT NMT 10/13/10 1835 10/24/10 1800 Weakly positive

1

Result - Based on relative qualitative visual observation of eluted dye sample

2

First water sample collected

3

Last water sample collected
90

Figure 1. Location of monitoring sites and general direction of groundwater flow in the vicinity of the Kamas Fish 

Hatchery, Samak, Utah.

RESULTS AND DISCUSSION

On October 12, 2010, at approximately 

1745, water was diverted from Beaver Creek 

into the sinkhole by breaching a rock berm that 

had been constructed around the sinkhole. 

Because no flow had been going into the 

sinkhole for an extended period of time, water 

was diverted into the sinkhole prior to the 

injection to “prime” the flow path. The rate of 

flow into the hole was adjusted to approximately 

200 gpm so that no pooling would occur. The 

following day, October 13, from about 1320 to 

1350, 300 grams of fluorescein dye were 
91

injected into the flow of water going into the 

sinkhole (fig. 2). Measured flow into the 

sinkhole at this time was 386 gpm, with no 

pooling. Streamflow loss was through large 

cobbles and boulders; no bedrock was exposed 

in the hole. Flow in Beaver Creek just above the 

sinkhole was measured to be 1,480 gpm and 

specific conductance of water in the creek was 

134 microsiemens per centimeter (microS/cm) at 

8.8 degrees Celsius. 

Figure 2. Injection of fluorescein dye into the Beaver 

Creek sinkhole.

On October 18, 5 days after dye injection, 

activated charcoal detectors were pulled from 

the north and south hatchery springs and from 

other springs and sites monitored during the test 

(fig. 1 and table 1) and extracted for the 

presence of dye. Detectors from both of the 

hatchery springs and their associated vaults 

(sites 4 – 7) were very positive for the dye, and 

dye was reported to be visual in the spring vaults 

(sites 5 and 7) the previous day, October 17 (Ted 

Hallows, Division of Wildlife Resources, oral 

commun.). Dye also was recovered from 

detectors placed in a vault into which 

groundwater is collected from other springs on 

the hatchery property (site 8). In the vicinity of 

the hatchery, dye was recovered from detectors 

placed in the outflow (site 11) from a private 

hatchery spring on the south side of Beaver 

Creek (site 3), from detectors placed in the 

outflow from several springs discharging along 

the valley margin in the lower part of the Lefthand Canyon drainage (site 13), about 1,000 feet 

northwest of the hatchery, and from detectors 

placed in the outflow from the main drainage 

(site 14). In addition, dye was detected in a 

spring located on the north side of the highway 

about 4,600 feet west of the hatchery (site 17). 

Dye was not detected downstream in Beaver 

Creek above the confluence with outflow from 

the hatchery (site 10) and was not present on 

detectors collected from a spring located near 

the Beaver Creek Inn (site 15). Left Fork Spring 

(site 12) also was monitored during the test 

although the altitude of the spring is about 40 

feet higher than the altitude of the sinkhole; 

samples from this site were negative. Water 

samples collected daily from the domestic-use 

wells (fig. 1, sites 1 and 2) also tested negative 

for the dye. 

Although dye injected into the sinkhole was 

detected on activated charcoal samples collected 

from both hatchery springs about 5 days after 

injection, samples analyzed from the automatic 

water sampler at the south hatchery spring 

degasser unit (site 7) showed that dye was 

present in the sample collected at 2200 on 

October 16 or about 3 days and 8 hours after 

injection. No dye was detected in the sample 

collected 6 hours prior, at 1600; thus, dye arrival 

at the south spring occurred sometime between 

1600 and 2200. The concentration of dye in the 

spring water peaked in the sample collected at 

1000 on October 17, or almost 4 days after 

injection, and dye was barely detectable in water 

from the spring in the last sample taken at 0600 

on October 30, indicating an overall transit time 

of the dye slug of about 17 days. Measured 

combined flow of both springs at the hatchery 

on October 13, the day of injection, was about 

1,500 gpm, and metered discharge of the south 

spring was about 745 gpm. By difference, flow 

from the north spring at this time was estimated 

to be about 755 gpm. During the tracer study, 

the metered flow of the south spring increased 

from about 745 gpm on October 12, prior to 

diversion of water into the sinkhole, to 775 gpm 

on October 16, just prior to first arrival of the 

dye, to almost 790 gpm on October 18, at about 

the time peak dye concentration occurred at the 

spring, or an overall increase of about 45 gpm.

Response of the south spring to the diversion of 

water into the sinkhole in relation to the dye 

breakthrough curve is shown in figure 3.
92

740

745

750

755

760

765

770

775

780

785

790

795

0

1

2

3

4

5

6

7

8

9

10

10/10/10 10/13/10 10/16/10 10/19/10 10/22/10 10/25/10 10/28/10 10/31/10 11/03/10

Discharge, in gallons per minute

Relative fluorescence units

Diversion into

sinkhole

October 12 1745

Dye injection

October 13 1320

Peak dye concentration

October 17 1000

Figure 3. Comparison of discharge and dye 

breakthrough at the south hatchery spring.

The specific conductance of water from the 

north and south springs also changed in response 

to the diversion of water into the sinkhole during 

the tracer study, decreasing from 432 and 443 

microS/cm on October 18 to 429 and 438 

microS/cm on October 24, respectively. On 

November 13, when observed flow into the 

sinkhole had dropped substantially and flow at 

the south hatchery spring had decreased to 745 

gpm, specific conductance of the north and 

south spring waters increased to 440 and 454 

microS/cm, respectively. Although these 

changes are small in relation to the substantial 

difference in specific conductance between 

water from the creek and the spring, taken in 

conjunction with changes in other parameters 

and similarity to historical measurements 

documenting changes in specific conductance 

during snowmelt runoff, they are considered 

significant. 

A bacteria sample collected from Beaver 

Creek at the sinkhole showed a total coliform 

count of more than 200 cells per 100 ml of 

sample and a fecal coliform count of 53 cells per 

100 ml of sample. The bacteria sample collected 

from the south spring source vault (site 6) 

showed a total coliform count of about 16 cells 

per 100 ml of sample and a fecal coliform count 

of 1 cell per 100 ml of sample. Although 

substantially reduced in numbers and 

considerably smaller in size (1 to 3 microns) 

than the spores that cause whirling disease 

(about 30 microns), the presence of these 

bacteria in the spring water indicates nonetheless 

that the permeability of and rate of groundwater 

movement through the aquifer are high enough 

for transporting particles or suspended materials 

(turbidity) to the springs from focused (point) 

surface-water sources such as the sinkhole.

Mississippian-age carbonate rock units crop 

out on the flank of the adjacent mountains and 

dip generally to the southwest beneath the valley 

of Beaver Creek in the vicinity of the hatchery. 

Well log data indicate, however, that the 

overlying unconsolidated deposits are as much 

as 150 feet thick along the Beaver Creek 

drainage (Hurlow, 2002). The log of the sampled 

domestic well only about 300 feet northwest of 

the sinkhole (site 1) indicates no bedrock at a 

depth of 200 feet. The upper 30 feet of these 

unconsolidated deposits consist primarily of 

alluvial-fan and debris-fan gravels, sands, and 

boulders, which can have relatively high 

hydraulic-conductivity values that are similar to 

those in fractured (non-cavernous) carbonate 

rocks (Heath, 1989). In addition, on the basis of 

the transit time of the dye from the sinkhole to 

peak concentration at the south hatchery spring, 

and assuming a linear flow path of about 3,800 

feet, average groundwater velocity would have 

been on the order of 1,000 feet per day, a value 

that is comparatively slow for conduit flow 

through karst aquifers, where average travel 

times of about 5,800 feet per day have been

documented from dye-tracer tests (Worthington, 

2007).

CONCLUSIONS

Results of the tracer study have shown not 

only that a hydrologic connection exists between 

the sinkhole and the hatchery springs but that 

groundwater also discharges at a number of 

other springs in the vicinity of and downgradient 

from the hatchery, both north and south of 

Beaver Creek. This distributary network of 

discharge points indicates diverging or possibly 

sub-parallel groundwater flow paths, rather than 

a pattern of converging flow paths to one 

principal spring such as Left Fork Spring and 

typical of many karst flow systems. The 

sinkhole along Beaver Creek probably lies on 

the margin or boundary of the hatchery springs 

groundwater basin that defines a weak divide 

from which water can flow in multiple 

directions. Variations in flow of the hatchery 
93

springs in comparison with the amount of water 

diverted into the sinkhole, relative changes in 

specific conductance of water from the hatchery

springs when compared with the low specific 

conductance of water diverted into the sinkhole, 

and most significantly, the detection of dye at 

springs other than the hatchery springs indicate 

that only part of the total amount of water 

diverted into the sinkhole during the tracer study

discharged from the hatchery springs and that 

most of the water discharging from the springs 

originates from other sources. The amount of 

flow discharging from other springs in the 

vicinity of the hatchery that also tested positive 

for dye more than accounts for the volume of 

water diverted into the sinkhole from the creek 

during the study. Given the distribution of 

discharge points, groundwater travel time, and 

geology of the study area, groundwater 

movement between the sinkhole and the 

hatchery and other springs may be at least in 

part, along preferential flow paths within the 

unconsolidated valley-fill deposits, in 

conjunction with flow along fractures in the 

underlying carbonate bedrock where near the 

surface.

 

ACKNOWLEDGMENTS

The author would like to acknowledge the 

assistance of Murray Stevens and Taigon 

Worthen with regard to the collection of water 

samples from their wells during the dye study. 

Ted Hallows, Kamas Fish Hatchery manager 

also provided assistance throughout the dye 

study. The analysis of dye samples by Steven 

Johnson with the Edwards Aquifer Authority is 

greatly appreciated. Finally, reviews of the 

manuscript by Bert Stolp, U.S. Geological 

Survey, and Dan Aubrey and Ed Fall, Utah 

Division of Water Resources, helped to improve 

readability and technical accuracy.

REFERENCES

Heath, R.C., 1989, Basic ground-water hydrology: 

U.S. Geological Survey Water-Supply Paper 

2220, 84 p.

Hurlow, H.A., 2002, The geology of the KamasCoalville region, Summit County, Utah, and its 

relation to ground-water conditions: Utah 

Geological Survey Water Resource Bulletin 29, 

plates 1-5. 

Spangler, L.E., Tong, M., and Johnson, W., 2005, A 

multi-tracer approach for evaluating the transport 

of whirling disease to Mammoth Creek fish 

hatchery springs, southwestern Utah, in U.S. 

Geological Survey Karst Interest Group 

Proceedings, Rapid City, South Dakota, 

September 12-15, 2005: U.S. Geological Survey 

Scientific Investigations Report 2005-5160, p. 

116-121.

Worthington, S.R.H., 2007, Ground-water residence 

times in unconfined carbonate aquifers: Journal 

of Cave and Karst Studies, v. 69, no. 1, p. 94–

102.

 

 
94

Using Geographic Information System to Identify Cave Levels and 

Discern the Speleogenesis of the Carter Caves Karst Area, Kentucky

By Eric Peterson1, Toby Dogwiler2, and Lara Harlan1

1

Department of Geography-Geology, Illinois State University, Normal, IL 61790 2

Southeastern Minnesota Water Resources Center, Department of Geoscience, Winona State University, 

Winona, MN 55987

Abstract

Cave level delineation often yields important insight into the speleogenetic history of a karst system.

Various workers in the Mammoth Cave System (MCS) and in the caves of the Cumberland Plateau Karst 

(CPK) have linked cave level development in those karst systems with the Pleistocene evolution of the 

Ohio River. This research has shown that speleogenesis was closely related to the base level changes 

driven by changes in global climate. The Carter Caves Karst (CCK) in northeastern Kentucky has been 

poorly studied relative to the MCS to the west and the CPK karst to the east. Previously, no attempt had 

been made to delineate speleogenetic levels in the CCK and relate them to the evolution of the Ohio 

River.

In an attempt to understand cave level development in CCK we compiled cave entrance elevations 

and locations. The CCK system is a fluviokarst typical of many karst systems formed in the Paleozoic 

carbonates of the temperate mid-continent of North America. The CCK discharges into Tygarts Creek, 

which ultimately flows north to join the Ohio River. The lithostratigraphic context of the karst is the 

Mississippian Age carbonates of the Slade Formation. Karst development is influenced by both bedding 

and structural controls. We hypothesize that cave level development is controlled by base level changes in 

the Ohio River, similar to the relationships documented in MCS and the karst of the Cumberland Plateau

The location and elevation of cave entrances in the CCK was analyzed using a GIS and digital 

elevation models (DEMs). Our analysis segregated the cave entrances into four distinct elevation bands 

that we are interpreting as distinct cave levels. The four cave levels have mean elevations (relative to sea 

level) of 228 m (L1), 242 m (L2), 261 m (L3), and 276 m (L4). The highest level—L4—has an average 

elevation 72 m above the modern surface stream channel. The lowest level—L1—is an average of 24 m 

above the modern base level stream, Tygarts Creek. The simplest model for interpreting the cave levels is 

as a response to an incremental incision of the surface streams in the area and concomitant adjustment of 

the water table elevation. The number of levels we have identified in the CCK area is consistent with the 

number delineated in the MCS and CPK. We suggest that this points toward the climatically-driven 

evolution of the Ohio River drainage as controlling the speleogenesis of the CCK area.

INTRODUCTION

Surface rivers play an integral role in the 

formation of many karst systems. In fluviokarst 

settings, the formation of phreatic cave passages 

is thought to occur at, or just below, the water 

table; hence as the rivers incise, lower levels of 

conduits are formed at increasingly lower 

elevation (Ford and Williams, 2007; Palmer, 

1987). As a karst system evolves, subsurface 

drainage at the current base level becomes more 

efficient at draining the watershed and surface 

systems and upper cave levels go dry 

(Kaufmann, 2009). Just as the development of 

terraces represent periods of river stability in 

surficial fluvial settings, the formation of cave 

levels across a region provide an archive 

recording periods of base level stability

(Kaufmann, 2009; Palmer, 1987).

The Ohio River and its tributaries provide 

ample evidence of this phenomenon. The 

entrenchment of the Ohio River and its 

tributaries produced multiple levels in the 

Mammoth Cave System (MCS) (Granger and 

others, 2001; Palmer, 1989) and in the 

Cumberland Plateau region (CPK) (Anthony and 

Granger, 2004). In the Mammoth Cave region 

and the Cumberland Plateau, investigators have 

demonstrated that alternating periods of climate-
95

driven incision and aggradation of the Ohio 

River have strongly influenced the evolution of 

the those karst systems over the past 3-5 million 

years (Anthony and Granger, 2004; Granger and 

others, 2001). In both cases, pre-existing cave 

systems experienced significant vertical 

development and modification during the PlioPleistocene due to changes in erosional base 

level, which forced alternating periods of 

incision and aggradation in river valleys south of 

the glacial margin (Teller and Goldthwait, 

1991). Isostatic responses of the continental 

crust to the waxing and waning of ice sheets and 

the resulting transmission of glacially-derived 

sediment packages down the Ohio River valley 

also played a role in the rates of stream erosion 

and deposition (Granger and others, 2001; 

Potter, 1955).

The MCS has four main cave levels1

(Granger and others, 2001) which generally 

correlate with similar levels in the CPK 

(Anthony and Granger, 2004). These cave levels 

 1

In MCS and CPK the cave levels are lettered with 

level ‘A’ representing the upper-most level in the 

system. In CCK we designated our lowest level ‘L1’ 

and increased the number (i.e., ‘L2’, ‘L3’, ‘L4’) with 

each higher elevation.

all ultimately coincide with base level changes 

in the Ohio River and its predecessor drainages. 

The oldest and highest cave levels formed in the 

Pliocene during a period of extremely low rates 

of river incision and landscape denudation. This 

led to long-term stabilization of water table 

levels and the development of extensive and 

large conduit systems (e.g., Collins Avenue in 

the MCS). With the onset of Pleistocene 

glaciation and the evolution of the modern Ohio 

River drainage, base level stability ended and a 

sequence of rapid incision and aggradation led to 

the development of several new cave levels. We 

hypothesize that the Carter Caves Karst (CCK) 

in Carter County, northeastern Kentucky, 

developed due to a similar speleogenetic 

response to the reorganization of the Ohio River 

drainage. 

Geologic Context

The CCK is located about 40 km south of 

the Ohio River in northeastern Kentucky (Fig. 1)

and has a stratigraphic and geologic setting 

similar to that of the MCS. The system has been 

described by several authors including (e.g., 

McGrain, 1966; Tierney, 1985). More recently, 

Engel and Engel (2009) provide a thorough and 

Figure 1. Location of the study area, highlighting Carter Caves State Resort Park. In the aerial image on the 

left, the park boundary is outlined in white. Rivers are represented as blue lines. The location of Carter 

County is shown in the map in the upper right. The location of the state park within Carter County is shown 

in the lower right figure.
96

up-to-date discussion of the local geologic 

setting, including an updated synthesis of the 

area’s stratigraphy. Unlike the larger karst 

systems in western Kentucky or West Virginia, 

karstification in the Carter Caves area is 

constrained by a relatively thin sequence of 

karstifiable carbonates (Engel and Engel, 2009). 

The carbonate sequence is Mississippian (latest 

Osagean to late Chesterian times) in age and is 

sandwiched between Mississippian and Middle 

Pennsylvanian siliciclastics (Ettensohn and 

others, 1984).

Northeastern Kentucky has experienced a 

complex drainage evolution through the PlioPleistocene as the Teays River system was 

abandoned and reorganized into the Ohio River 

drainage (Andrews, 2006; Rhodehamel and 

Carlston, 1963; Teller and Goldthwait, 1991; 

Ver Steeg, 1946). Prior to the onset of 

Pleistocene glaciation, northeastern Kentucky 

was part of the Teays River basin (Janssen, 

1953). Currently, the CCK area is highly 

dissected and characterized by deeply-incised 

stream valleys that are graded to Tygarts Creek,

the regional baselevel. Tygarts Creek flows 

north through Carter and Greenup Counties 

toward its confluence with the Ohio River.

Locally, Tygarts Creek has a very low gradient 

of 0.0007 m/m.Tygarts Creek is currently 

incised through the carbonate sequence into the 

shales of the underlying Borden Formation 

(Engel and Engel, 2009; Tierney, 1985). Thus, 

the lower stratigraphic limit of surficial karst 

development has been reached. Within the study 

area, Horn Hollow Creek drains to Cave Branch 

which in turn flows to Tygarts Creek. These 

tributary valleys are steeply graded (average of 

0.053 m/m) and are underlain by the karstforming carbonates. The tributary streams are 

characterized by numerous small waterfalls, 

sinking streams, and numerous resurgent 

springs. In Horn Hollow the stream is diverted 

into the subsurface by several caves (during 

floods, some flow is diverted to normally dry 

surface channels). 

The CCK is a fluviokarst system and is 

comprised of the surface and subterranean 

drainage associated with Tygarts Creek and its 

tributaries. The karst system includes a number 

of watersheds tributary to Tygarts Creek. Engel 

and Engel (2009) attribute the consistent 

distribution of caves with the same stratigraphic 

units throughout the various watersheds as 

evidence of simultaneous karstification. Engel 

and Engel (2009) also note two morphologically 

distinct cave passage types. The first are large 

trunk passages, that are stratigraphically high 

and whose development is controlled both by 

structure and stratigraphy. The second passage 

type are smaller passages, that are 

stratigraphically lower in the bedrock section 

and are characterized by simple passage 

segments with morphologies indicative of 

incision-driven water table lowering. The 

development of these passages was strongly 

controlled by bedrock fractures. Engel and Engel 

(2009) interpret these passage types as 

representing at least two distinct periods of 

karstification.

Methodology

Our objective is to assess the feasibility of 

applying the incision driven model of 

speleogenesis derveloped for the MCS(Granger 

and others, 2001) and the Cumberland Plateau 

region(Anthony and Granger, 2004) to the CCK.

More specifically, we posit that if cave levels 

can be distinguished in the CCK system, it may 

be possible to link them to the evolution of the 

Ohio River drainage, and thus devise a model 

for the timing and style of speleogenesis in 

CCK. The evolution of those systems has been 

worked out using sophisticated and expensive 

geochemical analysis, particularly of 

speleothems and cave sediments. Herein, we 

attempt to extrapolate the results of those studies 

in combination with GIS-based analysis of 

remotely-sensed data and direct observations 

into a robust model for the speleogenetic history 

of the CCK. Such an approach may be attractive 

to workers deciphering systems that lack the 

geoarchives (i.e., well-understood and wide 

spread sediments) that exist in these other 

systems or the technical and financial resources 

required for the requisite analytical 

methodologies employed in those systems. 

Studies such as ours may also provide a 

framework and an impetus for employing 

sophisticated analytical methods such as those 

used in MCS and CPK. 
97

We employed a GIS (ESRI ArcMapTM 9.2) 

to visualize and analyze various data from the 

karst system. Although the use of GIS to study 

karst has grown in the last decade, many early

applications of GIS to karst focused on GIS as a 

database and management tool for information 

(e.g., Florea and others, 2002; Gao and others, 

2006; Ohms and Reece, 2002). More recently, 

the use of GIS has increasingly included 

sophisticated data analysis, geoprocessing, and 

modeling as a central aspect of the research 

endeavor. For example, GIS has been used to 

indentify sinkholes, faults, and fractures (Angel 

and others, 2004; Florea, 2005; Seale and others, 

2008), to model depressions (Yilmaz, 2007), to 

create virtual field trips through caves (McNeil 

and others, 2002), to model karst hazards 

(McNeil and others, 2002), to delineate karst 

watersheds (Choi and Engel, 2003; Glennon and 

Groves, 2002), and to identify critical source 

areas of contaminants (Dockter and Dogwiler, 

2010).

GIS Data Sources 

Digital topography, hydrography, a digital 

elevation model, and orthophotos layers for the 

study area were obtained from the Kentucky 

Geological Survey2

and the United States 

Geological Survey3

 2

web sites. The DEM for the 

study area has a spatial resolution of 30m. DEM 

accuracy is expressed in terms of root mean 

square error (RMSE) and is assessed based on 

ground control points, the National Map 

Accuracy Standards (NMAS), and the National 

Standard for Spatial Data Accuracy (NSSDA). 

The NMAS is the RMSE that bounds 90 percent

of the values, while the NSSDA is the RMSE 

that bounds 95 percent of the values. Based on 

these various methods of assessment the CCK 

DEM accuracy ranges from 3.74 m to 7.34 m. 

The hydrography data were derived from the 

USGS National Hydrography Dataset (NHD), 

which was created from 1:24,000 Digital Line 

http://www.uky.edu/KGS/gis/kgs_gis.htm

3

http://seamless.usgs.gov/

Figure 2. Histogram of all the cave entrance elevations in CCK. Cave levels were delineated based on where 

there was a high frequency of caves at one elevation with breaks on either side or a reverse in the frequency 

trend.
98

Graphs (DLG). The NHD accuracy is reported 

as 98.5 percent.

ESRI’s ArcCatalogTM was used to build a 

geodatabase for data collected in the field. Cave 

locations and descriptions were obtained from 

the Wittenberg University Speleological 

Society. These data represent a thorough and 

systematic reconnaissance of the study area and 

surrounding areas and represents most—if not 

all—of the discoverable karst features in the 

area. The Wittenberg data lacked elevations of 

the documented features. All elevation data used 

in our analysis were obtained from the DEM and 

compared to elevations directly determined for 

selected features and fixed reference locations in 

the field using a combination of a differential 

GPS, an electronic altimeter, and an analog 

altimeter (Gorecki, 2008). Based upon 

temporally repeated measurements of fixed 

reference locations within the study area we 

estimate the electronic altimeter error at ±3.2 to 

4.3 m and the analog altimeter error at ±0.6 to 

0.9 m. The accuracy of the post-processed 

Figure 3. Box plots of the distance cave entrances are from streams. The ends of the boxes represent the 

25th and 75th percentiles with the solid line at the median; the error bars depict the 10th and 90th 

percentiles, and the dots represent the 5th and 95th percentile. The dotted line represents the mean.

Table 1. Summary of cave level elevations, lateral distance to the stream valley axis, and selected 

example caves from each level. Harlan (2009, p. 32-34) provides a detailed description and statistical 

summary of each of the cave levels. Engel and Engel (2009) provide additional description and context 

regarding the specific caves.

[All units are in meters]

Cave 

Level

# of 

Entrances

Mean / 

Mode 

Elevation

Elevation 

Range

Mean Distance 

to Stream ±

Std. Dev.

Example Caves

L4 25 276 / 270 268 – 283 72 ±56 X Cave, Coon-in-the-Crack Cave

L3 49 261 / 260 249 – 266 41 ± 37 Saltpetre Cave, Rat Cave

L2 47 242 / 245 234 – 247 33 ± 37 Cool James Cave

L1 36 228 / 231 215 – 234 24 ± 15 Laurel Cave, Lake Cave
99

digital GPS data was limited at many locations 

by the steep topography and dense tree canopy. 

Most positions had a vertical accuracy of 1 to 5 

m (one standard deviation error). 

Cave levels were determined based on the 

elevations of the cave entrances. The karst 

features geodatabase was filtered to remove 

vertically oriented features, such as vadose pits 

and sinks. Although the pits and sinks generally 

represent a vertical connection between the 

surface and an underlying cave, the depth from 

the surface to the cave was not known.Thus, the 

pits and sinks were not incorporated into the 

cave level determination. The elevation of cave

entrances was determined based on the elevation 

of the corresponding DEM cell. The distribution 

of cave entrance elevations was analyzed 

statistically in ArcMap (described more 

thoroughly in Harlan, 2009). The ArcGISTM

algorithm ‘Natural Breaks Classifier’ was used 

to generate a histogram showing the frequency 

of cave entrances by elevation (Fig. 2). Cave 

levels were delineated based on where there was 

a high frequency of caves at one elevation with 

breaks on either side or a reverse in the 

frequency trend.

In addition to identifying the cave entrance 

elevations, we determined the shortest distance 

of each cave entrance to the channel of the 

current surface stream valley. This was 

accomplished by deriving a stream line network 

based on the DEM and then using the 

‘Euclidean Distance’ tool in ArcGISTM to 

determine the distance of each DEM cell from 

the nearest stream line. The resulting raster of 

values was queried for each cave entrance using 

the ‘Extraction’ tool to yield the distance values 

which were added as an attribute to the karst 

features geodatabase. 

RESULTS

Error Analysis

Forty-three field-collected elevations for 

cave entrances and the fixed reference locations 

were compared to elevations obtained from the 

DEM. The DEM provides slightly higher 

elevation values. The mean error between the 

field-collected and DEM-derived elevations is -

0.48 m, with a 95percent confidence interval of 

1.25 m. However, the DEM elevations are 

statistically similar to the field-collected 

elevations [t(43) = -0.19, p = 0.85]. The RMSE 

between the field-collected elevations and the 

DEM elevations is 3.96 m. This error is only 

slightly higher than the 3.74 m RMSE of the 

DEM, but is below the NMAS and the NSSDA 

values of 6.15 m and 7.34 m, respectively. Both 

the mean error and the RMSE are within the 

error associated with electronic altimeter and 

DGPS; only the mean error is within the error of 

the analog altimeter. Overall, the data indicate 

that the DEM provides acceptable estimates of 

elevations.

Cave Levels

The location of 157cave entrances were 

analyzed in this study. Based on our analysis of 

the distribution of cave entrances by elevation 

we have delineated four cave levels (Figure 2, 

Table 1). We have denoted these levels as L1 

(mean elevation of 228 m above sea level), L2 

(x ̄ = 242 m), L3 (x ̄ = 261 m), and L4 (x ̄

DISCUSSION

= 276 

m). The mean lateral (horizontal) distances of 

entrances in each cave level to the nearest 

stream valley axis are shown in table 1. In 

general, L4 cave entrances were furthest from 

the streams, as would be expected if the crosssectional (normal to flow) valley shape is 

approximated by the classic ‘V-shaped’ valley 

of fluvial origin. Because the slope of valley 

walls varies significantly, there is also a 

reasonable expectation that this metric will show 

significant overlap in the lateral distances 

ranges. Nonetheless, analysis of variance 

(ANOVA) suggests that the cave levelsare a 

significant predictor of the lateral cave entrance 

to valley axis dimension [F(3, 156) = 8.78, p = 

<0.001]. 

Figure 2 shows our delineation of cave 

levels in the CCK system. The histogram 

contains many breaks and admittedly lends itself 

to a number of interpretations. However, an 

ongoing graduate project at Illinois State 

University is working to refine the delineation of 

cave levels presented here, and the preliminary 

results (Jacoby and Peterson, 2010) support our 

current interpretation. The development of larger 

trunk passages in levels L4 and L3 seem to be 

controlled by subtle changes in dip of the 

bedrock (Engel and Engel, 2009) and likely predate the onset of glaciation to the north of the 
100

study area the led to the reorganization of the 

Ohio and Teays River Drainage systems. Several 

of the upper level passages in the study area 

occur at similar vertical elevations and may be 

truncated remnants of a formerly integrated cave 

system. The lower cave levels in the CCK 

system are controlled by a combination of 

stratigraphic and structural influences and 

correlate strongly to modern-day surface stream 

patterns (Engel and Engel, 2009). 

The Horn Hollow Valley

We will focus our discussion of cave 

hydrology and geomorphology in CCK on the 

Horn Hollow Valley portion of the system 

(Figure 1). Horn Hollow has numerous caves, 

sinking streams, springs, sinks, and pits that are 

all well-studied and documented in the literature 

(Angel, 2010; Dogwiler and Wicks, 2004; Engel 

and Engel, 2009; Hobbs and Pender, 1985; 

McGrain, 1966; Ochsenbein, 1974; Tierney, 

1985) . The upper section of Horn Hollow is 

largely under-drained by an active cave system 

(variously referred to as Boundary Cave or 

Upper Horn Hollow Cave). The surface stream 

channel in this section of the valley is poorly 

maintained and ill-defined indicating a paucity 

of flow events large enough to inundate the 

active cave system and flow across the surface. 

In several places this section of surface stream is 

occupied by large blocks of limestone displaying 

anastomoses, scallops, and other dissolutional 

features associated with caves. It is likely that at 

least some of this portion of the surface stream is 

a former L3 cave that has been hydrologically 

abandoned and subsequently unroofed. 

Several caves higher up in the stratigraphic

section occur in the valley flanks. Some, such as 

Fudge Ripple Cave, are fairly near the contact 

with the siliciclastic units that overlie the 

carbonate sequence. Fudge Ripple Cave and 

another cave—Volcano Cave—appear to be 

examples of phreatic passages that have been 

overprinted with a vadose signature formed as 

waters have cut through the passage floors 

seeking pathways to lowering water tables. 

Stratigraphically, and in terms of elevation, 

these caves represent L3 and L4. In numerous 

places pits and sinks dot the hillslopes along the 

valley walls. Currently, the hydrologic function 

of these caves and pits is to direct water 

vertically down toward the modern phreatic 

zone. 

Dye tracing and water chemistry data 

(Angel, 2010), confirm that Bowel Spring (L2), 

in the central part of the Horn Hollow Valley, is 

a resurgence point for water flowing from 

Volcano Cave and Fudge Ripple Cave through 

Boundary Cave—hydrologically spanning 

several cave levels. From Bowel Spring the flow 

alternates from the surface to the subsurface 

through Cobble Crawl Cave, Horn Hollow Cave, 

New Cave, and H2O caves. Thus, in the lower 

part of Horn Hollow Valley, it is possible to 

explore several of the active L2 and L1 caves. 

H2O Cave and New Cave are phreatic tubes that 

meander along bedding planes and drain 

significant amounts of water during large flow 

events. H2O Cave (L1) emerges from Horn 

Hollow as a waterfall along the contact between 

the St. Louis limestone and the Borden Shale. 

As such, H2O Cave is formed at the 

carbonate/siliciclastic contact that forms the 

lower stratigraphic boundary of cave 

development in the region.

Comparison to Regional Karst Systems

In Mammoth Cave, Palmer (1987) and 

Granger and others (2001) identify four levels 

centered around 150 m, 167 m, 180 m, and 200 

m. The number of levels within the MCS

corresponds well to the CCK area, but there is 

an absolute difference of ~80 m between the 

levels of the two systems. We assume this 

difference in absolute elevations is a function of 

regional dips. However, the relative elevation 

differences between individual levels in each 

area are roughly comparable. Additionally, the 

three lowest levels in the MCS are also in the 

Ste. Genevieve Limestone, which is correlative

between the Mammoth Cave and CCK areas.

In the MCS the upper levels (Level A, 200m 

and Level B, 180 m) formed in the Pliocene and 

early Pleistocene due to slow valley deepening 

and aggradation, while the lower levels (Level 

C, 167 m and Level D, 150) developed during 

the Pleistocene glacial intervals during periods 

of base level stability (Palmer, 1987). Using 

cosmogenic26Al and 10Be dating, Granger and 

others (2001) determined that Levels A and B 

were both formed prior to 3.25 Ma and constrain 

the formation of Levels C and D as prior to 1.39 

Ma and 1.24 Ma, respectively. 

The CPK also has four levels (Anthony and 

Granger, 2004). Cosmogenic26Al and 10Be 
101

analysis demonstrates that the upper-most level 

(Level 1) was formed between 5.7 and 3.5 Ma, 

the second level was formed between 3.5 and 

2.0 Ma, the third level was formed between 2.0

and 1.5 Ma, and the fourth level was formed 

after 1.5 Ma. Thus, levels one and two formed in 

the Pliocene and levels three and four formed in 

the Pleistocene. 

MCS, CPK, and CCK are geographically 

close (within 300 km of one another), contain 

many of the same stratigraphic units, and are 

ultimately controlled by the base flow of the 

Ohio River. Thus, it is reasonable to hypothesize 

that CCK may share a similar history of cave 

development. However, unlike the Green River 

and the Cumberland Rivers which flowed west 

into the Old Ohio River in pre-Glacial times 

(Granger and others, 2001; Teller, 1973), 

northeastern Kentucky was part of the southern 

branch of the Teays River drainage that flowed 

from eastern North Carolina toward 

northwestern Ohio and Indiana (Hansen, 1995; 

Janssen, 1953). Whereas, the Green and 

Cumberland joined their master streams south of 

the glacial margin, the Teays drainage 

downstream of Kentucky was overrun by 

advancing ice sheets and flow was impounded 

south of the glacial margin (Andrews, 2006; 

Teller, 1973). It is difficult to ascertain precisely 

what effect these events had on karst 

development in CCK and how its progression 

may have differed from the other two karst 

systems. 

Nonetheless, we believe that enough 

similarities in cave level sequences, bedrock 

geology, and relative elevations exist between 

the three systems to pose some preliminary 

hypotheses regarding the development of the 

CCK system. Certainly, these hypotheses would 

benefit from future geochronology studies of 

CCK sediments and additional geomorphic field 

work in the study area. The L4 and L3 trunk 

passages in CCK, such as Saltpetre Cave and the 

upper level of Laurel Cave likely correlate with 

the upper cave levels in MCS and CPK and 

represent Pliocene or early Pleistocene karst 

development. These passages contain fine- to 

coarse-grained silt and sand deposits that Engel 

and Engel (2009) suggest are fluvial in origin. 

These sediments may be suitable for cosmogenic 

or paleomagnetic analysis. 

The Ohio River initially occupied its current 

course approximately 1.4 Ma and drove a rapid 

incision event that is attributed to the formation 

of MCS level D (Granger and others, 2001). 

After 1.24 Ma, the incision and aggradation 

history of the Ohio River becomes more 

complicated and Granger and others (2001)

attribute the relative instability of the river level 

to the lack of well-defined levels below level D. 

During this time period in northeastern 

Kentucky, it is possible that Tygarts Creek 

incised at times well into the siliciclastics 

underlying the carbonates—leaving the CCK 

hydrologically abandoned. Whereas, L4 and L3 

cave entrance elevations are tightly distributed 

across narrow distributions, cave entrance 

elevations in L2 and L1 are more broadly 

distributed. Thus, the “noise” in the L2 and L1 

distributions may represent the complex base 

level evolution of the Ohio River drainage over 

the last 1.24 Ma. 

CONCLUSIONS

The number of levels within the CCK shows 

that the area has experienced changes in the 

elevation of the water table. We posit that the 

upper-level trunk passages in the CCK may 

represent the remnants of a more extensive karst 

system that developed in the Plio-Pleistocene 

during a period of relatively slow landscape 

denudation prior to the abandonment of the 

Teays River network and the development of the 

Ohio River drainage. The lower level caves in 

the CCK system likely formed during periods of 

base level stability during the wax and wane of 

the Pleistocene ice sheets. 

We propose that accepted models for the 

Plio-Pleistocene development of the Mammoth 

Cave and Cumberland Plateau karst systems are 

appropriate starting points for deciphering the 

history of cave level development in the CCK 

area. Additional geomorphic analysis of the 

system, including geochronologic analysis of the 

cave sediments, could provide important insight 

into the demise of the Teays drainage and 

development of the modern Ohio River. 

ACKNOWLEDGMENTS

The authors would like to thank the 

Wittenberg University Speleological Society for

providing the karst features database, the 
102

Kentucky Department of Parks for access to 

CCSRP, and Kimberly Gorecki Boland who 

helped to collect the altimeter and GPS control 

point elevations. 

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104

Surface Denudation of the Gypsum Plain, West Texas and 

Southeastern New Mexico

By Melinda G. Shaw, Kevin W. Stafford, and B.P. Tate

Department of Geology, Stephen F. Austin State University, 1901 N. Raguet, Nacogdoches, TX 75962

Abstract

The Castile Formation crops out in Eddy County, New Mexico and Culberson County, Texas, 

covering more than 1500 square kilometers. The region contains more than ten thousand reported surficial 

karst features, including closed depressions, sinking streams and caves, with surface karren dominating 

exposed rock surfaces. A standard gypsum tablet study was conducted over the region for two years to 

delineate the rate of surface denudation across Castile outcrops of the Gypsum Plain. Fifty five sites were 

monitored with triplicate gypsum tabs that were deployed for four month intervals. Rates of surface 

denudation were calculated based on mass loss of gypsum tabs and compared with weather data collected 

at three control locations within the study area. Surface denudation rates exceeded 50 centimeters per 

thousand years (cm/ka) in some areas with an average denudation rate of more than 30 cm/ka. Greatest 

rates of denudation occurred during late summer and fall, in association with the monsoon season, while 

the lowest rates occurred during late spring and early summer. The Castile Formation is bounded by the 

Guadalupe Mountains to the northwest, Delaware Mountains and Apache Mountains to the west and 

southwest, Glass Mountains to the southeast and Rustler Hills to the east; orographic effects of these 

features contribute to a complex pattern of surface denudation associated with shifting seasonal climate 

patterns. True denudation within the gypsum plain varies from models developed from standard tablet 

studies due to variations in surficial deposits; however, standard tablet studies provide a quantitative 

measurement of the rate at which the gypsum plain is evolving.

INTRODUCTION 

Most gypsum and anhydrite rocks originate 

as evaporitic formations within somewhat 

restricted marine basins or epicontinental sea 

environments. The Castile Formation evaporites 

are exposed or shallowly buried over 

approximately 1800 km2 of west Texas and 

southeastern New Mexico. These rocks crop out 

in Culberson County, Texas and Eddy County, 

New Mexico along the western edge of the 

Delaware Basin and are commonly referred to 

as the Gypsum Plain (Hill, 1996) (Figure 1). The 

region is located in the semi-arid southwest on 

the northern edge of the Chihuahuan Desert; 

annual precipitation averages 26.7 cm with the 

greatest rainfall occurring as monsoonal storms 

in the late summer. Annual temperature averages 

17.3°C with an average annual minimum and 

maximum of 9.2°C and 25.2°C respectively 

(Sares, 1984).

The Gypsum Plain within the Delaware 

Basin is an area of low relief 12 – 40 km wide 

and 90 km long and is characterized by shallow, 

eastward-trending valleys and broad, low 

divides. Topographic features of the plain 

include scattered conical and circular hills rising 

15 – 20 meters above the plain, the “Castile 

Buttes”, linear solution subsidence troughs 

trending eastward to northeastward, and 

numerous karst features such as sinkholes, caves 

and dissolution pits (Stafford and others, 2008b). 

Gypsum is one of the most soluble rocks at or 

near the earth’s surface and is one of the few 

earth materials that can transition between its 

hydrated state (gypsum, CaSO4●2H2O) and its 

dehydrated state (anhydrite, CaSO4) with 

relative ease.

The extent of the exposed Castile Formation 

is bounded by the Guadalupe Mountains to the 

northwest, Delaware Mountains and Apache 

Mountains to the west and southwest, Glass 

Mountains to the southeast and Rustler Hills to 

the east (Figure 1). The thickness of the 

formation varies from 0 – 540 meters, depending 

on both depositional and dissolutional factors 

(Hill, 1996). Surface exposures are influenced 

and modified by several major streams; Hay 

Hollow and the Delaware River in the northern 

section, Salt Creek in the central section and 
105

Cottonwood Creek in the southern section 

(Figure 2).

Most karst systems develop within 

carbonate, sulfate or halide rock assemblages, 

with carbonate rocks (limestone and dolostone) 

being the most familiar and well documented. In 

the United States, studies of karst systems 

associated with sulfate rocks (gypsum and

anhydrite - collectively referred to as gypsum 

karst), have primarily focused on geohazards

Figure 1. Regional overview of the Castile Formation 

outcrop area in relation to the Delaware Basin, defined by 

the boundary of the Capitan Reef. Inset shows location of 

expanded region and the Delaware Basin in relation to 

other major seas of the Permian, including: OB –

Orogrande Basin; VB – Val Verde Basin; and MB –

Midland Basin (adapted from Stafford, 2008c).

and engineering potential. Internationally,

gypsum karst has been the focus of extensive 

speleological studies and worldwide some of the 

most exotic and interesting karst features are 

found within these sulfate systems (e.g. Johnson 

and Neal, 2003; Klimchouk, 2007). The high 

solubility of calcium sulfate rocks enables 

development of extensive karst features much 

more rapidly than in carbonate rocks and can be 

correlated to the evolution of karst surface 

features such as sinkholes, arroyos, and caves. 

Karst development within the Gypsum Plain is 

diverse and widespread with over 10,000 

surficial karst features documented for the 

region (Stafford and others, 2008a).

Figure 2. Digital Elevation Model of general study location 

with outline of Castile Formation. Map shows location of 

weather stations (numbered), major streams within the 

study area and geomorphological features adjacent to the 

Gypsum Plain.

GEOLOGIC SETTING

The Permian Basin formed as an area of 

rapid Mississippian-Pennsylvanian subsidence 

in the foreland of the Ouachita foldbelt. During 

this time, the basin was situated close to the 

western margin of Pangaea and lay about 10° 

north of the paleo-equator (Scotese and 

Langford, 1995). The region was part of the 

southern flank of the North American craton 
106

with predominantly carbonate sedimentation. 

During the late Paleozoic, as a result of 

Pennsylvanian Ouachita-Marathon orogeny, the 

Permian basin region was subdivided into a 

series of smaller basins and platforms including 

the Delaware (northwestern), Midland 

(northeastern), Val Verde (southeastern) and 

Marfa (southern) basins (Figure 1) (Scholle, 

2004). As basin partitioning commenced, 

topographical relief between the basins and 

adjacent platforms was accentuated by 

differential sedimentation and regional 

subsidence. Original platform and shelf areas, 

structural highs, accumulated much higher rates 

of sediment deposition than low-lying basins. 

By the end of the Permian, basins that began 

with relatively shallow depths eventually 

evolved to deep basins with water depths in 

excess of 500 meters (Leslie and others, 1996).

Basin filling began in the Pennsylvanian and 

by the mid Permian, the north and east parts of 

the Permian Basin had been infilled with clastic 

sediments. Connection with marine 

environments to the west became restricted or 

closed along the eastern parts of the Permian 

Basin and saline brines began to form (Kirkland 

and others, 2000). As seas continued to regress, 

circulation of briny waters in the Delaware 

Basin became more restricted and the depression 

was filled by the Castile Formation (Anderson 

and others, 1972). The Castile Formation 

originally consisted of approximately 5% 

calcite, 45% anhydrite and 50% halite, but 

migration of Tertiary groundwater along the 

western edge of the basin dissolved and removed 

most halite, resulting in intrastratal brecciation 

(Anderson and others, 1978). These evaporitic 

sediments consist of banded layers of gypsum, 

anhydrite and calcite with small lenses of halite 

and were deposited within a steep-walled basin 

with restricted circulation, coupled to the 

Permian Sea by a narrow channel known as the 

Hovey Channel. As seas regressed, circulation 

of briny waters became more restricted and 

although there may have been some drop in 

basinal water levels, most of the Castile was 

deposited in deep water below the wave base 

(Anderson and others, 1972). Evidence to 

support this conclusion includes the absence of 

shallow water sedimentary structures in most 

intervals and the presence of fine scale 

laminations consisting of alternating anhydrite 

and darker organic matter and calcite. These 

laminations have been interpreted as annual 

varve sequences, although some studies have 

questioned this characterization (Anderson and 

others, 1972). Regardless of origin or 

characterization, darker laminae probably 

represent a periodic influx of fresher water and 

the development of plankton blooms; anhydrite 

layers probably represent more restricted, 

evaporitic condition.

Sedimentation and precipitation of evaporite 

layers continued over a period of 260 Ka and 

coincided closely with termination of reef 

growth around the Delaware Basin margin. 

Closure of the Delaware Basin was most 

probably controlled by tectonic uplift primarily 

on the west side of the basin. This uplift 

continued throughout Ochoan time, causing 

truncation of lower Castile members and nondeposition of upper members on the western side 

of the basin (Horak, 1985).

METHODS

A standard gypsum tablet study (e.g. Nance, 

1993; Klimchouk and others, 1996) was 

conducted over the region for two years to 

delineate the rate of surface denudation across 

Castile outcrops of the Gypsum Plain. Fifty-five 

sites were monitored with triplicate gypsum tabs 

that were deployed for four month intervals. 

Data was collected during the time period 

November, 2005 – November, 2007 to 

determine the rate of surface denudation of the 

Gypsum Plain within the Delaware Basin. Fiftyfive sample locations were selected based on 

geomorphic attributes and land access. Blocks of 

homogeneous gypsum were extracted from a 

quarry site within the southern part study area, 

cored and sliced into discs 5 cm in diameter and 

1 cm thick (Figure 3). Each of the discs were 

abraded with coarse sand to remove rough edges 

and rinsed to remove dust. Discs were placed in 

an oven at 35°C for 48 hours to remove surface 

and interstitial water, then weighed and tagged 

for field placement.
107

Figure 3. Example of typical gypsum tablet used in study.

At each sample location, elevation and 

coordinates were taken using a portable GPS 

unit. Three tabs were placed directly on the 

ground surface at each sample location in areas 

to minimize mechanical erosion approximately 

10 meters apart and left for four months. Upon 

return, new tabs were placed and weathered tabs 

were retrieved. Weathered tabs were cleaned and 

dried in an oven at 35°C for 48 hours, then 

weighed and measured. Surface denudation was 

calculated using the loss of tab volume in 

centimeters per thousand years (cm/kyr).

The four month intervals for data collection 

were divided based upon distinct weather 

patterns within the region: November to March, 

Winter Season; March to July, Summer Season; 

and July to November, Monsoon Season. Three 

weather stations were installed within the study 

area to collect precipitation, temperature, wind 

direction and wind speed to determine the effect 

of those parameters on both dissolution and 

denudation rates of the exposed Castile 

Formation and mechanical erosion associated 

with the seasonal and temporal distribution of 

weather events. Station 1 was installed in the 

northwestern section near the Guadalupe 

Mountains, Station 2 in the central eastern area, 

and Station 3 in the far southeastern part of the 

study area (Figure 2).

RESULTS AND DISCUSSION

Surface denudation rates generated from the 

dataset shows seasonal variation based on the 

availability of moisture. Within each of the 

seasons, higher denudation rates are found in 

proximity to major stream systems: Hay Hollow 

and Delaware River in the north, Salt Creek in 

the central, and Cottonwood Creek in the south. 

Increased atmospheric moisture proximal to 

streams results in daily condensation from 

diurnal temperature variance inducing greater

dissolution than regions distal to surficial water 

supplies. Among the seasons, dissolution rates 

are higher in the Monsoon Season, correlating 

with increased precipitation across the study 

area. Composite figures representing average 

denudation rates for each of the seasons and an 

annual average are found in Figure 4.

Precipitation patterns within the study area 

vary seasonally and spatially: Winter Season is 

generally dry; Monsoon Season moist; and in 

general, higher denudation rates are generally 

reported in the north and can be directly 

correlated with an increase in moisture and 

precipitation. General trends also show higher 

surface denudation rates in the Summer and 

Monsoon Season in proximity to topographic 

features such as the Guadalupe Mountains and 

Delaware Mountains in the northern part of the 

study area. The exception to this would be the 

Winter Season, when freezing temperatures 

could temper liquid precipitation events in the 

north. Cumulative precipitation during the study 

period shows the variance of those trends both 

seasonally and spatially (Figure 5).

As a result of semi-arid climate and little 

relative humidity, diurnal temperature 

fluctuations within the study area tend to be 

greater than those in humid climates (Sares, 

1984). The northern section of the study area 

experienced greater fluctuations in diurnal 

temperatures; these fluctuations, coupled with 

readily available moisture in stream channels 

and springs, produced more effective 

evaporation / condensation cycles and 

accelerated denudation (Figure 6). The annual 

average ranges of temperatures throughout the 

study area were 9.9°C and 27.9°C, respectively. 

Dominant wind and weather patterns generally 

travel from west to east, with a strong 
108

Figure 4. Map of rates of denudation Winter (November – March), Summer (March – July) and 

Monsoon (July – November) for each season and an annual average. Castile Formation is outlined 

in white and study sites for tablet placement in black. Note: Contour intervals are 3cm/kyr.
109

Table 1. Average precipitation rates (cm) within 

seasons and sections of the study area.

southeastern component present during the 

Summer and Monsoon Season (Figure 7). 

Orographic barriers of mountain ranges on the 

northwestern and western edge of the Castile 

Formation and Delaware Basin play a role in 

channeling wind, storm systems and surface 

flow (Figure 2). In addition, although the 

topographic expression within the plain is 

minimal, a correlation can be made between 

elevation and denudation; points of higher 

elevation tend to denude at a greater rate than 

points of lower elevations. 

Winter Season

During the study period, annual average 

precipitation in the Winter Season was 3.1 cm; 

2.7 cm in the north, 3.4 cm in the central region 

and 3.3 cm in the south (Table 1). Temperatures 

during this season ranged from an average low 

of 2.9°C to an average high 

of 17.5°C. Predominant wind 

patterns were generally from 

the west during the Winter 

Season, with minimal input 

from the Gulf Stream (Figure 

7). In the northern section of 

the study area, warm air 

masses from the Chihuahuan 

Desert were channeled 

between the passes in the 

Guadalupe and Delaware 

Mountains, along the front of 

the Guadalupe Mountains. 

This pattern was repeated in 

the southern section, with 

warm desert air channeled 

between the Delaware and 

Apache Mountains. Wind 

patterns in the central section 

were dominated by cooler air 

masses moving over the 

Guadalupe Mountains from the northwest. The 

highest denudation rates are in proximity to 

drainage arteries and associated with moisture 

from the various regional springs that supply 

water to the atmosphere and surface. Overall, 

seasonal denudation rates are lowest within the 

winter, ranging from 5-40 cm, and correlating 

with the low rates of precipitation throughout the 

study area.

Summer Season

Precipitation rates within the Summer 

Season, March to July, show a four-fold increase 

with an annual average precipitation of 12.1 cm; 

17.5cm in the north, 10 cm in the central region 

and 8.9 cm in the south (Table 1). Temperatures 

during this season ranged from an average low 

of 7.7°C to an average high of 27.8°C. Wind 

patterns shifted with each section reporting a 

strong influence from the southeast (Figure 7). 

Warm air still flowed up from the desert, 

channeled between mountain passes in the 

northern and southern sections, with the central 

section still dominated by northwestern air 

masses channeled over the Guadalupe 

Mountains. The highest denudation rates are 

again focused along drainage arteries and 

associated with moisture from the various 

regional springs that supply moisture to the 

atmosphere and surface. Seasonal denudation 

Precipitation Rates (cm)

Season North Central South Average

Winter 2.7 3.4 3.3 3.1

Summer 17.5 10.0 8.9 12.1

Monsoon 27.6 16.4 11.9 18.6

Section 

Total 47.7 29.7 24.1 33.8

Figure 5. Cumulative precipitation (cm) through the duration of the 

study period showing increased precipitation rates moving northward 

across the Gypsum Plain. 
110

Figure 6. Temperature variation within the 

study area showing minor latitudinal variations. 

rates are somewhat higher in the summer and are 

more evenly distributed throughout the study 

area with values ranging from 5-50 cm. The 

difference in spatial distribution can be 

correlated with increasing rates of precipitation 

and humidity, and warmer temperatures 

throughout the study area.

Monsoon Season

The Gypsum Plain receives over 50% of its 

annual precipitation during the Monsoon 

Season. During the study period, annual average 

precipitation was 18.6 cm; 27.6cm in the north, 

16.4 cm in the central region and 11.9 cm in the 

south (Table 1). Temperatures during this season 

ranged from an average low of 8.9°C to an 

average high of 27.2°C. Wind patterns showed 

the greatest shift during this season, with a 

strong N/NE component predominant in the 

northern section of the study area (Figure 7). Air 

masses from the southeast, reported in the 

central and southern section of the study area 

could have redirected eastern moving air masses 

to the north of the Guadalupe Mountains, 

allowing wind patterns to shift in the northern 

section. Higher denudation rates are more evenly 

distributed across the study area, and the 

increase in moisture accelerates denudation 

along stream channels and arroyos. Seasonal 

denudation rates are somewhat higher in the 

summer and are more evenly distributed 

throughout the study area with values ranging 

from 24-80 cm. The difference in spatial 

distribution can be correlated with increasing 

rates of precipitation and humidity, and warmer 

temperatures throughout the study area.

CONCLUSION

Karst systems associated with sulfate rocks, 

both surficial and below surface, host some of 

the most exotic and interesting speleological 

features. Gypsum karst offers a medium for 

geochemical, engineering and petrochemical 

studies, many of which are of interest both 

within the United States and internationally. The 

high solubility of gypsum enables the 

development of extensive karst features much 

more rapidly, creating many opportunities for 

current and future studies of landscape 

evolution.

Within the Castile Formation, surface 

denudation rates vary both seasonally and 

spatially, and are directly correlated with 

available atmospheric and surficial moisture. 

Dissolution during the Winter Season is lowest; 

warm, dry air masses flowing from the 

Chihuahuan Desert reduce available atmospheric 

moisture. During the Summer and Monsoon 
111

Figure 7. Rose diagrams showing seasonal wind direction variations across the study area (n~80,000 per 

sampling period). Note: Rose diagrams are displayed in standard azimuth coordinates with north to the 

top of the page.

Seasons, southeastern winds from the Gulf of 

Mexico bring warm, moist air increasing 

atmospheric moisture, precipitation rates and 

relative humidity, thus inducing greater 

dissolution. Spatially, higher denudation rates 

are found in proximity to major stream systems; 

increased atmospheric moisture proximal to 

streams results in daily condensation from the 

diurnal temperature variance.

Precipitation rates can vary seasonally and 

spatially, during the Winter Season, cool, dry air 

masses from the northwest and southwest bring 

in little moisture, and the lowest precipitation 

rates are reported during this season. Summer 

and Monsoon Seasons report greater 

precipitation rates, with the influence of 

southeastern winds bringing moist, humid air 

from the Gulf of Mexico. Spatially, higher 

denudation rates are generally reported in the 

north and can be directly correlated with an 

increase in moisture and precipitation. The 

orographic affect of the Delaware and 

Guadalupe Mountains can influence moist air 

masses and induce precipitation.

True denudation within the Gypsum Plain 

varies from models developed from standard 

tablet studies due to variations in surficial

deposits; however, standard tablet studies 

provide a quantitative measurement of the rate at 

which the gypsum plain is evolving. The 

evolution of this landscape and variations of 

karst features within the Castile Formation are 

an indication that these features and rates of 

denudation are more complex than can be 

explained by a standard tablet study. Rates of 

incision of stream systems, precipitation, wind 

patterns and exposure of previously buried 

features by surface denudation all contribute to 

the geomorphic evolution of the Gypsum Plain.

ACKNOWLEDGMENTS

The authors are grateful to the numerous 

individuals who made this study possible. 
112

Generous land access to private ranches in Texas 

was provided by Jack Blake, Draper Brantley, 

Stanley Jobe, Lane Sumner and Clay Taylor. 

John Jasper and Jim Goodbar provided 

assistance with karst studies on BLM (Bureau of 

Land Management) public land in New Mexico. 

The authors are thankful for the useful and 

insightful comments provided by Wesley A. 

Brown and Andy Grubbs which helped to 

improve this manuscript. This research was 

partially funded through grants from the New 

Mexico Geological Society (NMGS), the 

American Association of Petroleum Geologists 

(AAPG), the Geological Society of America 

(GSA), with support from the National Cave and 

Karst Research Institute (NCKRI). 

REFERENCES

Anderson, R.Y., Dean, W.E., Kirkland, D.W. and 

Snider, H.I., 1972, Permian Castile varved 

evaporite sequence, West Texas and New 

Mexico: Geological Society of America Bulletin, 

v. 83, p. 59-85.

Anderson, R.Y., Kietzke, K.K., and Rhodes, D.J., 

1978, Development of dissolution breccias, 

northern Delaware Basin and adjacent areas: 

New Mexico Bureau of Mines and Mineral 

Resources Bulletin 159, p. 47-52.

Hill, C.A., 1996, Geology of the Delaware Basin, 

Guadalupe, Apache and Glass Mountains: New 

Mexico and West Texas: Permian Basin Section 

Society of Economic Paleontologists and 

Mineralogists, Midland, Texas, 480 p.

Horak, R.L., 1985, Trans-Pecos tectonism and its 

affect on the Permian Basin, in Dickerson, P.W., 

and Muelberger, W.R., eds., Structure and 

Tectonics of Trans-Pecos Texas: Midland, 

Texas, West Texas Geological Society, p. 81-87.

Johnson, K.S. and Neal, J.T., 2003, Evaporite Karst 

and Engineering/Environmental Problems in the 

United States, Norman, Oklahoma Geological 

Survey: 353 p.

Kirkland, D.W., Rodger, E.D, Dean, W.E., 2000, 

Parent Brine of the Castile Evaporites (Upper 

Permian), Texas and New Mexico, Journal of 

Sedimentary Research, v. 70, no. 3, p. 749-761.

Klimchouk, A., 2007, Hypogene Speleogenesis: 

Hydrogeological and Morphometric Perspective: 

Carlsbad, National Cave and Karst Research 

Institute, Special Paper No. 1, 106 p.

Klimchouk, A., Cucchi, F., Calaforra, J.M., Aksem, 

S., Finocchiaro, F., and Forti, P., 1996, 

Dissolution of gypsum from field observations: 

International Journal of Speleology, v. 25, no. 3-

4, p. 37-48.

Leslie, A.B., Kendall, A.C., Harwood, G.M., Powers, 

D.W, 1996, Conflicting indicators of paleodepth 

during deposition of the Upper Permian Castile 

Formation, Texas and New Mexico, Geological 

Society, London, Special Publications, v. 116, p. 

79-92.

Nance, R., 1993, Application of the standard tablet 

method to a study of denudation in gypsum 

karst, Chosa Draw, southeastern New Mexico, 

MS Thesis, University of Northern Colorado, 

Greely, Colorado, 82 p.

Sares, S.W., 1984, Hydrologic and geomorphic 

development of a low relief evaporite karst 

drainage basin, southeastern New Mexico [M.S. 

Thesis]: Albuquerque, University of New 

Mexico, 123 p.

Scholle, P.A, Goldstein, R.H., and Ulmer-Scholle, 

D.S., 2004, Classic Upper Paleozoic Reefs and 

Bioherms of West Texas and New Mexico, New 

Mexico Institute of Mining and Technology, 

Socorro, NM, 166 p.

Scotese, C.R. and Langford, R., 1995, Pangaea and 

the Paleogeography of the Permian, The Permian 

of Northern Pangaea, vol. 1, p. 3-19.

Stafford, K.W., Rosales-Lagarde, L., and Boston, 

P.J., 2008a, Castile Evaporite Karst Potential 

Map of the Gypsum Plain, Eddy County, New 

Mexico and Culberson County, Texas: A GIS 

Methodological Comparison: Journal of Cave 

and Karst Studies, v. 70, no. 1, p. 35-46.

Stafford, K.W., Nance, R., Rosales-Lagarde, L., and 

Boston, P.J., 2008b, Epigene and hypogene 

gypsum karst manifestations of the Castile 

Formation: Eddy County, New Mexico and 

Culberson County, Texas, USA: International 

Journal of Speleology, v. 37, no. 2, p. 83-98.

Stafford, K.W., Ulmer-Scholle, D.S., RosalesLagarde, L., 2008c, Hypogene Calcitization: 

Evaporite Diagenesis in the western Delaware 

Basin, Carbonates and Evaporites, v. 23, no. 2, 

2008, p. 89-103.
113

KARST PROGRAM UPDATES

National Cave and Karst Research Institute: 

Growing Capabilities and Federal Partnerships

By George Veni

National Cave and Karst Research Institute, 400-1 Cascades Avenue, Carlsbad, NM 88220 USA 

Abstract

The National Cave and Karst Research Institute’s (NCKRI) growth has accelerated since NCKRI’s 

report to the last Karst Interest Group (KIG) meeting in 2008. This is the result of three primary factors. 

First, NCKRI’s staffing has increased. While the addition of an Education Director and Advancement 

Director is numerically small, it is significant for a growing organization and the positions are critical to 

building key NCKRI programs. Second, NCKRI is buying and using significant research equipment. The 

most important purchase to date is an electrical resistivity survey system, which is being applied to solve 

field problems and in a NCKRI research program to refine and enhance the resistivity methodology. 

Third, NCKRI has moved into its newly constructed headquarters. While some rooms await additional 

funds for completion, all of the public spaces are complete and the building is functional for NCKRI’s 

core needs. NCKRI Headquarters provides a synergistic effect with growing staff and equipment by 

opening new opportunities for continued growth. For example, exhibits and classroom/meeting space 

bring in visitors, new partners, and new sources of revenue. Small conferences hosted by NCKRI are 

being scheduled and perhaps NCKRI will host a future KIG meeting.

NCKRI’s federal partners are also growing and strengthening. In March 2009, President Obama 

signed a bill that removed a legislative impediment to cooperative programs between NCKRI and federal 

agencies. In May 2010, NCKRI, Bureau of Land Management, National Park Service, U.S. Fish and 

Wildlife Service, and U.S. Geological Survey gave a joint briefing to the Department of the Interior 

(DOI) on the status of cave and karst research and management challenges in the U.S. One outcome of the 

meeting was an initiative to create a DOI-NCKRI Memorandum of Understanding (MOU). Currently in 

development, this MOU is expected to serve as an umbrella agreement to facilitate NCKRI working with 

and assisting DOI agencies. As the nation struggled with a major economic crisis during the last three 

years, NCKRI has continued to grow and will soon be positioned to launch forward when economic 

prosperity returns.

INTRODUCTION 

The National Cave and Karst Research 

Institute (NCKRI) was created by the U.S.

Congress in 1998 in partnership with the State of 

New Mexico and the City of Carlsbad. Initially 

an institute within the National Park Service,

NCKRI is now a non-profit 501(c)(3) 

corporation that retains its federal, state, and city 

partnerships. Federal and state funding for 

NCKRI is administered by the New Mexico

Institute of Mining and Technology (aka New 

Mexico Tech or NMT). Funds not produced by 

agreements through NMT can be accepted 

directly by NCKRI.

NCKRI’s enabling legislation, the National 

Cave and Karst Research Institute Act of 1998, 

16 U.S.C. §4310, identifies NCKRI’s mission as 

to:

1) further the science of speleology;

2) centralize and standardize speleological 

information;

3) foster interdisciplinary cooperation in 

cave and karst research programs;

4) promote public education;

5) promote national and international 

cooperation in protecting the 

environment for the benefit of cave and 

karst landforms; and
114

6) promote and develop environmentally 

sound and sustainable resource 

management practices.

Boston and Veni (2008) outlined some of 

the challenges that faced NCKRI as a young 

institute that was newly reorganized as a nonprofit corporation, as well as some opportunities. 

This paper reports on NCKRI’s progress over 

the past three years in its programs and federal 

partnerships. Additional details can be found in 

NCKRI (2008, 2009, 2010).

PROGRAM EXPANSION 

NCKRI’s staffing is increasing. The addition 

of an Education Director and Advancement 

Director is numerically small, but significant for 

a small organization with only one part-time and 

six full time employees. Likewise, NCKRI is 

buying critical research equipment and has 

moved into its newly constructed headquarters. 

All of these factors are major steps toward

building NCKRI’s key programs. Status 

summaries on these programs are provided 

below.

Education Program

Dianne Gillespie was hired as NCKRI’s first 

Education Program Director in June 2009. Her 

first 6 months were spent on a coast-to-coast 

research expedition across the U.S. She visited

64 sites including show caves, government 

agencies, universities, museums, parks, and 

private organizations focused on caves, karst, 

and related natural resources science. Through 

her travels she discovered that only 10 locations 

had a cave education specialist working with 

school groups. Additional funding and quality 

educational materials were common needs. 

Many people requested that NCKRI develop 

training for informal educators, focused on 

explaining cave development and karst. Many 

also asked NCKRI to serve as a clearinghouse 

for cave and karst education information.

Ms. Gillespie has since developed an 

Education Plan for NCKRI aimed to reach 

people of all ages and backgrounds. The plan is 

currently being reviewed by the NCKRI Board 

of Directors and approval is anticipated in May 

2011, at which time its details will be released.

NCKRI recognizes that effective education 

requires effective outreach. Education Program 

activities thus also include overhauling 

NCKRI’s website (www.nckri.org), and 

establishing the Institute on Facebook and 

Twitter. Cooperative relationships are also being 

developed with cave education programs at the 

American Cave Conservation Association, 

Carlsbad Caverns National Park, and Ohio State 

University. 

NCKRI is also connecting with people and 

organizations locally to nationally. For example, 

NCKRI sponsors a Boy Scouts of America 

Venture Crew, leads and teaches various cave 

programs, and serves on a federal White Nose 

Syndrome committee. 

Academic Program

Cave and Karst Studies at New Mexico Tech 

is NCKRI’s academic program and is taught 

through NMT’s Earth and Environmental 

Sciences Department. Dr. Penelope Boston is 

the program’s director. She teaches a variety of 

regular and special topic caves and karst 

courses, several in collaboration with other 

NMT faculty, including:

• Cave and Karst Systems

• Cave and Karst Laboratory

• Advanced Topics in Speleohydrology

• Moonmilk Research

• Fundamentals of Geobiology

• Extraterrestrial Dissolutional Landforms

Dr. Boston has supervised numerous Ph.D., 

Masters, and undergraduate senior theses on 

cave and karst research. Recent topics include: 

• Cave micrometeorological modeling 

• Salt Basin Karst Aquifer, Texas and New 

Mexico

• Sulfuric acid caves and sulfur springs of 

Tabasco, Mexico

• Origin of Moonmilk

• Hydrology and paleoclimatology of Snowy

River, Fort Stanton Cave, New Mexico

• Origin and geomicrobiology of fumarolic 

ice caves, Mount Erebus, Antarctica

Advancement Program

NCKRI is currently financed primarily 

through federal and State of New Mexico 

appropriations. These funds are considered seed 
115

monies to establish NCKRI and support its core 

operations. For long-term stability, growth, and 

success, NCKRI must increase and diversify its 

sources of revenue.

On 25 October 2010, Ann Dowdy became 

NCKRI’s first Advancement Director. She 

administers NCKRI’s Advancement Program to 

develop additional funds and support. This 

assistance will be used to build programs and 

partnerships at all levels to enhance the 

Institute’s activities and goals. Fundraising for 

NCKRI is especially challenging and requires 

close coordination with the Education Program. 

Potential donors must first understand what karst 

is, overcome common misconceptions about 

caves, and learn how caves and karst are 

important to daily life.

Ms. Dowdy has been on the job only a short 

time but is already making a big impact. She is 

developing a membership program for NCKRI, 

has written grant requests, and established 

NCKRI on fundraising websites with some large 

fundraising programs and potential donors and 

partners. She has also established a small 

revenue stream through NCKRI supporters who 

load the GoodSearch toolbar on their Internet 

browsers, http://www.goodsearch.com/

toolbar/national-cave-and-karst-researchinstitute. GoodSearch automatically generates 

small corporate donations to NCKRI with each 

Internet search and through much online 

shopping. Future fundraising methods will 

include online giving, annual and corporate 

giving, endowments, bequests, special events, 

and additional appropriation and grant requests.

Research Programs

NCKRI has basic and applied research 

programs but has not yet hired a director for 

either. Nonetheless, substantial research is 

underway.

Dr. Penelope Boston conducts most of 

NCKRI’s basic research, such as:

• National Science Foundation study of 

carbonate pool precipitates;

• Study of iron-manganese deposits in caves

and desert varnish;

• Geomicrobiological investigation of Cueva 

de los Cristales in the Naica Mine, 

Chihuahua, Mexico;

• Evaluating the role of extraterrestrial lava 

tubes on the Moon and Mars in future space 

exploration with partners from the 

Massachusetts Institute of Technology and 

TechShot, Inc.; and

• NASA funded development of an infrared 

tunable acousto-optical laser spectrometer to 

characterize organic components and 

microbial life—ultimately for extraterrestrial

planetary use but being tested on minerals 

that include cave microbial materials.

Dr. Lewis Land, NCKRI’s Research 

Hydrogeologist, also conducts basic 

investigations, such as hydrologic 

characterization of the Roswell Artesian Basin 

and hydrogeologic mapping of the Southern 

Sacramento Mountains watershed in southeast 

New Mexico. Beginning in 2008, his work has 

become more applied.

NCKRI is purchasing high quality 

hydrogeological field research equipment, such 

as flow and multi-meters, to enhance its 

investigative abilities. The Institute’s SuperSting 

R8/IP 112-electrode electrical resistivity meter 

and survey-grade Topcon GR3 global 

positioning system have seen the most demand 

for use.

In 2008, two brine well cavities collapsed in 

the Carlsbad area creating sinkholes over 100 m 

in diameter and 40 m deep. These events raised 

concerns that a similar cavity at a brine well 

within the city limits might collapse, and 

generated interest in high-precision surveys of 

brine well and natural karstic cavities. NCKRI is 

being contracted by the State of New Mexico to 

conduct a resistivity survey of the potential 

collapse site in Carlsbad. In addition to such 

problem-solving geophysical investigations, 

NCKRI has started experiments that may refine 

the resistivity methodology and interpretation of 

its data.

NCKRI is also now conducting contract 

research. This is an important component of 

diversifying the Institute’s sources of income 

and expanding its programs. However, NCKRI 

will not function as a consulting company but in 

a manner akin to the U.S. Geological Survey. 

The Institute will work on contracted projects 

that may produce results with broad application 

that can be published in NCKRI’s and other 

professional literature. While budgetary needs 

may initially see the Institute work on some 

projects that lack broad application, those 
116

projects will nonetheless provide important data 

that NCKRI will include in future studies.

Headquarters Construction

Twenty-seven months to the day after 

groundbreaking, NCKRI staff moved into its 

new headquarters on 24 February 2011. This 

1,609 m2 building contains a bookstore,

museum/exhibit hall, classrooms, laboratory, 

library, offices, and meeting space. NCKRI 

Headquarters is highly energy efficient, with 

innovative design features. Some of these 

features will serve nationally as models for 

green building techniques to prevent or

minimize adverse environmental impacts.

But NCKRI Headquarters is more than just a 

building. It is a true partner in advancing the 

Institute. This beautiful edifice provides:

• NCKRI legitimacy as a national 

organization in the eyes of many in ways 

impossible for our previous assemblage of 

temporary offices;

• Space to easily access our stores of books 

and research materials, as well as for 

volunteers, student interns, and visiting 

scholars to work with us;

• Museum space to better educate the public 

on the value of caves and karst;

• Educational space for classes, workshops, 

and small conferences like the Karst Interest 

Group (which is invited to have its 2014 

meeting at NCKRI); and

• Additional sources of funding through rental 

of rooms and bookstore sales.

An Opening Day Celebration is scheduled 

for 14 May 2011 as a community event that 

recognizes the many local and national partners 

who made the building and NCKRI a reality. 

While planned as a grand event, it will not be a 

“grand opening.” All public spaces in the 

building are finished, but the library, laboratory, 

and associated storage rooms wait for further 

funding. The Grand Opening will occur once the 

building is finished and all of the state-of-the-art 

museum exhibits are installed.

EXPANDING FEDERAL PARTNERSHIPS

The idea of NCKRI was born from the 

national cave and karst coordinators of the 

Bureau of Land Management, National Park 

Service, and U.S. Forest Service, who were all 

stationed in Carlsbad in the 1980s. They first 

envisioned a federal institute dedicated to 

studying and solving cave and karst problems 

for federal agencies. By the time Congress 

authorized NCKRI in 1998, its scope expanded 

to the current mandates.

In 2006, NCKRI reorganized as a nonprofit 

to maximize its flexibility to enter into

partnerships with other entities, raise funds, and

respond quickly to new opportunities. This 

change was a good thing, giving NCKRI the best 

of several worlds: national legitimacy though its 

federal partners and Congressional mandates; 

excellent administrative, legal, and technical 

support through its state partners at New Mexico 

Tech; a wonderful custom built headquarters 

through the City of Carlsbad; and the business 

agility achieved through its non-profit status. 

Yet NCKRI also felt diminished by the loss of 

its full federal status. That feeling wasn’t 

warranted and didn’t last long.

NCKRI’s federal connections have grown 

and strengthened over the past three years. The 

Institute’s Board of Directors has included 

representatives of the Bureau of Land 

Management, National Park Service, U.S. Fish 

and Wildlife Service, U.S. Forest Service, and 

U.S. Geological Survey, each working to 

support NCKRI in many ways.

The National Park Service (NPS) remains 

NCKRI’s primary federal partner. Federal funds 

are channeled to NCKRI, via New Mexico Tech, 

through the NPS’ Geologic Resources Division. 

The NPS provided architects whose attentive 

oversight dramatically improved the design of 

NCKRI Headquarters and assured its quality. 

The Institute’s website is hosted by the NPS and 

an exhibit for NCKRI’s museum was donated by 

Carlsbad Caverns National Park. In 2010, the 

Geologic Resources Division provided NCKRI 

with end-of-year funds to put the finishing 

touches on the public areas of the Institute’s 

Headquarters. NCKRI is also periodically 

consulted by the NPS on cave and karst issues, 

initiatives, and potential collaborative projects.

NCKRI also works regularly with other 

federal bureaus. NCKRI’s Education Director 

serves on a multi-agency committee on White 

Nose Syndrome and will co-teach a joint U.S. 

Forest Service-NCKRI-National Speleological 

Society workshop on cave and karst 

management in June 2011. The Institute is 
117

sponsoring a similar workshop which will be 

offered by the Bureau of Land Management 

hosted by NCKRI at NCKRI Headquarters in 

March 2011. I serve on a U.S. Fish and Wildlife 

Service recovery team for several endangered 

karst invertebrate species. NCKRI also 

continues to assist the U.S. Geological Survey 

whenever possible with the National Karst Map 

and in sponsoring these Karst Interest Group 

meetings.

Successful federal, state, and NCKRI 

teamwork was celebrated in March 2009 when 

President Obama signed a bill into law, 

removing a legislative impediment to 

cooperative programs between NCKRI and 

federal agencies. Building on that, about a year 

later NCKRI, the Bureau of Land Management, 

National Park Service, U.S. Fish and Wildlife 

Service, and U.S. Geological Survey gave a joint 

briefing to the Department of the Interior (DOI) 

on the status of cave and karst research and 

management challenges in the U.S. The meeting 

led to a cooperative project with the Bureau of 

Land Management. More significantly, it also 

started an initiative to create a DOI-NCKRI 

Memorandum of Understanding. The 

memorandum is currently in development, but it 

is expected to serve as an umbrella to facilitate 

NCKRI working with and assisting all DOI

bureaus. 

LOOKING AHEAD 

The U.S. is recovering from a severe 

economic recession that began three years ago. 

While many organizations have scaled back, laid 

off employees, and even gone out of business, 

NCKRI has continued to grow. NCKRI will 

continue further developing its administrative 

foundation for the next couple of years, but will 

now focus on building its education, 

advancement, and research programs. These will 

involve new and synergistic partnerships with 

federal agencies, and should blossom when 

national economic prosperity returns.

REFERENCES

Boston, Penelope J, and Veni, George, 2008, 

National Cave and Karst Research Institute: 

Partner for the USGS; in Kuniansky, Eve, ed., 

U.S. Geological Survey Karst Interest Group 

Proceedings, p. 12-12.

National Cave and Karst Research Institute, 2008, 

National Cave and Karst Research Institute 

Annual Report 2007-2008; National Cave and 

Karst Research Institute, Carlsbad, New Mexico, 

28 p.

National Cave and Karst Research Institute, 2009, 

National Cave and Karst Research Institute 

2008-2009 Annual Report; National Cave and 

Karst Research Institute, Carlsbad, New Mexico, 

28 p.

National Cave and Karst Research Institute, 2010, 

National Cave and Karst Research Institute 

2009-2010 Annual Report; National Cave and 

Karst Research Institute, Carlsbad, New Mexico, 

32 p.
118

The National Karst Map: an Update on its Progress

By David J. Weary and Daniel H. Doctor

U.S. Geological Survey, MS 926A, 12201 Sunrise valley Drive, Reston, VA 20192

Abstract

Over the last several years, the U.S. Geological Survey (USGS) has worked towards compiling a new 

map of karst in the United States in cooperation with the National Cave and Karst Research Institute 

(NCKRI), the National Speleological Society (NSS), and with contributions by various state geological 

surveys. Our goal is to produce a national-scale map that will have utility across the spectrum of karst 

research in the United States. The new National Karst Map will replace previous paper maps with a 

digital, GIS-based product. The geographic extent of the compilation now includes the conterminous 

United States, Alaska, Hawaii, and Puerto Rico. Resolution of the polygon data range from 1:24,000 to 

1:1 million depending on the source data used. This map reflects the presence, at or near the earth’s 

surface, of relatively soluble bedrock (carbonates and evaporites) and late-Tertiary to recent volcanic 

flows representing areas of potential for the occurrence of karst or volcanic pseudokarst. The compilation 

is being refined by input from karst researchers, and experts from state geological surveys. In addition to 

publishing the GIS database, we plan to publish a large paper map, most likely at a scale of 1:5 million. In 

addition, a USGS fact sheet will be published, and provided online, showing the extent of American karst 

lands and briefly explaining the importance and diversity of karst in the United States. This will include a

downloadable small map graphic suitable for illustration of the distribution of America’s karst for 

presentation purposes. Based on our current data, karst in soluble rocks in the United States comprises 

about 13.5 percent of the land area. Of that, about 12.4 percent of America’s area is underlain by 

carbonate rocks, and 2.1 percent by evaporite rocks. These numbers understate the extent of shallow karst 

aquifers, which serve as an important source for drinking water and are often obscured on geologic maps 

by overlying, non-karstic strata or sediments. Similarly, areas with potential for collapse hazard due to 

soluble rocks at depth are also understated. An additional 1.3 percent of the nation is underlain by rocks 

likely to contain volcanic pseudokarst. Future work will entail delineating areas prone to development of 

other types of pseudokarst, such as piping caves in unconsolidated sedimentary units. Estimating the 

extent of these areas is extremely difficult as data on the locations of such features as well as the factors 

contributing to their formation and localization are scarce. 

Integration of additional map layers and characteristics into the map database is ongoing. These 

include the potential extent of buried soluble rocks and paleokarst zones, density of karst features (where 

databases exist), the extent of mineralized areas that may be linked to hypogenic karst processes, and 

framework properties of regional carbonate aquifers. Karst areas will be categorized in terms of their 

geologic and eco-geographic contexts. The map database contains links back to the original geologic 

stratigraphic units selected from the source data (chiefly digital state geologic maps) so that detailed 

lithologic information may be extracted for this purpose. Spatial boundaries of many karst areas align 

reasonably well with Level III ecoregions (using the Omernick Ecoregion System hierarchy) and the 

database will allow for selection and sorting by these ecoregions. We believe that this organization will 

help establish not only geographic references to karst regions, but also supply a framework allowing 

definition of individual regional-scale karst systems characterized by shared genesis and a coherent set of 

characteristics. In addition, this categorization helps organize karst areas into regions with similar 

geology, climate, hydrology, vegetation, and wildlife. It is hoped that these regions will provide 

boundaries useful for multivariate analyses of various questions about the process of karstification and 

development of karst ecosystems. 
119

OZARKS PLATEAU KARST

Karst Hydrogeology of the Southern Ozarks

By Van Brahana

University of Arkansas, Department of Geosciences, Fayetteville, AR 72701 

Abstract

 The regional karst aquifers of the Salem Plateau of the southern Ozarks are unique in many ways 

amongst midcontinent U.S. karst aquifer systems. The Ozark region of Missouri and Arkansas contains 

many more large springs (> 1 m3

/sec) than contiguous Paleozoic aquifers elsewhere, including 3 of the 10 

largest springs in the county. Eight first-magnitude springs which occur in southern Missouri reflect a 

much larger median basin size than spring basins further south in the Springfield Plateau of Arkansas. 

Studies within the past 20 years that have evaluated regional aquifer systems (e.g. Regional Aquifer 

Systems Analysis [RASA] and National Water Quality Assessment [NAWQA]) have tended to lump the 

thick Ordovician and older stratigraphic interval of predominantly dolomite and sandstone into the Ozark 

aquifer. Stratigraphically, this includes the interval from the base of the Chattanooga or equivalent units 

(Ozark confining unit) to the top of the Davis Formation (St. Francois confining unit), approximately 800 

to 900 m thick.

 Generalization is appropriate to a degree for some regional assessment of karst hydrogeology, but for 

many applied problems dealing with scales ranging from site-specific (hundreds to thousands of m2

) to 

basin (km2 to hundreds of km2

) scale, this coarse regionalization is not appropriate. Emerging data

indicate that within this thick interval there are strong, localized karst-flow systems that behave 

independently. Large-scale studies that do not differentiate the entire Ozark aquifer interval into distinct 

functional karst flow zones tend to oversimplify the hydrogeology, with a concomitant loss of information 

and understanding. This paper applies a more detailed scale methodology and evaluates the karst 

attributes of the full stratigraphic range and areal extent of flow distributions in the Ozark aquifer interval, 

and seeks to build a coherent, refined-scale conceptual model of the evolution of the karst hydrogeology 

of the Potosi Formation. 

 Data from numerous municipal wells indicate that the widespread, vuggy porosity, high permeability 

and the predominant lithology of the Potosi Formation resulted from dolomitization and geochemical 

alteration early during burial. Well yield and hydraulic testing of individual units likewise establish the 

Potosi as the most permeable and productive interval within the Ozark aquifer. The concentration of large 

springs along linear trends that coincide with geophysical anomalies suggests that tectonic reactivation of 

basement faults fractured overlying brittle carbonate cover into systematic orthogonal joint sets that 

facilitated vertical flow. Surface faults that appear to truncate the entire hydrostratigraphic section are 

consistent with this interpretation. The broad distribution of endangered organisms such as Amblyopsis

rosae and Cambarus aculabrum and Cambarus setosis also supports this interpretation. Dye tracing over 

distances of several tens of kilometers, water-level mapping, interbasin stream piracy, borehole flowmeter testing, geochemistry provide additional constraints that explain most of the anomalous 

hydrogeologic behavior that cannot be accounted for with a single-layer Ozark aquifer model.
120

Overview of The Nature Conservancy Ozark Karst Program

By Michael E. Slay, Ethan Inlander, and Cory Gallipeau

Ozark Highlands Office, The Nature Conservancy, 38 W Trenton Blvd., Suite 201, Fayetteville, AR 72701

Abstract

The Ozarks contain an underground wilderness of caves, springs and aquifers that over the millennia 

have formed in the carbonate bedrock of the region. Stretching from northern Arkansas, southern 

Missouri, and into eastern Oklahoma, subterranean habitats in this landscape harbor bats, salamanders, 

fish, crustaceans and other invertebrates; including at least 60 species found nowhere else on Earth. The 

porous and fractured nature of karst terrain makes it susceptible to pollution caused by incompatible land 

use, and often the animals that live in the caves are negatively impacted by human visitation and 

vandalism. Because of these threats, The Nature Conservancy created the Ozark Karst Program to work 

with agencies, partners, and landowners to protect these species and habitats. Since 1978, the 

Conservancy has acquired 20 caves, installed or repaired two dozen cave gates, and removed 230 tons of 

garbage near caves or sinkholes. Within the last 10 years, we began a thorough investigation of 

Arkansas’ cave fauna discovering over 15 new species and additional populations of known species. 

Information about these species and habitats was used to reassess and update karst conservation priorities

outlined for Arkansas in the 2003 Ozark Ecoregional Conservation Assessment Plan. We developed a 

Geographic Information System (GIS)-based index model that assesses site specific threat associated with 

57 karst species using threat indicators derived from 25 geospatially available datasets. Threat was 

assessed for 297 karst habitats (caves, springs, or seeps). The results of the threat modeling can be used 

to evaluate conservation priorities at site, species, or community levels. We are currently exploring how 

reserve selection theory may allow us to combine these threat indices with biological information (for 

example; abundance, richness and endemism) to generate new karst conservation focal areas within 

Arkansas.
121

Microbial Effects on Ozark Karst-Water Chemistry—

Medicinal Implications

By John E. Svendsen

Springs of Arkansas, 16101 LaGrande Drive, Suite 101, Little Rock, Arkansas 72223

Abstract

 Many of the springs emerging from the karst regions of the Ozark Plateau have reputed medicinal properties.

Promoters, proprietors and "restored visitors" claim these springs are endowed with unique healing properties and 

that each spring has particular health-giving attributes, yielding different cures and often requiring specific forms of 

usage. During the late 1800s and early 1900s a large number of resorts and health spas were developed at these 

healing springs to provide lodging and services to the many people who visited the springs with expectations of 

being cured through their use. Thousands upon thousands of testimonials suggest that many health-seekers did 

indeed find relief.

 This presentation will examine the role microbial processes may have had in imparting distinct medicinal 

properties to the groundwater giving rise to these healing springs. It has long been recognized that microbial activity 

underlies many geochemical processes that influence groundwater chemistry. Microbes play a critical role in the 

subterranean environments by processing organic matter and facilitating the dissolution of rock and deposition of 

minerals. Without such biogenic influences, water quality in distinct environments would be significant different, 

many of the known minerals on Earth would never have formed, and the planet would lack the vast array of 

chemical diversity needed to sustain complex life forms. The presence of specific microbes and their 

biogeochemical processing within aquifers likely gives rise to unique water-quality attributes that contribute to the 

reputed medicinal virtues of these healing springs. 
122

KARST AQUIFER SYSTEMS

An Integrated Approach to Recharge Area Delineation of Four Caves

in Northern Arkansas and Northeastern Oklahoma

By Jonathan A. Gillip1, Rheannon M. Hart1, and Joel M. Galloway2

1

U.S. Geological Survey Arkansas WSC, 401 Hardin Rd., Little Rock, AR 72211

2

U.S. Geological Survey North Dakota WSC, 821 E. Interstate Ave., Bismark, ND, 58503

Abstract

A study was conducted from 2004 to 2007 by the U.S. Geological Survey in cooperation with the 

U.S. Fish and Wildlife Service to assess the characteristics of the local recharge areas for four caves in 

northern Arkansas and northeastern Oklahoma that provide habitat for a number of unique organisms. An 

integrated approach was used to determine the hydrogeologic characteristics and the extent of the local 

recharge areas for Civil War Cave, Wasson’s Mud Cave, Nesbitt Spring Cave, and January-Stansbury 

Cave. This approach incorporated methods of hydrology, structural geology, geomorphology, and 

geochemistry. Continuous water-level and water-temperature data were collected at each cave for various 

periods to determine recharge characteristics. Field investigations were conducted to determine surficial 

controls affecting the groundwater flow and connections of the groundwater system to land-surface 

processes in each study area. Qualitative groundwater tracing also was conducted at each cave to help 

define the local recharge areas. These independent methods of investigation provided multiple lines of 

evidence for effectively describing the flow patterns in complex hydrologic systems.

Civil War Cave is located near the city of Bentonville in Benton County, Arkansas. Civil War Cave is 

developed in and recharged through the Springfield Plateau aquifer. The daily mean discharge for the

period of study was 0.59 cubic feet per second and ranged from 0.19 to 2.79 cubic feet per second. The 

local recharge area calculated for Civil War Cave was approximately 0.13 to 2.5 square miles using a 

water balance equation and approximately 1.8 to 3.8 square miles using a normalized base-flow method. 

Tracer tests indicate water entered Civil War Cave from across a major topographic divide located to the 

southwest.

Wasson’s Mud Cave is located near the city of Springtown in Benton County, Arkansas. Wasson’s 

Mud Cave is developed in the Springfield Plateau aquifer. No flow was observed although the water level 

did fluctuate. Because there was no detectable flow and tracer tests proved ineffective, the recharge area 

for Wasson’s Mud Cave could not be calculated. The immediate topographic recharge area consists of 

less than 0.01 square miles around the opening of the cave. The location of the cave indicates that the 

cave is associated with the shallow groundwater-system (Springfield Plateau aquifer) within the Flint 

Creek Basin. Therefore, the entire basin potentially contributes recharge to the cave.

Nesbitt Spring Cave is located near the city of Mountain View in Stone County, Arkansas. Nesbitt 

Spring Cave is developed in the Ozark aquifer and is recharged through the Springfield Plateau aquifer. 

The daily mean discharge for the period of study was 4.5 cubic feet per second and ranged from 0.39 to 

70.7 cubic feet per second. The calculated recharge area for Nesbitt Spring Cave using a water-balance 

equation ranged from 0.49 square mile to 4.0 square miles. Tracer tests generally showed a portion of 

water discharging from Nesbitt Spring Cave originates from outside the topographic drainage area.

January-Stansbury Cave is located near the town of Colcord in Delaware County, Oklahoma. 

January-Stansbury Cave is developed in and recharged through the Springfield Plateau aquifer. The daily 

mean discharge for the period of study was 1.0 cubic foot per second and ranged from 0.35 to 8.7 cubic 

feet per second. The calculated recharge area for January-Stansbury Cave using a water-balance equation 

ranged from approximately 0.04 to 0.83 square mile. Tracer tests generally showed water discharging 

from January-Stansbury Cave originates from within the topographic drainage area and from an area 

outside the topographic drainage area to the southwest.
123

Groundwater Piracy in Semi-Arid Karst Terrains

By Ronald T. Green1, F. Paul Bertetti1, and Mariano Hernandez2

1

Southwest Research Institute®, Geosciences and Engineering Division 6220 Culebra, San 

Antonio, TX 78238 

2

School of Natural Resources and the Environment, University of Arizona, Tucson, AZ 85721

Abstract

Adequate characterization of aquifer extent and properties is needed to make informed decisions 

regarding water resource management. Discharge balancing performed as part of a water budget analysis 

provides an opportunity to determine the size and extent of groundwater basins. A comprehensive 

assessment of the hydrogeology of the Edwards-Trinity aquifer was undertaken to determine the extent of 

groundwater basins in the western portion of the Edwards-Trinity aquifer, a karst limestone aquifer 

located in central Texas. Recharge was calculated using baseflow estimates calculated from river 

discharge measurements. Effective recharge rates were calculated by averaging discharge-determined 

recharge over regions that are believed to have uniform infiltration based on similar precipitation, 

geology, soil type, and vegetation. The water-budget analysis corroborated the assessment made using a 

potentiometric-surface map that groundwater basin boundaries extend beyond their overlying surface 

watershed footprint, thereby support the premise that groundwater piracy in the southern portion of the 

Edwards-Trinity aquifer where the saturated aquifer is thickest. 

INTRODUCTION

Adequate characterization of aquifer extent 

and properties is needed to make informed 

decisions regarding water resource management. 

Groundwater divides must be identified to 

accurately designate separate pools or basins 

within regional-scale aquifers as part of this 

characterization. Identifying groundwater 

divides can be problematic, especially when 

sufficient field data are lacking. Groundwater 

divides are typically defined using 

potentiometric maps; however, even with a 

reasonably representative potentiometric map, 

the boundaries of groundwater basins, and 

groundwater divides in particular, are not easily 

determined. For occurrences where 

potentiometric maps are either not available or 

not sufficiently accurate, groundwater basin 

boundaries are commonly approximated by 

vertically extrapolating the overlying surface 

watershed boundaries. This approximation is not 

always valid. For those instances where 

groundwater basin boundaries do not coincide 

with surface watershed boundaries, water 

resource assessments can be misleading, if not 

wrong, unless both surface watershed and 

groundwater basin boundaries are accurately 

characterized. In the case of karst aquifers, 

groundwater basins extending into adjoining 

surface watershed basins is referred to as 

groundwater piracy (White and White, 2001; 

White, 2006). 

Common methods used to identify aquifer 

boundaries and delineate karst aquifer basins 

include hydrogeological mapping, discharge 

balancing (i.e., water-budget analysis) and dye 

tracing (Schindel and others, 1997). This study 

uses discharge balancing to demonstrate 

groundwater piracy in an eight-county portion of 

the Edwards-Trinity aquifer in central Texas. 

The use of discharge balancing demonstrates 

how to identify groundwater divides within a 

regional aquifer for which existing 

hydrogeological data are lacking or relatively 

sparse. Discharge balancing is determined using 

measured discharge for gaining rivers, 

particularly at locations that are proximal to 

river headwaters. This information sheds insight 

on effective recharge, the size of the 

groundwater catchment area that discharges to 

the river, and ultimately, the area and 

approximate location of hydraulic boundaries 

which are interpreted to be groundwater divides.
124

BACKGROUND

The Edwards-Trinity aquifer is a regionalscale karst limestone aquifer. The study area is 

centered over the headwaters for several river 

watersheds. Springs and seeps in the study area 

occur mostly where rivers and streams have 

incised into the aquifer allowing for discharge at 

locations where permeable media overlying a 

confining layer are exposed at the surface. There 

are occasional artesian springs associated with 

local faulting, but most artesian springs are 

located to the south of the study area in the 

Balcones Fault Zone (Brune, 1975). Springs 

offer an important opportunity to act as gauge 

points for the upstream basin to measure key 

components of the water budget (White and 

White, 2001; White, 2006). In the absence of 

spring discharge measurements, river gauges 

downstream from the springs can be used to 

estimate spring discharge. This approximation is 

valid if the distance from the spring to the river 

gauging station is not excessive and if the river 

does not lose significant water in the reach from 

the spring to the gauging station. The use of 

river gauging as a surrogate for spring discharge 

measurement is necessary in the western 

Edwards Plateau because discharge at local 

springs has not been measured on a regular basis 

(Green and Bertetti, 2010).

SURFACE WATERSHED BASINS

The western Edwards Plateau is bisected by 

a surface-water divide between the surface 

watershed basins of the Colorado River and the 

Rio Grande (Figure 1). In general, the 

northeastern half of the western Edwards Plateau 

lies in the Colorado River surface watershed

basin and the southwestern half lies in the Rio 

Grande surface watershed basin. The western 

Edwards Plateau is further divided into several 

small contributing surface watershed basins 

including the Rio Grande-Amistad, Lower 

Pecos, Devils, Middle Colorado-Concho, 

Middle Colorado-Llano, Nueces, and Rio 

Grande-Falcon (Figure 1). 

The study area serves as the headwaters for 

numerous rivers, streams, and their tributaries. 

The most notable of these are the Devils, 

Nueces, Llano, San Saba, and Frio rivers. The 

upper reaches of these rivers typically flow 

intermittently. Perennial flow occurs

downstream, with the exception of rivers and 

streams that lose to the subsurface where they 

flow across recharge zones for aquifers such as 

the Edwards. River gain/loss surveys compiled 

by Slade and others (2002) provide quantifiable 

evidence of where groundwater is discharged 

into rivers (gain) and where rivers discharge to 

the subsurface (loss). These studies indicate that 

virtually all rivers gain in the study area except 

for the lower San Saba River and various 

reaches of the Pecos River.

GROUNDWATER BASINS

Groundwater basins in karst aquifers can 

only be unambiguously delineated using 

techniques such as dye tracing, cave surveying, 

water chemistry sampling and assessment, 

development of potentiometric maps, and the 

full use of local geology (White, 2006). In the 

absence of dye tracer, water chemistry, and cave 

survey information, an existing potentiometric 

map of the Edwards-Trinity aquifer is used to 

characterize groundwater basins in the study 

area. In this case, the groundwater basins for the 

study area are approximated in the same way 

that a surface water basin is determined, except 

that the potentiometric surface is used to locate 

groundwater catchment divides. 

The potentiometric surface map prepared by 

Kuniansky (1990) and reproduced by Barker and 

Ardis (1996) is used to delineate the 

groundwater basin in the western Edwards 

Plateau (Figure 2). The surface-water divide 

between the Colorado River and the Rio Grande 

watersheds is overlain on the potentiometric 

surface to compare the extent of the groundwater 

basins to the surface-water basins. Although the 

resolution of the map is limited, areas in Upton, 

Reagan, Sutton, Schleicher, and Edwards 

counties exhibit groundwater flow toward the 

Rio Grande from locations within the Colorado 

River surface watershed basin. 

RIVER DISCHARGE ANALYSIS

River discharge was analyzed as a means to 

corroborate groundwater basin delineation based 

on the 1990 potentiometric surface map. Water 

budgets for individual surface watersheds within 

the study area were analyzed as a means to 

balance discharge. The water budget for each of 

these sub-areas was assessed using available 

hydrogoelogical information. In the absence of 

spring discharge measurements, river gauging 
125

measurements were used to estimate discharge 

from the groundwater catchment areas. 

Fortunately, rivers in the study area have long 

records of discharge measurements.

River discharge has two principal flow 

components, baseflow and surface runoff. 

Baseflow is considered to be the groundwater 

contribution to stream flow and is interpreted to 

equal recharge (Arnold and others, 1995; White 

and White, 2001; White, 2006). Baseflow 

recession is the rate at which the stream flow 

diminishes in the absence of recharge. The 

discharged volume is equated to the amount of 

recharge to the aquifer that discharges to the 

river. 

Decomposition of stream hydrographs 

provides estimates of the baseflow and storm 

flow components to stream discharge. The 

fraction of river discharge attributed to baseflow 

was calculated for each river gauging station 

analyzed in this study. An automated baseflow 

separation and recession analysis tool, 

BASEFLOW, was used to estimate the amount 

of stream flow attributed to baseflow for each 

sub-area (Arnold and others, 1995; Arnold and 

Allen, 1999). The automated procedure predicts 

baseflow recession from the point on the 

hydrograph where it is assumed that all surface 

flow has ceased. For those instances when there 

were gaps in the gauging station time series, 

recession was analyzed for each individual time 

series segment. A single value for baseflow 

fraction for each river watershed was calculated 

by averaging the baseflow fraction for each time 

segment weighted by the number of days in the 

time series subset (Green and Bertetti, 2010). 

Baseflow fractions of river discharge for the 

watersheds in the study area are summarized in 

Table 1. 

River gauging station baseflow analyses

were used to estimate recharge to their 

respective watersheds. For example, an average 

discharge of 125.5 cfs (90,860 acre-feet/year) 

was measured at the Frio River at Concan 

gauging station from 1923 to 2010. The 

baseflow fraction for the Frio River was 

calculated at 0.75, thus the baseflow component 

to flow is 94 cfs. The surface drainage area 

upstream of Concan is measured by the U.S. 

Geological Survey to be 383 mi2 (245,120 

acres). This equates to a recharge rate of 3.36 

inch/year if uniformly averaged over the 383 mi2

drainage area. Similar analyses were performed 

for all watersheds in the western Edwards 

Plateau that had river gauging stations with 

sufficient discharge measurements. Recharge 

calculations for these sub-areas are summarized 

in Table 1 and plotted at their drainage area 

locations in Figure 3.

There is a notable contrast in calculated 

recharge rates between the headwaters of the 

Nueces and Frio river watersheds and the Llano 

River watershed. The recharge rates for the 

Llano River and its north-flowing tributaries are 

significantly less than those for the southflowing Nueces and Frio river watersheds and 

the east-flowing Guadalupe and Medina river 

watersheds. Inspection of precipitation measured 

for the western Edwards Plateau (Figure 4) 

suggests that this significant difference in 

calculated recharge rates between the north and 

south cannot be attributed to variations in 

precipitation observed across the eastern portion 

of the study area because isohyets trend in a 

north-south direction. 

Significant differences in calculated 

recharge rates between northern and southern 

portions of the study area are interpreted to 

indicate that groundwater catchment areas for 

the Nueces, Frio, Medina, and Guadalupe rivers 

extend farther north than the boundaries of the 

overlying surface watersheds. This feature is 

indicative of groundwater piracy. The

interpretation is consistent with the combined 

surface watershed and groundwater flow map 

that indicates that groundwater catchment 

boundaries do not coincide with surface 

watershed boundaries (Figure 2). Groundwater 

piracy from watersheds north of the Rio 

Grande/Colorado River watershed divide results 

in lower calculated recharge rates for the 

watersheds of the Llano River and its tributaries, 

and higher calculated recharge rates for the 

watersheds to the south and east (i.e., Nueces, 

Frio, Medina, and Guadalupe river watersheds). 

Actual recharge rates are thus greater than the 

low values to the north and east of the Rio 

Grande/Colorado River watershed divide and 

less than the high values to the south and west.

Groundwater piracy is more likely to occur 

where the affected aquifer is thick and deep. The 

Edwards-Trinity aquifer is significantly thicker 
126

in Sutton, Edwards, Real, Val Verde, and 

southern Kimble and Schleicher counties 

relative to the Edwards-Trinity aquifer to the 

north (Figure 5). Examples of areas from which 

groundwater appears to be pirated are basins 

associated with the North Llano River at 

Junction, the Llano River at Junction, and the 

West Nueces River at Brackettville. Areas which 

gain from pirated groundwater will have greater 

discharge into rivers and streams and higher 

rates of calculated recharge. Examples include

basins associated with the Frio River at Concan, 

the Nueces River at Laguna, Johnson Creek at 

Ingram, and the Medina River at Bandera.

The extent to which groundwater is pirated 

from one surface watershed to another is 

difficult to measure, but can be estimated using 

discharge analyses. Factors that influence 

recharge (i.e., soil and vegetation type, 

topography) are relatively uniform over the 

areas where groundwater piracy is thought to 

occur. The only variable is precipitation. Using 

this reasoning, recharge rates were averaged 

along lines of equal precipitation (isohyetals) to 

estimate a uniform rate of recharge for areas of 

uniform precipitation. Because the isohyetals are 

essentially north trending in study area, uniform 

recharge rates should also be north trending.

Recharge was averaged between the gauging 

stations on the Frio River at Concan and on the 

South Llano River at Junction and between the 

gauging stations on the Nueces River at Laguna 

and on the North Llano River at Junction (Figure 

3). The averages were weight averaged relative 

to area to conserve mass. In this manner, an 

average recharge rate of 1.39 inch/year was 

assigned to the combined watersheds of the 

upper Frio River and the South Llano River. The 

combined basins have a total area of 1,468,000 

acres and total recharge of 170,500 acre-ft. In 

this calculation, it is assumed that the area of the 

combined groundwater basins is the same area 

as the area of the combined surface watershed. 

The only differences are the locations of the 

mutual boundaries of the two systems.

Similarly, an average recharge rate of 1.21 

inch/year was assigned to the combined 

watersheds of the upper Nueces River and the 

North Llano River. These combined basins have 

a total area of 1,057,000 acres and total recharge 

of 107,000 acre-ft. It is also assumed that the 

external boundary of the combined watersheds 

for the upper Nueces River and the North Llano 

River is the same as the boundary of the 

combined groundwater basins.

Final recharge rates were calculated for the 

eight counties in the study by taking an areal 

average of all recharge rates calculated within 

that county. Recharge rates for each of the eight 

counties in the study area are plotted in Figure 6.

The recharge rates are relatively uniform within 

areas of uniform precipitation, with the 

exception of Menard County. The reason for this 

discrepancy is that the San Saba River, which 

runs through the middle of Menard County, is a 

losing river. Therefore, the basic premise for the 

river discharge computation method used 

elsewhere in the study area is violated. Instead 

of a recharge rate of 0.50 inches/year as 

indicated in Figure 6, the actual recharge rate 

should be closer to 1.30 inch/year if the 

assumption of uniform recharge over areas of 

uniform precipitation is valid.

The extent to which groundwater basins 

associated with the Frio and Nueces rivers 

extend north beyond the surface-water divide 

between the Rio Grande and the Colorado River 

is interpreted to be influenced by the saturated 

thickness of the Edwards-Trinity aquifer. The 

surface-water divide is overlain on a contour 

map of the saturated thickness of the EdwardsTrinity aquifer in Figure 5. As illustrated, the 

greatest potential for groundwater piracy in the 

study area is northern Edwards, eastern Sutton, 

most of Kimble, eastern Schleicher, and Kerr 

counties because the Edwards-Trinity aquifer is 

locally thicker than in northerly areas. The limit 

to which groundwater piracy extends to the 

north can only be definitively determined with 

field verification such as dye tracer results and 

water chemistry analysis. In the absence of field 

verification, the area of potential groundwater 

piracy is estimated by the 200-ft saturated 

thickness contour in the Edwards-Trinity aquifer

in Figure 5. 

SUMMARY

A hydrogeological assessment was 

undertaken to approximate the boundaries of 

groundwater basins in the western portion of the 

Edwards-Trinity aquifer, a karst limestone 

aquifer located in central Texas. Eight counties 

were included in the study; Crockett, Edwards, 
127

Kimble, Menard, Real, Schleicher, Sutton, and 

Val Verde. The assessment relied on waterbudget analyses of hydrologically distinct subareas in the study area. Recharge was calculated 

using baseflow estimates calculated from river 

discharge measurements. Effective recharge 

rates were calculated by averaging dischargedetermined recharge over regions likely to have 

uniform recharge owing to similar geology, soil 

type, vegetation, and precipitation. This average 

was weighted by area to conserve mass. The 

final step in the process was to average effective 

recharge rates over each of the eight counties in 

the study area. These calculations provide 

corroborating evidence that groundwater basins 

extend beyond the extent of their overlying 

surface watersheds, thereby supporting the 

premise of groundwater piracy in the southern 

portion of the Edwards-Trinity aquifer where the 

saturated aquifer is thickest. 

ACKNOWLEDGMENTS

The authors thank Sutton County 

Groundwater Conservation District for 

permission to use the information in this paper. 

Reviews by Gary Walter, Stuart Stothoff, and 

Van Brahana improved the quality of the paper.

Figure 1. Map of the counties (black labels) and nine watershed basins (blue labels) of the major rivers of the study area 

(outlined in red). The blue line denotes the watershed divide between the Rio Grande on the southwest and the Colorado River on 

the northeast.

Figure 2. Map of the Rio Grande-Colorado River surface-water divide overlying the groundwater potentiometric surface. Blue 

arrows are added to denote the direction of groundwater flow assuming isotropic flow. Base map is from Barker and Ardis 

(1996).
128

Figure 3. Locations of the U.S. Geological Survey gauging stations in and near the study area. Numbers in parentheses denote 

recharge rates in inches/year calculated using river discharge rates and corrected for base flow.

Figure 4. Map showing average annual precipitation (inch/year) for the study area (outlined in blue).
129

Figure 5. Contour map of the saturated thickness of the Edwards-Trinity aquifer (Kuniansky and Holligan, 1994) with the 

surface-water divide separating the Rio Grande watershed from the Colorado River watershed (blue line) and the extent of 

groundwater piracy estimated using the 200-ft saturated thickness contour of the Edwards-Trinity aquifer (green line).

Figure 6. Recharge rates (inch/year) by county calculated for river watershed basins. Thick blue lines denote contours for average 

annual precipitation (inch/year).
130

Table 1. List of U.S. Geological Survey river gauging stations used in the water budget analysis. Adjusted discharge 

rates have been corrected for baseflow. All stations are in Texas.

Station 

Number

River Drainage 

Area (acre)

Gauging 

Start Date

(month/yr)

Baseflow 

Fraction

Gross 

Discharge 

(inch/yr)

Adjusted

Discharge

(inch/yr)

08190000 Nueces River at 

Laguna 471,680 10/1923 0.71 3.03 2.15

08190500 West Nueces 

River near 

Brackettville

444,160 9/1939 0.25 0.67 0.17

08195000 Frio River at 

Concan 284,960 10/1923 0.75 4.45 3.36

08178880 Medina River at 

Bandera 209,920 10/1982 0.68 6.77 4.54

08166000 Johnson Creek 

at Ingram 72,960 10/1941 0.62 3.10 1.92

08145000 Brady Creek at 

Brady 376,320 6/1939 0.29 0.15 0.04

08144600 San Saba River 

near Brady 1,045,120 7/1979 0.58 0.59 0.34

08144500 San Saba River 

at Menard 721,920 10/1915 0.48 0.72 0.35

08150800 Beaver Creek 

near Mason 137,600 8/1963 0.42 1.20 0.50

08148500

North Llano 

River at 

Junction

584,960 10/1915 0.46 1.00 0.46

08150000 Llano River at 

Junction 1,183,360 10/1915 0.64 1.45 0.92

08128000

South Concho 

River at 

Christoval

226,560 3/1930 0.60 1.15 0.69

08130500 Dove Creek at 

Knickerbocker 139,520 10/1960 0.76 0.98 0.74

08128400

Middle Concho 

River at 

Tankersley

714,240 4/1961 0.21 0.17 0.04

08449400

Devils River at 

Pafford 

Crossing

2,535,040 1/1960 0.76 1.25 0.95

08412500 Pecos River 

near Orla 13,586,560 10/1937 0.66 0.08 0.05

08446500 Pecos River 

near Girvin 18,918,400 9/1939 0.77 0.03 0.02

08447000 Pecos River 

near Sheffield 20,384,000 10/1975 0.79 0.09 0.07

08447410 Pecos River 

near Langtry 28,352,000 1/1/1967 0.74 0.08 0.06
131

Table 2. Recharge rates calculated by county for the Edwards-Trinity aquifer

Crockett Edwards Kimble Menard Real Schleicher Sutton Val Verde

Area (acres) 1,796,480 1,356,800 800,640 576,640 448,000 839,040 929,920 2,068,480

Recharge 

rate (in/yr) 0.25 1.30 1.50 0.50 2.14 0.80 1.00 0.63

Recharge 

(acre-ft/yr) 37,427 146,987 100,080 24,027 79,893 55,936 77,493 108,595

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White, W.B. 2006. Fifty years of karst hydrology and 

hydrogeology: 1953–2003. In: Harmon R.S., 

Wicks C. (eds) Perspectives on karst 

geomorphology, hydrology, and geochemistry: a 

tribute volume to Derek C. Ford and William B. 

White. Geol Soc Am Spec Pap 404:139–152.

White, W.B. and E.L. White. 2001. Conduit 

fragmentation, cave patterns, and localization of 

karst ground water basins: The Appalachians as a 

test case: Theoretical and Applied Karstology, v. 

13-14, p. 9-23.
132

Ten Relevant Karst Hydrogeologic Insights Gained from 15 Years

of In Situ Field Studies at the Savoy Experimental Watershed

By Van Brahana

University of Arkansas, Department of Geosciences, Fayetteville, AR 72701 

Abstract

 The Savoy Experimental Watershed (SEW) was established nearly 15 years ago, and during that 

time, has been the focus of applied research, education, and outreach dealing with the practical problems 

of humans living on a mantled-karst landscape. This report is a synthesis of ten important insights that 

have been gained during the life of this facility. Although the selection of ten is arbitrary, and much 

additional information has been learned, these encompass illustrate the wide range of understanding that 

can be gained from a long-term research site. Each is documented with specific examples and references 

that will hopefully provide understanding applicable far beyond SEW. These insights are: 1) the purity 

of the lithology controls the nature of the karst that develops; 2) structural control of brittle-carbonate 

units is a major determinant of sub-basin boundaries; 3) divergent flowpaths in karst tracing experiments 

are common; 4) groundwater flow directions commonly are a function of groundwater levels, changing 

as water levels in the aquifer change; 5) groundwater quality in the shallow aquifers is dominated by 

mixing; 6) evapotransporative losses from the epikarst are by far the most dominant of the hydrologic 

budget outflows; 7) flow zones near to one another are frequently hydrologically isolated from each other; 

8) the environments and ecosystems of discrete groundwater flow zones exhibit surprising diversity and 

distribution of organisms; 9) transport of suspended particles and pathogens is a major water-quality 

concern driven by storm pulses; and 10) the actual complexity of seemingly simple karst systems makes 

prediction of hydrogeologic response outside the range of conditions for which conceptual models are 

developed a risky task.

INTRODUCTION

 The Savoy Experimental Watershed (SEW) 

was established nearly 15 years ago, and during 

that time, has been the focus of applied research, 

education, and outreach dealing with the 

practical problems of humans living on a 

mantled-karst landscape. Twenty theses and 

dissertations (table 1) and a comparable number 

of technical papers (selected references) have 

been completed describing the karst 

hydrogeology, groundwater transport 

phenomena, geochemical processing of 

nutrients, hydrologic budgets, animal production 

impacts on karst waters, and development of 

tools and techniques appropriate to karst studies 

in general. Numerous classes from the 

Hydrogeology Program at the University of 

Arkansas have supplemented the research 

infrastructure at SEW, and the site currently has 

more than 100 permanent installations including 

weirs, flumes, wells, piezometers, lysimeters, 

rain gages, geophysical arrays, and a weather 

station.

 This report is a synthesis of the major insights 

that have been gained during the life of this 

facility, and although the selection of ten 

is arbitrary, each is documented with specific 

examples and references providing 

understanding that will be applicable far beyond 

SEW. The objective of this paper is to use 

insights gained in these years of targeted 

research to characterize the nature of mantled 

karst. The scope of this report is limited to 

recent studies at the site and to the mantled karst 

of the southern Ozarks.

 The karst of the SEW is not immediately 

apparent to most people. The setting is a 

dissected plateau with steep, dry valleys and few 

sinkholes. In most of the area, regolith covers 

the bedrock, leaving a thin, rocky soil that masks 

the carbonate bedrock beneath. SEW is typical 

of the Springfield Plateau province of the Ozark 

highlands. Most of the site (>70%) is covered in 

second- and third-growth forests, and the 

remaining land is in pasture. The karst 

groundwater system at SEW is underdrained by
133

 Table 1. Unpublished theses and dissertations focusing on the Savoy Experimental Watershed.

 [Full reference annotation is provided in the list of selected references; M.S., Masters of Science; Ph.D., Doctor of Philosophy; GEOL, Geology; 

CVEG, Civil Engineering; SOIL, Soils; GEOG, Geography; CENG, Chemical Engineering; ENDY, Environmental Dynamics; BIOL, Biology]

Author Date Degree Program Focus of Study

Al Qinna 2004 Ph.D. SOILS Soil water and solute transport in karst 

Al Rashidy 1999 M.S. GEOL Hydrogeologic characterization of basin 1; karst

Curtis

Dixon

Hamilton

Hobza

Laincz

Leh

Little

Parse

Pennington

Phelan

Stanton

Ting

Ting

Unger 

Wagner

Whitsett

Winston

Woodstrom

2000

2001

2001

2005

2011

2008

2000

1995

2010

1999

1993

2002

2005

2004

2007

2002

2006

1999

Ph.D. 

Ph.D.

M.S.

M.S.

Ph.D.

M.S.

M.S.

M.S.

M.S.

M.S.

M.S.

M.S.

Ph.D.

M.S.

M.S.

M.S.

M.S.

M.S.

CVEG

SOILS

GEOL

GEOL

ENDY

BAEG

GEOL

GEOG

GEOL

GEOG

GEOL

CENG

CENG

GEOL

GEOL

GEOL

BIOL

GEOG

Advective transport model encorporating lineations

Neuro-fuzzy modeling to predict groundwater vulnerability

Survival of E.coli in stream and spring sediments

Swine lagoon impact on karst water quality

Nitrogen transport and cycling in karst interflow zone

Rainfall-runoff mechanisms in pasture-dominated watershed

Dominant processes; flow and quality; basin 2

Geomorphic assessment of role of regolith in karst

Drainage basin analysis using hydrograph decomposition

GIS and 3-D visualization; static modeling; mantled karst

Processes and controls affecting Boone-St. Joe aquifer

Europium-tagged E. coli as karst tracer

Bacterial transport and storage using lanthanide tracers

Structural controls; boundaries; flow modeling

Assessment of waste-storage effectiveness using isotopes

Sediment and bacterial tracing in mantled karst

Nitrogen sources and processing in aquatic ecosystems

Land use effects on spring water quality

chert-rich carbonate-rock aquifers that have been 

selectively dissolved to form an open network of 

caves, enlarged fractures, bedding planes, 

conduits, sinking streams, and springs. Flow in 

these aquifers is typically rapid, flow directions 

are difficult to predict, interaction between 

surface and groundwater is typically extensive, 

and processes of contaminant attenuation that 

characterize many other groundwater settings 

are typically absent (White, 1988; Ford and 

Williams, 2007; Palmer, 2007). 

DISCUSSION

Insight 1—Lithologic Purity of 

Carbonates Controls Developing Karst

 Field reconnaissance by numerous researchers 

(Stanton, 1993; Fanning, 1994; Brahana, 1995; 

Al Rashidy, 1999; Little, 2000) at SEW and 

throughout the Springfield Plateau indicates the 

best cave development occurs in rocks that are 

fairly pure carbonates. The St. Joe, Pitkin, and 

upper Boone host some of the most karstified 

rocks in the Ozarks of northwest Arkansas, a 

fact reflected in not only in cavern development, 

but in bedding-plane anastomoses, enlarged 

joints, in driller’s reports, and in wireline 

geophysical logs. The bulk of the middle and 

lower Boone Formation, on the other hand, 

typically has as much as 70% chert and 

insolubles. Whereas caves in the Boone 

Formation are common, their size and form 

typically are small and limited, “pocket caves” 

as designated by Fanning (1994). The chert and 

clay plug and occlude nascent porosity and 

permeability in Boone carbonate. After floods 

and storms recharge the shallow karst aquifers, 

flow recedes, velocities decrease, and the 

suspended sediment load is deposited. Zones 

previously most permeable become flow 

restrictions and dams. The Illinois River graben 

that forms the western boundary of basin 1 at 

SEW (figure 1) provides an example. This fault 

is identified by displacement of the St. JoeChattanooga contact, an escarpment and the 

appearance of springs along its base at high 

flow, and a zone of low permeability. 

Observations indicate that the fault serves as a 

groundwater dam localizing the occurrence of 

Langle and Copperhead Springs. The impurities 

in the Boone source the weathering products of 

the regolith. This regolith is responsible for 

slow release of recharge that sustains epikarst 
134 

springs and the low flow of base-level springs, 

and is the reason that typical karst landforms 

such as dolines, karst windows, ponors, uvalas, 

and estavelles are not visible here.

 Continuous chert layers in the lower Boone 

Formation serve as confining layers, perching 

recharge from above (figure 2), and confining 

phreatic flow moving upward in joints within the 

underlying St. Joe Formation. The major role 

played by the chert layers is well-displayed at 

the two major drains from basin 1. Both Langle 

and Copperhead Springs resurge from small 

caves at the base of the lowermost continuous 

chert layer in the Boone Formation (Al Rashidy, 

1999). Downcutting of the Illinois River has 

allowed vadose entrenchment of these two baselevel spring systems.

Insight 2—Structural Deformation of 

Brittle Carbonate Rocks Defines Flow 

Boundaries and Flow Routes

 Structural geology and tectonic setting are 

well-known controls on groundwater in many 

karst areas (Brahana et al., 1988; Palmer, 2007) 

Nearly flat-lying formations of the Springfield 

Plateau in the southern Ozarks are thought to 

have been significantly impacted by reactivated 

basement faulting associated with the Ouachita 

orogeny. Although SEW lies more than 150 km 

north of the Ouachita Mountains, the evidence 

that basement faults were active tectonic forces 

in early Mississippian time is reflected in 

exposures of olistoliths within 5 km of SEW 

(Chandler, 2001). Utilizing borehole 

geophysical logs , field reconnaissance, and 

surveying, Unger (2004) delineated the top of 

the Chattanooga Shale.

1

Table footnote

Figure 1. Boundary of the Savoy Experimental Watershed, showing delineation of surface watersheds 

numbered 1-6, major springs within or draining the site, and major stream valleys. These alluvial valleys 

only have flowing water after intense precipitation; they are normally dry. 

1

2

3

4

5

6
135

Figure 2. Continuous chert layers perch flow and 

facilitate dissolution of the soluble limestone 

directly above the chert at SEW.

This provided an estimate of dip of the lower 

confining unit, which, like a tilted, 

impermeable block, controls flow directions 

toward springs (Bolyard, 2007). Tracing 

experiments coupled with wireline geophysical 

logs documented the tilted-block model which 

has allowed more consistent delineation of 

spring basins (Brahana, 1997; Ting, 2002; 

Ting, 2005). Stanton (1993) likewise showed 

the strong structural control of the SEW and the 

region with a transect of wireline geophysical 

logs across Clear Creek Valley, about 5 km east 

of SEW. These logs identified a previously 

unmapped graben with a downward 

displacement of about 30 meters in the valley. 

 Cavers in the region recognize that vertical 

shafts and pits are much more dominant in the 

area of the Buffalo National River (Tennyson et 

al., 2008), where the St. Joe Formation occurs 

about 70 meters higher in that area than it does 

at SEW. This is consistent with uplift along 

reactivated basement faults, delineated from 

geologic mapping and from gravity studies. 

The movement on the basement faults is 

thought to have created strongly orthogonal 

joint sets in the overlying, brittle carbonate 

rocks such as the St. Joe Formation (figure 3). 

Cave passages in this part of the Ozarks are 

strongly joint-controlled. Sinking streams such 

as Spanker and Spavinaw Creeks in northwest 

Arkansas are typically dry downstream from 

the intersection of basement fault with the 

streambed. The implications of this strong 

structural control include the need for careful 

design of water-quality sampling tasks in order 

to gather representative data, which in most 

cases means the springs.

Insight 3—Divergent Pathways Are 

Typical in Karst Tracing Experiments

Surface streams in the Ozarks typically are 

dendritic, yet dye tracing in karst aquifers 

commonly indicates dispersive flow—with one 

input point, and multiple recovery points (Aley, 

2002). Of the more than 25 tracing experiments 

conducted at SEW, it has been uncommon to 

have single recovery points. 

Figure 3. Orthogonal joints in brittle carbonate rocks 

serve as pathways for shallow groundwater flow. 

This site in the overflow channel for Crystal Lake, 

about 25 km north of SEW, developed as a subsoil 

karst that was subsequently removed by surfacewater erosion diverted into the overflow channel. 
136

The apparent explanation for this is based on 

multiple lines of evidence. This interconnected, 

multi-transmissive suite of flow routes (fast-flow 

along joints, bedding planes, and conduits) 

typically operates under confinement in the 

phreatic zone. The dynamic nature of pointsource recharge inputs, the turbulent mixing 

along fast-flow pathways, and the remarkably 

large transmissivities of the interconnected voids

establishes aquifer flow that can mimic a system 

of multiple and integrated estavelles, pipes, and 

parallel plates. High permeability allows flow to 

follow minor head variations, with the result that 

if voids are connected, the tracer will follow the 

path of least resistance and disperse. In 

groundwater systems, unlike unconfined flow in 

a stream channel at land surface, threedimensional confinement along the flowpath 

creates the necessary head variations. 

Depending on the outlet architecture of the 

recovery points and the hydrologic conditions at 

the time of the test, dispersive flow is the norm 

rather than the exception. 

Figure 4. Tracer studies at SEW include fluorescent 

dyes (such as Rhodamine WT shown here), 

conservative halogen solutes, lanthanum-series 

tagged clays, europium-tagged E. coli, and thermal 

pulses. 

Insight 4—Karst Groundwater Flow 

Paths Are Commonly Related to Water 

Levels—They Change with Stage and 

Their Resurgences May Be Covered

 Dye tracing and continuous flow monitoring

at weirs under varying hydrologic conditions has 

shown that spring-basin size is related to 

groundwater levels in basin 1. Langle and 

Copperhead Springs are known to be outlets 

from the same base level spring system; Langle 

is the underflow spring, with a slightly lower 

outlet elevation (~ 2 cm lower based on totalstation surveying). Copperhead is the overflow 

spring, with greater permeability at higher 

elevations. For both springs, fast-flow 

transmissivity is maximum along a set of 

systematic orthogonal joints in the St. Joe 

Formation. Langle Spring has higher base flow 

owing to its lower resurgence point, but 

Copperhead Spring has greater flood flow 

because of wider openings at higher elevations 

(figure 5). Also of relevance is the large number 

of dye traces undertaken that have not produced 

positive results with input locations close to 

these major resurgences. We now believe that 

buried spring outlets resurge into the Illinois 

River alluvium, and that our input points were 

isolated from the main flow paths by continuous 

chert layers. Ongoing tracer testing is 

evaluating this hypothesis.

 

Figure 5. Conceptual model of the LangleCopperhead Spring system reflecting differences in 

flow between springs under different water levels.
137

Insight 5—Groundwater Quality in 

Shallow Karst Aquifers Is Dominated 

by Mixing

 Geochemical processes documented in the 

shallow karst flow systems of SEW include 

dissolution, mixing, redox reactions, ion 

exchange, denitrification, volatilization, and 

precipitation. Of these, the advective transport 

and fast flow zones of the aquifer favor mixing 

as the most dominant of the processes observed, 

although in specific slow-flow zones any of the 

others may be important as well (Laincz, 2011; 

Winston, 2006). 

 Nutrient (nitrate) processing was 

characterized using stable isotopes, dissolved 

organic carbon concentration, and 

bioavailability, and concentration of reactive 

(nitrate) versus conservative (chloride) species

(Laincz, 2011). Additionally, concentration and 

isotopic composition of NO3

- was used to 

determine the extent of denitrification and 

immobilization of nitrate in the interflow zone. 

Study results show that as much as 30% of 

nitrate moving through the interflow zone can be 

microbially processed within this zone, and the 

level of processing is highly dependent upon 

flow-path and hydrologic conditions. 

Bioavailability of DOC in the interflow is 

increased relative to the focused-flow zone 

under high-flow conditions. Nitrogen and 

oxygen stable isotope data for nitrate suggest 

denitrification is occurring in the interflow zone. 

This appears to be a potentially important zone 

for nitrate attenuation in karst settings, and is an 

important finding for mantled karst areas where 

animal production is a major land use. 

Insight 6—Of the Hydrologic-Budget 

Parameters of Epikarst Springs, 

Evapotranspiration Losses Are Largest

The measurement of interflow through three 

springs in basin 1 (Tree, Woodpecker, and Red 

Dog) provides insight into a component of the 

karst hydrologic budget that is seldom quantified

(Brahana et al., 2005; Leh, 2008; Laincz, 2011). 

From 15 through 20 July 2005, representing the 

active plantgrowth period, and from 12 through 

17 December 2005, representing the plant 

dormant period, very precise measurements of 

discharge were made in early morning and midafternoon. Ten repetitions were averaged for 

each measurement, and data for diurnal and 

seasonal intervals were averaged (table 2).

 Results indicate that 1) diurnal fluctuation 

of these epikarst springs is notable during the 

growing season and minimal during the dormant 

season when deciduous trees have lost their 

leaves and transpiration is minimal. This loss of 

water from the shallow ground-water system 

during the growing season is interpreted to be 

evapotranspiration from the groundwater 

system, and ranges from 5 to 17 mL/s over the 

course of a diurnal cycle for each spring; 2) 

during the dormant season, flow losses between 

morning and afternoon are 1 mL/sec or less; 3) 

extreme low-flow conditions are manifest 

differently in different spring basins, and these 

provide an understanding of the flow 

mechanisms that may be active in karst settings; 

4) the wide range of hydrogeologic response in 

epikarst springs to identical stresses indicates 

that our models of these systems are likely 

grossly oversimplified.

Insight 7—Closely Spaced Flow Zones in 

Karst May Not Be Hydraulically 

Interconnected—Be Careful with Well 

Data

 One of the most surprising observations we 

have gained studying the karst of SEW and the 

Ozarks is the degree of anisotropy and 

heterogeneity present in the bedrock aquifers. 

The observation is based on numerous surfacecased open-hole wells, some as near to one 

another as 10 m, that show little or no hydraulic 

interconnection. Pumping and aquifer testing of 

these wells, many of which are more than 50 m 

deep, reveal hydraulic conductivites that are less 

than 1 cm/d. Approximately half of the open 

boreholes at SEW show this phenomenon 

(Pennington, 2010; Al Rashidy, 1999). This 

indicates that secondary karst development 

along bedding planes or other near-horizontal 

fractures is localized and not pervasive 

throughout the region. Drilling and aquifer 

testing the St. Joe in the Clear Creek graben area 

identified by Stanton (1993) indicated that, 

contrary to expectations, no karstification within 
138

the borehole had occurred. Other boreholes are 

closely interconnected, but this insight is 

especially valuable when interpreting data 

collected from wells in these settings. Because 

the well borehole intersects only a point, and 

they may or may not be connected to other 

components of the integrated flow system, data 

and interpretations therefrom should be carefully 

considered before flow considerations are made. 

Insight 8—Most Karst Environments and 

Ecosystems Are Surprisingly WellPopulated

In addition to microbial fauna which have 

been studied at SEW (Hamilton, 2001; Whitsett, 

2002; Ting, 2003 & 2005; Winston, 2006; 

Laincz, 2011; fig. 6), benthic invertebrates have 

been characterized by Mike Slay (written 

commun. The Nature Conservancy, 2010), and 

macro-invertebrates captured in passive traps set 

in springs and wells have been described by 

Mitchell (this volume). Combined assemblages 

total more than 30 genera, including cave 

amphpods (Stygobromus ozarkensis), troglybitic 

isopods (Caecidotea sp.) in springs, and 

copopods in wells isolated from surface water. 

This diversity and numbers of sensitive 

organisms observed indicate the ecosystems in 

this karst environment support a wide variety of 

stygofauna.

Insight 9—Understanding Transport of 

Suspended Sediments and Pathogens 

in Karst Requires a Different Tool Set

 In the rapid-flow part of the saturated 

aquifer, suspended sediments and bacteria are 

transported more slowly than dissolved tracers, 

averaging about 2 m/hr during conditions of low 

flow, as compared to as great as 800 m/hr for 

conservative solutes. Average velocity may be 

misleading, however, because it appears the 

suspended particles are deposited along the 

subsurface flow path, and do not move until 

increased recharge from storms increases 

velocity enough to scour, resuspend, and retransport the particles. Below a threshold 

ground-water velocity, turbulent flow does not 

occur and suspended particles are not in transit. 

This is consistent with observations and 

measurements at the spring resurgences; 

increases in stage, and E.coli, tagged clay and 

total suspended solids concentrations are rapid 

and occur around the same time, then decrease 

quickly. Ting (2002 & 2005) was able to utilize 

field-scale testing and verify that unique tagging 

of clays and microbes was a valuable tracing 

tool in this environment. 

Figure 6. This image is from a borehole televiewer in 

an uncased well below the regional confining layer 

(Chattanooga), showing thriving colonies of 

unidentified facultative bacteria.

Insight 10—Even the Simplest of Karst 

Systems Are Complex, and Predictions 

Should Be Offered with Great Care

The complexities of flow in karst terranes 

are well-documented throughout the world 

(Palmer, 2007), and SEW is no different. 

Continuous monitoring of two springs at the 

distal end of the groundwater basin 1 provides 

an illustrative example. Langle Spring, the 

underflow part of the focused flow system, is 

about 3 centimeters lower than Copperhead, the 

overflow spring. During a 33-day period 

starting in December 1997, five major storms 

perturbed the stability of the temperature of 

these springs (fig. 7). These produced an 

general overall cooling in the discharge waters 

of Langle, although the trend was by no means 

linear or gradual. Water from Copperhead, on 

the other hand, showed both warming and 

cooling trends, numerous abrupt reversals in 

heating and cooling, as well as many more 

fluctuations of about 1o C in the interval from 

December 13 through December 20. These 

diurnal variations were ultimately traced to a 
139

Table 2. Seasonal and diurnal variation of spring 

discharge in three epikarst springs in basin 1.

[units of discharge in mL/sec; am, 7:00 am measurement; 

pm, 3:00 pm measurement; ∆, difference between am and 

pm; growing, July; dormant, December]

\ Spring Tree

Season/Time\

WoodPecker Red Dog

growing/am 216 105 99

growing/pm 199 93 94

growing/∆ 17 8 5

dormant/am 103 92.7 78.9

dormant/pm 102 93.0 78.0

dormant/∆ 1 <1 <1

leaky, ephemeral surface pond that commonly 

stored water after periods of intense rainfall. 

The pond was exposed to solar heating during 

clear weather following storms, and the heated 

water leaked into the karst aquifer and as it 

mixed with water from other sources, it imparted 

the diurnal thermal signature to Copperhead 

Spring. The signature was not obvious at 

Langle, owing to its much longer distance from 

the leaky pond. The wider ranges of temperature 

in Copperhead Spring suggest that this part of 

the system is more open to surface water, and 

less thermally isolated than Langle. Cold water 

from the storm of 25 December obviously had 

an impact on the thermal regime of both springs, 

but the slope of the decrease at Langle supports 

the hypothesis that much of the water from the 

northern part of the spring system is more 

insulated from surface effects. Near identical 

temperatures from the two springs on 8 

December 1997 is interpreted to be caused by 

point-source input to each spring from fractures 

nearby the orifices. This effect is repeated on 25 

December, and is thought to be a strong 

indication of the temperature of the precipitation 

at the time of coincident temperatures.

 Documentation by continuous long-term 

monitoring gives us confidence that system 

responses cannot be easily simplified in what we 

initially conceptualized to be a non-complex 

karst system. Predictive numerical modeling 

does not appear to be appropriate, although 

hypothesis testing using models is definitely 

valuable and meaningful. 

Figure 7. Temperature of two springs, Langle and Copperhead, that resurge from the St. Joe Limestone at the distal 

end of the ground-water flow path in Basin 1. Temperature variations are natural tracers that show the complexity 

of input sources and thermal equilibration of this overflow/underflow spring.
140

SUMMARY

 The ten topics that are presented in this paper 

represent an arbitrary selection of studies 

conducted in mantled karst at SEW. The main 

concepts and insight are not new, but we feel the 

level of understanding gained contributes 

significantly to the overall knowledge base of 

karst science. The establishment of long-term 

research sites is neither trivial nor inexpensive, 

yet the opportunities to quantify watershed 

system response in situ to a wide range of 

natural stresses is essential to ultimately 

reaching true understanding of the processes that 

dominate this kind of setting. Time demands 

and economics will always influence applied 

research in karst, but the existence of sites such 

as SEW provides scientists the opportunity to 

develop new tools and to more fully explore 

quantifiable aspects of a truly complex system.

ACKNOWLEDGMENTS

 Insight described herein is based primarily 

on the work of many graduate students, field 

hydrogeology classes, colleagues, and visiting 

karst researchers, and I have tried to cite the 

importance of their many contributions. 

Unfortunately, limited space has constrained the 

full description of all who have contributed, and 

I encourage the interested to seek out published 

technical reports at the SEW website, which can 

be accessed at:

http://www.uark.edu/depts/savoyres/index.html

Reviews by Jozef Laincz, Justin Mitchell, and 

Phil Hays significantly improved this 

manuscript, and are gratefully acknowledged.

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141

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Highlands: Hydrological Processes, Wiley 

Interscience, DOI: 10.1002/hyp7021, 15 p.

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Orndorff, H.O., 2002, Movement of nitrate 

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aquifer in northwestern Arkansas: unpublished 

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mantled karst of the Savoy Experimental 

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biogeochemistry: Unpublished M.S. thesis, 

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Experimental Watershed, northwest Arkansas:

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142

Traps Designed to Document the Occurrence of Groundwater 

Stygofauna at the Savoy Experimental Watershed, Washington 

County, Arkansas 

By Justin R. Mitchell

Department of Geosciences, 113 Ozark Hall, University of Arkansas, Fayetteville, AR 72701

Abstract

The capture and documentation of stygofauna (obligate aquatic cave dwellers) in the Springfield 

Plateau of the Arkansas Ozarks has been typically limited to the physical access of caves, springs, and 

sinkholes. This project has expanded the potential search area for stygofauna to include wells through the 

use of specially designed traps. It was hoped that this study could find an unknown population of cave 

crayfish; however no Cambarus species were observed. Baited traps were set in three springs and seven 

wells at the Savoy Experimental Watershed, a University of Arkansas Experimental Station located in 

Washington, County Arkansas. Traps were checked three times over a three week period in November 

2010. These traps were proven effective in actively capturing various zooplankton species, lotic leaches, 

and the epigean crayfish Orconectes neglectus. Cyclopoid copepods (small shrimp-like crustaceans)

were repeatedly captured in two deep wells, whereas five deep wells produced no faunal specimens. 

Cyclopoid copepods as well as five other epigean (surface dwelling) aquatic taxa were captured in the

three springs sampled. The use of these combination traps has demonstrated an effective active technique 

for documenting aquatic organisms occurring in the groundwater of wells and surface springs at the 

Savoy Experimental Watershed.

INTRODUCTION

Six distinct species of stygobitic crayfish of 

the genus Cambarus (Decapoda: Cambaridae) 

are known to exist in the Ozark Plateau region of 

Arkansas, Missouri, and Oklahoma (Koppelman 

and Figg, 1995). These unpigmented sightless 

crayfish with long slender claws are the largest 

invertebrate members of Ozark cave ecosystems 

(Koppelman and Figg, 1995). The Benton 

County cave crayfish, Cambarus aculabrum, is a 

unique aquatic resident of Northwest Arkansas’ 

cave systems and a federally listed endangered 

species (Graening and others, 2006) (Fig. 1).

 The documented range of this troglobitic 

crustacean is limited to four locations within the 

Springfield Plateau of the Arkansas Ozarks, all 

in Benton and Washington counties (Graening 

and others, 2006). However, due to the geology 

of the extensive karst environment in which 

these crayfish are found, they are likely to occur 

elsewhere in the ground water of the region. 

Except for caves and springs at the distal end of 

karst flow systems, researchers are physically 

too large to access much of the hypothetical 

habitat of these animals. Therefore, specimens 

are almost always observed in caves and springs, 

explaining the common name of cave crayfish.

Figure 1. Cambarus aculabrum, The Benton 

County Cave Crayfish, Photo by Michael E. 

Slay, The Nature Conservancy.

 One proven historical method of 

documenting Cambarus in locations inaccessible 

for direct human exploration was through the 

use of specially designed traps lowered into 

wells (Purvis and Opsahl, 2005). This method 

was implemented in Georgia, U.S.A. and 
143

successfully extended the range of a karst 

dwelling Cambarus species. 

Similar traps were manufactured for this 

study and modified from the original design to 

potentially capture a variety of stygofauna in 

addition to crayfish. Numerous other organisms 

exhibiting exclusively subterranean aquatic life 

styles are known to occur in the Ozarks. These 

karst aquifer residents include cavefish, 

salamanders, snails, isopods, amphipods, and 

other invertebrates (Elliott, 2007). This report 

describes the effectiveness of these modified 

traps and investigates the existence of 

stygofauna at a rural location in Washington 

County, Arkansas (Figure 2).

Figure 2. Deep well 72 and surrounding 

landscape, Savoy Experimental Watershed, 

October 2010. 

Study Area

 The site for this study was the Savoy 

Experimental Watershed (SEW) located 24 

kilometers west of the University of Arkansas 

campus in Fayetteville, Arkansas. Situated 

within the Springfield Plateau of the Ozarks, the 

SEW is an approximately 1250-hectare property 

that serves as a scientific research area for the 

University of Arkansas’ Departments of Animal 

Science, Agronomy, Geosciences, and Poultry 

Science, along with various state and federal 

entities (Brahana and others, 2001). The 

landscape is hilly with well dissected plateaus 

cut by dry valleys. Vegetation is composed of 

mixed deciduous forests and grassy pastures. 

Limited cattle grazing, confined swine 

production, and poultry operations occur within

discrete parts of the property, associated with 

Animal Science research. Most of the SEW, 

however, is undeveloped second-growth forest.

Figure 3. Location of the Savoy Experimental 

Watershed and the distribution of Cambarus 

aculabrum with other cave crayfish populations 

in adjacent states-. Cave crayfish distributions 

modified from Graening and others, (2006).

The SEW is located in northern Washington 

county and is relatively close to documented 

Cambarus populations (Fig. 3) Although no 

Cambarus species are known from SEW, a cave 

isopod Caecidotea sp., and a cave amphipod

Stygobromus ozarkensis occur in Copperhead 

Spring, indicating suitable habitat for stygobitic 

organisms (Graening and others 2005; Slay,

2000). A Caecidotea species was also observed 

in Dead Cow Spring. A total of twenty-nine 

unique taxa have been collected in four SEW 

springs using a D-ring kick net following 
144

substrate disturbance (Slay, 2000). The taxa 

richness of the four springs studied was found to 

range from fourteen to twenty-two distinct 

species per spring (Slay, 2000). The 

biodiversity discovered previously in the SEW 

springs justifies continued investigations of 

groundwater habitats and their inhabitants 

occurring within the property.

The SEW is characterized as a mantled karst 

terrain and drains into the adjacent Illinois 

River. The cave-forming carbonate rock 

underlying the area is generally covered in a 

layer of regolith of variable thickness, consisting 

of insoluble chert and clays. Beneath the lower 

Mississippian Boone and St. Joe limestone 

formations, exists a layer of upper Devonian 

shale, the Chattanooga formation, which acts as 

a regional confining layer, preventing the 

vertical movement of groundwater (Brahana,

1997). Dissolution of the limestone occurs at 

this contact zone, along joints, and within 

overlying bedding planes influenced by the 

presence of insoluble chert (Brahana, 1997). 

This chemical weathering process enlarges paths 

for the movement of water, forms caves, and 

provides potential habitat for stygofauna. 

The existence of numerous deep test wells 

and springs on one contiguous tract of land with 

easy access was convenient and ideal for this 

study. The wells are in close proximity to 

natural springs and caves and can be visited in a 

relatively short period of time. The deep wells at 

SEW were drilled 30.5 meters to 80.5 meters 

deep. Geophysical logs of these wells have 

provided valuable information about the 

subsurface geology and hydrology at SEW

(Stanton and others, 1998) (Fig.4). Most deep 

wells penetrate through the regolith, Boone, 

St.Joe, and Chattanooga formations and 

terminate in the Ordovician dolomite beneath, 

accessing two karst aquifers separated by the 

Chattanooga confining layer. Fluid resistivity 

and fluid conductance logs indicate bedding 

plane openings in the Boone formation and 

voids in the contact zone between the St. Joe 

limestone and Chattanooga shale (J. Van 

Brahana, 2010, Personal Communication, 

Department of Geosciences, University of 

Arkansas, Fayetteville, Arkansas; Keys, 1990). 

These cavities facilitate the storage and 

movement of ground water and are possible 

habitats for Ozark stygofauna.

Figure 4. Geophysical log of Deep Well 72 at 

the SEW-. Well log modified from Stanton and

others, (1998).

Materials and Methods

 The working principle of the traps used is 

simple and non lethal: the bait attracts organisms 

to the trap, allows organisms to enter, and 

restricts an easy exit. The traps used for this 

study were made from polycarbonate sports 

drink containers similar to the design used by 

Purvis and Opsahl. However, for this study, 

improvements were made in the construction 

and modifications were made to facilitate the 

capture of smaller organisms including 

zooplankton (Fig.5). Only one specific type of 

bottle would work for this new design owing to 

the ribbed shape of the container. The top 6 cm 

of the bottle was cut off and slowly heated until 

it contracted and became malleable. The hot 

upper portion of the bottle was then pressed into 

the lower portion of the bottle and lodged 

between the ribs. This technique created a 

compression fitting between the upper and lower 

portions. No glue was required for attachment 

which allowed for easy opening and closing of 

the trap in the field. Small 4 mm holes to 

facilitate air and water flow were melted in the 

bottom and upper portion of the trap with a 

soldering iron. Holes were also melted in the 

sides for the attachment of three 15 cm nylon 

leader lines, used to lower and retrieve the trap. 
145

Leader lines were attached using plastic cable 

ties. Four 15 g glass weights were glued to the 

sides using 100% silicone adhesive. Instead of 

using lead sinkers, glass weights were used to 

avoid introducing lead into the ground water. A 

large plastic cable tie was threaded through 

holes in the bottom and left protruding out of the 

bottom. This forced the trap to fall or lean on its 

side instead of resting flat on its bottom. A 

custom designed conical polyester plankton net 

was then attached to the base with a cable tie. 

The plankton net was easily removed and 

refitted in the field. The traps were baited with 5

g of raw chicken liver placed in small polyester 

bags and fastened with a small cable tie to the 

base of the trap.

Traps were placed in seven deep wells and 

allowed to sink to the bottom. Traps were also 

set in three springs. Two springs sampled, 

Copperhead Spring and Langle Spring, emerge 

at the contact zone of the Boone and the St. Joe 

limestone formations (Al Rashidy, 1999). The 

other spring sampled, Red Dog Spring, is an 

epikarst spring that emerges from the contact 

zone between the regolith and the Boone 

formation. Traps were set at all 10 locations 

October 27, 2010 and checked November 6, 10, 

and 22, 2010. There was not any significant rain 

during the sampling period and the springs 

flowed at typical low flow discharges (Al 

Rashidy, 1999).

Each time a trap was checked, large 

organisms were removed from the bottle and a 

sample was collected from the plankton net. 

The net was detached from the trap, inverted 

over a collection container, and the contents 

rinsed with 70% ethanol. The specimens present 

were killed and preserved in this process. These 

samples were later examined in a lab using a 

dissection microscope. Organisms were 

identified to a basic taxonomic level using 

Pennak (1989), sorted, and preserved in 70% 

ethanol for later species identification. 

Figure 5. Combination Crayfish/ Zooplankton 

Trap

RESULTS

Aquatic invertebrates were captured in all 

three springs and two of the seven deep wells. 

Five wells did not yield any faunal organisms. 

The target species, Cambarus aculabrum, was 

not found at any location. The stygobitic 

Caecidotea sp. and Stygobromus ozarkensis

were not captured. However, the trap was able 

to capture zooplankton in both wells and 

springs. Cyclopoid copepods were captured in 

two wells: well 70 (2 individuals) and well 72

(27 individuals). Cyclopoid copepods were also 

captured in the three springs: 9 from Copperhead

Spring, 223 from Langle Spring, and 248 from 

Red Dog Spring. Ostracods were found in two 
146

springs: 174 in Red Dog Spring and 3 in Langle

Spring. Samples from Red Dog Spring also 

yielded 95 Daphnia and 1 Hydra. In addition to 

zooplankton, larger invertebrates were also 

captured in the springs. 80 gastropods (snails) 

and 3 Hirudinea (leeches) were collected in Red 

Dog Spring. Orconectes neglectus, the epigean 

crayfish commonly known as the ringed crayfish

(Pflieger, 1996), was captured once in 

Copperhead Spring and once in Langle Spring.

DISCUSSION

The well sampler originally designed by 

Purvis and Opsahl (2005) appears to be a useful

tool for sampling non-typical groundwater 

habitats (i.e. wells), and the trap modifications 

presented here allow the trap to capture aquatic 

organisms of a variety of sizes. Cambarus 

aculabrum is a federally listed endangered 

species and any discoveries of its existence or 

the existence of other stygofauna outside of 

known locations would be useful for 

conservationists and beneficial for the species. 

This project serves as a model study that could 

be reproduced, improved, and implemented in 

other wells and springs in the region or 

elsewhere in karst environments around the

world. Due to the negative anthropogenic 

impacts on cave habitats and ground water 

quality throughout the region, it is important to 

locate, document, and protect stygobitic species 

for their continued survival (Graening and 

others, 2006). Cave animals are sensitive to the 

water quality, chemistry, geology, and 

hydrology of their unique habitat. The presence 

(or absence) of some sensitive species can 

indicate changes in water quality over time 

(Palmer, 2007). The Purvis and Opsahl trap, 

along with the modifications, can potentially 

help ecologists with population and distribution 

studies, cave fauna inventories, and aid in new 

discoveries. Little characterization of 

organisms found in karst aquifers has been 

undertaken and more extensive work needs to be 

done to identify, understand, and protect the 

stygofauna of the Ozarks and beyond. Such 

studies translate into good stewardship of our 

precious groundwater resources. 

The copepods captured in deep well 72 and 

deep well 70 are of some significance and 

possibly cave forms due to their apparent 

isolation from the surface (Michael E. Slay,

2010, Personal Communication, The Nature 

Conservancy, Fayetteville, Arkansas). Well 72 is 

172 feet deep and at the time of this study the 

water level was 91 feet beneath the surface. Well 

70 is 103 feet deep and the water level was 7 

feet from the surface. The wells are always 

capped except for brief intervals of water-quality 

sampling and trap insertion and removal. The 

capture of these copepods on two separate

occasions in these wells during a period of base 

groundwater flow indicates the existence of 

groundwater invertebrates living deep below the 

surface, relatively far from any associated 

surface orifice at SEW. Plans call for specimens 

collected in this study to be sent to a copepod 

taxonomist for species identification.

Even though two crayfish were captured, 

larger crayfish, most likely Orconectes 

neglectus, appeared to be able to feed on bait 

and escape from this trap. One insecurely 

fastened bait bag was completely removed from 

a trap in Copperhead Spring (Fig.6) Another bait 

bag from a trap in Langle Spring was found torn 

with the bait removed. This problem has been 

considered and a new escape prevention 

modification consisting of a collapsible mesh 

tube attached to the inside of the trap’s entrance 

has been designed. This modification is 

currently being tested for functionality in the 

field. 
147

Figure 6. Copperhead Spring, SEW, October 

2010

ACKNOWLEDGMENTS

Thanks to Mike Slay of The Nature 

Conservancy (TNC) for suggesting the plankton 

net addition to create this combination trap. 

Slay provided support, materials, and assistance 

with identification of the collected specimens for 

this project and he facilitated permission from 

The Arkansas Game and Fish Commission to 

proceed with this project. Thanks to Jerry Anne 

Mitchell for sewing the plankton nets and bait 

bags, Jimmy Mitchell Sr. for suggesting glass 

for non-toxic weights, Erin Timmons-Mitchell 

for her continuing support of my education and 

scientific endeavors, and Professor J. Van 

Brahana for his continuing encouragement, 

assistance, and support for this project.

REFERENCES

Al Rashidy, S. 1999. Hydrogeologic controls of 

groundwater in the shallow mantled karst 

aquifer, Copperhead Spring, Savoy 

Experimental Watershed, Northwest 

Arkansas. Unpublished M.S. Thesis, 

Department of Geosciences, University of 

Arkansas, Fayetteville.

Brahana, J. V. 1997. Rationale and methodology 

for approximating spring-basin boundaries 

in the mantled karst terrane of the 

Springfield Plateau, Northwestern Arkansas. 

The Engineering Geology and 

Hydrogeology of Karst Terranes. Beck & 

Stephenson eds. Balkema, Rotterdam. p.77-

82.

Brahana, J. V., T. L. Phelan, P. D. Hays, T. J. 

Sauer, R. K. Davis, and J. H. Cole. 2001. 

Savoy Experimental Watershed and field 

research facility- a long-term karst research 

site. University of Arkansas, Fayetteville. 

poster.

Elliott, W. R. 2007. Zoogeography and 

biodiversity of Missouri caves and karst. 

Journal of Cave and Karst Studies. v. 69, no. 

1, p. 135-162.

Graening, G. O., M. E. Slay, and J. R. Holsinger. 

2005. Annotated checklist of the amphipoda 

of Arkansas with an emphasis upon 

groundwater habitats. Journal of the 

Arkansas Academy of Science. v. 59, p. 80-

87.

Graening, G. O., M. E. Slay, A. V. Brown, and J. 

B. Koppleman. 2006. Status and distribution 

of the endangered Benton cave crayfish, 

Cambarus aculabrum (Decapoda: 

Camaridae). The Southwestern Naturalist. 

v. 51, no. 3, p. 376-439

Harrod, J. 2006. Saving the last great places 

under earth. The Nature Conservancy. 

http:/www.nature.org/wherewework/northameric

a/states/Arkansas/preserves/art24994.html. 

accessed October 2010.

Keys, W. S. 1990. Techniques of waterresources investigations of the United States 

Geological Survey, Chapter E2 Borehole 

Geophysics Applied to Ground-Water 

Investigations. U.S. Department of the 

Interior, U.S. Geological Survey. Denver, 

Colorado. p. 9-22.
148

Koppelmann, J. B. and D. E. Figg. 1995. Genetic 

estimates of variability and relatedness for 

conservation of an Ozark cave crayfish 

species complex. Conservation Biology. v. 

9, no. 5, p. 1288-1294.

Palmer, A. N. 2007. Cave Geology. Cave Books. 

Dayton, Ohio. 

Pennak, R. W. 1989. Fresh-Water Invertebrates 

of the United States Protozoa to Mollusca 3rd

Ed. John Wiley & Sons, Inc. Wiley –

Interscience Publication. New York.

Pflieger, W. L. 1996. The Crayfishes of 

Missouri. Missouri Department of 

Conservation. Jefferson City, Missouri. 

Purvis, K. M. and S. P. Opsahl. 2005. A novel 

technique for invertebrate trapping in 

groundwater wells identifies new 

populations of the troglobitic crayfish, 

Cambarus cryptodytes, in Southwest 

Georgia, USA. Journal of Freshwater 

Ecology. v. 20, no. 2, p. 361-365.

Slay, Michael. 2000. Aquatic faunal inventories 

of four springs on the Savoy Experimental 

Watershed. Unpublished Karst 

Hydrogeology term paper, Department of 

Biological Sciences, University of Arkansas, 

Fayetteville.

Stanton, G. and Students. 1998. Geophysical log 

data collected by the Summer Field 

Hydrogeology class, Department of 

Geosciences, University of Arkansas and the 

U.S. Geological Survey Southeast Region 

Logger Service. in Al Rashidy, S. 1999. 

Hydrogeologic controls of groundwater in 

the shallow mantled karst aquifer, 

Copperhead Spring, Savoy Experimental 

Watershed, Northwest Arkansas. 

Unpublished M.S. Thesis, Department of 

Geosciences, University of Arkansas, 

Fayetteville.
149

Assessment of Sinkhole Formation in a Well Field in the Dougherty 

Plain, near Albany, Georgia

By Debbie Warner Gordon

U.S. Geological Survey, 3039 Amwiler Rd, Suite 130, Atlanta, GA 30360

Abstract

To assess the causes of sinkhole formation and determine if groundwater withdrawals are affecting 

the frequency of sinkhole development, the U.S. Geological Survey, in cooperation with the Albany 

Water, Gas, and Light Commission, conducted hydrologic studies at an Upper Floridan aquifer well field 

southwest of Albany, Georgia, during 2009–2010. The Upper Floridan aquifer is the main water-supply 

source in southwestern Georgia and consists of limestone that is overlain by 30 to 60 feet (ft) of sand and 

clay (undifferentiated overburden). Borehole data, including geophysical, lithologic, and video, indicate 

two zones of high secondary permeability (dissolution features) within the Upper Floridan aquifer at the 

well field. These zones are cavernous and may be vulnerable to collapse. Loading from ponds excavated 

during well-field construction and water-level fluctuations due to well-field pumping increase the 

potential for sinkhole development. 

Hydrologic data, including groundwater levels, production-well pumping records, and rainfall, were 

compared to subsurface data to assess the formation of sinkholes at the well field. From October 2003 

(initiation of well-field pumping) through December 2009, 23 sinkholes developed at the well field. Most 

of the sinkholes formed in or adjacent to storage ponds located along the western part of the well field 

(fig. 1). These manmade ponds increase the load on the underlying sediments, which may increase the 

potential for sinkhole formation (Tihansky, 1999). A cluster of eight sinkholes southwest of the ponds 

formed while a nearby production well was pumping, so the well was removed from production. Other 

sinkholes formed farther away from the ponds and may be linked hydrologically to high-permeable zones 

in the aquifer. 

Sinkholes may form in response to changes in water levels relative to the top of the limestone aquifer 

or relative to the top of a cavity in the limestone. A comparison of water levels in well 12L382 during 

October 2003–December 2009 to the depths to the top of the Upper Floridan aquifer and top of the upper 

permeable zone shows the relation of sinkhole formation to groundwater levels in the Upper Floridan 

aquifer (fig. 1). The relation between groundwater level fluctuations and sinkhole formation at the well 

field is not completely understood. In 2003 and 2005, water levels were almost 20 ft above the top of the 

aquifer for most of the year, yet two sinkholes formed during each of those years. During 2007, water 

levels oscillated above and below the top of the Upper Floridan aquifer and 12 sink holes formed; 

however, no sinkholes formed during 2008, even though water levels and pumping patterns were similar 

to those during 2007. Ongoing monitoring of groundwater levels, with the possible addition of pond 

stage, may provide additional insight into factors controlling sinkhole formation at the well field. 
150

Figure 1. The number of sinkholes that formed in the Albany well field, Dougherty County, Georgia, daily depth to 

water in well 12L382, and the depth to the top of the Upper Floridan aquifer and upper permeable zone at well 

12L382, October 2003–December 2009.

REFERENCE CITED

Tihansky, A.B., 1999, Sinkholes, west-central Florida, in Galloway, Devin, Jones, D.R., Ingebritsen, S.E., eds., 

Land subsidence in the United States: U.S. Geological Survey Circular 1182, p. 121–140.
151

Detecting Karst Conduits through their Effects on Nearby Monitoring 

Wells

By Fred Paillet and Terryl Daniels

Geosciences Department, University of Arkansas, 113 Ozark Hall, Fayetteville AR 72701 

Abstract

Karst aquifers transmit flow through high-permeability conduits that comprise a very small 

fraction of the porosity of the total rock mass. These conduits are very difficult to locate and characterize 

by the random drilling of boreholes. The characterization of conduits at the University of Arkansas 

Savoy wellfield is especially difficult because the conduits are developed along vertical joints in a 

formation overlain by competent chert beds that mask the surface expression of the subsurface solution 

openings. In this study we examine the conditions under which water level fluctuations and transient 

vertical flow in observation wells can indicate the presence of conduits in the surrounding rock mass. 

Slug tests in Savoy wells 69 and 70 indicate total well transmissivity in the range of 0.1 to 0.3 ft2

/day. 

We adopt a model where one or more bedding plane openings with T = 0.1 ft2

/day transmit flow to 

boreholes from hydraulic head increases of about a meter in a nearby karst conduit, with a linear increase 

over 24 hours followed by an exponential decline thereafter. The model (Paillet, 1998) uses an infinite 

bedding plane of specified T and S embedded in low-permeability rock and connected to a reservoir 

subject to fluctuating water levels. The model uses the slug test response integrated over time to compute 

water level and flow in the observation well as the water level in the rock surrounding the bedding plane 

is driven by hydraulic head changes in the karst conduit. The observation borehole water level reflects 

the full conduit head change modulated by the significant flow losses incurred in moving water into or out 

of the well bore. Model runs show that the low transmissivity of the bedding plane causes the response 

in the observation well to lag behind the head changes in the nearby conduit whenever bedding plane T is 

less than 100 ft2

/day. When bedding plane T = 0.1 ft2/day and we use S = 10-4, a conduit head change of 

1.5 m over 24 hours produces water level fluctuations that are about 20% of the conduit increase (readily 

measurable), and produces vertical flow in and out of casing storage of about 0.003 gpm (below the 

resolution of the HP flowmeter but measurable with dilution methods). If there are two such bedding 

planes and only one connected to the conduit, the model predicts very little change in the response, with 

outflow into the unconnected bedding plane having almost no effect on the response. If there are two 

bedding planes connected to two different conduits, the observation well response will give the Tweighted average of the heads in the two conduits. If, however, an unconnected plane has significant 

transmissivity (for example, T = 10 ft2

/day; probably not the case at Savoy), virtually all of the inflow 

driven by head changes in the karst conduit goes into that other more permeable “thief” zone. In that case 

there would be no measurable change in water level in the observation well, while the flow between 

bedding planes would be too small to measure with the HP flowmeter and could only be detected by 

dilution methods. On the basis of these results we propose a program at Savoy where we use dilution 

methods to identify the depth and T of inflow zones in observation wells, and then water level 

fluctuations after rainfall events to infer water level fluctuations in karst conduits to which those planes 

are connected. If we suspect that two karst conduits and two or more bedding planes are involved, we 

can use dilution methods to estimate flow in the observation well during response to rainfall events to 

estimate the head changes in the two separate conduits. We believe that this model approach will allow 

us to use monitoring well responses to infer hydraulic head changes in the karst conduits in the vicinity of 

observation wells even when the well does not actually penetrate the conduit opening. 

REFERENCE

Paillet, F.L., 1998, Flow modeling and permeability estimation using borehole flow logs in heterogeneous fractured 

formations, Water Resources Research, 34, 997-1010.
152

Spring Hydrology of Colorado Bend State Park, Central Texas

By Kevin W. Stafford, Melinda G. Shaw, and Jessica L. DeLeon

Department of Geology, Stephen F. Austin State University, 1901 N. Raguet, Nacogdoches, TX 75962 

Abstract

Karst development in Ellenburger carbonates near Colorado Bend State Park in central Texas exhibits 

complex polygenetic origins, with porosity development dominated by an early hypogene phase that has 

subsequently been overprinted to varying degrees by epigene processes. Quarterly physicochemical and

continuous thermal monitoring analyses of eight springs in the study area indicate that modern 

groundwater flow paths are highly variable. Springs exhibit patterns that range from shallow, distributed 

recharge into diffuse-flow dominated systems, to focused recharge into well-connected conduit systems, 

to deep-circulation systems that equilibrate with bedrock. All springs, except Sulphur Spring, exhibit 

physicochemical characteristics indicative of proximal epigenic groundwater flow through Ellenburger 

carbonates, while Sulphur Spring shows elevated temperature and dissolved-ion concentrations indicative 

of longer groundwater flow paths through deeper strata. The polygenetic nature of karst development in 

the Colorado Bend State Park has created an enhanced porosity structure which forms a complex modern 

groundwater flow network.

INTRODUCTION 

Colorado Bend State Park (CBSP) is located 

on the northern edge of the Texas Hill Country 

on the flank of the Llano Uplift (Figure 1, 2). 

Here Ordovician Ellenburger carbonates crop 

out along a highly entrenched segment of the 

Colorado River, immediately upstream from 

Lake Buchanan. Proximal to the river, numerous 

springs discharge including 

subaqueous springs, springs within a 

few meters of the river and springs that 

discharge hundreds of meters from the 

river (Figure 1). Most springs 

discharge with normal epigenic karst 

chemistries; however, one spring in the 

region, Sulphur Spring, discharges with 

a slightly thermal component, an 

elevated sulfate content and easily 

discernable odor of hydrogen sulfide.

CBSP is located approximately 180 

kilometers northwest of Austin, Texas 

in San Saba and Lampasas Counties. 

The park covers 21.6 square kilometers 

including a seven kilometer long 

stretch of the Colorado River. The area 

is located along the boundary between 

subtropical steppe climate and 

subtropical subhumid climate, with 

average annual temperature of 20oC 

and minimum and maximums of 8oC 

and 30oC respectively (Estaville and Earl, 2008). 

Annual precipitation averages 30 cm, with most 

precipitation occurring during Spring (March –

May) and Fall (September – November). 

More than 400 karst features have been 

identified within CBSP and surrounding 

properties, including more than 100 physically 

mapped caves. Most caves exhibit 

characteristics of complex, polygenetic origins, 

Figure 1. Map of study site showing location of springs, 

outline boundary of CBSP in yellow and approximate 

location of study site with reference to the state of Texas.
153

including many that exhibit classic hypogene

MSRF (Morphologic Suite of Rising Flow) 

(Klimchouk, 2007) characteristics with varying 

degrees of epigenic overprinting, while other 

caves are more purely epigene in origin. DeLeon 

(2010) showed that at least two thirds of caves

developed within the CBSP region exhibit 

hypogene origins with variable degrees of 

epigene overprinting, while less than one fourth 

of all caves showed clear dominance of epigene 

origins. Eight known perennial springs discharge 

subaerially within the park, while two other 

springs discharge upstream within the 

commercially operated Sulphur Springs fish 

camp (Figure 1). Stream mapping conducted by 

Mitchell and others (2011) suggests that at least 

ten additional springs also discharge 

subaqueously within the Colorado River in this 

region. The polygenetic nature of karst 

development with abundant springs 

that discharge from subaqueous

position in the Colorado River to 

more than sixty five meters above 

river level suggests a complex 

hydrogeologic system that has not 

completely equilibrated with current 

climatic and geomorphic conditions.

GEOLOGIC OVERVIEW

The Llano Uplift, and greater 

Texas Hill Country, is dominated by 

Precambrian (~1.0 bya - ~1.2 bya) 

basement rocks which form a large 

structural dome overlain 

unconformably by Cambrian and 

Ordovician clastics and carbonates 

(Figure 2) (Sellards and others, 

1932). Ordovician strata are 

unconformably overlain by 

Carboniferous rocks that are 

subsequently overlain 

unconformably by Cretaceous strata 

that compose the northern extension 

of the Edwards Plateau (Rose, 1972). 

Precambrian strata were emplaced as 

part of the Grenville orogenic event, 

while these and overlying Paleozoic 

rocks in the study area were modified 

by tectonism associated with the 

Ouachita Orogeny, primarily 

resulting in minor tilting and faulting 

(Standen and Ruggiero, 2007). 

Throughout the Cenozoic, all strata within the 

region have seen additional brittle deformation 

and minor tilting as a result of uplift of the 

Edwards Plateau and down-dropping of the gulf 

coastal plain, with intense faulting along the 

Balcones Fault Zone, approximately 80 

kilometers east of the study area (Collins, 1995).

Karst development within the study area is 

largely limited to Ordovician Ellenburger 

carbonates, including cave and spring 

development in all three Ellenburger units, 

oldest to youngest — Tanyard, Gorman and 

Honeycutt Formations. The Tanyard Formation 

is ~170 meters thick and consists of fine- to 

coarse-grained, irregularly bedded dolomite 

deposited as high-energy, restricted, subtidal 

facies, including common ooidic zones and 

cryptalgal laminae (Kerans, 1990). The Gorman 

Figure 2. Spatial distribution of Llano Uplift aquifers in relation 

to project study area (red box) with ~10X vertically exaggerated 

cross section (modified from Preston and others, 1996).
154

Formation is ~130 meters thick and consists of 

micro-granular dolomite associated with 

deposition in a low-energy, restricted-shelf 

environment, including common macrofossils, 

distributed zones of intense burrowing and 

cryptalgal laminae, as well as zones indicative of 

subaerial exposure that include rip-up clasts and 

siliciclastic sediments (Kerans, 1990). The 

Honeycutt Formation is ~40 meters thick and 

consist of thinly interbedded limestones and 

dolomites that were deposited in an open, 

shallow-water shelf environment, including 

common structures indicative of periods of brief 

subaerial exposure (e.g. desiccation cracks, ripup clasts) and structures indicative of variable 

current energies (e.g. ooids, current ripples, 

cryptalgal laminae) (Kerans, 1990).

Within the study area, three minor aquifers 

collectively referred to as the Llano Uplift 

aquifers are developed in the Paleozoic rocks, 

including from bottom to top: Hickory aquifer, 

Ellenburger-San Saba aquifer, and Marble Falls 

aquifer (Preston and others, 1996). The aquifer 

system dips gently into the subsurface away 

from the central Llano Uplift, dipping mainly to 

the north in the study area. The Hickory aquifer 

is developed in the Cambrian Hickory 

Sandstone, which is underlain by Precambrian 

basement rocks. The Ellenburger-San Saba 

aquifer includes all three formations of the 

Ellenburger Group plus the San Saba member of 

the underlying Wilberns Formation. The 

Ellenburger-San Saba aquifer is 

compartmentalized in regions due to local and 

regional block faulting, with significant 

solutional overprinting and is in some regions 

locally connected hydrologically with the 

overlying Marble Falls aquifer (Preston and 

others, 1996). The Marble Falls aquifer is 

developed in corresponding Carboniferous 

limestone, which exhibits highly variable 

permeability due to well-developed secondary 

porosity. All three of the Llano Uplift aquifers

show gradual increases in total dissolved solids 

in the down-dip direction away from the main 

Llano Uplift dome, with deep, distal components 

containing total dissolved solids greater than 

10,000 mg/L (Preston and others, 1996).

In the CBSP area, Ellenburger strata are

exposed at the surface and most associated 

springs exhibit normal, epigene karst 

chemistries, while one spring exhibits 

anomalous characteristics. The seven springs 

that exhibit normal epigene karst chemistries, 

include Bear Spring, Gorman Cave Spring, 

Gorman Falls Spring, Gorman Spring, Lemon 

Spring, McLarrin Spring and Well House Spring

(Figure 1, Table 1). Sulphur Spring, as the name 

implies, exhibits anomalous patterns. Bear 

Spring discharges from a solutionally widened 

fracture in the Honeycutt Formation that is 

located four meters above and twenty five 

meters away from the river. Gorman Cave 

Spring is associated with Gorman Cave and 

discharges from the Gorman Formation. Gorman 

Cave Spring discharges directly into the 

Colorado River through alluvial sediments, but 

the spring is accessible for sampling in a stream 

passage in Gorman Cave, which is more than 

one hundred meters from the river and six 

meters above. Gorman Falls Spring is located 

immediately adjacent to the Colorado River with 

discharge from a decimeter-scale conduit in the 

Gorman 

Formation. 

Gorman Spring is 

located at the 

headwaters of 

Gorman Creek 

with discharge 

occurring as 

artesian flow 

from a vertical 

fracture more 

than a kilometer 

from the river 

and sixty five 

meters above

(England and others, 2010). Lemon Spring 

discharges horizontally from a thin bedding 

TDS pH Conductivity Sulphate

ppm stdev pH stdev mV stdev ppm stdev

Bear Spring 345 41 6.8 0.1 -8 9 1.3 0.4

Lemon Spring 386 139 6.8 0.2 -16 3 1.0 0.2

Gorman Cave Spring 325 88 6.6 0.2 -5 10 1.3 0.5

Gorman Falls Spring 344 26 6.7 0.2 -16 12 1.2 0.6

Gorman Spring 351 22 6.7 0.3 -9 19 0.7 0.4

McLarrin Spring 309 46 6.9 0.2 -17 13 0.4 0.4

Sulphur Spring 2067 235 6.6 0.1 -1 14 3.0 0.3

Well House Spring 361 8 6.8 0.1 -13 12 1.1 0.3

Table 1. Average physiochemical characteristics of springs with standard deviation (stdev) 

based on three month sampling.
155

plane in the Tanyard Formation approximately 

seventy five meters from the river and six meters 

above. McLarrin Spring discharges through a 

conduit in the Gorman Formation within twenty 

meters of the river and less than two meters 

above. Well House Spring has been encased and 

is used as the primary water supply for CBSP. 

Well House Spring discharges under artesian 

pressure from the Tanyard Formation 

approximately one hundred meters from the 

river and five meters above. Sulphur Spring 

discharges under artesian pressure from the 

Honeycutt Formation and is channelized into an 

artificial impoundment for recreation. Sulphur 

Spring is located fifty meters from the river and 

five meters above.

SPRING MONITORING

Springs in the CBSP area were monitored in 

order to evaluate spatial and temporal variations 

in groundwater discharge from the Ellenburger 

karst system along the Colorado River. Onset 

HOBO Pendant Temperature Data Loggers were 

installed at each of the springs to record thermal 

variations through the course of the study. At 

three-month intervals, starting in March 2009 

and concluding in June 2010, the temperature 

data loggers were downloaded and redeployed. 

At each of these intervals, chemistry of spring 

discharge was measured in the field with Oakton 

Portable meters (i.e. Oakton pH 300 and CON 

400 meters) in order to evaluate spring pH, 

conductivity and total dissolved solids. Also at 

each sample period, a water sample was 

collected for sulfate analysis, which was 

conducted in the Stephen F. Austin State 

University Geochemistry Lab with use of an 

Agilent 8453 UV-Vis Spectrophotometer 

through precipitation of sulfate ions as barium 

sulfate induced by the addition of barium 

chloride (methodology after Shaw, 2006).

Throughout the study, gaps in data exist for 

various springs. Gaps in thermal data collected 

with data loggers occur as a result of loss of 

loggers either from theft or as a result of being 

dislodged from springs and lost due to natural 

conditions. Gaps in data collected with handheld meters are the result of temporary 

inaccessibility to spring sites either as a result of 

bat hibernation (e.g. Gorman Cave Spring) or 

seasonal hunting activity (e.g. Bear Spring). 

However, in the case of all springs, except Bear 

Spring from which deployed data loggers were 

lost in every sample interval, sufficient data 

were collected to provide initial characterization 

of each springs based on temporal and spatial 

variability.

RESULTS

The seven springs with typical epigenic 

karst characteristics (i.e. Bear, Gorman, Gorman 

Cave, Gorman Falls, Lemon, McLarrin and Well 

House) exhibited average total dissolved solid 

values of 346 ppm, average pH values of 6.7, 

and average conductivities of -11.8 mV (Table 

1). Sulfate concentrations in these seven springs 

were more variable with an average of 0.98 

ppm, with slightly elevated concentrations for 

Gorman Spring, Gorman Falls Spring and Bear 

Spring and significantly lower concentrations in 

Gorman Cave Spring and McLarrin Spring 

(Table 1). In contrast, Sulphur Spring 

consistently exhibited sulfate levels with an 

average of 2.98 ppm and correspondingly 

elevated total dissolved solids averaging 2067 

ppm, elevated average temperature of 23.0°C, 

elevated average conductivity of -1.3 mV and a 

slightly lower pH of 6.6 (Table 1).

Temporal monitoring of spring temperature 

(Figure 3) coupled with monitoring of surface 

temperature and precipitation (Figure 4) showed 

greater variability amongst springs as expected 

from quarterly spot sampling with hand-held 

meters. Springs exhibited variable seasonal 

fluctuation (Figure 3, Table 2), with maximum 

fluctuation observed at Gorman Falls Spring and 

virtually no seasonal fluctuation at McLarrin and 

Sulphur Springs. Lags in seasonal fluctuation of 

one to two days occurred with Gorman Cave 

Spring and Gorman Spring, one to two weeks in 

seasonal fluctuation was observed in Lemon, 

Forman Falls and McLarrin Springs; while no 

data were collected for Bear Springs because of 

logger loss. Thermal shifts associated with 

precipitation events show that Lemon Spring, 

Well House Spring and Sulphur Spring waters 

thermally equilibrate completely with host rock 

and do not show thermal response to individual 

precipitation events. Daily thermal fluctuations 

are observed in Well House Spring because the 

data logger was installed in a parkwater 
156

overflow discharge pipe that is exposed to

diurnal heating that overprints spring behavior. 

Gorman Cave Spring, Gorman Spring, McLarrin 

Spring and Gorman Falls Spring, show varying 

degrees of thermal equilibration between 

precipitation recharge and discharge, ranging 

from one degree of 

equilibration to ten 

degrees of

equilibration,

respectively (Table 2).

DISCUSSION

Significant 

variability in chemical 

and thermal 

characteristics is 

observed in the 

springs within the 

CBSP study area. 

Physicochemical 

characteristics 

measured with 

portable meters 

suggest that all of the 

springs in the study 

area except Sulphur 

Spring are associated 

with relatively 

shallow flow through 

Ellenburger Strata, 

because they exhibit 

consistent values for 

temperature and ion 

activity. In contrast, 

Sulphur Spring 

exhibits elevated 

concentrations of 

dissolved ions, 

including significantly 

greater sulfate 

content, which 

combined with the 

mild thermal 

component, ~2°C 

higher than the other 

springs, suggests a 

deeper / longer 

relative flow path 

(Table 1, 2).

Thermal patterns provide insight into the 

relative flow paths of the groundwater systems 

associated with heat-transfer between water and 

conduit walls. Various researchers have made 

great progress towards characterizing karst 

Figure 3. Thermal spring data for springs in study. Note that gaps in data 

represent lost data loggers, data logger malfunction or inability to retrieve data 

during hunting seasons.
157

systems with temperature being used as a 

groundwater tracer (e.g. Long and Gilcrease, 

2009; Manga, 2001); however, the complexities 

of karst systems and the relative effectiveness of 

heat-transfer in different karst systems greatly 

complicate interpretation. Luhmann and others

(in press) have developed qualitative methods 

for comparing karst systems based on thermal 

patterns observed at spring discharge, in which 

effectiveness of heat-transfer exchange is 

evaluated. Storm hydrographs, coupled with 

thermal monitoring, can provide significant 

insight into the configuration of karst systems 

(Birk and others, 2006), but the installation of 

weirs and level loggers for monitoring of 

discharge are often 

not possible because 

of location and 

economics.

Thermal patterns 

analyses indicate that 

at Lemon Spring, 

Sulphur Spring and 

Well House Spring 

effective heat 

exchange occurs on 

the flow paths which 

is associated with 

distributed recharge. 

Well House and 

Sulphur Spring 

appear to be largely 

decoupled from 

surficial processes and 

represent deepcirculation flow paths 

with artesian 

discharge. Lemon 

Spring is a shallow karst system influenced by 

seasonal temperature fluctuations and likely 

dominated by matrix flow. Gorman Cave 

Spring, Gorman Falls Spring, Gorman Spring 

and McLarrin Spring all exhibit ineffective heat 

exchange along flow suggesting that these 

systems are controlled by localized recharge 

through well-integrated fractures and conduits. 

Gorman Falls Spring and McLarrin Spring show 

significant thermal response to individual storm 

events, indicating that flow velocities through 

these systems are sufficiently high to prevent 

thermal equilibration with conduit walls. 

Gorman Cave Spring and Gorman Springs show 

Annual 

Temperature (°C) Sesonal Variation Precipitation Event Response

avg min max

°C 

Change

lag time 

(days)

avg °C 

temp shift

residence 

time

Bear Spring n/a n/a n/a n/a n/a n/a n/a

Gorman Cave Spring 21.0 20.5 21.5 0.5 2.0 1.2 hours

Gorman Falls Spring 21.8 20.9 22.8 0.9 14.0 10.5 hours

Gorman Spring 21.8 21.5 22.1 0.3 1.0 4.8 hours

Lemon Spring 21.3 20.7 21.9 0.6 11.0 equilabrated days

McLarrin Spring 21.1 20.6 21.6 0.5 10.0 7.0 hours

Sulphur Spring 23.0 22.9 23.0 0.1 30.0 equilabrated weeks

Well House Spring 20.9 20.7 21.0 0.1 21.0 equilabrated weeks

Figure 4. Surface temperature with daily averages (black) and seasonal average 

(red) compared with precipitation events (blue).

Table 2. Thermal variability of spring discharge, including seasonal response and response to storm events with 

temperature variations significant enough to show responses in spring temperature. Note that no data is available for Bear 

Spring because the data logger was lost during all sampling periods.
158

greater thermal equilibration with individual 

storm events indicating longer flow paths than 

Gorman Falls Spring and McLarrin Spring.

CONCLUSION

Karst development in Ellenburger 

carbonates proximal to CBSP is of complex 

polygenetic origin, with current groundwater 

flow paths exhibiting highly variable behavior. 

Two springs, Sulphur and Well House Springs, 

are associated with deep-circulation flow paths 

associated with distal, distributed recharge. 

Lemon Spring is a shallow, karst system 

associated with shallow, distributed recharge 

and diffuse groundwater flow that is being 

focused as discharge along a bedding plane. 

Gorman Falls Spring and McLarrin Spring are 

associated with well-connected, conduits and 

localized recharge that rapidly respond to storm 

events and seasonal fluctuations. Gorman Cave 

Spring and Gorman Spring also exhibit wellconnected conduit systems, but localized 

recharge appears to be significantly distal to 

spring discharge based on the degree of 

equilibration between storm-event water and 

conduit channels. 

These four observed patterns of thermal 

response suggest that groundwater circulation 

within the study area likely occurs across

multiple horizons of groundwater flow that cross 

one another at different depths in the subsurface 

as flow ultimately directed to the potentiometric 

low imposed by the locally deep incision of the 

Colorado River. Variability in flow paths may 

be associated with varying degrees of epigenic 

overprinting of previously existing hypogene 

conduits, where unconfined groundwater flow is 

now occurring through solutional paths 

established during confined conditions.

In conjuction with thermal patterns, 

physicochemical characteristics indicate that all 

spring flow paths, except Sulphur Spring, are 

effectively limited to flow within the 

Ellenburger carbonates, even though thermal 

patterns indicate that the Well House Spring 

system is associated with deeper / longer 

circulation flow paths. Elevated temperature and 

dissolved ions in Sulphur Spring discharge, 

coupled with degassing of hydrogen sulfide, 

indicates that this deep-circulation system is 

likely associated with regional groundwater flow 

in contact with mineralized zones near to

underlying basement rocks. Reported 

occurrences of sulfide minerals are associated 

with lead mineralization proximal to basement 

rock (Barnes, 1956), which could provide the 

source for observed fluid chemistry at Sulphur 

Spring. 

Currently, investigations in the study area 

are attempting to delineate the stratigraphic and 

structural controls on local and regional karst 

development. Spring monitoring continues and 

isotopic studies are planned, including tritium 

analysis, in order to better delineate the age of 

water discharging from Sulphur Springs. Heavy 

metal and trace element analyses are planned to 

determine the source of sulfur associated with 

anomalous springs in the Llano Uplift and Texas 

Hill Country region.

ACKNOWLEDGMENTS

The authors are grateful to Colorado Bend 

State Park, and specifically Cory Evans and 

Kevin Ferguson, for their assistance and access 

to research sites within the study area. The 

authors also thank Wesley Brown, Phillip Hays

and Andy Grubbs for their constructive reviews 

of this manuscript which helped to improve it.

REFERENCES

Barnes, V.E., 1956, Lead Deposits in the Upper 

Cambrian of Central Texas, Report of 

Investigations No. 26; University of Texas, 

Bureau of Economic Geology, Austin, Texas, 68 

p.

Birk, S., Liedl, R., and Sauter, M., 2006, Karst spring 

response examined by process-based modeling; 

Groundwater, Vol. 44, No. 6, p. 832-836.

Collins, E.W., 1995, Structural framework of the 

Edwards Aquifer, Balcones fault zone, central 

Texas; Transactions – Gulf Coast Association of 

Geological Societies, Vo. 45, p. 135-142.

DeLeon, J.L., 2010, Characterizing the Temporal and 

Spatial Variability of Spring Discharge Along 

the Colorado River Near Colorado Bend State 

Park, Central Texas [Master’s Thesis]; Stephen 

F. Austin State University, Nacogdoches, TX, 

156 p.

England, J. Suter, A., and Stafford, K., 2010, Tufa 

diagenesis in a carbonate karst fluvial 
159

environment; Geological Society of America 

Abstracts with Programs, Vol. 42, No. 2, p. 88.

Estaville, L.E. and Earl, R.A., 2008, Texas Water 

Atlas; Texas A&M University Press, College 

Station, TX, 129 p.

Kerans, C., 1990, Depositional Systems and Karst 

Geology of the Ellenburger Group (Lower 

Ordovician), Subsurface West Texas, Reports of 

Investigations; University of Texas, Bureau of 

Economic Geology, Austin, TX, 63 p.

Klimchouk, A., 2007, Hypogene Speleogenesis: 

Hydrolgeologic and Morphometric Perspective; 

National Cave and Karst Research Institute, 

Carlsbad, NM, 106 p.

Long, A.J. and Gilcrease, P.C., 2009, A onedimensional heat-transport model for conduit 

flow in karst aquifers; Journal of Hydrology, 

Vol. 378, p. 230-239.

Luhmann, A.J., Covington, M.D., Peters, A.J., 

Alexander, S.C., Anger, C.T., Green, J.A., 

Runkel, A.C., and Alexander, E.C., in press, 

Classification of thermal patterns at karst springs 

and cave streams; Groundwater.

Manga, M., 2001, Using springs to study 

groundwater flow and active geologic processes; 

Annual Reviews in Earth and Planetary Sciences, 

Vol. 29, p. 201-228.

Mitchell, K., Dornak, S., and Stafford, K.W., 2011, 

Geochemical spatial variability of the Colorado 

River associated with karst springs, Colorado 

Bend State Park, Central Texas; Geological 

Society of America Abstracts with Programs, 

Vol. 43, No. 3, p. 10.

Presont, R.D., Pavilcek, D.J., Bluntzer, R.L., and

Derton, J., 1996, The Paleozoic and Related 

Aquifers of Central Texas, Report 346; Texas 

Water Development Board, Austin, TX, 76 p.

Rose, P.R., 1972, Edwards Group, Surface and 

Subsurface, Central Texas; University of Texas, 

Bureau of Economic Geology, Austin, TX, 198 

p.

Sellards, E.H., Adkins, W.S., and Plummer, F.B., 

1932, The Geology of Texas, v. 1, Stratigraphy; 

University of Texas, Bureau of Economic 

Geology, Austin, TX, 1007 p.

Shaw, M.G., 2006, Geologic Controls of Stream 

Water Composition in Cherokee, Smith and 

Rusk Counites, Texas [Master’s Thesis]; Stephen 

F. Austin State University, Nacogdoches, TX, 

193 p.

Standen, A. and Ruggiero, R., 2007, Llano Uplift 

Aquifers: Structure and Stratigraphy; Texas 

Water Development Board, Austin, TX, 78 p.
160

Analysis of Long-Term Trends in Flow from a Large Spring Complex 

in Northern Florida

By Jack W. Grubbs

U.S. Geological Survey, 2639 North Monroe Street, Suite A-200, Tallahassee, FL 32303 

Abstract

Nonparametric regression analysis of historic flow and rainfall data was used to estimate declining 

flows in a river draining a large spring complex in northern Florida, USA. The analysis indicated that 

flow declined by an estimated 23 percent from 1900 to 2009. The rate of decline appeared to increase 

over time, from about 0.8 cubic foot per second per year during the period from 1930-1970, to about 1.1 

cubic feet per second per year over the period from 1970-2009. The estimated decline for the period prior 

to 1980 is consistent with evidence indicating groundwater withdrawals to the east of the study area have 

diverted groundwater that formerly flowed toward the Ichetucknee River under predevelopment 

conditions.

INTRODUCTION 

The Ichetucknee River is a tributary to the 

Suwannee River and drains an area of karst 

topography in northern peninsular Florida, USA. 

Flow in the 5.2-mile long river is sustained by 8 

named springs and spring complexes that 

discharge groundwater from the highly 

transmissive Floridan aquifer system. The 

purpose of this paper is to describe a long-term 

trend of declining flow in the Ichetucknee River, 

and changes in the groundwater flow system in 

the vicinity of the river that have most likely 

contributed to this trend.

REGIONAL GROUNDWATER FLOW 

SYSTEM

Groundwater in the Floridan aquifer system 

generally flows toward the Ichetucknee River 

from the region north and east of the springs. 

This pattern of flow is evident in potentiometricsurface maps of conditions prior to the initiation 

of substantial groundwater withdrawals from the 

Floridan aquifer system in the late 1800s 

(predevelopement conditions), as well as 

conditions subsequent to groundwater 

development, in May 1980 (Johnston and others, 

1980; Johnston and others, 1981; figs. 1 and 2, 

respectively). These two potentiometric-surface 

maps were chosen because they are the only 

maps that had been published for the entire 

Floridan aquifer system, and therefore make it 

possible to evaluate patterns of the groundwater 

flow system over large areas that could affect 

the discharge of groundwater to the Ichetucknee 

River. The flow lines shown in figures 1 and 2 

are drawn so that (1) they are directed from 

areas of higher water levels toward areas of 

lower water levels, and (2) points of 

intersections between the flow lines and 

equipotential (contour) lines in the 

potentiometric surface form right angles. In 

addition to representing directions of 

groundwater flow, the flow lines can also be 

used to define the approximate location of 

important boundaries within the groundwater 

flow system, as well as the areas that contribute 

groundwater flow to springs and rivers.

A key feature in the regional groundwater 

flow system of the Floridan aquifer system in 

northern Florida is the groundwater flow line 

that runs roughly northwest to southeast and 

defines the boundary between groundwater 

flowing eastward (toward the Atlantic Ocean) 

and westward (toward the Suwannee, Santa Fe, 

and Ichetucknee Rivers) (figs.1 and 2). The 

boundary is formed by two flow lines. The first 

originates on a high, dome-shaped area of the 

potentiometric surface of the Floridan aquifer 

system, near the city of Valdosta, in southern 

Georgia. The second flow line originates on 

another high, dome-shaped area of the 

potentiometric surface of the Floridan aquifer 

system east of the Ichetucknee River, near 

Keystone Heights, Florida. These flow paths 

meet at one or more points where groundwater 

flow direction diverges and splits (stagnation 
161
162
163

point) into eastward- and westward-directed 

flow paths. In this paper, this divide is referred 

to as the northeastern flow-line boundary.

The location of the northeastern flow-line 

boundary is evident in the potentiometricsurface map representing predevelopment 

conditions, as well as in maps made subsequent 

to groundwater development. The configuration 

of the potentiometric-surface map corresponding 

to predevelopment conditions (Johnston and 

others, 1980) suggests that the northeastern 

groundwater flow boundary passed through an 

area near the eastern boundaries of Baker and 

Bradford Counties in Florida (fig. 1; Grubbs and 

Crandall, 2007, fig. 19, page 26) prior to 

groundwater development. This area coincides 

with the Trail Ridge, a physiographic feature 

with the highest elevation in the area between 

the Suwannee and the Atlantic Ocean.

The configuration of the May 1980 surface 

suggests that the northeastern groundwater flow 

boundary migrated westward after development 

of the groundwater flow system (fig. 2; Grubbs 

and Crandall, 2007). This movement is 

consistent with the historic patterns of 

groundwater development (2011, 

http://fl.water.usgs.gov/infodata/wateruse/counti

es.html), in which groundwater withdrawals 

have been substantialy larger to the northeast (in 

Duval and Nassau Counties) than in the four 

counties closest to Ichetucknee Springs 

(Columbia, Union, Baker, and Suwannee) (fig. 

3). Comparison of the two potentiometricsurface maps also indicates that groundwaterlevel declines between predevelopment 

conditions and May 1980 were negligible near 

the Ichetucknee River, but became more 

pronounced toward the northeast. For example, 

the maps indicate that groundwater levels fell by 

approximately 10-12 feet in Baker and northern 

Union Counties, approximately 25 feet in 

eastern Duval and Nassau Counties, and by as 

much as 90 feet near Fernandina Beach, where 

paper mills withdrew approximately 50 cubic 

feet per second (ft3

s

-1

) from the Floridan aquifer 

system in 2000 (Richard Marella, U.S. 

Geological Survey, written commun., 2010). 

This geographic pattern of groundwater-level 

decline and groundwater withdrawals suggest 

that larger withdrawals, by lowering 

groundwater levels over a larger area, shift the 

northeastern flow-line boundary and create a 

larger contributing area to capture greater 

amounts of groundwater recharge. Westward 

migration of the boundary has also resulted in a 

contraction of the area of westward-flowing 

groundwater toward the Suwannee, Ichetucknee 

and Lower Santa Fe Rivers (fig. 2). It should be 

noted that this area where the groundwater 

contributing area has changed (fig. 2) was 

originally delineated in figure 18 in Grubbs and 

Crandall (2007). The area shown in figure 2 

differs from that originally shown in figure 18 in 

Grubbs and Crandall (2007) because it reflects 

corrections that were made to the contour lines 

that were shown in the original figure.

The declines in groundwater levels of the 

Floridan aquifer system observed at individual 

wells are consistent with the differences seen in 

the predevelopment and May 1980 

potentiometric-surface maps. Data from selected 

wells (fig. 4) indicate that average-annual 

groundwater levels declined by 2 to 10 feet in 

wells east and west of the northeastern 

groundwater flow divide from 1960 to 1980, and 

by approximately 4 to 12 feet in these same 

wells from 1960 to 2009. Longer-term data from 

well 1 near the metropolitan area of Jacksonville 

indicated a decline of more than 30 feet from 

1930 to 2009. Kendall’s tau tests (Conover, 

1980) for negative correlations between 

groundwater-levels and time were significant for 

all of the time series shown in figure 4, with the 

highest levels of significance (p-value < 0.0001) 

associated with wells 1-5 (well 6 had a onesided p-value of 0.048). The computed values of 

Kendall’s tau also became progressively more 

negative (and increased in absolute value) for 

wells closer to Duval and Nassau Counties, 

indicating a closer correspondence between 

groundwater level and time as one moves from 

the Ichetucknee River toward these two 

counties.
164 
165 

TRENDS IN THE FLOW OF THE 

ICHETUCKNEE RIVER

Between December 1898 and September 

2009 the U.S. Geological Survey (USGS) made 

538 individual field measurements of flow at a 

cross section of the Ichetucknee River 

downstream from its major springs, where it 

passes under the bridge at U.S. Highway 27 

(USGS station 02322700). Comparison of timeseries plots of the flows measured at this cross 

section and rainfall totals accumulated at Lake 

City (National Weather Service coop station 

084731) for the 24 month period prior to (and 

including) the month during which each flow 

measurement illustrate the relation between river 

flow, rainfall, and time (fig. 5). As expected, 

measured flows generally increased with 

increasing antecedent rainfall totals, and 

decreased as these rainfall amounts decreased. 

For example, flows in the river generally 

increased as conditions became wetter from late 

1963 to late 1965, and subsequently decreased 

as conditions became drier from late 1965 to late 

1968. In addition, periods of extreme low flow 

coincided with periods of deficient rainfall in the 

mid 1950s, 2000-2003, and 2007-2009. 

Evidence for a possible relation between time 

and flow is suggested by noting that the average 

flow from 1930-1965 (346 ft3

s

-1) was higher 

than the average flow for the period after 1970 

(326 ft3

s

-1), even though the average 24-month 

rainfall accumulation for measurements made in 

the earlier period was lower than that of the later 

period (100 inches versus 111 inches, 

respectively). Similarly, the low flows during 

dry periods in 2000-2003 and 2007-2009 were 

generally lower than those during the dry period 

in the mid 1950s, even though the 24-month 

antecedent rainfall totals in the 2000-2003 and 

2007-2009 dry periods were much higher than 

the rainfall totals during the mid 1950s (fig. 5). 

These observations suggest that flow in the 

Ichetucknee River may have declined over time 

from causes that are unrelated to rainfall.

To assess the significance of the time trends 

in the flow in the Ichetucknee River while 

controlling for the effects of rainfall, a 

multivariate, locally-weighted scatterplot

smoothing, (LOWESS; Cleveland and others, 

1988) regression model was created to estimate 

flow in the Ichetucknee River at U.S. Highway 
166

27 as a function of two explanatory variables: 

time and 24-month antecedent rainfall. 

LOWESS regression was chosen because it 

requires no prior assumptions about the form of 

the relation between the response and 

explanatory variables, such as whether the 

relation is linear or nonlinear. The data used to 

fit the model were obtained by grouping the 

individual flow measurements by the year in 

which each measurement was made, and then 

computing the average of the flows, 24-month 

rainfall accumulations, and dates associated with 

each measurement made in a given year. This 

was done to obtain more reliable estimates of the 

model error and confidence limits for the flow 

values produced by the model. The overall 

model fit was highly significant (p-value < 

0.001). Nested F-tests indicated that, although 

both explanatory variables were significant at a 

confidence levels of 0.96 or better, the addition 

of the explanatory variable associated with the 

24-month rainfall accumulation was responsible 

for a greater degree of model improvement (pvalue < 0.007) than that associated with time (pvalue < 0.04).

The regression model made it possible to 

evaluate the magnitude and significance of the 

time trend over selected time intervals by using 

it to estimate river flow at the beginning and end 

of a given time interval (using the same 

antecedent rainfall amount for computing both 

estimates), then computing changes in flow over 

the interval as the difference between the two 

estimates. Table 1 shows the results from an 

analysis of changes in the Ichetucknee River 

flow under conditions of typical rainfall (24-

month rainfall accumulation was set equal to the 

median of 24-month rainfall accumulation 

values used to fit the regression model). In this 

analysis, the estimated flow declined by 92 ft3

s

-1

from 1900 to 2009 (table 1). The analysis also 

suggests that the stream-flow losses have 

probably accelerated over time. For example, the 

estimated slope of the trend for the period 

between 1930 and 1970 was about -0.8 cubic 

foot per second per year (downward), but 

steepened to about -1.1 cubic feet per second per 

year in the period between 1970 to 2009.

DISCUSSION

The estimated decline in the flow of the 

Ichetucknee River from 1900 to 1980 is 

approximately 60 ft3

s

-1 (table 1), with a 95 

percent confidence interval of 23 to 96 ft3

s

-1

. 

The configuration of the May 1980 

potentiometric surface and a more recent 

detailed study of the surface in the vicinity of the 

Ichetucknee River (Sepulveda and others, 2006) 

suggest that the area contributing groundwater to 

the river during this period was probably located 

within an area roughly comprising the southern 

half of Columbia County and perhaps half of 

Union County. The estimated total fresh 

groundwater withdrawals from the Floridan 

aquifer system in these two counties was 

13 ft3

s

-1 in 1980 (2011, 

http://fl.water.usgs.gov/infodata/wateruse/counti

es.html). Thus it seems unlikely that estimated 

decline in the flow of the Ichetucknee River 

from predevelopment conditions to 1980 can be 

explained entirely by groundwater withdrawals 

within the post-development groundwater 

contributing area to the river. The most likely 

alternative explanation is that estimated flow 

declines were caused by the westward migration 

of the northeastern groundwater flow boundary 

(and the concurrent contraction of the 

groundwater contributing area to the 

Ichetucknee River).

This explanation is also consistent with 

estimates of the magnitude of groundwater flow 

that was diverted from the river in response to 

the contraction of its contributing area, 

computed as the product of change in the size of 

the contributing area and groundwater recharge 

rate to the Floridan aquifer system. The 

potentiometric-surface maps suggest a plausible 

range of approximately 450 to 850 square miles 

for the reduction in the size of the groundwater 

contributing area to the Ichetucknee River from 

predevelopment conditions to May 1980. 

Recharge rates of approximately 1 to 1.8 inches 

per year over an area of this magnitude are 

sufficient to produce a flux of 60 ft3

s

-1

, and are 

also consistent with plausible rates of recharge 

(for example, Sepulveda, 2002) in this area 

where groundwater flow directions have been 

reversed. Therefore it seems likely that some (if 

not most) of the estimated reductions in flow in 

the Ichetucknee River from predevelopment 
167

conditions to 1980 resulted from changes in the 

groundwater flow system arising from 

withdrawals outside of the area that contributed 

groundwater to the river in and prior to 1980.

Table 1. Estimated Changes in Flow in the Ichetucknee River at USGS Station 02322700 for Typical Rainfall 

Conditions and Selected Time Intervals.

[Unit abbreviations: ft3s

-1

, cubic feet per second; ft3s

-1

yr-1, cubic feet per second per year]

Time 

Interval

Time Span 

(∆t),

in years

Estimated 

Decline in 

Flow (∆Q), 

in ft3

s-1

Standard 

Error of ∆Q, 

in ft3

s-1

Lower limit 

of 95% 

confidence 

interval for 

∆Q, in ft3

s-1

Upper limit of 

95% confidence 

interval for ∆Q, in 

ft3

s-1yr-1

Estimated 

Rate of 

Decline in 

Flow,

in ft3

s-1yr-1

1900-2009 119 92 20 52 132 0.84

1930-1970 40 31 11 10 53 0.78

1970-2009 39 42 11 20 65 1.08

1900-1980 80 60 18 23 96 0.74

1980-2009 29 32 11 10 55 1.12

REFERENCES

Cleveland, W.S., and Devlin, S.J., 1988, Locallyweighted Regression: An Approach to 

Regression Analysis by Local Fitting. J. Am. 

Statist. Assoc., Vol. 83, pp 596-610.

Conover, W. J., 1980, Practical Nonparametric 

Statistics. 2nd. ed. New York: Wiley

Grubbs, J.W., and Crandall, C.A., 2007, Exchanges 

of Water between the Upper Floridan Aquifer 

and the Lower Suwannee and Lower Santa Fe 

Rivers, Florida: U.S. Geological Survey 

Professional Paper 1656-C, 83 p.

Johnston, R.H., Healy, H.G., and Hayes, L.R., 1981, 

Potentiometric surface of the Tertiary limestone 

aquifer system, southeastern United States, May 

1980: U.S. Geological Survey Open-File Report 

81-486, 1 sheet.

Johnston, R.H., Krause, R.E., Meyer, F.W., Ryder, 

P.D., Tibbals, C.H., and Hunn, J.D., 1980, 

Estimated potentiometric surface for the Tertiary 

limestone aquifer system, southeastern United 

States, prior to development: U.S. Geological 

Survey Open-File Report 80-406, 1 sheet.

Sepulveda, A.A., Katz, B.G., and Mahon, G.L., 2006, 

Potentiometric Surface of the Upper Floridan 

Aquifer in the Ichetucknee Springshed and 

Vicinity, Northern Florida, September 2003: 

U.S. Geological Survey Open-File Report 2006-

1031, 1 sheet.

Sepulveda, Nicasio, 2002, Simulation of GroundWater Flow in the Intermediate and Floridan 

Aquifer Systems in Peninsular Florida: U.S. 

Geological Survey Water-Resources 

Investigations Report 02-4009, 130 p.
168

KARST MODELING

Modifications to the Conduit Flow Process Mode2 for MODFLOW-2005

By Thomas Reimann1, Steffen Birk2, and Christoph Rehrl2

1

Technische Universitaet Dresden, Institute for Groundwater Management, D-01062 Dresden, Germany 2

University Graz, Institute for Earth Sciences, Heinrichstraße 26, A-8010 Graz, Austria

Abstract

Generally, groundwater flow is assumed to be slow and therefore laminar. However, karst aquifers 

exhibit highly conductive features resulting from rock dissolution processes that cause strong 

heterogeneity and anisotropy. Groundwater flow can become turbulent within these structures. It is a 

challenging task to address these karst features in numerical models of groundwater flow. One approach 

is to couple the continuum model with a pipe network that represents the conduit system, for example the 

Conduit Flow Process Mode 1 (Shoemaker and others, 2008). These models exhibit an enhanced 

parameter demand as well as increased numerical efforts. In contrast, the Conduit Flow Process Mode 2 

(CFPM2) computes turbulent flow by adapting the hydraulic conductivity within the linear Darcy law 

used by the MODFLOW continuum model (Shoemaker and others, 2008). Consequently, this approach 

allows computation of turbulent flow within the continuum model. In doing so, CFPM2 reduces the 

practical as well as the numerical efforts for simulating turbulent conduit flow. 

Contrary to laminar flow, which is usually described by linear equations like the well known Darcy 

formula, turbulent flow can be described by nonlinear gradient functions. A universal formulation is:

∂h/∂x =β1q+β2qm (1)

where h is the hydraulic head, x is the spatial coordinate, q is the specific discharge, β1 as well as β2

are parameters related to fluid and rock properties, and m represents the constant exponent of a power 

law. Laminar flow is covered by the linear term in equation (1) whereas the nonlinear term describes 

turbulent flow by a power law. For turbulent flow conditions with high specific discharge the linear term 

in equation 1 becomes small and therefore can be neglected.

The original formulation within CFPM2 was for large pore aquifers where the onset of turbulence 

occurs at low Reynolds numbers (1 to 100) and not for conduits. Benchmark computations demonstrate 

that CFPM2 considers turbulent flow according to the universal power law with m = 1.5 (Reimann and 

others, 2011). In addition, the existing CFPM2 needs several time steps for convergence because of 

iterative adjustment of the hydraulic conductivity without additional control of closure. Modifications to 

the existing CFPM2 were made by implementing a generalized power function where the exponent m is 

user-defined (Reimann and others, 2011). This allows for matching turbulence in porous media as well as

pipes and eliminates the time steps required for iterative adjustment of hydraulic conductivity. The 

modified CFPM2 successfully reproduced simple validation test problems.

REFERENCES

Reimann T., Birk, S,, Rehrl, C., and Shoemaker, W. B., 2011, Modifications to the Conduit Flow Process Mode 2 

for MODFLOW-2005: Ground Water, (online early, March, 2011).

Shoemaker, W.B., Kuniansky, E. L., Birk, S., Bauer, S., and Swain, E. D., 2008, Documentation of a Conduit Flow 

Process (CFP) for MODFLOW-2005: U.S. Geological Survey Techniques and Methods 6-A24, 50 p.

This work was funded by the German Research Foundation (DFG) under grants no. LI 727/11-1 and 

SA 501/24-1 and the Austrian Science Fund (FWF) under grant no. L576-N21. The authors express their 

appreciation to Martin Sauter and Tobias Geyer from the University Göttingen as well as to Rudolf Liedl 

from the TU Dresden for thoughtful discussions. Sadly, we have to report that our colleague and friend 

Dr. Christoph Rehrl died on 21 March 2010. His dedication and person were an inspiration to us all.
169

Comparison of Three Model Approaches for Spring Simulation, 

Woodville Karst Plain, Florida

By Eve L. Kuniansky1, Josue J. Gallegos2, and J. Hal Davis3

1

U.S. Geological Survey, 2850 Holcomb Bridge Road, Norcross, GA 30092 

2

Graduate Student, Department of Geological Sciences, Florida State University, Tallahassee, FL 32306 3 U.S. Geological Survey, 2639 North Monroe Street, Suite A-200, Tallahassee, FL 32303

Abstract

Simulation of flow in karst aquifers with conduits is difficult because the presence and exact locations 

of conduits are frequently unknown. The conduits typically convey as much as 99 percent of the fast flow 

(Palmer, 2007). The Wakulla-Leon Sinks submerged cave system of the Woodville Karst Plain is an 

exception as many conduits are mapped. Extensive mapping of the conduit system has been accomplished 

over the past twenty years by cave divers of the Global Underwater Explorers as part of their Woodville 

Karst Plain Project. As a result of their pioneering work, this karst drainage system is believed to be 

moderately well characterized. Their maps served as the basis for three model approaches used to 

simulate the system. The three model approaches for simulation of the groundwater-conduit system are: 

(1) single continuum laminar flow only (Davis and others, 2010); (2) single continuum that allows for 

turbulent and laminar flow (model of Davis and others, 2010 with Conduit Flow Process mode 2; 

Shoemaker and others, 2008 as modified by Reimann and others, 2011); and (3) hybrid model—that 

consists of a single continuum coupled to pipe-flow network capable of both turbulent and laminar flow 

(Gallegos, 2011 with Conduit Flow Process mode 1; Shoemaker and others, 2008 ). 

The presentation of field applications of MODFLOW-2005 (Harbaugh, 2005) and the Conduit Flow 

Process (Shoemaker and others, 2008) examines spring discharge under various time discretization and 

flow conditions. Model approach 1,the subregional model of Davis et al. 2010 was a transient simulation 

from January 1, 1966, when sprayfield operations began through 2018, when effects of system upgrades 

should have occurred. The stress periods were mostly annual. Calibration data for water levels and spring 

discharge were available for November 1991 and May to early June 2006, plus tracer test datasets 

conducted in 2006 and 2007 by Hazlett-Kincaid, Inc (Davis and others, 2010). Model approach 2 uses 

Model 1 and was not calibrated. Model approach 3, the hybrid model, was calibrated to stressed steadystate conditions for the same calibration datasets as model 1 (Gallegos, 2011). Models (1) and (3) 

simulated observed average spring discharge at Wakulla and Spring Creek Springs within 10 percent and 

head residuals were within the calibration criteria of plus or minus 5 ft. Thus, for average spring 

discharge and tracer test time of travel, the models produce similar results and both are considered 

acceptable calibrations.

In order to observe differences between the model approaches, a transient period with daily 

springflow observations and a storm event was simulated (rising and falling of spring flow, no 

recalibration). For the 52 day storm event, which was simulated with daily time steps, none of the models 

simulated the observed peak discharge well, defined as within 10 percent of observed. Model approach 2 

used the identical input for MODFLOW-2000 from Davis and others (2010) and the transient recharge 

data prepared for the 52-day period, but with the implementation of the turbulence in standard 

MODFLOW as derived by Reimann and others (2011) using constant values of pipe diameter 20, lower 

critical Reynolds number of 3,000 and upper critical Reynolds number of 10,000. The greatest 

differences between the shape of the simulated storm hydrographs for Wakulla Springs and Spring Creek 

were between the hybrid model (3) and the continuum model (1). The continuum model with turbulence, 

model 2, had similar hydrographs as the hybrid model. Methods 1 and 3 were in error by approximately 

30 percent in simulation of Wakulla Spring peak discharge and method 2 by 10 percent. However, if the 

total volume of discharge for the 52 day period at Wakulla Springs is compared, model 3 matched 

observed (within 0.01 percent); model 1by 23 percent; and model 2by 16 percent. Model 1 over-simulated 
170

peak observed discharge and models 2 and 3 under-simulated peak observed discharge at Wakulla Spring. 

For Spring Creek, model 3 was in error by 75 percent at peak discharge; model 1 by 42 percent; and 

model 2 by 7 percent (considered a good match). Models 1 and 3 over-simulated peak observed for 

Spring Creek. However, if the total volume of discharge for Spring Creek is compared, model 3 was in 

error 22 percent; model 2by 27 percent; and model 1by 41 percent. The computation time required by the 

hybrid model (3) was more than 100 times longer than the single continuum model 1 (over 72 hours on a 

personal computer, which precludes using parameter estimation to improve model fit). This test of hybrid 

models is perhaps the largest with over 1,000 pipes and nodes. The hybrid approach requires that 2 

models are iteratively solved until both converge at each one-day stress period and the pipe network is 

very slow to converge for this application.

A recent study (Hill and others, 2010), comparing the hybrid model approach with the single 

continuum model approach for two other springs systems in Florida, concluded that the hybrid model 

more closely simulated observed transient spring discharge. This study concludes that for average 

conditions neither model approach 1 or 3 is distinctly better. For the storm hydrograph, only model 2 met 

the calibration criteria of matching peak spring discharge within 10 percent and better calibration with 

parameter estimation techniques was not possible for the hybrid model approach as a result of the long 

simulation time. However, the hybrid model (3) did outperform the continuum models (1 and 2) for 

simulation of total discharge and matched total discharge at Wakulla Springs and mathematically could 

mimic nature more accurately if fully calibrated. Thus, it is unclear if the extra effort required to use a 

hybrid model, both in data preparation and computation time, is justified based on these two apparently 

conflicting case studies. Two additional considerations of this study need to be examined further; the 

first- the lack of inclusion of turbulence in the single continuum model by Davis and others (2010) did not 

result in a greater peak discharge for Spring Creek than model 3 as it did for Wakulla Springs, which was 

expected since turbulent flow was simulated in the pipe network (Gallegos, 2011). The second was that 

change in temporal discretization for simulation of a storm event appeared to result in the need for 

recalibration of the models, indicating that the models are not suitable for simulation of storm 

hydrographs and should only be used for annual average conditions.

Note the U. S. Geological Survey version of CFP mode 2 that incorporates the new algorithm by 

Reimann and other (2011) is a beta version and full testing of the code is not complete at this time, but 

will be incorporated in version 2. 

References

Davis, J.H., Katz, B.G., and Griffin, D.W., 2010, Nitrate-N Movement in Groundwater from the Land Application 

of Treated Municipal Wastewater and Other Sources in the Wakulla Springs Springshed, Leon and Wakulla 

Counties, Florida, 1966-2018: U.S. Geological Survey Scientific Investigations Report 2010-5099, 90 p.

Gallegos, J.J., 2011, Modeling Groundwater Flow in Karst Aquifers: an evaluation of MODFLOW-CFP at a 

laboratory and sub-regional scale. Master’s Thesis, Florida State University,Tallahassee, Florida.

Harbaugh, A.W. 2005. MODFLOW-2005, the U.S. Geological Survey modular ground-water model -- the GroundWater Flow Process: USGS Techniques and Methods 6-A16.

Hill, M. E., Stewart, M. T. and Martin, A., 2010, Evaluation of the MODFLOW-2005 Conduit Flow Process: 

Ground Water, vol. 48, no. 4, pp. 549–559. 

Palmer, A.N., 2007, Cave Geology: Cave Books, Dayton, Ohio, 453 p.

Reimann, Thomas., Birk, Steffen., Rehrl, Christoph., and Shoemaker, W.B., 2011, Modifications to the Conduit 

Flow Process for MODFLOW-2005 Mode 2: Ground Water (online early, March 2011).

Shoemaker, W.B., Kuniansky, E. L., Birk, Steffen, Bauer, Sebastian, and Swain, E.D., 2008, Documentation of a 

Conduit Flow Process (CFP) for MODFLOW-2005, U.S. Geological Survey Techniques and Methods 

Book 6, Chapter A, 50p.
171

Synthesis of Multiple Scale Modeling in the Faulted and Folded Karst 

of the Shenandoah Valley, Virginia and West Virginia

By Kurt J. McCoy1, Mark D. Kozar2, Richard M. Yager3, George E. Harlow1, and David 

L. Nelms1

1

U.S. Geological Survey, 1730 East Parham Road, Richmond, VA 23228 

2 U.S. Geological Survey, 934 Broadway, Tacoma, WA 98402 

3 U.S. Geological Survey, 30 Brown Road, Ithaca, NY 14850

Abstract

Groundwater flow models were developed at the basin scale, 3,300 square miles (mi2); subbasin scale , 300 mi2

; and watershed scale, 10 mi2; in the Shenandoah Valley to evaluate 

geologic controls on groundwater flow patterns and to describe water budgets at various 

management scales. At basin scale, ground-water flow in the folded and faulted sedimentary 

rocks of the Shenandoah Valley is largely controlled by anisotropy along bedding. Refinement 

of the conceptual model at sub-basin and watershed scales highlights the role of faults, fault 

orientation, heterogeneity in carbonate unit aquifer properties, and epikarst to control the rates of 

direction of groundwater flow. A synthesis of these modeling efforts shows how scale-related 

changes in conceptual models provide hydrogeologic insight into differing geologic features of 

importance to water resource managers at varying scales.
172

A Hydrograph Recession Technique for Karst Springs with Quickflow 

Components That Do Not Exhibit Simple, Zero-Order Decay 

By Darrell W. Pennington1 and Van Brahana2

1

FTN Associates, Ltd., 124 W. Sunbridge Drive, Suite #3, Fayetteville, AR 72703 2

Geosciences, University of Arkansas, 113 Ozark Hall, Fayetteville, AR 72701

Abstract

Recession analysis of karst-spring hydrographs provides valuable information about the 

hydrodynamic performance of the system following storm events. Responses of karst aquifers to 

recharge events depend upon numerous controlling factors which generate a variety of observed recession 

forms. Hydrograph decomposition techniques using exponential curve fitting can identify up to three, 

sometimes more, straight-line segments. Observations of springs from karst regions throughout the world 

have shown that in some instances the intermediate slope (or recession coefficient) is steeper than the 

initial slope, producing a convex form on the quickflow curve. Inasmuch as discharge is proportional to 

head, the discharge is expected to decline most rapidly (i.e. steepest slope segment ) immediately 

following the peak when greater water volume provides driving force. Additionally, in semi-log space, 

the quickflow component is expected to decay in a linear manner, and the baseflow component is 

expected to decay in an exponential manner, producing a concave shape on the quickflow curve. Whereas 

numerous recession methodologies are described in the literature, almost all of these approaches are 

designed for springs that exhibit simple zero- or first-order decay and are not applicable to recession 

coefficients that increase with time. We revised the methodology of Mangin to evaluate two karst springs 

that exhibit concave and convex quickflow components located at the Savoy Experimental Watershed in 

the southwestern Ozark Plateaus. Three functions are used to simulate the hydrograph recessions: 1) 

Mangin’s linear equation for quickflow components that exhibit concavity, 2) a new, empirically derived 

equation, proposed herein, for quickflow components that exhibit convexity, and 3) Maillet’s exponential 

for baseflow components. Comparisons of results show that despite a similar hydrogeologic setting, these 

two springs have very different transmission behaviors that are attributed to differences in flow path 

properties and hydraulic connection with the epikarstic zone. Based on available monitored parameters, 

the recession form appears to be a function of antecedent saturation in aquifer components prior to 

rainfall, available recharge, and hydraulic control and interaction among the reservoir zones. 

INTRODUCTION 

The temporal and spatial variation of 

hydraulic properties in karst aquifers under different flow regimes makes quantification and 

prediction of groundwater movement through 

these systems using traditional field methods 

designed for flow in porous media difficult 

(Baedke and Krothe, 2001). The distribution 

and frequency of a high-velocity conduit 

network (quickflow) embedded in a slowvelocity rock matrix (baseflow) creates dualistic, 

heterogeneous, and anisotropic hydraulic 

processes (i.e. variable infiltration, flow, and 

discharge rates) that affect the overall hydrodynamic performance of the system (Atkinson, 

1977; Király, 2003). In these settings, well data 

gathered from conventional aquifer tests cannot 

be extrapolated far beyond the borehole owing 

to unknowns about the geometry, changes in 

aperture size, and interconnectivity of the conduits (Baedke and Krothe, 2001).

In contrast to wells, karst springs integrate 

flow over an area and show the response of the 

entire basin to recharge events (Padilla et al., 

1994; Baedke and Krothe, 2001). Analytical 

spring-hydrograph techniques provide realistic 

estimates of hydraulic parameters of karstic 

aquifers and reveal information regarding the 

void size and spatial configuration of the conduit 

network (Kovács and Perrochet, 2008). The 

principal influences of precipitation, storage 

conditions, basin characteristics (such as size 

and slope), drainage network density, geological 

variability, vegetation, and soil affect the lag 

time between input and output response as well 

as the output pattern; therefore, spring hydro-
173

graphs can show a wide range of forms (Ford 

and Williams, 2007). 

Karst-Spring Recession Analysis

Traditional karst-spring hydrograph decomposition techniques using exponential curve 

fitting can identify up to three, sometimes more,

straight-line segments (Figure 1). The reason 

for these segments is debated where one group 

of researchers propose that the segments 

represent discharge components from three 

individual (or parallel) reservoirs: conduit network, fracture network, and low-permeability 

matrix (Atkinson, 1977; Padillia et al., 1994; 

Shevenell, 1996). These three segments appear 

due to changes in hydraulic conductivities (K) 

where water levels in aquifer components with 

lower hydraulic resistance respond more quickly 

to recharge. Accordingly, transient flow through 

each of these reservoirs proceeds at rates controlled by K, causing drainage among the three

reservoirs to occur in an out-of-phase sequence 

that is easily recognizable on a hydrograph 

(White, 1988). 

Figure 1. Fitting of exponential segments to 

karst-spring recession limbs. The combination 

of each segment represents the master depletion 

curve (after Shevenell, 1996).

A major limitation to this argument is that it 

is based on Darcy’s Law, which is only applicable to laminar flow and is not valid for high 

velocity flows in karst systems that produce 

turbulence (Baedke and Krothe, 2001). In the 

interpretation of this theory, each exponential 

segment (or recession coefficient) is thought to 

represent the depletion of a particular reservoir 

component. In reality, these three reservoir 

components are not isolated from each other and 

drain concomitantly, and formulae designed to 

determine the K, transmissivity (T), and storage 

(S) of each reservoir component separately 

could result in gross miscalculations (Király, 

2003). 

Other researchers advocate that the straightline segments represent water flowing through a 

series of large conduits characterized by constrictive pathways or debris-filled passages, a 

concept referred to as reservoir-constricted flow 

(Bonacci 1987 and 1993). Halihan et al., (1998) 

numerically modeled the storm-event responses 

of a karst-spring system, Devil’s Ice Box in 

Missouri, characterized by reservoir-constricted 

flow. Results suggested that hydrograph forms 

were controlled by changes from phreatic to 

epiphreatic conditions in a conduit, in conduit 

geometry, or from multiple springs draining the 

same system. Models by Covington et al., 

(2009) suggest that recession hydrographs only 

reveal information about conduit geometry when 

reservoir constrictions are present along the flow 

path. Owing to these changes in geometry, flow 

is delayed through the conduits by different 

amounts of time, and the constrictions and voids 

of varying sizes produces a shape that may be

characteristic only for a particular spring 

(Baedke and Krothe, 2001). 

An unusual situation for recession curves 

occurs when the intermediate recession coefficient (α2) is steeper than the initial slope (α1), 

producing a convex form on the master 

depletion curve (see Figure 1). Inasmuch as 

discharge (Q) is proportional to head (H) [Q

H], the discharge is expected to decline the 

quickest (i.e. the slope to be steepest) 

immediately following the peak when greater 

water volume provides driving force (Bonacci, 

1993). Additionally, in semi-log space, the 

quickflow component is expected to decay in a 

linear manner, and the baseflow component is 

expected to decay in an exponential manner, 

producing a concave shape on the quickflow

curve (Padilla et al., 1994). 

Recession coefficients that increase with 

time have been observed in southern France 

(Bailly-Comte et al., 2010), Greece (Soulios, 

1991), the Dineric Karst (Mangin, 1975; 

Bonacci, 1987 and 1993), and, recently, in 

northwest Arkansas. Three models have been 

proposed for recession coefficients that increase 

with time: 1) Spring basins with two lithologies 
174

of contrasting permeabilities, which creates a 

lag-time for the inflow of water to the spring 

orifice (Soulios, 1991), 2) changes in storageoutflow relations created by temporarily flooded 

cave passages or ponor retention ponds that are 

drained by smaller diameter conduits (Bonacci, 

1987 and 1993), and 3) water exchange and 

pressure transfer between the matrix and conduit 

when antecedent water levels in the aquifer 

matrix are medium to high compared to those of 

the conduits prior to rainfall (Bailly-Comte et 

al., 2010).

Bonacci (1987 and 1993) proposed that

when the water levels in a cave chamber or a

ponor retention pond rise above the surrounding 

ground-water table in the aquifer, the capacity of 

the conduits draining these features is exceeded. 

At this point, the conduits are under pressure and 

discharge occurs according to a form of the 

Darcy-Weisbach equation. While the water 

levels in the flooded passages remain elevated 

above the surrounding groundwater table, the 

discharge of the conduit is independent of the 

hydraulic head (i.e. the discharge can only occur 

according to the conduit capacity where continued in-flow will cause the water level in the 

flooded features to rise without a corresponding 

increase in discharge). This situation will cause 

a flattening in the initial recession coefficient 

(α1). Once the flooded cave chamber, or ponor 

retention pond, is drained to a level equal to, or 

below, that of the water table, the discharge of 

the conduit is no longer under pressure, and 

spring flow will, once again, be proportional to 

the head (Q H). This transition can cause a 

break in the recession slope, producing an 

intermediate recession coefficient (α2)

characterized by a steeper slope with respect to 

the initial recession coefficient (α1), creating a 

quickflow curve with a convex form (Bonacci, 

1987 and 1993). 

Bonacci (1987) observed this behavior and 

hydrograph form in the discharge of springs 

draining flooded ponors in the Dinaric karst. In 

one case, the discharge of the spring ceased to 

increase beyond approximately 22 m3

/s, while 

the ponor retention pond continued to rise 

another 5 m above land surface. Mangin (1975) 

was the first to notice this phenomenon and 

called it a “trop-plein” hydrograph, meaning 

“too-full”. 

Bailly-Comte et al. (2010) observed a 

similar hydrograph form in the Véne Spring near 

Montpellier, France. Véne Spring is an 

ephemeral spring fed by the sinking of the 

Coulazou River into the karstified Aumelas 

Causse Aquifer. In that study, the spring, a 

nearby well with no known connection to karst 

features, and several conduits of a cave (Puits de 

l’aven) were monitored for stage responses to 

storm events. All of the monitoring stations had 

a direct hydraulic connection to the Coulazou 

River. 

Stage fluctuations between the well (representing the matrix flow regime (MFR) and

the ground-water table) and the springs and cave 

conduits (representing the conduit flow regime 

(CFR)) indicated that various recession forms 

were produced on the master depletion curve 

depending upon antecedent water levels in the 

MFR. When pre-storm water levels in the MFR 

were medium to high with respect to the water 

levels in the conduits, a convex form developed 

on the master depletion curve. The authors 

observed that following a storm event, focused 

recharge to the conduits would cause water 

levels to rapidly increase to a stage above that of 

the MFR. At some time after the increase in 

stage, the spring hydrograph would peak and 

enter into a linear decrease while the stage of the 

well would continue to rise for a period of time. 

When the water level in the conduits equaled the 

water level in the matrix, an inflexion point 

occurred on the hydrograph, after which time the 

discharge decreased exponentially. The authors 

suggested that while the stage in the conduits 

exceeded the water level in the well, the conduits were recharging both the spring and the 

MFR. When the water levels in the MFR and 

CFR became equal, there was an abrupt modification within the karst drainage network 

where the flow regime transitioned from a CFR 

to a MFR. This transition manifested as an 

intermediate recession coefficient that was 

steeper than the initial recession coefficient 

owing to water exchange and pressure transfer

(Bailly-Comte et al., 2010). 

Study Location

The Savoy Experimental Watershed (SEW) 

is a University of Arkansas property that encompasses about 1250 ha in the mantled-karst of 

the Springfield Plateau in the southwestern 
175

Ozark Plateaus (Figure 2). SEW is a long-term, 

integrated, and multidisciplinary field laboratory 

where processes, controls, and hydrologic and 

nutrient–flux budgets for surface water, soil, and 

shallow ground-water environments may be 

evaluated in a well-characterized and highlyinstrumented setting. 

Figure 2. SEW site location map with property 

(red line) and watershed boundaries (black lines 

with sequential numbering) delineated. Basin-1 

is the focus of this paper. 

The karst hydrogeology in this area includes 

carbonate aquifers covered by a thin, rocky soil, 

and a variable thickness of regolith. The carbonate rocks have been dissolved to form an 

open network of caves, enlarged fractures, 

bedding planes, conduits, sinkholes, swallets, 

sinking streams, and springs. Flow in these 

aquifers is typically rapid, flow directions are 

difficult to predict, interaction between surface 

and ground water is typically extensive, and 

surface watersheds and ground-water basins do 

not always coincide (Brahana et al., 1999)

SEW Conceptual Model

The previous studies at the SEW established 

a three-component conceptual model consisting 

of the interface (soil-regolith) zone, interflow 

(soil-epikarst) zone, and focused flow (phreatic) 

zone. The focused flow system is embedded in 

a block matrix characterized by low porosity 

(1%), low hydraulic conductivity (6 x 10-6 ft/d) 

and a low storage capacity, where matrix 

contribution to flow is considered to be 

negligible. 

Permeability contrasts between these zones 

and within the aquifer are the primary control 

for ground-water flow, with structural features, 

such as joints and fractures, forming a secondary 

control. Recharge occurs during precipitation 

events, where fast-flow conduits govern groundwater movement that is highly variable and 

exhibits appreciable surface-groundwater interaction. Surface-water and groundwater basins 

do not coincide, and interbasin transfer is 

common. An example of this is spring recharge 

zones changing with naturally varying water 

levels. Basin-1 is under-drained by two major 

springs, Langle and Copperhead (Figure 2), 

which issue from the phreatic zone and form an 

underflow/overflow system. Copperhead, the 

overflow spring, resurges at an elevation of 

314.21 m mean sea level (msl) with an estimated 

spring-basin area of 0.30 km2

, and is characterized by a flashy discharge (0.85 to 379.16 l/s 

over the period of record), a low-flow discharge 

of 0.91 l/s, and temperature and conductivity 

profiles indicative of shallow flow paths with a 

high degree of surface-ground water interaction. 

Langle Spring, the overflow spring, resurges at

an elevation of 314.05 m msl with an estimated 

spring-basin area of 1.40 km2

, and is characterized by a lower discharge range (1.13 to 89.48 

l/s over the period of record), a low-flow 

discharge of 2.01 l/s, and temperature and 

conductivity profiles indicative of deeper flow 

paths with less surface-ground water interaction 

than Copperhead (Al-Rashidy, 1999; Brahana et 

al., 1999). 

Approach

From October 1, 2007 through 2008, 

continuous monitoring of water levels at 

Copperhead and Langle Springs was reinitiated 

using HOBO® U20 submersible data loggers 

equipped with pressure transducers. The 

pressure transducers record the absolute pressure 

values, which consist of atmospheric pressure 

and water head. Barometric effects were compensated for by using one additional data logger 

as a reference station placed near land surface.

After compensation, the pressure sensor has a 

typical error of 2.1 cm and a maximum error 6.5 

cm. These stations were programmed to record 

measurements every 15 minutes. In addition to 

the data loggers, a weather station located at the 

SEW (Figure 2) was programmed to record precipitation measurements every 10 minutes. The 

weather station consists of a tipping bucket rain 
176

gauge with a 0.254 cm resolution and a 

Campbell Scientific® CR10X Measurement and 

Control System. 

Recession Methodology

Numerous recession methodologies are 

described in the literature, and Brodie and 

Hostetler (2005) provide a detailed summary of 

published techniques. Almost all of the proposed approaches are designed for recessions

that exhibit zero- or first-order decay and are not 

applicable to recession coefficients that increase 

with time. Review of available literature has revealed only one proposed method for recession 

limbs that exhibit a convex form. Bonacci 

(1993) proposed a system of functions for 

simulating hydrographs with recession coefficients that increased with time. However, 

these equations required either the collection (or 

assumption) of hydrological and climatological 

parameters, such as evapo-transpiration and soil 

moisture retention as a function of time. These 

parameters are seldom, if ever, available or costeffective for most spring-basin studies. A new 

method is proposed herein requiring only discharge data to simulate the recession. This 

methodology relies on an adaptation of the 

recession functions published by Mangin (1975) 

as described in Padilla et al. (1994) to handle 

convexity and requires no assumed variables.

Mangin (1975) proposed that, during the

recession of a karst spring, the discharge at time 

(Qt) can be expressed with the following 

formula: 

 

where Ψt is a function that translates the effects 

of recharge through the unsaturated zone to the 

spring, modified by transport through the 

phreatic zone; and Φt is described by Maillet’s 

formula (1905):

where Qt

bf is the baseflow discharge at time t;

qo

bf is max baseflow discharge as extrapolated 

from time of interest (ti) – when flow transitions 

from quickflow to baseflow – at the onset of the 

recession as shown in Figure 3; t is the time increment for which the equation is being evaluated; and α is the baseflow (or recession) 

coefficient (Ford and Williams (2007):

where qtf is the discharge at time final (tf), 

which equals the end of the recession period; 

and ∆t is the difference in time between qo

bf and 

qtf.

Maillet (1905) observed that recession 

curves from basins consisting of porous media 

during dry periods followed first-order decay 

that could be treated linearly in semi-log space. 

Although equation 2 was developed for 

homogeneous and isotropic media, numerous 

studies have successfully applied forms of this 

equation to karstified and fissured media 

(Atkinson, 1977; Bonacci, 1987 and 1993; 

Padilla et al., 1994). 

The parameter Ψt is an empirical function 

that simulates quickflow and is expressed by the 

following formula for a recession limb with a 

concave form (Padilla et al., 1994): 

where Qt

qf is the quickflow discharge at time t; 

q0

qf is the max quickflow discharge, which is the 

difference between peak discharge (Q0) and q0

bf

at time t equals 0 as shown in Figure 3; n is the 

reciprocal of ti; and ε is the slope of the trend 

line that best fits the curve created by the 

function Z (t):

Equation 5 is only applied to the baseflow 

discharge between times t0 and ti.

For a recession with a convex form, the 

parameter Ψt is expressed using the following 

formula:

The coefficient ε defines the degree of concavity 

exhibited in the quick-flow component in terms 

of inverse time (1 – nt). For the case of no coneqn. (1) 

eqn. (2) 

eqn. (3) 

eqn. (4) 

eqn. (5) 

eqn. (6) 
177

cavity, ε equals 0, and the term (1 + εt) equals 1 

and can be eliminated from equation 4. The 

elimination of the ε coefficient will produce a 

linear curve that has a convex shape in semi-log 

space. 

Total discharge at any time t (Qt) is the 

summation of the quickflow and baseflow 

discharges for the time increment of 

interest.

Figure 3. Karst-spring recession limb analysis 

using Mangin’s (1975) approach. The shaded 

area represents baseflow and the unshaded area 

quickflow runoff (after Padilla et al., 1994).

Integration of the areas under the recession 

curves with respect to t provides the volume of 

water in the spring basin available for discharge 

(Padilla et al., 1994). The initial volume of 

stored water at t equals 0 available for discharge 

during the quickflow component (V0

qf) of a recession limb (between t = 0 and ti) with a concave quickflow depletion is calculated by the 

formula: 

The initial volume of water stored at t equals 

0 available for discharge during the quickflow 

component of a recession limb (between t = 0 

and ti) with a convex quickflow depletion curve 

is calculated by the formula.

The initial volume of stored water available 

for discharge for the baseflow component (VD) 

from the onset of the recession (between t = 0 

and ∞) is calculated from the formula:

Equation 9 is known as the dynamic volume, 

and it represents the volume of water that would 

drain from the spring if the hydrograph was 

allowed to go from peak discharge to zero. To 

determine that actual volume of water that is 

discharged during the recession period (V0

bf), 

equation 9 must be evaluated between t equals 0 

and time final (tf):

Summation of the V0

qf and V0bf provides the 

total volume of water discharged during the recession to be determined, and summation of V0

qf

and VD provides the thoretical storage volume.

RESULTS

The functions presented above were applied to 

42 recession hydrographs of Copperhead (14) 

and Langle (28) Springs. Both springs displayed recessions that were characterized by 

quickflow curves that exhibited concavity (8 for 

Copperhead and 16 for Langle), by quickflow 

curves that exhibited convexity (3 for Copperhead and 4 for Langle, or by recessions that 

exhibited no quickflow period (3 for Copperhead and 1 for Langle). In instances of no 

observed quickflow, only Maillet’s exponential 

was used to simulate the recession. 

eqn. (7) 

eqn. (8) 

eqn. (9) 

eqn. (10) 
178

For simulations that produced more than 20 

percent average error (7 for Langle), extrapolating the curves beyond the influence of small 

deviations in flow to model the observed discharge was necessary, and results of these recessions are not discussed in this paper. A common reason that the functions failed to adapt to 

observed discharge was attributed to recession 

limbs characterized by more than two inflexion 

points, which separated the recession into more 

than three straight-line exponential segments in 

semi-log space. Despite the complexity of recession forms observed at Langle Spring, this 

spring was more sensitive to small-magnitude 

recharge events than Copperhead Spring producing a greater number of storm-event hydrographs to analyze. 

Graphical representations of four recession 

analyses are shown in Figure 4. A summary of 

the main parameters that characterized the recessions are presented in Table 1, and a summary of volume determinations are presented in 

Table 2. 

DISCUSSION

The various recession forms observed at 

the SEW provide insight into the antecedent 

water levels/saturation prior to rainfall of, 

transmissive and storage properties of, and 

hydraulic control of and interaction between the 

three recognized flow components. 

Recession limbs that can be 

characterized by a single exponent imply that 

there is a linear re-lation between hydraulic head 

and flow rate that is a function of the volume of 

water held in storage (Ford and Williams, 2007). 

The lack of a quickflow component indicates 

that the vadose zone (soil and epikarst) is 

effective in mod-erating infiltration and that the 

saturated zone (phreatic) is responsible for 

observed discharge response (Padilla et al., 

1994). Exponential re-cessions tend to occur 

under low precipitation amounts (average of 38 

mm), over long durations (average of 1.5 d), and 

during the dry season (October through February 

for this study). Under these circumstances, 

recharge is primarily autogenic in that rainfall is 

diffusely distributed through soil macropores 

and down many fissures across the epikarst 

zone. Low average total and dynamic volume 

determ-inations further support this 

interpretation. The high average percent drained 

volumes (52% in 3.6 days for both springs) 

indicate that the phreatic zone is characterized 

by highly trans-missive elements. 

Recession limbs that exhibit a quickflow 

curve indicate that recharge reaches the phreatic 

zone with less moderating affect from the 

vadose zone than exponential declines (Padilla 

et al., 1994). During heavy rains (average of 81 

mm over 2 days) and/or when soil moisture and 

epikarst saturation is high, the infiltration 

capacity of the unsaturated zone is exceeded and 

inflow to the saturated zone occurs via preferential, vertical pathways. The porosity 

contrast between the epikarstic (typically 20%) 

and unweathered portions of the endokarstic 

(average of 1% at SEW) zones creates a 

permeability contrast that promotes a lateral 

component of flow (Ford and Williams, 2007). 

At the SEW, this interflow discharges from 

epikarstic springs into a sinking stream, providing an additional input of concentrated flow 

to the phreatic conduit network that occurs as 

allogenic recharge. 

For quickflow components that exhibit 

con- cavity, the ε term is a measure of the 

infiltration rate through the unsaturated zone 

(Padilla et al., 1994). Low ε values, such as 

those of Langle Spring (0.70) suggest that 

infiltration through the vadose zone is slow. 

This observation is further supported by the 

average quickflow period (ti) of 37 h. 

Contrastingly, Copperhead Spring has a fairly 

high ε value (1.5) and an average ti of 53 h. In 

this timeframe, Copper-head Spring drains an 

average total volume that is one order of 

magnitude higher than Langle Spring. These 

parameters suggest that Copper-head Spring 

receives a greater percentage of flow from 

concentrated and highly transmissive zones with 

a direct hydraulic connection to the vadose zone.

Despite the influence of the un-saturated zone on 

Copperhead, most of the total volume of water 

supplied to Copperhead and Langle Springs over 

the course of the recession is derived from the

baseflow component (86% and 93%,

respectively). The dominance of the baseflow 

recession suggests that the phreatic zone 

provides the primary control on discharge from 

these two springs. Whereas the average 

dynamic volume determinations are similar for

each spring, Copperhead releases 66% of the
179

available water from storage in about nine days 

compared to 31% for Langle over the same time 

period. The high drainage rage from Copperhead suggests that the spring system is characterized by a better developed and more transmissive conduit network with shallower flowpaths. 

Quickflow components that exhibit convexity occur under similar input conditions as 

those of convex quickflow curves (an average of 

81 mm over 2 days). Additionally, concave and 

and convex quickflow curves develop under 

similar flow regimes. Parameters that were not 

monitored as a part of this study that could have 

assisted in understanding the provenance of the 

convex quickflow curve include rainfall intensity and antecedent saturation in the inter-face 

and interflow zones. 

Figure 4. Hydrographs of two recessions (1 concave on top row, 1 convex on bottom row) from 

Copperhead and Langle Springs selected for analyses.
180

Table 1. Average Values of Recession Metrics by Recession Type

Recession 

Type

Spring Q0

(l/s)

q0

bf

(l/s)

α

(day-1)

q0

qf 

(l/s)

ti 

(day)

ε

(day-1)

Avg. 

Error

(%)

Concave CH 150 56.6 0.10 95.7 2.21 1.5 10.51

L 23.8 14.2 0.06 9.91 1.56 0.70 7.17

Convex CH 76.5 14.2 0.15 62.3 3.19 - 14.17

L 31.7 10.8 0.07 20.7 4.85 - 7.09

Exponential CH 40.2 - 0.24 - - - 5.74

L 6.23 - 0.24 - - - 4.34

Table 2. Average Values of Volume Determinations by Recession Type.

Recession 

Type

Spring Vqf

(l .105)

%VT Vbf

(l .105)

%VT VT

(l .105)

VD

(l .105)

Percent 

Drained

Concave CH 35 14 218 86 254 384 66

L 5 7 65 93 70 287 24

Convex CH 119 62 72 38 151 251 60

L 46 33 95 67 140 244 57

Exponential CH - - - - 117 185 63

L - - - - 9 22 41

 Discharge through a conduit depends on the 

amount of available recharge (catchment 

control) and on the hydraulic capacity of the 

conduit (hydraulic control) (Palmer 1991). If 

instant-aneous inflow exceeds the hydraulic 

capacity of the conduits draining the epikarst 

zone, the ex-cess water must be stored in the 

epikarst or dis-charged from the system via 

lateral flowpaths. 

Evidence of exceeding hydraulic capacity in 

the recessions is shown in the longer ti periods 

where the average percent volume drained over 

the course of the recession increases to 62% for 

Copperhead and 33% for Langle, indicating 

greater availability of water in the epikarst zone 

for draining conduits. The quickflow of Langle 

is 40 hours longer than Copperhead but drains a 

much lesser volume (7.3 x 106 l/s), indicating a 

less developed and transmissive conduit network 

with deeper flowpaths. Despite similar input 

rainfall amounts, Langle and Copperhead have 

an average dynamic volume that is one order of 

magnitude less than recessions with concave 

quickflow curves, which suggests that excess 

water is drained through other routes, such as 

overland flow or additional spring outlets 

observed along the escarpment in basin-1 that 

only operate under extreme flood events. 

CONCLUSION

A hydrograph decomposition method is 

presented in this paper that permits quantitative 

analysis of karst-spring recession hydrographs 

that exhibit multiple modes of decay. This 

methodology relies on an adaptation of the 

recession functions published by Mangin (1975) 

as described in Padilla et al. (1994). It is 

appropriate for springflow recession curves 

displaying convexity and requires no assumed 

variables, unlike the method of Bonacci (1993). 

Two karst springs located at the SEW that 

exhibit multiple recession forms were evaluated 

using this methodology. Comparisons of results 

show that despite a similar hydrogeologic 

setting, these two springs have very different 

transmission behaviors that are attributed to 

differences in flowpath properties and hydraulic 

connection with the epikarstic zone. Interpretations from analyses are consistent with 

observations from previous investigations of 

these two spring systems. Based on available 

monitored parameters, the recession form 

appears to be a function of antecedent saturation 

in aquifer components prior to rainfall, available 

recharge, and hydraulic control and interaction 

among the reservoir components. The proposed 

model incorporates and refines parts of 
181

previously published models. Whereas Bonacci 

(1987 and 1993) attributed the hydrograph form 

to conduit capacity along a reservoir-constricted 

flowpath, interpretations from this study suggest 

that the form could also be generated from 

instantaneous inflow from the epikarstic zone 

exceeding the endokarstic conduit capacity. 

This scenario is more likely to occur during the 

wet season when vadose saturation is high and 

the infiltration capacity can easily be exceeded 

producing focused flow. 

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182

Numerical Evaluations of Alternative Spring Discharge Conditions for 

Barton Springs, Texas, USA

By William R. Hutchison and Melissa E. Hill 

Texas Water Development Board, 1700 North Congress Avenue, Austin, Texas 78711-3231

Abstract

Numerical model simulations were performed to evaluate spring discharge conditions over a wide 

range of seven-year climatic conditions at Barton Springs, Texas, USA. The purpose for these predictive 

scenarios was to evaluate the effects of starting heads and alternative pumping conditions. Emphasis was 

placed on quantifying the relative frequency of discharges at or below 0.31 cubic meters per second 

(m3

/s), which is equivalent to the estimated historic minimum monthly discharges that occurred during 

the seven-year drought-of-record from the early 1950s. Each of the model scenarios included 342 sevenyear simulations based on tree-ring data extending from 1648 through 1995. Results from the predictive 

scenarios were evaluated in terms of: i) average simulated discharges under intermediate starting head 

conditions and ii) low simulated discharges under various starting head and pumping conditions. Results 

from the predictive scenarios for the average simulated discharges indicate that under average 

precipitation conditions, the seven-year average simulated discharge can range from 0.90 to 2.57 m3

/s. 

Based on these results, 59 percent of the seven-year simulations had an average discharge above the 

long-term historical average discharge (1.5 m3

/s). Discharges at or below 0.31 m3/s occur at a relative 

frequency of 5 percent under the low starting head condition, using an annual average pumping quantity 

of 8.4 million m3

/y, which is comparable to current groundwater withdrawal estimates, but decreases to 0 

percent using this same pumping dataset under intermediate and high starting head conditions. 

Conversely, the relative frequency for discharges at or below 0.31 m3

/s under a high pumping scenario 

(20.1 million m3

/y) ranged from 4 to 17 percent regardless of starting head conditions. These results 

indicate that the relative frequency of low discharge events (at or below 0.31 m3

/s) using pumping 

quantities of 8.4 million m3

/y, or below, and with the pumping distribution applied in this analysis, will be 

highly dependent on the starting head conditions going into a seven-year drought. However, significant 

increases from current annual average pumping quantities would likely increase the relative frequency of 

low discharge events during drought conditions regardless of starting head conditions due to capture of 

natural discharge.

INTRODUCTION

Barton Springs is located in central Austin 

within Travis County, Texas, USA, fig. 1 and is 

within the jurisdiction of the Barton 

Springs/Edwards Aquifer Conservation District. 

It consists of several vents emanating from the 

underlying Edwards (Balcones Fault Zone) 

Aquifer, which is comprised of the soluble 

Cretaceous-age Edwards Group and the 

Georgetown Formation (Scanlon and others, 

2001). Recharge to the aquifer occurs from rain 

falling over the outcrop area and from streams 

that cross the outcrop area of the aquifer 

(Scanlon and others, 2001). Long-term average 

monthly discharge for Barton Springs from 

January 1917 to January 2007 was 1.5 m3

/s 

(Slade and others, 1986; Hunt, 2009). The 

minimum average monthly discharge of 0.31 

m3

/s occurred in both July and August of 1956 

during the seven-year historical drought-ofrecord (Slade and others, 1986). Discharges 

from the springs provide habitat for the federally 

listed endangered Barton Springs Salamander 

(Eurycea sosorum) (U.S. Fish and Wildlife 

Service, 2010). 

The Texas Water Development Board is 

mandated by the Texas Legislature to develop 

numerical groundwater availability models for 

the major and minor aquifers of the state. The 

models typically utilize the MODFLOW 

groundwater flow simulator developed by the 

United States Geological Survey, and are 

typically large regional models intended to 

simulate regional flow conditions. The model for 
183

the Barton Springs segment of the Edwards 

(Balcones Fault Zone) Aquifer is a sub-regional 

model, with relatively smaller cell dimensions, 

that was developed to estimate groundwater 

availability and simulate spring discharges under 

future pumping scenarios and drought 

conditions. 

Figure 1. Location of Barton Springs.

As part of an evaluation of spring flow 

conditions, the Barton Springs/Edwards Aquifer 

Conservation District requested that the Texas 

Water Development Board estimate 

groundwater pumping for drought conditions 

under average monthly spring discharges of 

0.31, 0.25, 0.20, 0.14, and 0.08 m3

/s. 

APPROACH

Model Recalibration

The existing groundwater availability model 

for the Barton Springs segment of the Edwards 

(Balcones Fault Zone) Aquifer (Scanlon and 

others, 2001) was calibrated using data from 

1989 through 1998, which did not include the 

drought-of-record period in the 1950s. 

Therefore, in order to increase the confidence in 

the ability of the model to simulate spring flow 

during historic drought conditions, the model 

calibration period was extended and the model 

was recalibrated. Details of the recalibration 

effort are provided in Hutchison and Hill (2011). 

The calibration period for the recalibrated model 

was January 1943 through December 2004 (744 

monthly stress periods), with a steady-state 

stress period (stress period 1) preceding the 

transient simulation for a total of 745 stress 

periods. The steady-state stress period did not 

represent a true pre-development condition, but 

it provided a stable initial head solution to 

initialize the transient simulation.

Although groundwater elevation 

measurements, and estimates for discharge and 

pumping were available for the historic droughtof record, stream gage data that would be useful 

for estimating recharge rates were not available. 

Therefore, recharge estimates for the extended 

calibration period were extrapolated from 

recharge estimates in Scanlon and others (2001) 

using 84 regression relationships developed for 

each month (12 total) and recharge zone (7 total) 

in the model. The recharge zones of Scanlon and 

others (2001) coincide with the outcrop area and 

stream locations. In addition, a relationship to 

estimate recharge from precipitation based on 

current month rainfall and rainfall in the 

preceding two months was developed in an 

effort to account for antecedent conditions. 

In the recalibrated model, hydraulic 

conductivity was estimated within the same nine 

zones as used by Scanlon and others (2001). 

However, Scanlon and others (2001) assumed 

that the hydraulic conductivity is isotropic, 

whereas in the recalibrated model, hydraulic 

conductivity is assumed to be anisotropic. 

Significant anisotropy was expected given the 

geologic setting, and the estimated hydraulic 

conductivity values were consistent with that 

conceptualization. 

The model of Scanlon and others (2001) was 

developed with MODFLOW-96 (Harbaugh and 

McDonald, 1996) and the recalibrated model 

was developed with MODFLOW-2000 

(Harbaugh and others, 2000). Statistics from the 

recalibrated model for estimated/measured 

discharges and measured groundwater elevations 

are summarized in Table 1. A hydrograph 

showing the match between estimated/measured 

discharges versus simulated discharges for 

Barton Springs is provided in fig. 2. 
184

Table 1. Calibration statistics for discharges and groundwater elevations.

Figure 2. Hydrograph of estimated/measured discharges versus simulated discharges for Barton Springs using the 

recalibrated model.

Predictive Simulations Using Recalibrated 

Model

Predictive simulations using the recalibrated 

model consisted of 15 scenarios that involved 3 

different starting head conditions (low, 

intermediate, and high) and 5 pumping scenarios 

with annual pumping quantities of 4.7; 5.5; 6.7; 

8.4; and 20.1 million m3

/y. Current pumping is 

estimated to be approximately 8.4 million m3

/y. 

Starting head conditions for the predictive 

simulations were extracted from the recalibrated 

model by evaluating groundwater storage at the 

end of each stress period of the calibration 

period. Based on this review, February 1957 was 

selected as the low starting head condition, June 

1992 was selected as the high starting head 

condition, and January 2004 simulated heads 

were selected as the intermediate starting head 

condition. 

Calibration Statistic

Calibrated Model 

Simulated Discharges

(m3/s)

Calibrated Model 

Simulated Groundwater Elevations

(m)

Minimum Residual -1.45 -58.44

Maximum Residual 1.84 79.00

Absolute Residual Mean 0.38 9.59

Standard Deviation of Residuals 0.48 13.62

Range of Measured Values 3.51 141.48

Standard Deviation/Range 0.14 0.096

Absolute Residual Mean/Range*100 11 7

Sum of Squared Residuals 6.31 × 103 1.37 × 106
185

Groundwater pumping quantities and 

distribution for the 5 pumping scenarios were 

based on pumping estimates from the 

recalibrated model for 1982, 1987, and 2002. A 

factor of 1.25 and 3 was applied to the 2002 

dataset to achieve 2 additional well datasets. The 

1982 and 1987 amounts and distributions from 

the model’s well package were used because the 

annual pumping quantities for these years are 

lower than the current pumping quantities. The 

1.25 factor was applied to the 2002 dataset 

because groundwater withdrawals in the model 

for that year are lower than the estimates of 

current pumping provided by the Barton 

Springs/Edwards Aquifer Conservation District. 

The larger factor (3) was applied to the 2002 

dataset to account for potential increases of 

groundwater withdrawals. The resulting annual 

average pumping quantities for the 5 well 

datasets were 4.7; 5.5; 6.7; 8.4; and 20.1 million 

m3

/y. 

Each of these 15 scenarios included 342 

seven-year simulations extending from 1648 

through 1995 that were based on reconstructed 

rainfall based on the composite of six post-oak 

tree-ring chronologies for South Central Texas 

(Cleaveland, 2006). Rainfall data from the treering dataset were used to estimate recharge from 

rainfall and streamflow. The first seven-year 

simulations were run using rainfall data from 

1648 to 1654, followed by rainfall data from 

1649 to 1655, and so forth with the last sevenyear period ranging from 1989 to 1995. 

Precipitation from the tree-ring dataset was 

estimated on an annual basis. Monthly rainfall 

estimates from these annual estimates were 

developed based on a lookup table that utilized 

the monthly distribution of rainfall during the 

calibration period. For example, if the annual 

rainfall for a given year in the tree-ring dataset 

expressed as a percent of long-term average (for 

example 84 percent) matched the annual rainfall 

of a year in the calibration period the monthly 

precipitation for that year were used to develop 

monthly precipitation estimates for that year of 

the simulation. If an exact match was not 

identified, the next closest match was selected 

and scaled to match the percentage based on the 

reconstructed rainfall values. 

Results from the predictive simulations 

included spring discharge estimates under a 

wide variety of starting head conditions and 

pumping scenarios. These results can be used to 

quantify the relative frequency and magnitude of 

low discharges under a range of pumping 

scenarios. A more complete description of the 

predictive simulations is presented in Hutchison 

and Hill (2011). 

RESULTS

Results from the predictive simulations were 

evaluated in terms of: i) average simulated 

discharges under intermediate starting head 

conditions and ii) low simulated discharges 

under various starting head and pumping 

conditions. Results for the average simulated 

discharges were useful to evaluate the sensitivity 

of simulated discharges to increases in pumping, 

and to quantify discharge variation over a wide 

range of possible precipitation conditions over a 

seven-year period. Results related to low

simulated discharges under various starting 

heads and pumping scenarios were used to 

quantify the relative frequency of low discharge 

events, and evaluate how starting head 

conditions and increased pumping could affect 

the occurrence of undesirably low discharges.

Average Simulated Discharges Under

Intermediate Starting Head Conditions

As expected, increases in pumping resulted in 

lower simulated discharges. As described in 

Bredehoeft and Durbin (2009), one effect of 

groundwater pumping is the capture of natural 

discharge from the groundwater flow system. In 

this case, under pre-development conditions, 

groundwater is recharged from precipitation and 

streamflow, and discharged at Barton Springs. 

Pumping increases result in groundwater storage 

depletion (which can be observed by declining 

groundwater elevations), and decreased 

discharges. The effects of increased pumping on 

the simulated seven-year average discharge 

using intermediate starting head conditions are 

summarized in fig. 3. Also, note that simulated 

spring discharge is equal to or above the longterm average of 1.5 m3

/s when pumping is below 

about 10 million m3

/y. 
186

Figure 3. Effects of increased pumping on the simulated seven-year average discharge using intermediate 

starting head conditions.

An example for the variation of spring flow 

under the range of seven-year precipitation 

conditions for one of the five pumping scenarios 

(8.4 million m3

/y) under intermediate starting 

head conditions is shown in fig. 4. Note that 

each data point represents a seven-year average 

simulated discharge and under a particular 

seven-year period based on the tree-ring dataset. 

It can be seen that under average precipitation 

conditions, the seven-year average simulated 

discharge can range from 0.90 to 2.57 m3

/s. 

Moreover, 59 percent of those average sevenyear simulated discharges (342 total) are at or 

above the long-term historical average discharge 

(1.5 m3

/s). 

Low Simulated Discharges Under Various 

Starting Head and Pumping Conditions

Results from the predictive scenarios were 

evaluated in terms of low simulated discharges 

to be responsive to the request of the Barton 

Springs/Edwards Aquifer Conservation District. 

The relative frequency for discharges at or 

below each of those listed in the request (0.31; 

0.25; 0.20; 0.14; and 0.08 m3

/s) were calculated 

for each of the 15 scenarios, fig. 5. Simulated 

discharges for Barton Springs at or below 0.31 

m3

/s, which are equivalent to the estimated 

minimum monthly discharges during the 1950 to 

1956 drought-of-record, occurred at a relative 

frequency of 5 percent under the low starting 

head condition (essentially a seven-year drought 

that follows a seven-year drought), and an 

annual average pumping quantity of 8.4 million 

m3

/y. The significance of starting head 

conditions is apparent as the relative frequency 

decreases to 0 percent using the same pumping 

scenario under intermediate and high starting 

head conditions, fig. 5.

The relative frequency of simulated 

discharges for Barton Springs at or below 0.31 

m3

/s under low starting head conditions varies 

from 4 to 17 percent using the 4.7; 5.5; 6.7; 8.4 

and 20.1 million m3

/y pumping scenarios. 

However, the relative frequency of simulated 

discharges at or below 0.31 m3

/s decreases to 0 

percent under intermediate and high starting 

head conditions using the 4.7; 5.5; 6.7; and 8.4 

million m3

/y scenarios, whereas the relative 

frequency of simulated discharges at or below 

0.31 m3

/s using the largest pumping dataset 

(20.1 million m3

/y) varies from 4 to 6 percent 

under intermediate and high starting head 

conditions. This suggests that, under the highest 

pumping scenario, simulated discharges become 

more sensitive to pumping relative to starting 

head conditions as more of the natural discharge 

is captured (Hutchison and Hill, 2009).

Throughout this analysis, the pumping 

distributions did not vary from those originally 

assigned in the datasets extracted from the 

recalibrated model. Moreover, we did not 

evaluate the effects on simulated discharges that 

may result due to alternative pumping 

distributions. Results presented here will likely 

differ under alternative pumping distributions, 

particularly if pumping were to increase in the 
187

Figure 4. Plot showing the variation of precipitation versus simulated discharges for the 8.4 million m3/y pumping 

scenario. Fifty-nine percent of the seven-year average simulated discharges are at or above the long-term historical 

average discharge (1.5 m3

/s). 

Figure 5. Curves for annual average pumping quantities (million cubic-meters per year) versus the relative 

frequency (percent) for 0.31, 0.25, 0.20, 0.14, and 0.08 cubic-meters per second spring discharge with low starting 

head conditions.
188

Figure 5 continued. Curves for annual average pumping quantities (million cubic-meters per year) versus the 

relative frequency (percent) for 0.31, 0.25, 0.20, 0.14, and 0.08 cubic-meters per second spring discharge with 

intermediate starting head conditions (top) and high starting head conditions (bottom).
189

vicinity of the springs which would result in 

more efficient capture of the natural discharge.

SUMMARY AND CONCLUSIONS

Results from the predictive simulations, 

evaluated in terms of average simulated 

discharges, under intermediate starting head 

conditions demonstrate that the seven-year 

average discharge decreases with increases in 

pumping, but only the largest pumping dataset 

(20.1 million m3

/y) is below the long-term 

historical average discharge of 1.5 m3

/s. A 

graphical representation, of the seven-year 

percent average precipitation versus the sevenyear average discharges, using intermediate 

starting head conditions, show that simulated 

discharges range from a low of 0.90 m3

/s to a 

high of 2.7 m3

/s. Additionally, 59 percent of 

those average seven-year simulated discharges 

(342 total) are at or above the long-term 

historical average discharge (1.5 m3

/s). 

The predictive simulations for low discharges 

under low starting head conditions and various 

pumping scenarios resulted in relatively higher 

frequencies (4 to17 percent) for simulated 

discharges at or below 0.31 m3

/s compared to 

those using intermediate (0 to 6 percent) or high 

starting head conditions (0 to 4 percent). 

Conversely, the relative frequency for 

discharges at or below 0.31 m3

/s under the high 

pumping scenario (20.1 million m3

/y) ranged 

from 4 to 17 percent regardless of starting head 

conditions. In conclusion, these results indicate 

that the relative frequency of low discharge 

events (at or below 0.31 m3

/s) using pumping 

quantities of 8.4 million m3

/y or below, and with 

the pumping distribution applied in this analysis, 

will be highly dependent on the starting head 

conditions at the beginning of a seven year 

drought. However, significant increases from 

current annual average pumping quantities, 

would likely increase the relative frequency of 

low discharge events during drought conditions 

regardless of starting head conditions due to 

capture of natural discharge. 

ACKNOWLEDGMENTS

The authors thank Mr. Kirk Holland, Dr. 

Brian Smith, and Mr. Brian Hunt from the 

Barton Springs/Edwards Aquifer Conservation 

District for providing data and for fruitful 

discussions pertaining to the conceptualization 

of groundwater flow in the Barton Springs 

springshed. The authors also graciously thank 

Dr. Robert Mace and Ms. Cindy Ridgeway from 

the Texas Water Development Board for 

providing review comments that improved the 

overall quality of this manuscript.

REFERENCES

Bredehoeft, J. and Durbin, T., 2009, Ground water 

development-the time to full capture problem: 

Ground Water, vol. 47, no. 4, p. 506-514. 

Cleaveland, M.K., 2006, Extended chronology of 

drought in the San Antonio area, revised report: 

University of Arkansas, 29 p. 

http://www.gbra.org/documents/studies/treering/

TreeRingStudy.pdf

Harbaugh, A.W., Banta, E.R., Hill, M.C., and 

McDonald, M.G., 2000, MODFLOW-2000, The 

U.S. Geological Survey modular ground-water 

model – user guide to modularization concepts 

and the ground-water flow process, U.S. 

Geological Survey Open-File Report 00-92, 121 

p. 

Harbaugh, A.W. and McDonald, M.G., 1996, User’s 

documentation for MODFLOW-96, an update to 

the U.S. Geological Survey modular finitedifference ground-water flow model, U.S. 

Geological Survey Open-File Report 96-485, 56 

p. 

Hunt, B., 2009, Written communication, June 19. 

Austin, Texas: Barton Springs/Edwards Aquifer 

Conservation District.

Hutchison, W.R. and Hill, M.E., 2011, Recalibration 

of the Edwards Balcones Fault Zone (Barton 

Springs segment) Aquifer Groundwater Flow 

Model Draft Report: Texas Water Development

Board, variously p.

Hutchison, W.R. and Hill, M.E., 2009, GAM Run 09-

019 Draft Report: Texas Water Development

Board, 28 p.
190

Scanlon, B.R., Mace, R.E., Smith, B., Hovorka, S., 

Dutton, A.R., and Reedy, R., 2001, Groundwater 

availability of the Barton Springs segment of the 

Edwards aquifer, Texas: Numerical simulations 

through 2050: Bureau of Economic Geology, 36 

p.

Slade, R.M. Jr., Dorsey, M.E., and Stewart, S.L., 

1986, Hydrology and water quality of the 

Edwards Aquifer associated with Barton Springs 

in the Austin area, Texas, U.S. Geological 

Survey Water-Resources Investigations Report 

86-4036, 117 p. 

U.S. Fish and Wildlife Service, 2010, Species Profile, 

http://ecos.fws.gov/speciesProfile/profile/species

Profile.action?spcode=D010
191

FIELD TRIP GUIDE

Geology and Karst Landscapes of the Buffalo National River Area, 

Northern Arkansas

By Mark R. Hudson,1 Kenzie J. Turner,

1 and Chuck Bitting2

1

U.S. Geological Survey, Box 25046, Mail Stop 980, Denver, CO 80225

2

National Park Service, 402 N. Walnut, Suite 136, Harrison, AR 72601

Abstract

The Buffalo River watershed of northern Arkansas contains abundant karst features and associated 

karst hydrologic systems developed within Mississippian and Ordovician stratigraphic sequences. Cherty 

limestone of the Lower to Middle Mississippian Boone Formation is the predominant host of karst 

features and corresponds to the Springfield Plateau karst aquifer. Sinkholes are best developed at the top 

of the Boone Formation where it is overlain by Batesville Sandstone. Springs are concentrated at the base 

of the Boone Formation, with the largest karst systems localized within structural lows whose recharge 

may include interbasin transfer from the adjacent Crooked Creek topographic basin. A lower karst 

aquifer, the Ozark aquifer, is recognized where structural highs expose lower parts of the Ordovician 

Everton Formation within the Buffalo River watershed. Both modern karst systems and late Paleozoic 

lead-zinc mineralizing fluids were influenced by lithologic and structural characteristics of the bedrock, 

and modern karst systems may locally reoccupy ancient fluid pathways.

INTRODUCTION

Extensive karst landscapes have developed 

within the Ozark Plateaus region due to their 

favorable geology and ample precipitation in a 

temperate climate. These landscapes host a 

variety of important and sometimes unique 

natural resources and ecosystems. Karst 

landscapes are well expressed at Buffalo 

National River, a park administered by the 

National Park Service (NPS) in northern 

Arkansas (fig. 1) that was established in 1972 to 

preserve this 214-km-long free-flowing river. 

The watershed for the Buffalo River and 

adjacent areas contains abundant caves, 

sinkholes, losing streams, and springs that are 

linked in karst hydrologic systems. During the 

last 15 years, geologic mapping at 1:24,000 

scale by the U.S. Geological Survey and the

Arkansas Geological Survey has advanced 

knowledge of the lithologic and structural 

characteristics that have potentially influenced

karst development in this region (figs. 1 and 2). 

Stops on this field trip (fig. 3) will examine the 

geologic setting of karst features in the westcentral part of the Buffalo River watershed.

GEOLOGIC, PHYSIOGRAPHIC, AND 

HYDROLOGIC SETTINGS

Geologically, the Buffalo River watershed 

lies within the southern flank of the Ozark dome 

(fig. 1), a late Paleozoic uplift developed in the 

foreland to the Ouachita orogenic belt. The 

watershed includes part of the Boston Mountains 

and Springfield Plateau physiographic provinces 

that are held up by Pennsylvanian sandstone and 

shale and by Mississippian cherty limestone, 

respectively. The river incises Ordovician rocks 

that hold up the Salem Plateau farther to the 

northeast (Purdue and Miser, 1916). 

Erosion by the Buffalo River and its 

tributaries has enhanced bedrock exposures of 

an approximately 1,600 ft (490 m) thick 

stratigraphic sequence (fig. 4; Appendix) of 

carbonate and clastic sedimentary rocks (Purdue 

and Miser, 1916; McKnight, 1935) that were 

deposited predominantly in shelf-marine, nearshore-marine, or fluvial environments. The 

distribution of carbonate lithologies determines 

the potential for karst development. Ordovician 

rocks are alternating dolostones, limestones, and 

carbonate-cemented sandstones, as typified by 

the heterogeneous Middle Ordovician Everton 

Formation (Suhm, 1974) that crops out widely in 

lower elevations of the Buffalo River valley.
192

Figure 1. Location of study area within northern Arkansas, overlapping the western part of Buffalo National River. 

Field trip stops are approximately located. U.S. Geological Survey geologic maps in area (Hudson, 1998; Hudson and 

others, 2001, 2006; Hudson and Murray, 2003, 2004; Hudson and Turner, 2007, 2009; Turner and Hudson, 2010) are 

listed by publication number. Lower regional map illustrates geological and selected physiographic provinces of 

Arkansas and adjacent areas. The Ozark Plateaus region includes Salem and Springfield Plateaus and Boston 

Mountains physiographic provinces.
193

Figure 2. Geologic map for west-central part of Buffalo River watershed and adjacent areas compiled from published maps listed in fig. 1 as well as McMoran (1968) 

and Lucas (1971). Locations for field trip stops and dye tracer tests are also shown. Other abbreviations include: KLF, Kyles Landing fault; BHM, Big Hurricane Mine; 

CBC, Chilly Bowl Cave.
194

Figure 3. Regional road map of northern Arkansas with Buffalo National River (shaded area) and locations of field-trip stops.
195

Thin, but hydrologically significant, shale beds 

are present within both the Ordovician Powell 

Dolomite and St. Peter Sandstone. The exposed 

Ordovician rocks comprise the upper part of the 

Ozark aquifer of Adamski and others (1995).

Mississippian rocks are predominantly limestone 

but also have significant thickness of shale and 

sandstone in their upper part. The Lower to 

Middle Mississippian Boone Formation is 

divided into the basal St. Joe Limestone Member 

(Kinderhookian to Osagean) and the main body 

of the Boone (Osagean to Meramecian); together 

these form the Springfield Plateau aquifer of 

Adamski and others (1995). Upper 

Mississippian Pitkin Limestone is present in 

parts of the watershed. Pennsylvanian rocks are 

dominantly sandstone and shale, with carbonatecemented sandstone and limestone restricted to 

the Prairie Grove Member of the Hale Formation 

and lower Bloyd Formation, respectively. Due to 

their significant thickness of low-permeability 

shale, the Upper Mississippian and 

Pennsylvanian strata were included in the 

Western Interior Plains confining hydrogeologic 

unit of Adamski and others (1995). Regional 

unconformities (fig. 4) separate the Ordovician, 

Mississippian, and Pennsylvanian sequences;

uppermost strata of both the Ordovician and 

Mississippian sequences are truncated northward 

beneath these two unconformities. Quaternary 

surficial deposits are present as colluvial 

wedges, landslides, and as a series of riverterrace deposits that are linked to the history of 

modern landscape development. 

Paleozoic strata were deformed by a series 

of faults and open folds (fig. 2) that developed in 

three late Paleozoic deformation phases 

(Hudson, 2000; Hudson and Cox, 2003). The 

oldest deformation, D1, is aerially restricted to 

the >4-km-long, north-northwest-trending Kyles 

Landing fault and an associated fold that were

active in latest Mississippian-earliest 

Pennsylvanian time (Hudson and Murray, 2003).

A widespread deformation phase, D2, formed 

most of the west- to west-northwest-striking 

normal faults, northeast-striking dextral strikeslip faults, and variably oriented monoclinal 

folds throughout the area. These structures acted 

in a coordinated fashion to accommodate northsouth extension during flexure of the continental 

margin as it was being loaded beneath the 

Ouachita orogenic belt to the south, probably 

during Middle-Late Pennsylvanian time 

(Hudson, 2000). In the final deformation phase, 

D3, preexisting structures were locally 

reactivated in north-south shortening related to 

final docking of the Ouachita orogenic belt.

Regionally in northern Arkansas, north-south 

shortening of D3 is recorded in calcite-twin 

deformation in limestones (Chinn and Konig, 

1973) and by a pervasive north-trending joint set 

(Hudson and others, 2010).

Late Paleozoic lead-zinc mineralization was 

widespread in the Ozarks region and has been 

linked to a northward flow of mineralizing 

brines out of the adjacent Arkoma Basin during 

the Ouachita orogeny (Leach and Rowan, 1986).

Leach and Rowan (1986) suggest that regional 

flow of ore-bearing fluids was focused along 

basal Cambrian clastic strata, but in northern 

Arkansas lead-zinc deposits are hosted by 

Middle Ordovician and Lower Mississippian 

strata near faults (McKnight, 1935), 

demonstrating that faults acted to divert 

mineralizing fluids up section (Erickson and 

others, 1988). Late Pennsylvanian to early 

Permian ages for the mineralization are 

indicated by isotopic dating (Brannon and 

others, 1996) and paleomagnetic studies (for 

example, Pan and others, 1990). Commonly 

associated with lead-zinc deposits (McKnight, 

1935), collapse breccias demonstrate that 

ancient episodes of carbonate dissolution predate 

modern karst development.

Karst Features

In the modern landscape, carbonate rocks in 

the Buffalo River region were subjected to 

dissolution as capping Pennsylvanian and 

Mississippian sandstones and shales were eroded 

from the area. In particular, the thick limestone 

interval of the Mississippian Boone Formation 

developed karst systems from the interchange of 

surface and groundwater via recharge from 

sinkholes and losing-stream segments and 

discharge at springs. Dye tracer tests (fig. 2) 

indicate subsurface transport of groundwater

occurs rapidly (1000’s of ft/day) within the

Springfield Plateau aquifer (Mott and others, 

2000), consistent with transport via solution 

conduits. Many cavern systems are known 

within the Boone Formation. An unpublished 
196

Figure 4. Representative stratigraphic columns for Paleozoic rocks for the eastern and western parts of the study 

area. Regional hydrogeologic units follow Adamski and others (1995) and local karst aquifers recognized in this study 

are indicated. FM.–Formation
197

NPS inventory of caves within the park 

demonstrates that Boone Formation is the 

dominant host of caves within the area (fig. 5).

Discharge occurs at springs in both the Buffalo 

River and adjacent Crooked Creek watersheds 

and recharge areas for springs in the Buffalo 

River watershed locally extend beyond 

topographic divides of surface drainages (Mott 

and others, 2000). An inventory of springs 

within the Mill Creek subbasin (Hudson, 1998) 

of the Buffalo River watershed demonstrates a 

strong frequency maximum at or near the basal 

contact of the St. Joe Limestone Member of the 

Boone Formation (fig. 5). Although these 

characteristics indicate that the Boone Formation 

is the main karst aquifer in the watershed, the 

largest spring in the Park (Mitch Hill Spring) 

discharges from the lower part of the Ordovician 

Everton Formation, another common host of 

caves (fig. 5). Lower parts of the Ordovician 

stratigraphic sequence are exposed by several 

structural highs within the watershed. 

Comparison of springs and connecting dye 

traces to the underlying geology in these areas 

indicates that a limestone-bearing interval in the 

lower part of the Everton Formation forms a 

second karst aquifer discharging into the Buffalo 

River watershed.

Field Trip Stops

This field trip will provide an overview of 

the west-central Buffalo National River region 

and will include key stops to illustrate probable 

stratigraphic and structural controls on karst 

systems. It will address the effects of both a 

spatially variable Paleozoic stratigraphy and late 

Paleozoic faulting and folding on the location of 

major springs, caves, and recharge basins within 

two stacked karst aquifers (fig. 4). Most stops 

will be near roads (fig. 3), although other 

interesting features will be noted that may be 

reached on foot with more time. All location 

coordinates for trip stops are given relative to 

the North American Datum of 1927 to match 

1:24,000-scale topographic maps. This trip is 

designed to be easily completed in a single day.

Stop 1 – Gaither Mountain (36°10.464'N, 

93°12.474'W)

This overlook serves as an introduction to 

the physiographic setting of the southern Ozark 

Figure 5. Frequency of caves and springs 

within stratigraphic units of part of the study 

area. (A) Pie diagram of cave entrances within 

park boundaries of Buffalo National River from 

unpublished NPS inventory. Description of 

Silurian limestone can be found in Turner and 

Hudson (2010). (B) Relative frequency of 

springs versus stratigraphic position from parts 

of Ponca, Jasper, Hasty, Gaither, and Harrison 

quadrangles (from Hudson, 1998).
198

Plateau region. The southern Ozark Plateau 

region contains three regional geomorphic 

surfaces. The highest, Boston Mountains, is held 

up by Pennsylvanian strata, mostly sandstone 

and shale. The intermediate Springfield Plateau 

is held up by Lower to Middle Mississippian 

limestone. The Salem Plateau is held up by 

Cambrian and Ordovician strata, predominantly 

dolostone with lesser limestone and sandstone.

The overlook stop is sited within Lower 

Pennsylvanian (Morrowan) strata of the Bloyd 

Formation (figs. 2, 3, and 4) with a notable thick 

cross-bedded sandstone member (middle Bloyd 

sandstone) that forms prominent bluffs 

throughout the region (fig. 6). The view to the 

east gives an overview of the Springfield Plateau 

that is developed on top of the Mississippian 

Boone Formation. A cluster of mountains 

(Sulphur, Boat, and Pinnacle Mountains) are 

capped by Pennsylvanian strata and represent 

remnants of the Boston Mountains plateau that 

have not yet been lost to erosion. Looking east, 

the surface drainage drains northward into 

Crooked Creek. A topographic divide to the 

southeast separates the Crooked Creek watershed 

from the deep valley of the Buffalo River to the 

south that cuts through Pennsylvanian and 

Mississippian strata into Ordovician strata. Dye 

tracer studies have documented interbasin transfer 

of groundwater from the southern Crooked Creek 

topographic watershed into the Mill Creek 

subbasin of the Buffalo River watershed (Mott 

and others, 2000).

Figure 6. Photograph of thick cross-bedded 

sandstone of middle part of the Pennsylvanian Bloyd 

Formation present at Stop 1. This unit holds up a 

significant part of the high Boston Mountains

geomorphic surface in the region.

Stop 2 – Fitton Spring (36°5.273'N, 93°14.001'W)

Springs within karst systems in the 

Buffalo River watershed reflect the discharge 

of groundwater, and larger springs may be 

associated with cavern systems. Fitton Spring 

(figs. 7, and 8) represents the discharge point 

for Fitton Cave, the longest known cave 

system in Arkansas; cave mapping has 

identified more than 12.8 km (8 mi) of 

mapped passages with multiple vertical levels 

spanning a 120 m (393 ft) elevation range. The 

deepest cave passages have been mapped to 

within about 150 m (500 ft) distance of Fitton 

Spring, and dye traces have confirmed the 

intervening hydrologic connection.

A series of dye tracer tests conducted by T. 

Aley for the NPS in the late 1990’s 

demonstrated that Fitton Spring and nearby Van 

Dyke Spring gather their recharge from 

upstream segments of Cecil Creek. Traces 

detected at Fitton Spring received input from the 

Bartlett Cove branch and from slopes northwest 

of the spring. Van Dyke Spring received input 

from traces farther west in the Cecil Creek basin.

Geologically, the Fitton Spring and Cave 

system are fully hosted within Mississippian 

Boone Formation and the spring emerges from 

the upper part of the basal St. Joe Limestone 

Member. Three geologic factors may have 

contributed to development of this large karst 

system. (1) Facies of the main body of the 

Boone Formation limestone in this part of the 

watershed contain less insoluble chert than 

typical for northern Arkansas, facilitating 

connectivity of caverns. (2) An uncommonly 

thick section of Batesville Sandstone and 

Fayetteville Shale overlies Boone Formation in 

the Cecil Creek area. These units contain 

disseminated pyrite whose oxidation may have 

provided chemically aggressive waters that 

aided limestone dissolution. Such pyrite 

oxidation also provides a source of sulfate for 

gypsum cave formations that are common in 

highest passages within Fitton Cave, closest to 

the contact with the overlying formations. (3) 

The Fitton Cave karst system is localized at the 

base of a structural trough bounded by two westnorthwest-trending monoclines. The association 
199

Figure 7. Geologic map of the northeastern Ponca quadrangle (Hudson and Murray, 2003) and northwestern Jasper 

quadrangle (Hudson and others, 2001) with location of stops 2 and 3. Map unit symbols are keyed to Fig. 4 and 

Appendix. Dye tracer tests were conducted by T. Aley in the late 1990’s for National Park Service.
200

Figure 8. For Stop 2, (A) Map view of survey lines for 

Fitton Cave passages, from unpublished mapping 

data gathered by NPS. (B) View to north of Fitton 

Spring, the main discharge for Fitton Cave that emits 

from the uppermost part of the St. Joe Limestone 

Member of the Mississippian Boone Formation. Within 

Fitton Cave, gypsum cave formation types include (C) 

polycrystalline flowers, (D) needles, and (E) crusts 

and pendants.

of large karst systems with structural troughs in 

the Boone Formation has been documented

elsewhere in the Buffalo River watershed 

(Hudson, 1998) and demonstrates that these low 

points in the Springfield Plateau aquifer gather 

groundwater where they are perched above lesspermeable Ordovician rocks at the surface.

Stop 3 – Batesville sinkhole (36°4.782'N, 

93°14.430'W)

Field observations indicate that all carbonate 

lithologies in the Buffalo River region undergo 

dissolution and thus have potential to develop 

shallow surface depressions. However, sinkholes 

of sufficient size and depth to have expression 

on 1:24,000-scale topographic maps are strongly 

concentrated at or near the contact of the 

Batesville Sandstone with the Boone Formation.

This stop (fig. 7) examines a sinkhole (fig. 9A) 

developed at the edge of a local Batesville 

Sandstone topographic flat where the overlying 

Fayetteville Shale was removed by erosion.

Batesville Sandstone is interpreted to have 

sufficient strength to hold up topography until 

caverns in the underlying Boone Formation 

enlarge to sufficient size to cause collapse and 

form large sinkholes. In contrast, solution 

cavities preserved on lower slopes of Boone 

Formation are mostly smaller, in part because 

they are infilled by residual chert blocks 

weathered from higher in the formation. An 

additional factor in creation of large sinkholes in 

the stratigraphic interval of the upper Boone 

Formation-Batesville Sandstone may be the 

presence of pyrite in the Batesville Sandstone 

and overlying Fayetteville Shale whose 

oxidation may have increased acidity of 

descending waters. Oxidized pyrite framboids 

are common (fig. 9B) and form red spheres

distributed through the tan, parallel laminated, 

very fine-grained sandstone blocks of Batesville 

Sandstone.

Within the sinkhole, vertical conduits 

developed within limestone of the underlying 

Boone Formation are elongated along northeasttrending joints, one of several dominant regional 

joint sets (fig. 10) (Hudson and others, 2010).

Such apparent local control of dissolution 

features by joints is prevalent at outcrop scale 

throughout the Buffalo River region, although at 

regional scales the intersection of multiple joint 

trends probably leads to near-isotropic fracture 

permeability. 
201

Figure 9. For Stop 3, (A) photograph of sinkhole 

developed in Batesville Sandstone and (B) oxidized 

pyrite (pyr) framboids standing in relief on weathered 

surface of Batesville Sandstone block (1.9-cm 

diameter coin for scale).

Figure 10. Rose diagram of joints measured within 

Buffalo River mapping areas.

Stop 4 - Ozark (36°3.876'N, 93°9.535'W)

This stop illustrates both Ordovician 

Everton Formation and multiple Quaternary 

terrace deposits of the Buffalo River (fig. 11).

The Middle Ordovician Everton Formation is a 

heterogeneous mix of dolostone, sandstone, and 

limestone that were deposited in tidal flat and 

barrier island environments, and have been 

divided into several members (Suhm, 1974).

Although grouped within the Ozark aquifer of 

Adamski and others (1995), the Everton 

Formation contains significant sandstone 

intervals that act as local aquitards. Across the 

river at this stop, Briar Bluff exposes the thick 

Newton Sandstone Member overlying an 

interval containing thin beds of stromatolitic 

limestone (figs. 12A and B). Local dissolution is 

evident in the limestone interval below the 

Newton Sandstone Member. This boundary is 

used to separate informal upper and lower map 

units of the Everton Formation and represents 

the top of a proposed lower carbonate aquifer 

within the watershed.

Sandy deposits form a terrace level about 

20 ft (6 m) above the river base level that is 

extensive along the Buffalo River. An older 

terrace deposit is preserved as an interval of 

cobbles and capping sand 80 ft to 110 ft (24 m 

to 33.5 m) above the river base level that is 

inset into Ordovician bedrock to the south. The 

cobbles are subrounded to well rounded and 

include clasts of sandstone and chert derived 

from a mixture of Ordovician, Mississippian, 

and Pennsylvanian bedrock sources. Remnants 

of similar cobble deposits at similar heights 

above the current river base level have been 

mapped for many kilometers both upstream 

and downstream from this site (fig. 11). 

Terrace deposits provide a record of valley 

incision and thus the lowering of the ultimate 

base level for karst springs discharging into the 

watershed. Studies are underway to better 

characterize the style and rate of Buffalo River 

valley incision and how this may have 

influenced karst development.
202

Figure 11. Geologic map modified from the eastern Jasper quadrangle (Hudson and others, 2001) and western 

Hasty quadrangle (Hudson and Murray, 2004) with location of stops 4 and 5. Map unit symbols are keyed to fig. 4 

and Appendix. 
203

Figure 12. Photographs of (A) Briar Bluff exposing 

Ordovician Everton Formation, with designation of 

upper and lower parts. Top of lower part placed below 

Newton Sandstone Member at uppermost 

stromatolitic limestone beds. See lower right box for 

location of (B) close view of stromatolitic limestone 

beds with local dissolution conduits. Note 

hemispherical casts of algal heads in lower bed.

Stop 5 - Welch Bluff (36°3.841'N, 93°7.617'W)

This stop (fig. 11) displays a well-exposed 

normal fault zone (fig. 13A) in a bluff on the 

opposite side of the Buffalo River. The fault 

zone forms the southern margin of the Braden 

Mountain graben, part of an east-striking graben 

system that continues eastward for about 29 km 

(18.3 mi). This graben system accommodated 

north-south D2 extension that developed during 

late Paleozoic flexure of the southern Ozarks 

region in the foreland of the Ouachita orogenic 

belt. At Welch Bluff, the contact between the 

Ordovician Everton Formation and the basal, 

reddish St. Joe Limestone Member of the 

Mississippian Boone Formation rises 20 m 

(66 ft) above the river across two fault splays to 

form the southern footwall of the fault. On the 

hanging wall, this Boone-Everton contact is 

dropped below river level adjacent to the fault 

but it rises above the surface farther to the 

northwest. The southeast-dipping strata in the 

hanging wall (fig. 13B) are an example of an 

extensional relay ramp caused by a decrease of 

the vertical displacement of faults toward the 

western termination of the graben. Farther east, 

throw on the southern fault of the Braden 

Mountain graben is as much as 120 m (393 ft).

In the bluff, the fault zone is composed of two 

fault planes that dip 74°N with a 3-m- (10-ft-)

wide intervening fault block in which thin 

limestone beds are folded about an axis parallel 

to the fault strike. This fold geometry and the 

preservation of several high-rake slickensides 

on these fault planes indicate a dip-slip, normal 

displacement for the main movement on this 

fault zone.

Figure 13. For Stop 5, (A) View looking east of South 

Braden Mountain fault zone with two fault splays that 

drop contact between the St. Joe Limestone Member 

(Mbs) of the Boone Formation and the Everton 

Formation (Oe) below river level on the northern 

hanging wall block. Mb—main body Boone Formation. 

(B) View to west of gently southeast-dipping strata in 

hanging wall of fault zone. 
204

Fault zones commonly coincide with 

topographic lineaments in the Buffalo River 

area. In several instances, parallel alignments of 

small sinkholes are coincident with faults, 

particularly where both affect the stratigraphic 

interval near the Boone Formation-Batesville 

Sandstone contact. These observations suggest 

that karst dissolution is enhanced near fault 

zones due to an increased fracture density and 

associated fracture permeability. Fault zones 

also localized late Paleozoic fluid flow 

associated with lead-zinc mineralization in 

northern Arkansas, as illustrated by ore extracted 

from Big Hurricane Mine (McKnight, 1935) 

located about 17 km (10.7 mi) to the east along 

the same graben system (fig. 2). These relations 

demonstrate the potential for modern cavern 

systems located near fault zones to have had 

hypogene, late Paleozoic origins, as has been 

documented at Chilly Bowl Cave (Tennyson and 

others, 2008) located about 9 km (5.7 mi) 

northeast of the stop (fig. 2). 

Stop 6 – Yardelle Spring (36°4.115'N, 93°0.156'W) 

Yardelle Spring discharges from the St. Joe 

Limestone Member of the Mississippian Boone 

Formation (figs. 14 and 15) at the bottom of a 

structural trough framed by a northeast-trending 

down-to-southeast monocline on the west, the 

northeast-trending Shaddock Branch anticline to 

the east, and the west-northwest-trending downto-north Yardelle monocline to the south (figs. 

14 and 16). The spring emerges just above the 

basal contact with sandstone of Ordovician 

Everton Formation. Dye tracer studies

conducted by the National Park Service 

demonstrate that Yardelle and the nearby 

Yardelle Branch Spring receive water from sites 

across the topographic divide within the Clear 

Creek watershed to the north. Similar interbasin 

transfer of groundwater discharging from 

springs in structural lows within the Springfield 

Plateau aquifer (Boone Formation) has also been 

documented in areas to the east and west (fig. 2).

Due to their location within a perched, 

unconfined karst aquifer, we interpret that the 

groundwater gradient linked to these lowelevation discharge points allows these 

associated springs to gather recharge from long

distances that can extend into the adjacent 

watershed. In the Mill Creek subbasin to the 

west, such interbasin recharge has practical 

resource management concerns because 

increased nutrient loads are introduced into the

springs from agricultural lands within the 

adjacent Crooked Creek watershed (Mott and 

others, 2000).

Stop 7 – Mitch Hill Spring (36°0.895'N, 

92°57.044'W) 

Mitch Hill Spring (fig. 14), located near Mt. 

Hersey, is the largest spring within the park with 

an average flow of about 2 cubic feet per 

second. It is sourced from the lower part of the 

Ordovician Everton Formation and thus 

represents discharge from a lower karst aquifer 

within the Buffalo River watershed. In contrast 

to springs within the Boone Formation that 

respond quickly to precipitation, Mitch Hill 

Spring has more constant flow with a modest 

artesian head (fig. 17A) that suggests the source 

aquifer is partially confined. Geologically, the 

lower part of the Everton Formation and 

underlying Powell Dolomite are brought to the 

surface in a structural high by the northwesttrending down-to-northeast Cane Branch 

monocline. Powell Dolomite is an argillaceous 

dolostone (McKnight, 1935) that is interpreted 

to act as a lower confining unit forcing 

groundwater discharge at Mitch Hill Spring 

(Hudson and others, 2005; Turner and others, 

2007). Nearby the spring to the west, McKnight

(1935) reported that zinc carbonate ore was 

extracted at Old Granby mine from the 

lowermost Everton Formation as early as 1875.

Sandstone ledges directly above the Mitch Hill 

Spring horizon commonly contain silicified 

collapse breccias (fig. 17B), giving indirect 

evidence for karst dissolution in underlying 

strata. However, the strong silicification and the 

observation that breccias are cut by joints that 

probably formed during late Paleozoic orogeny 

(Hudson and others, 2010), indicates that the 

breccias are probably ancient and likely formed 

due to dissolution associated with flow of leadzinc mineralizing fluids. Thus, the modern karst 

system of Mitch Hill Spring is in part coincident 

with, and may have reused, ancient fluid-flow 

pathways.
205

Figure 14. Generalized geologic map from Turner and others (2007) based on geologic data from Hudson and Murray (2004) and 

Hudson and others (2006). Also shown are Davis Creek subbasin of Buffalo River watershed, springs, and dye traces conducted by 

National Park Service. 
206

Figure 15 (left). View looking west of Yardelle Spring, 

discharging from St. Joe Limestone Member of Boone 

Formation.

Figure 16 (above and right). Images from three-dimensional geologic model (Turner and 

others, 2007) of the Hasty and Western Grove quadrangles (area of fig. 14). (A) Overview of 

model, looking to north. (B) View to northeast with layers above the St. Joe Limestone 

Member of the Boone Formation removed. Boxes represent dye trace input and detection 
207

points with connecting lines. (C) View to northwest showing potential dye trace path from 

sink in lower part of the Everton Formation, in Davis Creek below Yardelle Spring through 

graben toward Mitch Hill Spring. (D) View to west of potential groundwater flow path across 

graben between lower Everton Formation on graben flanks via St. Joe Limestone Member of 

Boone Formation within graben. (E) View to northwest showing end of potential dye trace at 

Mitch Hill Spring at upwarp of underlying Powell Dolomite at Cane Branch monocline.

Dye tracer studies from the upper part of the 

immediately adjacent Davis Creek drainage 

demonstrate a transition of groundwaters and 

surface waters from the upper Springfield 

Plateau aquifer (Boone Formation) to the lower 

Everton aquifer, but this transfer is complicated 

by intervening faults and folds. Downstream 

from Yardelle Spring, Davis Creek can go dry 

where the lower part of the Everton Formation is 

brought to the surface. Dye traces from this 

losing stream segment emerge from Mitch Hill 

Spring, but the flow path must cross an 

intervening graben (fig. 14). A threedimensional geology model (Turner and others, 

2007) allows visualization of likely pathways for 

groundwater flow across the graben (fig. 16).

This model demonstrates likely paths of 

groundwater flow from the lower Everton 

Formation through down-faulted Boone 

Formation within the graben, and back into the 

lower Everton Formation.

Figure 17. Photographs of (A) Mitch Hill Spring, and 

(B) Breccias within silicified sandstone layers that lie 

above Mitch Hill Spring horizon. North-trending joints 

that cut breccia are probably late Paleozoic in age.

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APPENDIX

Geologic Map Unit Descriptions (modified from 

Hudson and others, 2001; Hudson and Murray, 

2003, 2004; Hudson and others, 2006; Hudson 

and Turner, 2009). 

Qal Alluvial deposits (Quaternary)—

Unconsolidated sand and gravel. Active-channel 

gravel deposits are composed of subangular to 

rounded Paleozoic rock clasts of mixed lithology 

along drainages, and interspersed with bedrock 

exposures (not mapped). Low terraces of lightbrown, fine sand are locally present adjacent to 

creeks. As thick as 10 ft

Qty Younger terrace and channel alluvial deposits 

(Quaternary)—Unconsolidated sand and gravel of 

Buffalo River. Terrace deposits are principally lightbrown fine sand and have smooth upper surfaces that 

are about 20 ft above the base-flow level of river. 

Channel gravel is subangular to rounded and composed 

of Paleozoic sandstone, chert, dolostone, and limestone; 

unmapped bedrock exposures are interspersed along 

channel. Deposits are as thick as 20 ft

Qto Older terrace and alluvial deposits 

(Quaternary)—Unconsolidated gravel, sand, and silt 

deposits adjacent to Buffalo River. Deposits are 

principally a lag of brown-weathered, subrounded to 

rounded cobbles of Paleozoic sandstone and chert in 

a reddish-brown sandy to silty matrix. Fine sand, silt, 

and clay locally overlie cobble deposits. Deposits are 

80–120 ft above base-flow level of river. Deposits are 

as thick as 20 ft 

Qc Colluvial deposits (Quaternary)—

Unconsolidated deposits of subrounded to angular 

blocks as large as 20 ft in diameter, commonly in an 

orange-brown silty clay matrix. Blocks are mostly 

derived from the basal sandstone of the upper part of 

the Bloyd Formation (IPbu) or the Cane Hill Member 

(IPhc) of the Hale Formation. Deposits have fan-like 

morphology and were mapped where sufficiently 

thick to mask typical ledge-flat topography of 

underlying bedrock. Smaller, thinner colluvial 

deposits elsewhere were not mapped. Thickness 

uncertain but probably more than 10 ft 

Bloyd Formation (Lower Pennsylvanian, 

Morrowan)—Interbedded sequence of sandstone, 

siltstone, shale, and thin limestone beds separated 

into lower and upper intervals. As much as 400-ft 

thick

IPbu Upper part—Dominantly sandstone with 

interbedded siltstone and shale. Upper part of 

sequence contains dark-gray to black shale and 

siltstone beds that form topographic flats interbedded 

with sandstone beds that form ledges. Upper 

sandstone intervals are 5–20 ft thick and vary from 

orange-brown, fine to coarse grained with local 

quartz pebbles, and medium to thick planar bedded to 

crossbedded to tan or olive, fine to very fine grained, 

ripple cross-laminated to planar bedded. Base of unit 

is crossbedded sandstone as thick as 80 ft that forms 

prominent cliffs in area. This basal sandstone is 

white to light-brown, fine- to medium-grained quartz 
210

arenite that has a sharp erosional base and is 

commonly a composite of several tabular and troughcrossbed sets. Sandstone contains local 

concentrations of white quartz pebbles and casts of 

wood fragments. Rocks of the upper part of the 

Bloyd Formation were originally assigned to the 

Winslow Formation by Purdue and Miser (1916). 

Zachry (1977) concluded that the basal sandstone 

was a time-equivalent unit with the Woolsey Member 

of the Bloyd Formation farther west and designated it 

informally as the “middle Bloyd sandstone.” 

Thickness is 200–300 ft

IPbl Lower part—Sequence of predominantly 

dark-gray to black shale and siltstone with thin beds 

of sandstone and limestone; poorly exposed under 

moderate to steep hillslopes. Siltstone is medium to 

dark gray, fissile, and moderately bioturbated. 

Medium beds of rippled, very fine grained 

sandstone are locally exposed as ledges. Coarse 

bioclastic limestone is reddish gray and weathers 

dark brown but is poorly exposed and mostly 

observed in loose blocks in lower part of sequence. 

Thickness 90–120 ft

Hale Formation (Lower Pennsylvanian, 

Morrowan)—Interbedded sequence of sandstone, 

siltstone, shale, and thin limestone. Thickness 100–

180 ft

IPhg Prairie Grove Member—Brown to reddishbrown, fine- to medium-grained, thick-bedded, 

calcite-cemented sandstone. Locally contains quartz 

pebbles in its base and thin interbeds of fossiliferous 

limestone. Beds are planar or crossbedded, and 

crossbeds may have bi-directional dips. Weathered 

sandstone forms rounded surfaces. Sandstone forms 

steep hillslopes or ledges that are commonly covered 

by slope debris from overlying units. Thickness 10–

60 ft

IPhc Cane Hill Member—Interbedded sequence of 

shale, siltstone, and sandstone. Most of unit is poorly 

exposed dark-gray fissile shale and thin-bedded 

siltstone that form gentle to moderately steep slopes.

Upper part locally contains rippled, thin-bedded, very 

fine and fine-grained sandstone intervals as thick as 5 

ft. Lower part includes 10-ft-thick sandstone interval 

that changes downward from olive-brown, very finegrained, thin-bedded sandstone with ripple crosslamination to reddish-brown, medium- to thickbedded, very fine-grained to medium-grained 

sandstone with trough crossbeds. Lower sandstone 

contains sparse casts of wood fragments and 

conglomerate lenses with quartz pebbles and angular 

to subrounded clasts of shale, siltstone, and 

sandstone. Unit unconformably overlies the Pitkin 

Limestone. Thickness 80–160 ft

Mp Pitkin Limestone (Upper Mississippian, 

Chesterian)—Generally medium- to dark-gray, fetid 

limestone. Limestone in medium to thick wavy beds 

varies from micrite at base to coarse grained and 

locally oolitic near top. Limestone beds may contain 

abundant crinoids, brachiopods, corals, and bryozoan

Archimedes. Basal contact with the Fayetteville Shale 

is conformable, although rarely exposed. The Pitkin 

generally forms a prominent ledge or steep slope.

Thickness 40–100 ft

Fayetteville Shale (Upper Mississippian, 

Chesterian)—Fine-grained sandstone and siltstone 

of the Wedington Sandstone Member that grades 

downward into main body of black, slope-forming 

shale. Thickness 250–390 ft

Mfw Wedington Sandstone Member—Brown to 

light-gray, well-indurated, calcite-cemented 

sandstone and siltstone. The Wedington caps a steep 

slope and is commonly separated from the overlying 

Pitkin Limestone by a topographic bench. Sandstone 

is very fine grained and is present in thick to thin 

planar beds with internal parallel laminations and 

locally developed low-angle crossbeds. Sandstone 

grades downward into siltstone beds that are ripple 

cross-laminated and bioturbated. The Wedington 

grades downward into main body. Thickness 20–40 ft 

Mf Main body—Below the Wedington Sandstone 

Member, the Fayetteville Shale is poorly exposed 

black shale that locally contains thin, rippled beds of 

olive-brown siltstone near its top. The lower part of 

the Fayetteville crops out along stream gullies where 

it consists of black fissile shale that may contain 

medium- to light-gray, fetid septarian concretions as 

large as 2 ft in diameter. The Fayetteville Shale is 

susceptible to landslides. Thickness 230–350 ft 

Mbv Batesville Sandstone (Upper Mississippian, 

Chesterian)—Fine-grained to very fine grained, 

light- to medium-brown, calcite-cemented sandstone 

with sparse interbedded limestone. Thin to medium 

beds are typically parallel laminated; low-angle 

crossbeds common in upper part of unit. Sandstone 

commonly contains burrows on bedding plane 

surfaces. Sandstone breaks into thin flat blocks. One 

or more discontinuous, 1- to 3-ft-thick, medium- to 

dark-gray, fetid, fossiliferous limestone beds are 

locally interbedded with sandstone; limestone beds 

contain crinoids and brachiopods. Both sandstone and 

limestone beds may contain 2- to 10-mm-diameter 

oxidized pyrite framboids that weather to reddishbrown spheres. Where stripped of the overlying 

Fayetteville Shale, the top of the Batesville forms 

topographic flats. Unit hosts sinkholes formed by 

collapse into dissolution cavities in the underlying 

Boone Formation. Thickness 40–60 ft
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Boone Formation (Upper to Lower 

Mississippian)—Limestone and cherty limestone of 

main body that grade into the basal St. Joe Limestone 

Member. The Boone Formation is a common host of 

caves and sinkholes. Total thickness is 380–400 ft 

Mb Main body (Upper to Lower Mississippian, 

Meramecian to Osagean)—Medium- to thickbedded, chert-bearing bioclastic limestone.

Limestone is light to medium gray on fresh surfaces 

and generally coarsely crystalline with interspersed 

crinoid ossicles. A 1- to 3-ft-thick bed of oolitic 

limestone is common in upper part of unit. Beds of 

dense, fine-grained limestone are present in upper 

one-third of unit. Beds are typically parallel planar to 

wavy, but channel fills are locally present in lower 

part of unit. Chert content varies vertically and 

laterally within the Boone and forms lenticular to 

anastomosing lenses. Chert-rich horizons are 

generally poorly exposed but produce abundant float 

of white-weathered chert on hillslopes. Chert in 

uppermost part of unit contains prominent 

brachiopod molds. Thickness 310–350 ft 

Mbs St. Joe Limestone Member (Lower 

Mississippian, Osagean to Kinderhookian)—Thinand wavy-bedded bioclastic limestone containing 

ubiquitous 3- to 6-mm-wide crinoid fragments in a 

finely crystalline matrix. Limestone is commonly 

pink to red on fresh surfaces due to hematite in 

matrix, but color varies depending on hematite 

concentration. Lenticular red to pink chert nodules 

are locally present in upper part. Contact with 

overlying main body of Boone Formation is 

gradational and marked by increase in chert and 

thicker beds upward. Middle to lower part of St. Joe 

Limestone Member contains greenish-gray shale 

interbeds. Base of St. Joe is a 0.3- to 8-ft-thick, very 

fine to fine-grained, moderately sorted, tan sandstone 

containing phosphate pebbles. Total thickness of 

member is 30–50 ft

Of Fernvale Limestone (Upper Ordovician)—

Thick- to massive-bedded, coarsely crystalline 

bioclastic limestone. Limestone is light gray to pink 

on fresh surfaces and contains abundant 3- to 10-mmwide cylindrical to barrel-shaped crinoid ossicles in a 

coarsely crystalline matrix. Thickness 10–20 ft

Opl Plattin Limestone (Middle Ordovician)—

Thin- to medium-planar-bedded, fine-grained, dense 

limestone locally interbedded with calcarenite.

Characteristically, dense, very fine grained, mediumgray limestone that breaks with conchoidal fracture 

and weathers to very light gray, thin, tabular blocks.

Locally in upper parts there are interbeds of 

laminated, tan to light-gray, fine-grained to very fine 

grained, limey to dolomitic calcarenite as thick as 10 

ft; calcarenite has fetid odor when freshly broken.

Thickness 10–100 ft

Osp St. Peter Sandstone (Middle Ordovician)—

Very fine grained, grayish-yellow sandstone 

interbedded with blue-green siltstone, shale, and 

sparse dolostone. Sandstone is calcite-cemented 

arenite containing well-rounded quartz grains. Upper 

part of unit, as thick as 25 ft, is sandstone interbedded 

with blue-green fissile shale, thin-bedded siltstone, 

and sparse 0.5- to 3-ft-thick beds of brown-gray, fine 

crystalline dolostone; shale, siltstone, and dolostone; 

this interval is commonly marked by zone of slumped 

sandstone blocks on hill slopes. Main part of unit 

contains a sequence of 1- to 2-ft-thick sandstone beds 

overlying a massive basal sandstone ledge as thick as 

20 ft. Bioturbation is common in most sandstone beds 

and includes cylindrical burrows (Skolithos) 

perpendicular to bedding that weather to distinctive 

straw-like forms. Basal sandstone is crossbedded, 

contains inclined fractures, and discordantly overlies 

Everton Formation. Thickness ranges from 0–70 ft

Everton Formation (Middle Ordovician)—

Interbedded dolostone, limestone, and sandstone 

sequence, divided into upper and lower parts. Unit is 

250–450 ft thick

Oeu Upper part—Interbedded dolostone, sandstone, 

and limestone. Limestone is finely crystalline, light 

gray, thick bedded, and locally fossiliferous, and is 

present in intervals as thick as 20 ft. Dolostone is 

light to dark gray, finely to medium crystalline, 

laminated, and medium to thick bedded commonly 

containing sandstone stringers. Sandstone is arenite 

composed of well-sorted, well-rounded, fine to 

medium quartz grains. Sandstone is light tan to white 

in planar, medium to thick beds and is variably 

cemented by dolomite or calcite; sandstone locally 

grades into sandy dolostone or limestone. Limestone 

interbedded with limy sandstone and lesser dolostone 

at top of unit comprises Jasper Member of Glick and 

Frezon (1953). Dolostone and dolomitic sandstone 

underlying Jasper Member make up most of upper 

part of Everton. Base of upper part of Everton is 

marked by base of Newton Sandstone Member 

(McKnight, 1935), a planar, medium-bedded quartz 

arenite composed of well-sorted, well-rounded, fine 

to medium quartz grains. Newton Sandstone Member 

is 10–90 ft thick. Upper part of Everton Formation is 

180–200 ft thick 

Oel Lower part—Interbedded dolostone, sandstone, 

and limestone. Limestone is restricted to upper part of 

unit and is interbedded with dolostone and sandstone. 

Limestone is light gray, finely crystalline, thin to 

medium bedded, commonly laminated, and contains 

sparse stromatolite hemispheres as large as 1 ft in 

diameter. Dolostone, present throughout unit, is light 
212

gray to grayish brown, finely to medium crystalline, 

thin to thick bedded, and commonly laminated, and it 

commonly contains stromatolites. Sandstone is lighttan to white quartz arenite, composed of well-sorted, 

well-rounded, fine to medium grains, variably 

cemented by dolomite or calcite. Sandstone intervals, 

composed of planar, thin to medium beds as thick as 6 

ft, are present throughout unit. Sandstone also forms 

lenses and stringers within limestone and dolostone 

intervals. Light- to dark-brownish-gray dolostone and 

sandy dolostone interspersed with sandstone are 

predominant in lower part of unit. Collapse breccia 

with clasts of planar laminated dolostone and 

dolomitic sandstone is present. Unit is 140–250 ft thick 

Op Powell Dolomite (Lower Ordovician)—Finegrained, thin-bedded, argillaceous dolostone. 

Yellowish gray on fresh surfaces. Weathers to thin, 

platy blocks. As much as 65 ft of Powell is exposed 

along the Buffalo River where it is exposed in the 

eroded upthrown side of Cane Branch monocline near 

Mt. Hersey (Glick and Frezon, 1953). Regionally, 

formation thickness ranges from 40–200 ft